Altered Volcanic Rocks

Erratum

There are several instances of a typographic error inFigure 4.2 (page 79). The references to 'Y/Zr! in thecaption and 'Y/Zr' in labels on the diagram areincorrect. They should be 'Zr/Y' in all cases.

Altered Volcanic RocksA guide to description and interpretation

Cathryn Gifkins

Walter Herrmann

Ross Large

Published by the Centre for Ore Deposit Research

University of Tasmania, Australia

UTAS

Published by CODESCentre for Ore Deposit Research,

University of Tasmania,

Private Bag 79,

Hobart, Tasmania, Australia 7001

An ARC Special Research Centre

© Centre for Ore Deposit Research, 2005

National Library of Australia Cataloguing-in-Publication Data

Gifkins, Cathryn.Altered volcanic rocks : a guide to description and interpretation.

Bibliography.Includes index.ISBN 1 86295 219 1.

1. Rocks, Igneous. 2. Hydro thermal alteration. I.Herrmann, Walter, 1951- . II. Large, Ross R. III.University of Tasmania. Centre for Ore Deposit Research.IV. Title.

552.2

another Pongratz Production 2005

Copy editing: Im'press: clear communicationIndex: Word Wise and Im'press: clear communication

Printed in Australia by the Printing Authority of Tasmania

Ill

I CONTENTS

PREFACE v.viiACKNOWLEDGEMENTS viiiINTRODUCTION ix

1 | ALTERATION IN SUBMARINE VOLCANIC SUCCESSIONS 1

1.1 Submarine volcanic successions 1Volcanic facies 1Volcanic facies associations 2Evidence for submarine environment of emplacement 2

1.2 Alteration in submarine volcanic successions 2Devitrification 4Alteration processes 4Characteristics inherited from volcanic facies 6

1.3 Geology of the Mount Read Volcanics 7Stratigraphy of the Mount Read Volcanics 9Submarine facies associations and architecture 10Post-depositional alteration processes 11Mineral deposits and prospects 11

1.4 Geology of the Mount Windsor Subprovince 12Stratigraphy of the Seventy Mile Range Group 12Submarine facies associations and architecture 13Post-depositional alteration processes 14Mineral deposits and prospects 14

2 | DESCRIBING ALTERED VOLCANIC ROCKS 15

2.1 Frequently asked questions 152.2 Alteration nomenclature 19

Mineral-based alteration nomenclature 19Compositional alteration nomenclature 20Generic alteration nomenclature 20Descriptive nomenclature — alteration facies 20

2.3 Alteraction facies — the recommended method 222.4 Alteration mineral assemblage 23

Tools for mineralogical determination 242.5 Alteration intensity 25

Qualitative estimates of alteration intensity 25Quantitative estimates of alteration intensity 26An integrated approach to alteration intensity 33

2.6 Alteration data sheets 36

3 | COMMON ALTERATION TEXTURES AND ZONATION PATTERNS 37

3.1 Alteration textures 37Replacement textures 37

iv | CONTENTS

Infill textures 41Dissolution textures 41Static recrystallisation textures 52Dynamic recrystallisation textures 52Deformation textures 52

3.2 Pseudotextures 54Pseudoclastic textures 54False polymictic texture 63False matrix-supported texture 63False coherent textures 63

3.3 Alteration distribution 633.4 Alteration zonation patterns 64

Regional diagenetic zones 64Regional metamorphic zones 64Regional, deep, semi-conformable altered zones 66Local contact metamorphic or hydrothermally altered halos 66Local hydrothermally altered halos around ore deposits 67Vein and fracture altered halos 67

3.5 Overprinting relationships and timing of alteration 69Method 70Overprinting textures 70

4 | GEOCHEMISTRY OF ALTERED ROCKS 73

4.1 Lithogeochemistry 73Sampling and analytical methods 73Closure in composition data 78Chemostratigraphy 79

- Mass transfer techniques 81Rare-earth-element geochemistry related to alteration 87

4.2 Mineral chemistry 87Principles 87Applications 88

4.3 Stable isotopes 92Theoretical background 92Isotopic applications in alteration studies 92

5 | SEAFLOOR-AND BURIAL-RELATED ALTERATION 97

5.1 Alteration related to sea-floor processes and burial 97Physical conditions 98Definitions 98

5.2 Hydration 98Palagonite 99Perlite 100

5.3 Diagenesis (glass to zeolite facies) 102Diagenetic minerals 102Diagenetic zones 105Genesis of diagenetic minerals and zones 108

5.4 Regional metamorphism (zeolite to amphibolite facies) 115Transition from diagenesis to regional metamorphism 115Burial metamorphism 115Burial metamorphic facies 115Burial metamorphic zones 115Zeolite facies 116Genesis 116

5.5 Diagenetic alteration in the Hokuroku Basin 118Geological setting 118Alteration facies and zones 119Genesis of altered zones 120Data sheets 122

CONTENTS | V

5.6 Diagenetic alteration in the Mount Read Volcanics 128Geological setting 128Alteration fades and zonation 128Genesis of alteration fades 128Data sheets 133

6 | SYNVOLCANIC INTRUSION-RELATED ALTERATION 139

6.1 The role of intrusions in generating hydrothermal systems 140Subseafloor regional hydrothermal systems 140

6.2 Regional altered zones assodated with intrusions 141Recharge zones 141Discharge zones 141Deep, semi-conformable altered zones 142Altered zones as part of a regional hydrothermal system 147

6.3 Altered zones within intrusions 148 'Deuteric alteration 148Hydrothermal alteration 148

6.4 Contact altered halos around intrusions 149Contact altered zones 149Genesis of contact altered zones 153

6.5 Contact altered zones associated with the Darwin Granite 154Geological setting 155Alteration fades and zonation 155Genesis of the alteration system 156Data sheets 157

7 | LOCAL HYDROTHERMAL ALTERATION RELATED TOVHMS DEPOSITS 163

7.1 Common features of VHMS deposits 1637.2 Hydrothermal alteration halos associated with VHMS deposits 164

Footwall alteration pipes 164Stratabound altered footwall zones 166Altered hanging wall zones 167Chemical reactions and mass changes 168Alteration box plot trends in altered footwall zones 169The genesis of footwall alteration pipes 170Significance of pyrophyllite and kaolinite in VHMS systems 174Metamorphism of altered zones 174

7.3 The spectrum of volcanic-hosted deposits and associated alteration patterns 174Hydrothermal alteration related to the spectrum of deposits 176

7.4 Comparisons between Archaean, Palaeozoic and Cainozoic VHMSalteration systems 178

Australian Palaeozoic VHMS alteration halos 178Japanese Cainozoic VHMS alteration halos 179Canadian and Australian Archaean VHMS alteration halos 179Comparisons : 180

7.5 Hellyer: a massive elongate polymetallic lens 181Geological setting 182Alteration fades and zonation 182Ore genesis 183Data sheets 184

7.6 Rosebery: a polymetallic sheet-style deposit 194Geological setting 194Alteration facies and zonation 195Genesis of the ore lenses and alteration system 195Data sheets 196

7.7 Western Tharsis: a hybrid Cu-Au VHMS deposit 202Geological setting 202Alteration facies and zonation 202

Vi | CONTENTS

Ore genesis 203Data sheets 204

7.8 Henty: a volcanogenic gold deposit 212Geological setting 212Alteration fades and zonation 212Ore genesis 213Data sheets 214

7.9 Thalanga: a polymetallic sheet-style deposit 221Geological setting 221Alteration facies and zonation 222Ore genesis 222Data sheets 223

7.10 Highway-Reward: a pipe style Cu-Au VHMS deposit 232Geological setting 232Alteration facies and zonation 232Ore genesis 232Data sheets 233

| FINDING ORE DEPOSITS IN ALTERED VOLCANIC ROCKS : 241

8.1 Principles of discriminating between diagenetic, hydro thermaland metamorphic alteration facies 241

Diagenetic facies 241Metamorphic facies 242Hydrothermal alteration facies 242

8.2 Exploration vectors and proximity indicators 243Mineral zonation 243Major element lithogeochemistry 243Alteration indices 244Mass change vectors 245Mineral chemistry vectors 245Isotopic vectors 246

| REFERENCES 251| INDEX 271

I vii

PREFACE

Altered volcanic rocks is principally for hands-on geologists,our fortunate colleagues who practise in mineral explorationand mining geology, and the students who may in the futureplay in those professional fields.

We began designing and writing this book in mid 2001after struggling for several decades to come to terms with avariety of alteration styles in ancient submarine volcanicsuccessions. We realised that although a large number ofcompany and research geologists were working on similarrocks there was no existing text to help guide us through thecomplexity of altered volcanic rocks. The so-called volcanicrocks we deal with in ancient volcanic successions and aroundore deposits frequently bear little resemblance to their freshcounterparts, which are studied in undergraduate igneouspetrology and volcanology courses. It is typically only withlong experience that geologists develop the confidence andskills to be comfortable working with altered volcanic rocks,to interpret the original volcanic facies, unravel complexalteration histories and determine their significance in termsof mineral deposit prospectivity, particularly in ancient anddeformed successions.

The topic and content of the book were inspired byproblems that we have faced, and in many cases overcome,while working on industry-related volcanic facies, alterationgeochemistry and economic geology research projects,particularly in the Mount Read Volcanics. Many of the ideaspresented in this book come from the results of CODESresearch projects, which have been run in collaborationwith industry partners and the Australian Research Council(ARC) over the last 15 years. In particular, AMIRA-ARCLinkage project P439 (Studies of VHMS-related alteration:geochemical and mineralogical vectors to ore) provided an

enormous amount of data, case studies and expertise. Someof the results of this project have previously been published asa special issue in Economic Geology (Gemmell and Herrmann,eds., A special issue devoted to alteration associated with volcanic-hosted massive sulfide deposits, and its exploration significance,August 2001, v. 96, no. 5).

We were encouraged by the wide acceptance and successof the CODES publication by Jocelyn McPhie, MarkDoyle and Rod Allen (1993) Volcanic textures: a guide to theinterpretation of textures in volcanic rocks, which has been amajor factor in improving the description and interpretationof volcanic facies over the last decade. The advance we havemade in Altered volcanic rocks is to integrate observationsand data on volcanic facies and textures with alterationmineralogy and geochemistry at both regional and local scalesin order to provide a multidisciplinary method for the studyand discrimination of different alteration types: diagenetic,metamorphic and hydrothermal alteration.

We hope that this book will help to equip geologistsworking in altered and deformed successions with the skilland confidence to interpret the original volcanic facies andencourage the use of altered rocks as discriminants andvectors in mineral exploration. This book may not provideall the answers, but if it gives readers the courage to tackle thestudy of altered rocks, embrace the problems and pursue theanswers it will have been worthwhile.

Cathryn C. GifkinsWally HerrmannRoss R. Large

viii |

| ACKNOWLEDGEMENTS

While preparing this book, we were fortunate to have valuablesupport, assistance and advice from many people.

We extend our sincere thanks to those people whosediscussions and/or reviews of various chapters have helpedshape this book. Chapters were peer-reviewed by Stuart Bull,David Cooke, Mark Doyle, Kim Denwer, Allan Galley, BruceGemmell, Anthony Harris, Jocelyn McPhie, Andrew Rae,Mike Solomon and Fernando Tornos. Valuable discussionswere also held with Ron Berry, Stuart Bull, Jocelyn McPhie,Phil Robinson and Mike Solomon.

Although samples and photographs used herein areprincipally from the authors' collections, we also made use ofhand specimens and thin sections from the School of EarthSciences rock catalogue at the University of Tasmania, andsamples and photographs from colleagues. Thank you to thosepeople who contributed: Sharon Allen, Stuart Bull, Kate Bull,Tim Callaghan, Cari Deyell, Bruce Gemmell, George Hudak,Karin Orth and Jocelyn McPhie. We also thank Izzy vonLichtan, Curator at the School of Earth Sciences, for her helpin finding and returning hundreds of catalogue samples.

Andrew McNeill very kindly provided a long projectionof the Rosebery ore lenses. Tim Callaghan assisted with corespecimens and whole-rock geochemical data from MountJulia. Jon Huntington and Melissa Quigley at CSIROprovided HyMap® images of the Mount Lyell field.

We are infinitely grateful for the hard work of theproduction team. Karin Orth and Simon Stephens helped withsample preparation. Mike Blake and Karin Orth assisted withphotography. Rose Pongratz and Izzy von Lichtan preparedthe bibliography and checked references. June Pongratzprovided expert drafting, design and desktop publishing, andwas incredibly tolerant of the endless revisions. Final editingwas by Impress: clear communication and indexing by WordWise and Impress: clear communication.

We also appreciate our families, friends and colleagueswho have been very understanding of our commitment tothis project over the last three years. Thank you for yoursupport and patience.

r

I ix

INTRODUCTION

This book is about the processes and products of alterationin submarine volcanic successions, although many of theconcepts presented here can be applied to altered volcanicrocks from almost any environment. Its emphasis is onhydrothermal alteration associated with volcanic hostedmassive sulfide (VHMS) deposits.

Few volcanic rocks in submarine settings are entirelyunaltered, and in hydrothermal environments all rocksare altered to some degree. Recognising, describing andinterpreting altered volcanic rocks is not always easy, butthe results can have important implications for volcanology,petrology and ore deposit studies, and can improve andaccelerate success in mineral exploration. Determining pre-alteration characteristics and discriminating between primaryvolcanic, magmatic and secondary alteration features requiresknowledge of the alteration processes and their products.

Valuable base-metal, gold and silver deposits existin a variety of modern and ancient submarine volcanicsuccessions. Many of these deposits are surrounded by, orspatially related to, extensive altered zones that record thepassage of mineralising hydrothermal fluids and fluid-wallrock reactions. Research into the textural, mineralogical andcompositional effects of alteration around VHMS depositshas shown that they can be quantified and used as effectiveexploration tools for discriminating deposit styles and guidingexploration towards mineralised zones.

An introduction to alteration

Guilbert and Park (1986) defined alteration as any change inthe mineralogical composition of a rock brought about byphysical or chemical means, especially by the interaction withhot or cold aqueous solutions or gases. Alteration typicallyencompasses mineralogical changes and changes in the rocktexture and composition. Components of rocks, includingore metals, can be dissolved, replaced or recrystallised. Newminerals may precipitate and isotopic ratios may change.Porosity and permeability may be reduced or increased.Primary volcanic textures are overprinted, and may bedestroyed and replaced by new 'false' textures, or enhanced.The resulting altered rock is described as the 'alteration fades'(e.g. Riverin and Hodgson, 1980).

Thus, alteration involves complex modification of a rock.Furthermore, a rock may undergo several episodes of syn- topost-depositional alteration, not all of which are related tomineralising hydrothermal systems. Each alteration episodeis influenced by the existing texture and composition of therock, and may also overprint and modify that texture andcomposition. As a result the characteristics of altered rocksare highly variable. In ancient volcanic rocks it is a challengeto determine host volcanic facies, unravel complex alterationprocesses and interpret their significance in terms of mineralprospectivity. That challenge is the focus of this book.

How the book is organised

Altered volcanic rockshzs two main themes, which are organisedinto eight chapters: (1) it describes the basic principles behindrecognising and describing altered volcanic rocks; and (2) itdiscusses the different alteration processes that are commonin submarine volcanic successions and their products.

Chapter 1 introduces the concepts of alteration insubmarine volcanic successions and summarises the mainalteration processes and volcanic facies. It outlines the regionalgeology of two of the most productive Australian submarinevolcanic successions: the Cambrian Mount Read Volcanicsin western Tasmania and the Cambro-Ordovician MountWindsor Subprovince in Queensland. This book principallyemploys examples from these two successions, and includesdescriptions of other ancient submarine volcanic successionsfor comparison.

Chapter 2 discusses alteration nomenclature, mineralogy,intensity and indices, and the principles of alteration facies.It proposes an integrated multidisciplinary approach todescription and classification. The main elements of alterationfacies — mineral assemblage, intensity, texture, distribution,zonation and timing — are described in Chapters 2 and 3.Chapter 4 outlines geochemical methods used in alterationstudies and their applications. It emphasises whole-rocklithogeochemistry, mineral chemistry and stable isotopeanalysis.

Chapter 5 concentrates on regional alteration stylesincluding hydration, diagenesis and metamorphism associatedwith burial, Chapter 6 on intrusion-related alteration styles,

X I INTRODUCTION

and Chapter 7 on hydrothermal alteration and mineralisationassociated with VHMS deposits. We present short case studiesfor these different alteration styles, emphasising hydrothermalalteration associated with a variety of VHMS depositsincluding Rosebery, Hellyer, Henty and Thalanga. Thesecase studies incorporate pictorial data sheets that present themineralogical, textural and compositional characteristics ofeach of the main alteration facies. They combine volcanicfacies analysis and alteration mineral assemblages, textures,intensity and geochemistry to interpret the features ofdifferent alteration styles.

Chapter 8 outlines the methods for discriminatingbetween the products of mineral deposit-related hydrothermalalteration and other alteration processes, and identifyingfavourable altered zones for mineral exploration. It alsodiscusses geochemical vectors that may guide explorerstowards mineralised rock within these altered zones.

Significance of altered volcanic rocks to mineralexploration

Economic geologists are particularly interested in alterationbecause hydrothermal mineral deposits are commonly hostedby altered rock. Hydrothermally altered zones around mineraldeposits provide much larger targets for mineral explorationthan the deposits themselves. The mineral assemblages,and in some cases the chemical composition, of the alteredrocks may provide indications of the proximity of an oredeposit, and thus vectors towards mineralised rock. Inaddition, mineralogical, textural and compositional studiesof alteration facies can provide important constraints onthe timing, physical and chemical conditions, and originsof hydrothermal systems and related mineralisation (Barnes,1979). The texture and distribution of alteration facies canalso be used to infer changes in porosity, permeability andfluid pathways in the host succession. The results of alterationstudies are commonly incorporated into ore deposit modelsused in mineral exploration. Thus, the identification andinterpretation of alteration facies is, and should be, a routinepart of exploration for hydrothermal mineral deposits.

.

11

1 | ALTERATION IN SUBMARINE VOLCANICSUCCESSIONS

This chapter describes submarine volcanic successions,the common processes of alteration that occur in thesesuccessions, and provides two examples of ancient submarinevolcanic successions, that have been variably altered andmineralised: the Mount Read Volcanics and the MountWindsor Subprovince.

In submarine volcanic environments, the coincidenceof magmatic fluids, heat and abundant seawater generateshydrothermal convection. As a consequence, submarinevolcanic successions may host important hydrothermalmineral deposits, commonly referred to as volcanic-hostedmassive sulfide (VHMS) deposits.

VHMS deposits are a significant source of zinc, copper,lead, silver and gold, and continue to be a target for base-metalexploration. They range in size from less than one milliontonnes to over 200 million tonnes, and commonly containhigh metal grades. For example, the Hellyer deposit in westernTasmania produced 16.2 Mt at 13.9% Zn, 7.1% Pb, 0.4%Cu, 168 g/t Ag and 2.5 g/t Au in its nine years of operation.VHMS deposits occur mainly in submarine rift environ-ments particularly back arc and mid ocean rifts; however,they can occur in a variety of other submarine environmentsincluding continental rifts, oceanic basins or plateaux, andarc-continent or continent-continent collision zones. They areone of the few classes of ore deposits that exist throughout thegeological record from early Archaean to Recent.

1.1 | SUBMARINE VOLCANICSUCCESSIONS

Submarine volcanic successions are significantly different fromsubaerial volcanic successions, as the processes of eruption,transport, emplacement, and post-emplacement alterationmay be strongly affected by the presence of water. Typically,submarine volcanic successions comprise a wide variety ofcoherent and volcaniclastic facies intercalated with mixedprovenance and non-volcanic sedimentary facies (Fig. 1.1).The volcanic facies may be derived from intrabasinal, extra-basinal or basin-margin eruptions in submarine or subaerialsettings. Eruption styles may be effusive or explosive and theproducts may remain in situ or be redeposited or reworked

by sedimentary processes. In addition, volcanic units may beemplaced into the succession as synvolcanic intrusions.

This section summarises the main volcanic facies thatoccur in submarine volcanic successions. For a more detaileddiscussion of submarine volcanism, volcanic textures, faciesand their interpretation, readers are referred to McPhie et al.(1993) Volcanic textures: a guide to the interpretation of texturesin volcanic rocks.

Volcanic facies

For descriptive purposes, volcanic facies are divided into twomain textural types: coherent and volcaniclastic. Coherentfacies consist of solidified magma and are commonlycharacterised in volcanic rocks by aphyric (fine grained orglassy) or porphyritic textures, where porphyritic refers toevenly distributed euhedral crystals (phenocrysts) in a fine-grained or glassy groundmass (McPhie et al., 1993).

Volcaniclastic facies are those composed mainly of volcanicparticles (Fisher, 1961). Volcanic particles are crystals, crystalfragments, shards, pumice clasts, scoria clasts and densevolcanic clasts, which may be produced by primary volcanic(pyroclastic and autoclastic) or sedimentary (weathering anderosion) processes. Volcaniclastic facies include a spectrumof facies: primary volcanic facies, syneruptive volcanicfacies generated by coeval eruptions and deposited fromsedimentary processes, and volcanogenic sedimentary faciesthat show evidence of temporary storage and reworking priorto deposition (McPhie et al., 1993).

Primary volcaniclastic facies result from volcanic processesof clast formation, transport and deposition and includehydroclastic, pyroclastic and autoclastic facies. Hydroclasticfacies is a general term for facies, typically comprising blockyglassy particles, produced by magma-water interactions,whether by explosive steam generation or by non-explosivequench fragmentation of magma (Fisher and Schmincke,1984; Hanson, 1991). Pyroclastic facies comprise volcanicparticles (pyroclasts) that were generated by explosiveeruptions and deposited by primary volcanic processes, byfallout, flow or surge. Autoclastic facies comprise volcanicparticles generated by in situ non-explosive fragmentation oflava or magma (autobrecciation and quench fragmentation).

2 | CHAPTER 1

Autobrecciation occurs when the more viscous parts of amoving lava respond in a brittle fashion to locally higherstrain rates, and fragment into blocky clasts (Fisher, I960).Quench fragmentation occurs in situ where hot lava ormagma comes into contact with water, ice or water-saturatedsediment (Rittmann, 1962; Pichler, 1965; Yamagishi, 1987).The resulting autoclastic deposits - autobreccia, hyaloclastiteor peperite - typically comprise dense blocky or splinteryclasts, but they may be pumiceous and have fluidal shapes(Fisher, 1960; Pichler, 1965; Busby-Spera and White, 1987;Gifkins et al., 2002).

Syneruptive volcaniclastic fades are composed dominantlyof unmodified volcanic clasts that were fragmented byvolcanic process such as explosive eruptions, autobrecciationor hydration, but were transported and deposited bysedimentary processes (McPhie et al., 1993; McPhie andAllen, 2003). They can occur directly from eruption whenclasts bypass initial deposition as primary deposits and aredelivered directly to sedimentary transport and depositionsystems, such as subaqueous eruption-fed water-supportedgravity currents or water-settled fall (e.g. White, 2000;McPhie and Allen, 2003). They may also occur indirectlyby rapid remobilisation and redeposition during or shortlyafter the eruption (Fisher and Schmincke, 1984; Cas andWright, 1987; McPhie et al., 1993). Unconsolidated volcanicdebris may be remobilised by: the slumping and sliding ofgravitationally unstable rapidly accumulated clastic debris;explosive eruptions; local uplift; syn-depositional faulting;and extrusion and intrusion of magma.

Volcanogenic sedimentary fades (epiclastic volcanic, Fisher,1960) contain volcanic particles derived from the post-eruptive erosion and reworking of pre-existing volcanic faciesand may include a significant proportion of non-volcanicparticles (McPhie et al., 1993).

In submarine volcanic successions, volcaniclastic facies aredominated by in situ autoclastic and syneruptive volcaniclasticfacies where the particles were derived from either autoclasticfragmentation or explosive eruption. Most volcanic and non-volcanic clastic deposits were emplaced by water-supporteddensity currents (i.e. high- and low-concentration turbiditycurrents, debris flows and grain flows) and as fallout fromsuspension in the water column.

Evidence for submarine environment ofemplacementVHMS deposits occur in submarine volcanic successions,thus exploration for new deposits is restricted to submarinesuccessions. However, there are few volcanic or sedimentaryfacies that unequivocally constrain the host depositionalenvironment.

A subaqueous setting (marine or lacustrine) may beinterpreted based on the presence of: water-supported mass-flow deposited facies; hemi-pelagic, biogenic, biochemicaland chemical sedimentary facies; pillow lavas; and quenchfragmented volcaniclastic facies. Also seawater-related dia-genetic alteration facies (e.g. widespread albite alterationfacies) can suggest a submarine environment.

Without fossil evidence, differentiating between marineand lacustrine settings is difficult as few facies are restrictedto either environment. Facies with tidal and wave tractionalstructures, such as bimodal-bipolar ripples, are submarine,whereas lacustrine settings may be indicated by the presenceof evaporites. Hummocky cross stratification is more commonin, although not restricted to, marine settings. Carbon-oxygenisotope signatures of carbonates can be used to support marineor lacustrine environments.

Although bedforms, sedimentary structures and somesedimentary deposits help us to interpret a marine environ-ment, they are of little help in constraining the water depth.Water depth may be an important consideration for mineralexploration as recent research suggests that Au-rich VHMSdeposits are restricted to shallow water environments (e.g.Hannington et al., 1999; Hannington and Herzig, 2000;Herzig et al., 2000). Shallow water environments are typicallydominated by the tractional processes of tidal and waveaction and result in characteristic sedimentary structuresand bedforms. In contrast, deep water environments, belowstorm wave base, generally lack tractional currents: sedimentdistribution and deposition mainly occurs through theactions of turbidity currents, debris flows and the processof suspension sedimentation. Water depth and depositionalsetting may be more accurately constrained by the presenceof fossiliferous limestone or sedimentary facies that containmarine fossils intercalated in the volcanic succession.

Volcanic facies associations

A facies association is a collection of facies that are spatially,mineralogically, compositionally or texturally related, andthat may also be genetically related (Cas and Wright, 1987).There are three common types of volcanic units representedby facies associations in submarine volcanic successions(Fig. 1.1): lavas, synvolcanic intrusions (cryptodomes, sillsand dykes) and syneruptive volcaniclastic facies. Lavas andsynvolcanic intrusions comprise associations of coherent andautoclastic facies. The syneruptive volcaniclastic facies can bedivided into two principal categories, those dominated bynon- to poorly-vesicular blocky lava clasts and related to thesubmarine emplacement of lavas and lava domes, and othersthat contain abundant pumice or scoria clasts produced byexplosive eruptions. In addition, there is a wide variety ofvolcanogenic sedimentary facies.

1.2 | ALTERATION IN SUBMARINEVOLCANIC SUCCESSIONS

After emplacement, volcanic facies are commonly subjectedto a variety of alteration processes (Fig. 1.2). Alteration occurswhen existing components become unstable under changingphysical and chemical conditions, and alter to more stableminerals. Volcanic glass, which is the main component ofmany volcanic facies, is a metastable solid with the structureof a liquid (Carmichael, 1979). It is undercooled to the pointwhere extreme viscosity has prevented crystallisation. As aresult, volcanic glass readily devitrifies to minerals that aremore stable under surface conditions; generally clay minerals,zeolites, carbonates, feldspar, quartz and oxides (Carmichael,1979; Henley and Ellis, 1983; Fisher and Schmincke, 1984;

ALTERATION IN SUBMARINE VOLCANIC SUCCESSIONS | 3

FIGURE 1.1 | Facies model of a submarine basin in which a variety of coherent and clastic volcanic facies are intercalated with sedimentary facies. The volcanicfacies include primary coherent and autoclastic facies, syneruptive and post-eruptive volcaniclastic facies derived from submarine and subaerial eruptions. Many ofthe volcanic facies associations are laterally discontinuous. Common facies associations represent (A) lavas and lava domes composed of coherent and autoclasticfacies; (B) synvolcanic sills and cryptodomes; (C) syneruptive volcaniclastic facies derived from explosive and effusive submarine eruptions; (D) volcanogenicsedimentary or resedimented volcaniclastic facies derived from pre-existing deposits; (E) syneruptive volcaniclastic facies derived from subaerial explosive eruptions;(F) mixed provenance sedimentary facies; and (G) marine sedimentary facies.

FIGURE 1.2 | Facies model showing the distribution of different styles of altered zones in a submarine volcanic succession that hosts VHMS deposits. See Figure

1.1 legend for the patterns denoting volcanic and sedimentary facies.

4 | CHAPTER 1

Friedman and Long, 1984; Cerling et al., 1985). Alterationof volcanic glass involves not only devitrification, but changesin texture, composition, porosity and permeability, and maysimultaneously affect both the chemistry and circulation ofpore fluid in the volcanic succession (Noble, 1967; Dimrothand Lichtblau, 1979; Fisher and Schmincke, 1984; Noh andBoles, 1989; Torres et al., 1995).

Understanding alteration requires a range of skills thatinclude recognising alteration minerals, textures, paragenesis,distribution, zonation, intensity, mineralogical and chemicalchanges associated with alteration, pathways and mechanismsfor fluid migration, and fluid origin. These characteristics arerelated to the alteration processes and to the characteristics ofthe host volcanic succession.

Devitrification

The cooling history of volcanic facies may involve primarycrystallisation and later devitrification. Primary or high-temperature crystallisation refers to crystallisation of magmaresulting in phenocrysts, microcrysts and microlites. Incontrast, devitrification refers to crystallisation of glass at lowtemperatures (i.e. below the glass transition temperature:Lofgren, 1971a). High-temperature devitrification accom-panying first cooling may produce spherulites, lithophysaeand micropoikilitic or snowflake texture (e.g. Lipman, 1965;Anderson, 1969; Lofgren, 1971b; Bigger and Hanson,1992; McArthur et al., 1998) and is not considered to bealteration. Low-temperature devitrification results in thegradual conversion of glass to fine-grained granular crystallineaggregates, which may happen over time as a result ofalteration during changing physical conditions or in responseto interaction with fluid. Devitrification may be accompaniedby changes in whole-rock composition (Lipman, 1965;Lofgren, 1971a; Friedman and Long, 1984).

the presence of fluid (seawater, magmatic fluid or a mixture ofboth). There are gradations from isochemical metamorphismto metasomatism with increasing compositional changes.The different alteration processes, hydration, diagenesis,metamorphism and local hydrothermal alteration, are all partof this continuum in submarine volcanic successions (Fig.1.3).

The effects of each alteration process may be difficultto distinguish. Hydrothermal alteration, diagenesis andmetamorphism can result in similar mineral assemblagesand textures. In addition, in many cases, different alterationprocesses, such as diagenetic and hydrothermal alteration,are contemporaneous and their products may be inseparable(Iijima, 1974, 1978; Ohmoto, 1978; Reyes, 1990; Utada,1991;Paradisetal., 1993).

In Chapters 5, 6, and 7 of this book, the commonalteration processes are grouped into those related to burial,intrusions and VHMS deposits (Fig. 1.4). Thus, burial-related alteration styles include hydration, diagenesis andburial metamorphism. Alteration styles associated withintrusions are hydrothermal alteration within intrusions,contact metamorphism and hydrothermal alteration, andregional hydrothermal alteration. Included below is a briefintroduction to each of the common alteration processes thatoperate in submarine volcanic settings.

Hydration of volcanic glass

Hydration of glass involves the absorption of externalwater into glass and modification of the glass structure,either during cooling or at ambient temperatures (Ross andSmith, 1955; Friedman and Long, 1984). Hydration doesnot directly produce new minerals, but can form perliticfractures or palagonite in basaltic glass and it can facilitatesubsequent alteration (see Chapter 5). Compositional changes

Alteration processes

Alteration may result from regional or local processes. Itcan occur as a result of the interaction with hydrothermalfluid, as a result of changing physical (mainly temperatureand pressure) conditions during burial, in association withthe emplacement of intrusions, or a combination of all theseprocesses. Submarine volcanic facies, especially glassy facies,are readily altered during hydration, diagenesis, hydrothermalalteration, metamorphism and tectonic deformation.

Hydrothermal alteration is defined as the alteration ofrocks or minerals by the reaction of hydrothermal fluidwith pre-existing solid phases (Henley and Ellis, 1983).Hydrothermal fluid is a hot aqueous solution or gas, with orwithout demonstrable association with igneous processes.Hydrothermal alteration usually results in significant changesin rock texture, mineralogy and composition.

Alteration is either metasomatic or isochemical. Meta-somatism involves changes in mineralogy, texture andcomposition, whereas isochemical alteration (or meta-morphism) involves mineralogical and textural changesonly. In submarine volcanic successions, almost all alterationinvolves some degree of metasomatism, which is facilitated by

FIGURE 1.3 | This cartoon depicts the continuum between isochemicaland hydrothermal alteration and shows the alteration processes common insubmarine volcanic successions. They are positioned based on the relativedegrees of chemical exchange for each process.

accompanying hydration include gains in H2O, and minorlosses in silica and alkalis (Noble, 1967; Friedman and Long,1984; Mungall and Martin, 1994).

Diagenesis

Diagenesis encompasses the changes that occur in responseto changing temperature and pressure during burial.During diagenesis of volcanic facies, significant textural andmineralogical changes can be produced by precipitation ofcement, dissolution and replacement of original components,especially glass, and compaction (Fisher and Schmincke,1984; Marsaglia and Tazaki, 1992; Torres et al., 1995). Intheory, diagenesis in submarine volcanic successions is ametasomatic process involving minor chemical exchangebetween the host facies and trapped modified seawater at lowtemperatures (<250°C). Transitions between diagenesis andmetamorphism, and diagenesis and hydrothermal alterationhave not been rigorously defined and are discussed in Sections5.4 and 6.2.

Regional metamorphism

Regional metamorphism involves pervasive, mainly iso-chemical, mineralogical and textural changes in response toincreasing pressure and temperature (Yardley, 1989). Duringmetamorphism, H2O and CO2-bearing fluids are generatedby dehydration and decarbonation reactions (Rose and Burt,1979).

Contact alteration associated with intrusions

Contact alteration refers to the changes caused by the rise intemperature of the host rock immediately surrounding an

intrusion, which may be accompanied by the circulation ofheated pore fluids around and within the intrusion. Contactalteration can be isochemical (i.e. contact metamorphism)or metasomatic (i.e. hydrothermal alteration). Contactmetamorphism typically results in recrystallisation of existingminerals or components and minor remobilisation of elements(Yardley, 1989). The effects of hydrothermal alterationmay include major changes in texture, mineral assemblageand whole-rock composition on a scale of centimetres tokilometres.

Hydrothermal alteration related to VHMS deposits

Two styles of hydrothermal alteration are commonly relatedto VHMS mineralisation: (1) local alteration halos aroundore deposits; and (2) regional hydrothermally altered zones.Regional hydrothermally altered zones are commonly spatiallyand genetically associated with large intrusions and hencein this book are discussed in Chapter 6 - intrusion-relatedalteration styles.

Local hydrothermally altered halos around VHMSdeposits result from the reaction between the host facies andthe mineralising hydrothermal fluid (Sangster, 1972; Franklinet al., 1975; Riverin and Hodgson, 1980; Green et al., 1983;Urabe.et al., 1983). These altered halos are commonly zoned,reflecting changes in the composition, pH and temperatureof the hydrothermal fluid with time, or the extent of reactionwith the host facies (Rose and Burt, 1979; Lydon and Galley,1986; Schardt et al., 2001).

The nature of altered halos around VHMS deposits dependson the host volcanic facies, host-rock composition, timing ofthe hydrothermal alteration relative to the emplacement ordeposition of facies, structures, fluid pathways, distributionpattern of the ore, and chemical and physical conditions of thehydrothermal fluid (Large, 1992). Thus altered halos aroundVHMS deposits exhibit a wide variety of geometries, sizes,

FIGURE 1.4 | Common alteration processes and their products.


I CHAPTER 1 - "alteration mineral assemblages, intensities, compositions, andzones (Chapter 7).

Thick, extensive (up to 20 km), pervasive sub-horizontalsemi-conformable altered zones, referred to as regionalhydrothermal alteration zones or deep semi-conformablealteration zones (Section 6.2), have been recognised in somevolcanic successions hosting VHMS deposits (Gibson et al.,1983; Galley, 1993). Many of these regional hydrothermallyaltered zones are spatially associated with, and possiblygenetically related to, intrusions (Galley, 1993; Brauhart etal., 1998). They are interpreted to result from large-scaleconvection of seawater through permeable volcanic successionsat elevated geothermal gradients (Spooner and Fyfe, 1973;Munha and Kerrich, 1980; Baker, 1985). Reactions betweenthe volcanic successions and modified seawater have involvedNa-, Si-, Ca-Fe-, K-, or Mg-metasomatism, and the leachingof ore-forming metals (Munha and Kerrich, 1980; Gibson etal., 1983; Baker, 1985; Galley, 1993).

Syntectonic hydrothermal alteration

Hydrothermal alteration can also be synchronous withtectonic deformation: syntectonic hydrothermal alteration. Inthis case, the hydrothermal fluid may be modified seawater,magmatic water or volatiles released during metamorphism,or a combination of these, and migrates principally alongfaults and shear zones. Contemporaneous deformationmay modify or destroy pre-existing textures and create newtextures or foliations.

Many VHMS deposits, such as Rosebery, Hercules andMount Lyell (western Tasmania), have been affected bylater tectonic deformation and modified by syntectonichydrothermal fluids (Walshe and Solomon, 1981; Khin Zawand Large, 1992). Although syntectonic hydrothermal fluidswere not responsible for the formation of these VHMS ores,they can be critical to the subsequent formation of other stylesof ore deposits in the submarine volcanic successions, such asmesothermal gold deposits. Detailed discussion of syntectonichydrothermal alteration is not dealt with in this book.

Characteristics inherited from volcanic facies

Volcanic components and facies with different compositionsand textures behave differently during the initial stages oflow-temperature alteration. Some components react morerapidly than others and their composition may influencethe composition of the alteration mineral assemblage. Forexample, Marsaglia and Tazaki (1992), in their study ofdiagenetic trends in volcaniclastic sandstones of the Izu-Bonin Arc, found that black mafic crystalline fragments wereunaltered, brown intermediate to mafic glassy fragmentsshowed evidence of dissolution, and colourless rhyoliticfragments were altered to clay minerals. These differenceswere due to the variable reactivity of the components and theproportions of volcanic glass to crystalline facies.

Generally glass is the most reactive component, followedby olivine —> magnetite, titanomagnetite and ilmenite —*pyroxene and amphibole —> biotite —* Ca-plagioclase —*microcline, sanidine and orthoclase —* quartz, apatite, rutile

and zircon (Browne, 1978; Reyes, 1990). Alteration ratesfor different minerals vary considerably because of mineralstructure and composition. Silicate materials with an extensivecross-linked (e.g. tetrosilicate) structure react slowly; whereasthose silicates with poorly connected fabric tend to reactrapidly and uniformly (Casey and Bunker, 1990).

Felsic volcanic facies typically have a higher proportion ofglassy to crystalline facies than mafic facies. This is becausethe viscosity of silicic magmas (71-77% SiO2) inhibitsdiffusive crystal growth and thus produces thick bodies ofglass, whereas the low viscosity of basaltic magmas favourscrystallisation (Friedman and Long, 1984).

Glassy facies or facies that contain glassy fragments arelikely to be more rapidly altered, and to form different mineralassemblages, than those that are crystalline (Lee and Klein,1986). In addition, mafic glasses are more rapidly altered thanfelsic glasses (e.g. Whetten and Hawkins, 1970; Fisher andSchmincke, 1984; Friedman and Long, 1984). The rate ofalteration is related to the glass's viscosity, which in turn isa function of the composition (particularly SiO2 and H2O)and temperature (Friedman and Long, 1984). Increased SiO2

decreases the rate of alteration, whereas increased MgO, CaOand H2O increases the rate. Thus the higher SiO2 content ofrhyolitic glasses retards reaction (Hawkins, 1981).

The primary composition can influence the alterationmineralogy mainly because the dissolution of glass liberatesalkalis and silica, which are consumed by subsequentreactions. Hence, highly silicic volcanic facies result in thecrystallisation of Si- and Na-rich minerals, such as opal,quartz, tridymite, cristobalite and Na-zeolites (Sheppard etal., 1988). In contrast, mafic glasses, such as those on pillowrims, alter to Ca-, Fe-, Mg- and Mn-rich minerals such assmectites, phillipsite, oxides and chlorite.

Volcaniclastic facies, particularly pumice-rich facies,initially have very high porosities. In volcaniclastic facies, theinter- and intra-particle pore space controls the porosity andpermeability and thus grain size, type and sorting influencethe distribution of early alteration facies. Early diageneticalteration in well-sorted pumice breccias, although commonlypatchy, is pervasive. In poorly sorted polymictic volcaniclasticfacies the porosity and permeability are initially muchmore variable and diagenetic facies typically have complexdistribution patterns.

Coherent facies have much lower porosity and permea-bility, factors that are controlled by fractures produced byquenching, flowage and hydration. Alteration in coherentfacies typically progresses as fronts that move outward fromfractures into the less altered domains (e.g. the alteration ofperlite, Noh and Boles, 1989).

In addition, patchy or domainal alteration styles involcaniclastic facies may also be related to variations in thequenching and hydration of glassy clasts (Surdam, 1973;Boles and Coombs, 1977; Marsaglia and Tazaki, 1992).Hydrothermal experiments on rhyolitic glass indicate thatat high temperatures (>200°C), rhyolitic glass does notrecrystallise but instead acts as Na-K ion exchanger. Quenchedglass fixes K+, whereas slowly cooled glass fixes Na+. Thereforevariations in cooling history may explain why some glassyfragments alter more readily to particular minerals than others(Marsaglia and Tazaki, 1992).

6 | CHAPTER 1

1.3 | GEOLOGY OF THE MOUNT READVOLCANICS

Many of the examples of altered volcanic rocks and alterationsystems discussed in the following chapters come from theMiddle to Late Cambrian Mount Read Volcanics in westernTasmania. The Mount Read Volcanics are a submarinesuccession of rhyolitic to basaltic volcanic and intrusiverocks with variable proportions of intercalated sedimentaryrocks. They are interpreted as the products of post-collisionalvolcanism associated with arc-continent collision (Berry andCrawford, 1988; Crawford and Berry, 1992). The successionoccurs in a 200 x 20 km area that extends from Elliott Bay in


the south through Queenstown and Rosebery to Deloraine inthe north (Fig. 1.5). The volcanic succession was depositedin a series of troughs separated by areas of Proterozoicbasement (Corbett and Lees, 1987; Corbett, 1992; Crawfordand Berry, 1992). The Mount Read Volcanics host gold,silver and base-metal massive sulfide (VHMS) ore depositsat Hellyer, Que River, Rosebery, Hercules, Henty and MountLyell (Fig. 1.5). The mineral district is referred to as theMount Read province.

The Mount Read Volcanics have undergone diageneticand hydrothermal alteration, metamorphism, at least twophases of deformation, and intrusion by Cambrian andDevonian granites (Corbett and Lees, 1987; Corbett, 1992).

FIGURE 1.5 | Distribution of the principal lithostratigraphic units and major massive sulfide deposits in the central Cambrian Mount Read Volcanics, westernTasmania. Modified after Corbett (1992,2002).

8 I CHAPTER 1

FIGURE 1.6 | Stratigraphic correlation diagram showing the major lithostratigraphic units in the Mount Read Volcanics to the west (A) and east (B) of the Hentyfault. The sections are located at (1) Hellyer-Que River, (2) Pinnacles, (3) Hollway, (4) Mount Black, (5) Rosebery-Hercules, (6) White Spur, (7) Hall Rivulet Canal,(8) Murchison Gorge, (9) Henty, (10) South Henty, (11) Anthony Road, (12) Comstock-Lyell, (13) Lynchford, (14) Jukes-Darwin. Sections are modified after Fitzgerald(1974), Corbett (1979,1992, 2001), Cox (1981), Komyshan (1986), Coutts (1990), Allen (1991), Dugdale (1992), Waters and Wallace (1992), Jones (1993,1999),McKibben (1993), Herrmann and MacDonald (1996), McPhie (1996), White and McPhie (1996,1997), Callaghan (2001), Gifkins (2001) and Wyman (2001).

(A) To the west of the Henty fault, the Central Volcanic Complex interfingers with, and is conformably and disconformably overlain by, the Dundas and Mount CharterGroups of the western volcano-sedimentary sequences (Corbett and Lees, 1987; Corbett, 1992). Immediately overlying the Central Volcanic Complex is a varietyof small-volume sedimentary (Black Harry beds and Animal Creek greywacke) and volcanic units (rhyolite and pumice breccia). These units are overlain by theQue-Hellyer Volcanics, a succession of calc-alkaline to shoshonitic, intermediate to mafic lavas and volcaniclastic units (Corbett and Komyshan, 1989; Waters andWallace, 1992). The Que-Hellyer Volcanics host the Que River and Hellyer ore deposits, and extend via Sock Creek to Burns Peak and Pinnacles (i.e. the Brown'stunnel sequence). The Que River Shale overlies the Que-Hellyer Volcanics and is similar to mudstone in other lithostratigraphic units of the Mount Read Volcanics.The Southwell Subgroup overlies the Que River Shale and is lithologically similar to the White Spur Formation and the Rosebery hanging-wall volcaniclastic units,comprising quartz-bearing volcaniclastic mass-flow units interbedded with black mudstone and Precambrian basement-derived turbidites (Corbett, 1992; McPhie andAllen, 1992). Overlying the Southwell Subgroup is the Mount Charter Group in which the upper Mount Cripps Subgroup is a correlate of the Tyndall Group (Corbett,1992).

Although the primary textures, mineralogies and whole-rockcompositions have been modified to various degrees, volcanictextures are generally well preserved. Locally, two regionaltectonic cleavages have been recognised; however, the axialplanar S2 Devonian cleavage is the dominant cleavage. S2

strikes north, dips steeply and varies from a weak, spacedcleavage to an intense, pervasive, anastomosing cleavage inthe most strongly deformed rocks adjacent to faults and inphyllosilicate-rich altered zones.

The geology of the Mount Read Volcanics has beendescribed in detail by Campana and King (1963), Corbett(1981, 1986, 1992, 1994), Corbett and Lees (1987), Corbettand Solomon (1989), Pemberton and Corbett (1992), McPhieand Allen (1992) and Crawford et al. (1992).


B

(B) To the east of the Henty fault, the Eastern quartz-phyric sequence overlies the Sticht Range Beds, interfingers with the Central Volcanic Complex and isconformably overlain by the western volcano-sedimentary sequences (Farrell Slates). The southern Central Volcanic Complex is flanked to the west by the YolandeRiver Sequence, part of the western volcano-sedimentary sequences. To the east, it interfingers with the Eastern quartz-phyric sequence and is overlain by theTyndall Group (Corbett, 1992) and locally by andesite and basalt lenses that occur between Henty and Queenstown (Anthony Road andesite, Crown Hill andesite,Howards basalt, Spillway basalt, and Lynchford basalt). The Cambrian Murchison and Darwin granites intruded the succession in the Mount Murchison and MountDarwin areas (Corbett and Lees, 1987; Corbett, 1992). The Tyndall Group is the youngest lithostratigraphic unit. It extends north-south along the eastern margin of thesuccession where it unconformably overlies Tyennan basement, Sticht Range Beds, southern Central Volcanic Complex, Eastern quartz-phyric sequence, westernvolcano-sedimentary sequences, and the Darwin Granite (Corbett and Lees, 1987; White and McPhie, 1997). The Owen Conglomerate overlies the Mount ReadVolcanics both conformably and unconformably.

Stratigraphy of the Mount Read Volcanics

The stratigraphy of the Mount Read Volcanics can be dividedinto (Figs 1.5 and 1.6): Sticht Range Beds, Eastern quartz-phyric sequence, Central Volcanic Complex, western volcano-sedimentary sequences, and the Tyndall Group and correlates(Corbett, 1992). These lithostratigraphic units comprisecompositionally and texturally diverse coherent volcanic andvolcaniclastic facies intercalated with sedimentary rocks, whichare distinguished and mapped on the basis of the dominantfacies. The principal volcanic facies associations are lavas,synvolcanic intrusions and syneruptive volcaniclastic units(McPhie and Allen, 1992). Lavas and synvolcanic intrusions

are predominantly calc-alkaline rhyolites and dacites withlocally abundant andesites and basalts (Crawford et al.,1992). The volcaniclastic facies associations include a varietyof primary and secondary volcaniclastic facies includingthick extensive syneruptive pumice- and crystal-rich unitsand in situ and resedimented hyaloclastite (McPhie andAllen, 1992). Sedimentary facies include black mudstone,and graded, bedded sandstone of mixed volcanic and meta-sedimentary Precambrian basement provenance (McPhie andAllen, 1992).

Regional stratigraphic relationships between the litho-stratigraphic units are complex and laterally variable(Fig. 1.6). The Mount Read Volcanics are conformably and

10 | CHAPTER 1

unconformably overlain by the Owen Group, a thick sequenceof Late Cambrian-Early Ordovician siliciclastic, shallowmarine to fluvial conglomerate and sandstone (Corbett,1992; White, 1996).

VHMS deposits occur in a variety of volcanic facies inthree of the main lithostratigraphic subdivisions of theMount Read Volcanics (McPhie and Allen, 1992; Pembertonand Corbett, 1992; Waters and Wallace, 1992). In particular,VHMS deposits are interpreted to occur: (1) at the top of theCentral Volcanic Complex, close to large felsic volcanic centres(Rosebery, Hercules and Mount Lyell); (2) associated withproximal facies of andesite-dacite volcanoes in the westernvolcano-sedimentary sequences (Hellyer and Que River); and(3) in the base of the Tyndall Group (Henty and Comstock)(Corbett and Solomon, 1989; Halley and Roberts, 1997).

Sticht Range Beds

The Sticht Range Beds are a thin (<500 m) basal successionof interbedded basement-derived sedimentary rocks andvolcaniclastic rocks that occur along the eastern margin of theMount Read Volcanics (Fig. 1.5: Baillie, 1989).

Eastern quartz-phyric sequence

The Eastern quartz-phyric sequence is a 2.5 km-thicksuccession of quartz + feldspar-phyric lavas, synvolcanicintrusions and volcaniclastic units limited to the easternmargin of the Mount Read Volcanics (Fig. 1.5: Polya, 1981;Polya et al., 1986; Pemberton et al., 1991; Corbett, 1992).

Central Volcanic Complex

The 3 km-thick Central Volcanic Complex dominates thecentral part of the Mount Read Volcanics between MountDarwin and Mount Block (Fig. 1.5: Corbett, 1979). It consistsof feldspar-phyric rhyolitic and dacitic lavas, synvolcanicintrusions and pumiceous volcaniclastic units (Corbett, 1979,1992; Corbett and Lees, 1987; Corbett and Solomon, 1989;Pemberton and Corbett, 1992; Gifkins, 2001). Andesitesand basalts are locally intercalated with the felsic succession(Crawford et al., 1992). Quartz + feldspar-phyric intrusionsand tholeiitic basalt and dolerite dykes (Henty dyke swarm)occur throughout the northern Central Volcanic Complex(Corbett and Solomon, 1989; Crawford et al., 1992).

Western volcano-sedimentary sequences

The western volcano-sedimentary sequences include theYolande River Sequence, Dundas Group, Mount CharterGroup and Henty fault wedge sequence (Corbett, 1992).These sequences are thick (>3 km) mainly sedimentarysuccessions of quartz + feldspar-phyric volcaniclastic facies,mixed provenance sandstone and mudstone intercalatedwith rhyolitic, andesitic and basaltic lavas and synvolcanicintrusions, (Corbett and Lees, 1987; Corbett, 1989; McPhieand Allen, 1992). The volcaniclastic facies contain a diverse

range of clasts including quartz + feldspar porphyry, feldspar-phyric rhyolite, pumice, granite and massive sulfide clasts.

Tyndall Group and correlates

The Tyndall Group is the youngest lithostratigraphic unitin the Mount Read Volcanics. It extends along the easternmargin and locally along the western side of the succession(Fig. 1.5). The Tyndall Group varies in thickness from 350 to1300 m and comprises distinctive quartz + feldspar crystal-rich sandstone, volcanic breccia and volcanic conglomerateintercalated with minor rhyolitic welded ignimbrite, felsicto intermediate lavas and intrusions, and non-volcanicsedimentary rocks including limestone, mudstone andsandstone (White and McPhie, 1996, 1997).

Cambrian granites

Five Cambrian granitoids (commonly referred to as 'granites')have been recognised in western Tasmania: the Murchison,Darwin, Elliott Bay, Dove and Timber Tops granites (Leamanand Richardson, 1989). Cambrian granites may also occurat depth in a belt that extends along the eastern margin ofthe Mount Read Volcanics between Mount Darwin andMount Murchison (Large et al., 1996). They are typicallymedium grained quartz + K-feldspar + plagioclase + biotite +hornblende + apatite + zircon ± rutile granite or granodiorite(McNeill and Corbett, 1992). They intrude the westernvolcano-sedimentary sequences, Central Volcanics Complexand Eastern quartz-phyric sequence and are unconformablyoverlain by the Tyndall Group in the Murchison and Darwinareas (Corbett, 1992). They are interpreted to be subvolcanicintrusions genetically related to the host volcanic succession(Solomon, 1981).

Submarine facies associations and architecture

The Mount Read Volcanics were deposited in a predominantlybelow wave-base, moderate to relatively deep submarinesetting. This interpretation is supported by the presence oftrilobite and other marine fossils, fossiliferous limestone,turbidites and black pyritic mudstone in the sedimentaryfacies association (Jago et al., 1972; McPhie and Allen, 1992).The presence of massive sulfide ore deposits, very thickvolcaniclastic mass-flow units, hyaloclastite, peperite andpillow lava in the volcanic facies association are also consistentwith a subaqueous environment (Corbett, 1992; McPhie andAllen, 1992; Waters and Wallace, 1992).

The essential elements of the facies architecture inthe Mount Read Volcanics are a variety of volcanic andsedimentary facies associations that include lavas, lavadomes, synvolcanic intrusions and diverse volcaniclasticunits (McPhie and Allen, 1992). Lavas and sills occurseparately or in clusters in the succession (McPhie andAllen, 1992). The four common types of volcaniclasticfacies associations in the Mount Read Volcanics are: (1)very thick (tens of metres), massive to graded beds ofrhyolitic to dacitic pumice breccia; (2) very thick, massive

to diffusely stratified units of crystal-rich (feldspar, quartz,clinopyroxene) sandstone; (3) thick to very thick, massive tograded beds of polymictic volcanic conglomerate or breccia;and (4) massive or laminated shard-rich siltstone (McPhieand Allen, 1992). Many of these volcaniclastic facies containa high proportion of crystals, crystal fragments, shards andpumice clasts, which are interpreted to be juvenile pyroclaststransported by water-supported gravity flows (McPhie andAllen, 1992). The sedimentary facies association comprisesconglomerate, sandstone, interbedded turbiditic sandstoneand mudstone, mudstone, and fossiliferous carbonate (Selley,1997; Large et al., 2001a; McPhie and Allen, 2003). Thesefacies are of non-volcanic and mixed provenance, and includepelagic marine sediment, meta-sedimentary and ultramafic(bonninite, gabbro, peridote) rock fragments derived fromthe Precambrian basement, and volcanic clasts and crystals.

There are regional variations in the proportion of volcanicversus sedimentary facies, the types of volcanic facies and thedominant magma composition. Volcanic facies associationslocally dominate the stratigraphy (e.g. at Rosebery) where-as, elsewhere, volcanic facies are intercalated with, or sub-ordinate to, sedimentary facies (e.g. in the hanging wall atHellyer). Parts of the volcanic succession are dominatedby the products of effusive, intrabasinal eruptions, such asthe andesitic lavas and domes in the footwall of the Hellyerdeposit (McPhie and Allen, 1992). In contrast, other areasare dominated by volcanic facies generated by explosiveeruptions, such as the crystal and pumice-rich volcaniclasticunits of the White Spur Formation (McPhie and Allen, 1992;McPhie and Allen, 2003). The volcanic facies associations alsodisplay marked regional variations in composition. Rhyoliteand dacite dominate much of the succession (^90% of themapped area of the central Mount Read Volcanics: Gifkinsand Kimber, 2004); however, intermediate to mafic volcanicfacies are locally important at Hellyer and between Henty andQueenstown (Corbett, 1992; Crawford et al., 1992; Large etal., 2001a).

Post-depositional alteration processes

Formerly glassy or partly glassy volcanic rocks dominatethe Mount Read Volcanics. These rocks have textures andcompositions that reflect subsequent modification by avariety of processes, which include: hydration, diagenesis,hydrothermal alteration, regional and contact metamorphism,and deformation.


Two regional Cambrian diagenetic zones (albite zoneand epidote zone, Section 5.6) are preserved in the northernCentral Volcanic Complex (Gifkins, 2001). Locally, hydro-thermal alteration and mineralisation was synchronous withdiagenesis. In addition, altered halos developed around thicksynvolcanic intrusions and Cambrian granites (Eastoe et al.,1987; Large et al., 1996; Gifkins, 2001).

The Mount Read Volcanics were faulted during theMiddle to Late Cambrian and more extensively deformed,folded and faulted during the Early to Middle Devonian(Corbett and Lees, 1987; Crawford and Berry, 1992).Regional metamorphism to lower greenschist facies producedassemblages of quartz, albite, sericite, calcite, chlorite,tremolite-actinolite, K-feldspar, epidote and biotite, and wascontemporaneous with the Devonian deformation (Corbett,1981; Green et al., 1981; Walshe and Solomon, 1981;Corbett and Solomon, 1989). During the Late Devonian toEarly Carboniferous, contact metamorphism was associatedwith the intrusion of I- and S- type granites (Corbett andLees, 1987; Williams et al., 1989; Corbett, 1992).

Mineral deposits and prospects

In 2004, the Mount Read province contained two majoroperating base-metal-sulfide mines (Rosebery and MountLyell), one gold mine (Henty), and a number of exhaustedand smaller sub-economic deposits or prospects. Publishedresource estimates are listed in Table 1.1.

The Hellyer polymetallic massive sulfide deposit was aclassic mound-shaped ore body discovered by a combinationof geophysical, geological and geochemical explorationtechniques in 1982 (Sise and Jack, 1984). Productioncommenced in 1989 and mining was complete by 2000. Thehigh grade and simple geometry of the Hellyer ore body madeit a profitable operation although metal recoveries were lowdue to the fine grainsize of the sulfides.

The Que River polymetallic massive sulfide deposit was asmall deposit comprising five steeply dipping ore lenses (fourZn + Pb rich and one Cu rich). It was discovered in early 1974by airborne electromagnetic and soil geochemical exploration(Webster and Skey, 1979). Production from the Que Riverdeposit occurred from 1981 to late 1991.

Rosebery is the largest polymetallic massive sulfide depositin western Tasmania. It comprises at least 16 separate stackedore lenses over a strike length of 1.5 km. The deposit wasinitially discovered in 1893 when prospectors traced gold and

Table 1.1 | Tonnages and grades of massive sulfide deposits in the Mount Read province (Data from Mineral Resources Tasmania, Pasminco

Mining and Exploration Goldfields P/L, and Gemmell and Fulton, 2001: in situ values based on average metal prices in 2000).

Deposit

Hellyer

Que River

Rosebery

Hercules

Henty-Mount Julia

Mount Lyell field

tx106

16.2

3.1

32.1

2.7

2.2

311

Znwt%

13.9

13.5

14.7

15.9

-

0.04

Pbwt%

7.1

7.5

4.5

5.1

-

0.01

Cuwt%

0.4

0.6

0.58

0.4

-

0.97

Ag

g/t

168

200

146

159

-

7

Au

g/t

2.5

3.4

2.3

2.54

12.1

0.31

In situ valueUSS billion

3.99

0.81

7.76

0.71

0.23

6.83

Form

Massive lens

Stratabound sheet

Stratabound sheets

Stratabound sheets

Sheet-like

Disseminated

Status

Past producer

Past producer

Current mine

Past producer

Current mine

Current mine

in stratigraphic order, the Puddler Creek, Mount Windsor,

Trooper Creek and Rollston Range formations (Henderson,1986; Paulick and McPhie, 1999). These have a total thicknessof at least 14 km (Henderson, 1986) and generally young tothe south. The four formations respectively represent initialcontinent-derived sedimentation and minor rift-related maficvolcanism, succeeded by voluminous eruptions of crustallyderived rhyolitic magmas, abruptly followed by mixed mafic-felsic volcanism derived from subduction-modified mantle,and culminating in deep-water fine-grained sedimentation(Stolz, 1995). The stratigraphic relationships and lithologiesare summarised in Figure 1.8.


Rollston Range Formation

The uppermost formation of the Seventy Mile Range Group,the Rollston Range Formation, is poorly exposed, except inthe southern central part of the belt where it has a minimumthickness of 1 km. It consists of Early Ordovician, fossiliferous,thinly bedded sandstone and siltstone of largely volcanicprovenance. Minor intervals of felsic lava and volcaniclasticunits exist locally (Berry et al., 1992).

Submarine facies associations and architecture

Puddler Creek Formation

The Puddler Creek Formation is the oldest formation in theSeventy Mile Range Group and consists mainly of massive tolaminated lithic sandstone, greywacke and siltstone of mixedcontinental and volcanic derivation. Minor altered trachy-andesitic to trachytic coherent volcanic rocks are intercalatedwith clastic rocks in the upper few hundred metres. Thevolcanic rocks have geochemical signatures indicating analkali intraplate association related to lithospheric thinningand incipient back-arc basin development (Stolz, 1995). Theformation is up to 9 km thick in the western part of the beltand is partly stoped out by the Ravenswood Batholith in theeast.

The volcanic facies associations in the vicinity of the Highway-Reward, Liontown and Thalanga base-metal sulfide depositsare known from several detailed deposit scale studies (Hill,1996; Miller, 1996; Doyle, 1997; Paulick, 1999; Paulick andMcPhie, 1999; Doyle and McPhie, 2000). Recent regionalstudies of volcanic facies and lithostratigraphy contribute toan improved understanding of the volcanic facies associations(Simpson and McPhie, 1998; Simpson, 2001); however, muchof the regional data is currently confidential or unpublished.

In summary, the Mount Windsor and Trooper CreekFormations comprise deep submarine volcanic facies thatinclude pyroclasts, probably from both subaerial andsubmarine explosive eruptions. Lithofacies such as sparsemicrobialitic ironstones (Simpson and McPhie, 1998)indicate shallow marine settings for the Mount WindsorSubprovince.

Mount Windsor Formation

The Mount Windsor Formation is a 0.4 to 5 km-thicksuccession of subaqueous rhyolitic volcanic rocks dominatedby thick lavas, domes and high-level intrusions with sub-ordinate volcaniclastic breccias. Isotopic evidence (Stolz,1995) suggested the magmas were derived from the meltingof continental crust.

Trooper Creek Formation

The Trooper Creek Formation is a 0.5 to 2 km-thicksuccession of highly variable basaltic-andesitic, dacitic andrhyolitic coherent and brecciated volcanic rocks intercalatedwith abundant volcanogenic siltstone, and minor calcareousmeta-sedimentary rocks. It is internally heterogenous and thereare major lateral variations in the proportion of volcanic andsedimentary facies. Base-metal sulfide deposits and exhalativesiliceous ironstones occur at various stratigraphic levels (Fig.1.8: Duhig et al., 1992). Stolz (1995) suggested that theTrooper Creek Formation volcanic rocks were derived froma melted subduction-modified, sub-arc mantle wedge thatwas erupted during back-arc extension. Decreased volcanicactivity, or an increase in clastic sedimentation, appears tohave coincided with the change from exclusively rhyoliticvolcanism in the preceding Mount Windsor Formation.

FIGURE 1.8 | Simplified stratigraphic column for the Seventy Mile RangeGroup. Modified after Large (1992) and Paulick and McPhie (1999).

14 I CHAPTER 1

Table 1.2 | Tonnages and grades of massive sulfide deposits in the Mount Windsor Subprovince (data from Berry et al., 1992; Large,

1992: in situ values based on average metal prices in 2000).

Deposit

Thalanga

Highway-Reward

Liontown

Handcuff

Waterloo/Agincourt

Magpie

tx106

6.6

3.7

1.8

1

0.4

0.3

Znwt%

8.4

-

6.2

10

19.7

15

Pbwt%

2.6

-

2.2

0.4

2.8

2

Cuwt%

1.8

6.2

0.5

0.6

3.8

2

Agg/t

69

-

29

8

94

30

Au

g/t

0.4

1.5

0.9

0.2

2

1

In situ valueUS$ billion

1.02

0.47

0.18

0.13

0.13

0.06

Form

Stratabound sheet

Subvertical pipes

Stratabound sheet

Stratabound sheet

Stratabound sheet

Stratabound sheet

Status

Past producer

Current mine

Prospect

Prospect

Prospect

Prospect

Post-depositional alteration processes

Some of the least-altered felsic coherent facies of the SeventyMile Range Group commonly have relict spherulitic orperlitic textures due to devitrification and hydration (Paulickand McPhie, 1999; Doyle, 2001). Pseudoclastic textures,attributed to domainal devitrification and subsequentdiagenesis of coherent rhyolites, are prominent in the MountWindsor Formation at Thalanga and probably elsewhere.Zones of intense hydrothermal alteration comprising quartz+ sericite + chlorite + pyrite + carbonate assemblages partlyenclose the major sulfide deposits. They vary in style fromthe broadly stratabound zone extending laterally beneaththe Thalanga deposit, to the discordant concentric zonesenveloping the Highway-Reward sulfide pipes. Early dia-genetic and hydrothermal alteration facies were overprintedby regional deformation coeval with extensive intrusion ofMid-Late Ordovician gneissic granitoids (Berry et al., 1992).This deformation produced relatively low-pressure regionalmetamorphic assemblages that range from prehnite grade inthe east, to upper greenschist grade in the west, and a nearvertical axial planar cleavage. Subsequent intrusion of post-kinematic Siluro-Devonian plutons in the central and easternparts of the subprovince produced contact metamorphicaureoles with assemblages up to amphibole-hornfels grade.

Despite the multiple alteration processes, well preservedprimary volcanic textures that enable detailed interpretationsof facies associations occur away from zones of hydrothermalalteration and mineralisation (e.g. Simpson and McPhie,1998). Even in intensely hydrothermally altered rocks, the

existence of resistant primary components such as quartzphenocrysts allow volcanic facies interpretation (e.g. Paulickand McPhie, 1999).

Mineral deposits and prospects

The Mount Windsor Subprovince contains two major base-metal sulfide deposits and several small sub-economic depositsand historical prospects. Published resource estimates arelisted in Table 1.2. The known deposits are all in the TrooperCreek Formation. The two largest deposits, Thalanga andHighway-Reward, exist at the base and near the top of theformation respectively (Fig. 1.8).

A gossanous outcrop led to the 1975 discovery of theThalanga deposit and its eventual development for open pitand underground mining. Production between 1990 and1998 amounted to 4.7 Mt from an estimated total resourceof 6.6 Mt. Thalanga mine was not highly profitable, mainlybecause of ore dilution in underground mining and stabilityproblems related to the thin ore lenses.

The Highway-Reward Cu-Au deposit consists of twodiscordant, vertical pipe-like bodies of massive pyrite about200 m apart. Originally discovered in a surface road-metalscrape in 1953 (Beams et al., 1998), the Highway-Rewarddeposit has been the subject of intense but sporadicexploration. Open pit mining of small oxide and supergenehigh-grade Cu-Au resources occurred during the late 1980sand from 1997 to the present. The deeper hypogene parts ofthe sulfide pipes remain undeveloped.

I 1 5

2 I DESCRIBING ALTERED VOLCANIC ROCKS

This chapter addresses some of the common problems thatwe face when studying altered volcanic rocks. Recognisingand describing the characteristics of the altered rocks is animportant step towards understanding the processes ofalteration. Alteration involves complex modifications ofthe pre-existing rock and can encompass mineralogical,textural and compositional changes. Resolving these complexrelationships is dependent on a systematic multidisciplinarydescriptive approach incorporating aspects of volcanology, oredeposit geology, petrology and geochemistry. Unfortunatelyrelatively few studies adequately integrate these datasets.Studies of ore deposits generally describe the characteristicsof the host rocks (i.e. lithology, petrology, geochemistry andalteration) in separate sections of manuscripts. In many cases,particularly in unpublished company reports, the geochemicaldata and petrographic descriptions are in appendices,discouraging integration and interpretation.

The integration of physical or textural observations andgeochemical data is a powerful tool in the study of alteredrocks. The physical characteristics and immobile elementconcentrations of altered volcanic rocks can help to identifythe original rocks, where relict primary minerals and texturesare inconclusive (e.g. Paulick and McPhie, 1999; Barrett et al.,2001). Physical and chemical changes that occurred duringalteration may help to determine the degree of alteration (i.e.alteration intensity), the style of alteration (i.e. isochemicalversus metasomatic), and to discriminate between alterationprocesses such as diagenesis, metamorphism and hydrothermalalteration (e.g. Offler and Whitford, 1992; Gifkins and Allen,2001). In addition, this integrated approach can lead to thedevelopment of vectors to guide explorers toward ore deposits(e.g. Large et al., 2001c).

This chapter compares alternative schemes of alterationnomenclature and presents a multi-variable system fordescribing and naming alteration facies. The differentelements of this descriptive approach to nomenclature areexplained in detail in subsequent sections and chapters (i.e.alteration mineral assemblage in Section 2.4, alterationintensity in Section 2.5, alteration textures in Section 3.1and 3.2, distribution and zonation in Section 3.3, and timingin Section 3.5). It also explains alteration indices, and thephysical and geochemical techniques for determining theintensity of alteration. Alteration data sheets, which visually

combine mineralogical, textural and chemical data for alteredvolcanic rocks, are introduced. These alteration data sheets areused in Chapters 5, 6 and 7 to present examples of alterationfacies associated with various alteration processes and differentVHMS deposits.

2.1 | FREQUENTLY ASKED QUESTIONS

Most geologists are introduced to the basic principles ofhydrothermal alteration when they are students. However,they typically have a limited knowledge of how to recognise,characterise and interpret altered rocks. Common questionsare:• Was the rock altered?• What was the nature of the alteration?• Was the rock hydrothermally altered?• How do we name the altered rock?• What was the original rock?

To address these questions we need to make some simpleobservations, which include the recognition of primaryminerals and textures, alteration colour, mineral assemblage,texture and intensity, overprinting relationships, and alter-ation distribution patterns. These observations can be madeat a variety of scales: map, outcrop, hand-specimen and thin-section scales.

Was the rock altered?

Few volcanic rocks in submarine settings are entirely un-altered and most altered rocks are easily recognised as such.The most effective method of determining if a rock is alteredis by comparing it with other samples from the same unit.Observed differences in mineral assemblage, texture andcolour may indicate a spectrum from fresh, or least-altered, tosignificantly altered samples.

Some indicators of alteration in submarine volcanic faciesmay be:• absence of glass• colour differences• presence of abundant minerals that typically form during

alteration, such as clays, zeolites, chlorite, micas, kaolinite,

16 | CHAPTER 2

tourmaline, apatite, alunite, epidote, carbonates and quartzassociation between a distinctive mineral assemblage andsulfidespresence of halos around veins, faults, intrusions andmineralised rocklack of, or only partial preservation of, primary textureshardness of the rock: if not silicified, altered volcanic rockstend to be softer than unaltered volcanic rocks, which aretypically glassy or crystalline, hard and brittledegree of deformation: clay- or phyllosilicate-altered rocksare commonly more deformed than unaltered or least-altered rocks because deformation-related strain is typicallypartitioned into softer altered rocks.

What was the nature of the alteration?

Characterising the nature of the alteration can be challenging.Nevertheless, systematic descriptions of alteration mineralassemblages, alteration textures, preservation of relict mineralsand textures, patterns of distribution and overprintingrelationships combined with interpretations of alterationindices and compositional changes provide importantinformation for subsequent classification and geneticinterpretation. To ensure that the data are meaningful,systematic schemes for core logging and sample descriptionshould be employed. Figures 2.1 and 2.2 are examples wheredetailed observations in drill core and hand specimen haveestablished the characteristics of the altered rocks.

Was the rock hydrothermally altered?

Hydrothermal alteration can be discriminated frommetamorphism and diagenesis, which are typically regionallyextensive processes that result in weakly altered rocks andpreserve delicate volcanic textures. In contrast, hydrothermalalteration styles, especially those associated with mineral-isation, are local in their distribution, have variable intensity(from weak to intense) and generally destroy primarytextures.

Discriminating accurately between different alterationstyles (Section 8.1) requires knowledge of: the host rock;alteration intensity; distribution; timing; mineralogical,textural and chemical changes; and comparison with changesrelated to diagenesis, metamorphism and hydrothermalalteration, which have been documented in well-preserved,geologically young, submarine volcanic successions.

How do we name the altered rock?

Typically, rocks that are only weakly to moderately altered,in which primary textures and minerals can be easilyrecognised, are given precursor names (e.g. quartz-phyricpumice breccia). In contrast, rocks that are intensely altered,in which few primary textures or minerals are recognisable,are given alteration names (e.g. quartz-augen schist or massivechlorite rock). This is similar to metamorphic rocks wherelow-grade rocks are given precursor names — the prefix 'meta-'is assumed — and pervasively deformed and metamorphosedrocks are given metamorphic names.

What was the original rock?

Outcrops and hand specimens of ancient volcanic rocksrarely exhibit clear evidence of their modes of eruption andemplacement. In many cases, the best we may hope for isto recognise features that help distinguish coherent volcanicfacies from volcaniclastic facies.

The simplest approach to recognising the primary rockis to move out of the altered zone and examine unalteredrock. However, ancient volcanic successions rarely containunaltered rocks. As a result we rely on the preservation ofrelict textures and minerals in altered rocks to provide a guideto the interpretation of the primary volcanic facies. Relicttextures are original pre-alteration features that have not beendestroyed by alteration. Relict textures are most likely to bevisible in polished hand specimens with the aid of a hand lensand in thin sections cut parallel to the tectonic foliation.

There are a small number of volcanic, devitrificationand hydration textures, and components or structures thatusually survive diagenesis, moderate hydrothermal alteration,low-grade metamorphism and deformation - these areparticularly helpful in deciphering the primary volcanic facies.For example, porphyritic texture, spherulites, lithophysae,micropoikilitic texture, perlite, flow banding, columnarjoints and pillows are all characteristic of coherent volcanicfacies. Volcanic components such as pumice and scoria clasts,glass shards, accretionary lapilli and non-vesicular volcanicbombs or blocks, as well as bedding and cross stratification,are characteristics of volcaniclastic facies. For a more detaileddiscussion of volcanic, devitrification and hydration textures,components and structures that help to determine the hostvolcanic facies, readers are referred to McPhie et al. (1993)Volcanic Textures: a guide to the interpretation of textures in

volcanic rocks.

Primary crystals and crystal fragments are found in a widevariety of volcanic facies and can also be helpful indicators ofthe host volcanic facies. Whole crystals and crystal fragmentsin volcanic facies are mainly derived from porphyritic magmas.Crystals may be liberated from magmas during volcanicprocesses (explosive eruption or auto fragmentation) or bysurface sedimentary processes. The shape and distribution ofcrystals in an altered volcanic rock can be used as a guideto whether the primary facies was coherent or clastic. Inpyroclastic facies, angular and broken crystal fragmentsare much more common than whole euhedral crystals. Inautoclastic facies, whole crystals and clusters of jigsaw-fitcrystal fragments are common. In contrast, coherent volcanicfacies typically, but not necessarily, contain very few brokencrystal fragments. The distribution of crystals and crystalfragments in volcaniclastic facies may be random, relatedto size or density sorting, or concentrated in particularclasts, clusters or lenses. Crystal-bearing coherent facies areporphyritic; they contain evenly distributed euhedral crystalsin a fine-grained or glassy groundmass.

Although relict textures and primary crystals can be usedas a guide, the discrimination of coherent and volcaniclasticfacies in altered volcanic rocks is not trivial. In originallyglassy volcanic rocks, alteration may produce convincingpseudotextures such as pseudobreccia, false polymictictexture, false thin-bedded texture and pseudomassive texture(Section 3.2: Allen, 1988).

DESCRIBING ALTERED VOLCANIC ROCKS I 17

FIGURE 2.1 | Part of a drill core graphic log using a modified standard logging sheet, which incorporates volcanic and alteration fades descriptions.This drill core, EHP319, is from western Tasmania and includes a thick interval of Central Volcanic Complex rocks. Abbreviations: So = bedding, S, andS2 = tectonic foliations, LCA = long core axis, cc = calcite, chl = chlorite, fsp = feldspar, qtz = quartz, ser = sericite and gb = graded bedding.

18 | CHAPTER 2

FIGURE 2.2 I This sample description - for a rock sample from 245.2 m depth in drill core EHP319 - shows the main descriptive fields for alterationstudies. Abbreviations: S2 = regional cleavage, chl = chlorite, fg = fine grained, fsp = feldspar, hem = hematite, plag = plagioclase, py = pyrite, qtz = quartzand ser = sericite.

Crystal assemblages may reflect the primary volcaniccomposition; they are relicts of original magmatic mineralassemblages. For example, a rock containing abundant quartzcrystals was probably derived from a quartz-phyric magma,which was likely of rhyolitic composition (Table 2.1).

In cases of low temperature (<200°C), weak to moderateintensity alteration, the alteration mineral assemblage mayalso be a guide to the primary composition of the volcanicfades. Alteration minerals rich in Fe, Mg and Ca are commonin mafic volcanic rocks; K- and Na-rich minerals in felsicrocks. Typical alteration minerals in mafic rocks are chlorite,epidote, calcite, palagonite, zeolites, albite, micas, actinolite-tremolite and clays (Table 2.2). In contrast, common alterationminerals in felsic rocks are quartz, micas, feldspars, zeolites,cristobalite, opal and clays, especially montmorillonite andkaolinite (Table 2.2). At temperatures above 200°C and athigh water-rock ratios, the alteration mineral assemblageformed is less dependent on primary host composition andmore on the fluid composition, temperature, permeabilityand pressure (Browne, 1978; Henley and Ellis, 1983; Reyes,1990).

Generally, consideration of a combination of fieldrelationships, relict textures and mineral assemblages willenable interpretation of coherent versus clastic, and felsicversus mafic volcanic facies, in all but the most intenselyaltered volcanic rocks. Beyond that we must resort tolithogeochemical techniques (Chapter 4).

DESCRIBING ALTERED VOLCANIC ROCKS | 19

2.2 I ALTERATION NOMENCLATURE

A variety of approaches have previously been taken to theclassification of alteration and altered rocks, particularlyhydrothermal alteration associated with different stylesof mineral deposits. Common methods of alterationnomenclature are mineral based, compositional, generic, oruse terminology that reflects a combination of mineralogicaland textural characteristics (e.g. alteration facies). Discussionsof alteration nomenclature appear in Meyer and Hemley(1967), Rose and Burt (1979), Beane (1982), Titley (1982),Guilbert and Park (1986) and Thompson and Thompson(1996).

Mineral-based alteration nomenclature

Classifying altered rocks in terms of mineral assemblagewas discussed in detail by Creasey (1959). Mineral-basedclassification involves field and petrographic observations,in some cases supported by other analytical techniques (e.g.microprobe, X-ray diffraction and SWIR spectroscopy). It isbased on direct observations and provides the simplest non-genetic approach to naming alteration and altered rocks.

There are two levels of mineral-based alterationnomenclature: (1) terminology based on the dominant

TABLE 2.1 | Summary of the common volcanic rock compositions, their chemical classification (SiO2 content) and likely

phenocryst minerals. SiO2 contents for unaltered modern subduction-related volcanic rocks are from Ewart (1979).

Rhyolite

Dacite

Andesite

Basalt

>69 K-feldspar (orthoclase) ± quartz ± plagioclase ± biotite ± muscovite± amphibole ± pyroxene ± fayalite

± quartz ± biotite ± amphibole ± pyroxene63-69 Na-plagioclase

52-63 Na- or Ca-plagioclase + biotite ± quartz ± K-feldspar ± olivineor amphibole or pyroxene

<52 Ca-plagioclase + pyroxene ± olivine ± hornblende

TABLE 2.2 | Common alteration minerals that replace glass and magmatic minerals in volcanic rocks. Alteration minerals are fromSchwartz (1959), White and Sigvaldason (1962), lijima (1978), Hay (1978), Honnorez (1978), Brey and Schmincke (1980), Tucker(1987) and Utada (1991).

Silicic volcanic glass

Mafic volcanic glass

Magnetite, ilmenite and titano-magnetite

Pyroxene, amphibole, olivine and biotite

Plagioclase

Anorthoclase, sanidine and orthoclase

Quartz

Zeolites (mordenite, clinoptilolite, laumonite, analcime, heulandite),cristobalite, opaline silica, quartz, calcite, clays (montmorillonite, smectite,mixed-layer clays)

Palagonite, nontronitic clays, smectite, calcite, chlorite, epidote, Ca-richzeolites, Fe/Ti/Mn-oxides

Pyrite, leucoxene, titanite, pyrrhotite, hematite

Chlorite, illite, quartz, calcite, pyrite, anhydrite

Calcite, albite, adularia, wairakite, quartz, anydrite, chlorite, illite, kaolin,montmorillonite, epidote, sericite

Adularia, albite, sericite

Microcrystalline quartz

20 | CHAPTER 2

mineral; and (2) the use of the complete or abbreviatedalteration mineral assemblage. Some authors also use negativemineral-based names, such as K-feldspar-destructive alteration(e.g. Gustafson and Hunt, 1975).

The simplest method of alteration nomenclature uses thedominant or most recognisable mineral phase in the alteredrock. Examples of this are albitic, which is dominated byalbite; silicic, dominated by quartz; chloritic, dominated bychlorite; and sericitic, dominated by sericite. In addition,the terms chloritisation, sericitsation, silicification andcarbonitisation are common in VHMS literature in referenceto the processes of alteration (e.g. Sangster, 1972; Paradis etal., 1993). They, like the terms alteration and mineralisation,are widely misused (Solomon, 1999).

Deciding which mineral is dominant in an altered zoneis not always a straightforward task. Several minerals maybe obvious and their proportions may vary. In addition,common alteration minerals, such as sericite, can occur as thedominant mineral in several different mineral assemblagesthat have different origins, timing and economic significance.In fact, sericitic assemblages are probably the most abundantand widespread of all alteration assemblages. They are presentin aluminous rocks in nearly all types of hypogene alterationassociated with ore deposits (Meyer and Hemley, 1967).Along with sericite, carbonates, chlorite and quartz are amongthe most widespread alteration minerals (Meyer and Hemley,1967).

Alternatively, more detailed alteration mineral assemblagescan be used; either complete assemblages of all the visiblealteration minerals, or abbreviated assemblages of the mostabundant and distinctive minerals. Minerals are usually listedin order of decreasing abundance; thus a mixture of 60%sericite, 35% quartz and 5% pyrite becomes the sericite +quartz + pyrite alteration assemblage. This nomenclature hasthe advantage of clearly defining the alteration assemblage.However, some confusion may exist where mineral assemblagescontain identical or similar minerals in different abundances.For example, sericite + quartz + pyrite is easily confused withsericite + chlorite + quartz + pyrite.

Dana's Textbook of Mineralogy (Dana, 1957) definedsericite as 'fine scaly muscovite united in fibrous aggregatesand characterized by its silky lustre'. The term has sincebeen widely used to refer to all fine-grained pale-colouredmicas, and indeed almost any fine-grained aggregates ofpale-coloured layer-lattice minerals (Whitten and Brooks,1972), particularly in hydro thermally altered and low-grademetamorphic rocks. White mica is the preferred term to avoidthe ambiguity in sericite where compositional differences suchas sodic muscovite, muscovite and phengite, may be important(e.g. Yang, 1998). In this book, we always use sericite in theloose sense, referring to fine-grained pale-coloured micasof undetermined composition. We use the non-specificalternative term: white mica, where compositions are known(e.g. in discussions of mineral chemistry in Sections 4.2 and8.2).

Compositional alteration nomenclature

Chemical methods of assessing hydrothermally alteredrocks (i.e. lithogeochemistry) have been applied in mineral

exploration, especially around VHMS and porphyryCu deposits, leading to the classification of alteration bycompositional changes that occurred during alteration(Hemley and Jones, 1964). Examples of this include Na-metasomatism or soda-metasomatism, Mg-metasomatismand K-metasomatism (e.g. Hemley and Jones, 1964) or K-enriched alteration, Ca-enriched alteration and Mg-enrichedalteration (e.g. Elliott-Meadows and Appleyard, 1991), andNa-depleted alteration (e.g. Date et al., 1979, 1983).

There are several problems with compositional alter-ation nomenclature: (1) it becomes increasingly complicatedwhere more than one element is mobilised during alteration,which is almost always the case in the alteration of volcanicrocks; (2) it cannot be applied in the field, as it requires adetailed knowledge of the addition and removal of elements;and (3) the character of the chemical alteration can onlybe accurately determined if a least-altered protolith can beunequivocally identified (Section 4.1).

Generic alteration nomenclature

A number of generic terms, such as advanced argillic,intermediate argillic, phyllic or sericitic, potassic, propylitic,skarn and greisen, have been applied to common alterationmineral assemblages or groups of assemblages (Table 2.3:Meyer and Hemley, 1967). These terms are widespread inthe geological literature, however they are not always clearlydefined or uniformly applied by different authors, and areless precise than alteration assemblages. Many workers applygeneric terms based on the occurrence of indicator mineralsrather than complete alteration assemblages, with the resultthat the terms are not always distinguishable in their usage(Rose and Burt, 1979). To apply these terms rigorously,alteration mineral assemblages for specific host rocks need tobe identified and correlated.

Generic alteration nomenclature tends to reflect detailedwork on altered rocks associated with particular deposittypes or geothermal systems, specifically porphyry, skarn,mesothermal vein and epithermal deposits (Table 2.4). Ineach case, the generic classification conveys a sense of themineralogical composition and implies knowledge of alterationprocesses and environment of formation. Although genericclassification of altered rocks surrounding an ore deposit canbe useful, the reliance on understanding the environmentof formation can cause problems and may incorrectly implygenetic processes. Thus, generic classification is best avoidedduring the early stages of recognition, description, mappingand interpretation of altered rocks in favour of a more rigorousand descriptive classification.

Descriptive nomenclature — alteration facies

The term alteration facieswas first proposed by Creasey (1959)in an attempt to standardise the subdivision of hydrothermallyaltered rocks in a similar manner to metamorphic facies,which are assemblages of co-existing metamorphic mineralsthat characterise particular pressure and temperature regimesduring metamorphism (Yardley, 1989). Creasey's concept ofthree chemically and mineralogically distinctive alteration


TABLE 2.3 | Generic alteration terms based on common alteration mineral assemblages. Modified after Creasey (1959), Meyer and Hemley (1967),Lowell and Guilbert (1970), Rose (1970), Gustafson and Hunt (1975), Rose and Burt (1979), Beane and Titley (1981), Guilbert and Park (1986), Beane(1982), and Thompson and Thompson (1996). Forsimplic,*, skarn implies a limestone or dolomite host rock.

Argillic

Advanced argillic

Intermediateargillic

Phyllic (or sericitic)

Sericitic (or phyllic)

Propylitic (orsaussuritization)

Potassic

Greisen

Skarn

Calcic skarn(or tactite)

Kaolinite (or halloysite, metahalloysite or dickite) +montmorillonite ± sericite (or muscovite) ± chlorite

Pyrophyllite + kaolinite (or dickite) ± quartz ± sericite± andalusite ± diaspore ± alunite ± topaz ± zunyite ±enargite ± tourmaline ± pyrite ± chalcopyrite ± hematite

Chlorite + sericite ± kaolinite ± montmorillonite ± illite-smectite ± calcite ± epidote ± biotite ± pyrite

Sericite + quartz + pyrite ± biotite ± chlorite ± rutile ±leucoxene ± chalcopyrite ± illite(Note: K-feldspar absent)

Sericite + quartz + pyrite ± K-feldspar ± biotite ± calcite± dolomite ± chlorite ± andalusite ± chloritoid ± albite ±pyrrhotite

Epidote (or zoisite or clinozoisite) + chlorite + albite± carbonate ± sericite ± montmorillonite ± septachlorite± apatite ± anhydrite ± ankerite ± hematite ± pyrite ±chalcopyrite

K-feldspar (orthoclase) + biotite + quartz ± magnetite± sericite (or muscovite) ± albite ± chlorite ± anhydrite ±apatite ± rutile ± epidote ± chalcopyrite ± bornite ± pyrite

Muscovite (or sericite) + quartz + topaz ± tourmaline± fluorite ± rutile ± cassiterite ± wolfranite ± magnetite ±zunyite ± K-feldspar

Porphyry Cu, high-sulfidation epithermal, low-sulfidation epithermal, geothermal


Porphyry Cu, high-sulfidation epithermal

Porphyry Cu

Porphyry Cu, low-sulfidation epithermal,geothermal, VHMS , sediment hosted massivesulfide


Porphyry Cu

Porphyry Cu, porphyry Sn

Pyroxene + garnet + wollastonite ± epidote (or zoisite) Porphyry, skarn± actinolite-termolite ± vesuvianite ± pyrite ± chalcopyrite± sphalerite

Magnesian skarn Forsterite + diopside + serpentine + talc ± actinolite-tremolite ± calcite ± magnetite ± hematite ± chalcopyrite± pyrite ± sphalerite

Retrograde skarn

Jasperiod

Calcite + chlorite ± hematite ± pyrite

Quartz + pyrite + hematite

Porphyry, skarn

Porphyry, skarn

Sedimented-hosted Au, VHMS

fades — propylitic, argillic and potassium silicate facies —was abandoned for a wide variety of generic and non-genericterms.

Subsequently, Riverin and Hodgson (1980) proposed thatalteration facies be used as a descriptive term to refer simplyto altered rocks that could be identified during the courseof mapping or in hand specimen. They described a 'spottedfacies' characterised by a well-developed spotted texture dueto large, strongly altered, cordierite porphyroblasts, and a'silicified facies' that lacked spots and was typically grey incolour and siliceous in appearance.

More recently, the concept of alteration facies hasbeen expanded to incorporate other descriptive elements,particularly alteration mineral assemblages (e.g. Gibson et al.,1983; Elliott-Meadows and Appleyard, 1991; Paradis et al.,1993; Tiwary and Deb, 1997; Brauhart et al., 1998; Gifkinsand Allen, 2001). Examples are 'mottled quartz-epidote' and

'silicifi cation alteration facies' described by Gibson et al. (1983)in the Amulet Rhyolite of Noranda, Canada, and 'domainalfeldspar-quartz-sericite', 'fracture-controlled chlorite-sericite'and 'stylolitic chlorite-sericite-hematite alteration facies'described by Gifkins and Allen (2001) in a regional study ofalteration in the Mount Read Volcanics, western Tasmania.

The advantage of characterising alteration in terms ofalteration facies is that it is a purely descriptive scheme inwhich the basic criteria used to classify the alteration can berecognised and established in the field or in hand specimen.More importantly, by using a combination of textural andmineralogical terms, alteration facies convey the generalappearance of an altered rock. Also, the descriptive variablesin the alteration facies provide information that is critical tosubsequent genetic interpretations of the alteration process(e.g. diagenetic, metamorphic, or hydro thermal).

22 | CHAPTER 2

TABLE 2.4 | Examples of different alteration nomenclature (i.e. dominant mineral, abbreviated mineral assemblage, compositional and generic terminology)

applied to altered rocks in a variety of ore deposit environments.

VHMS depositsSilicicChloriticSericiticAlbiticCarbonate

Porphyry depositsKaoliniticPyrophylliticKaolinitic

SericiticFeldspathic

Biotitic

Chloritic

Epithermal depositsSilicicAl unite

K-mica or kaoliniteChloriticSericitic

Sediment-hosteddepositsSilicicSilicic

TourmalineCarbonateSericiticAlbitic

Quartz + sericite + pyrite ± chlorite ± K-feldsparChlorite + pyrite + sericite ± quartzSericite ± quartz ± chlorite ± pyriteAlbite + sericite ± quartzDolomite/siderite/ankerite ± quartz ± sericite ± chlorite± pyrite

Kaolinite + montmorillonite ± sericite + chloritePyrophyllite + kaolinite ± quartz ± sericiteKaolinite + chlorite + sericite ± montmorillonite ± illite-smectite ± calcite ± epidote ± biotiteSericite + quartz + pyrite ± chlorite ± biotiteK-feldspar ± biotite ± quartz ± sericite ± albite ±anhydrite ± epidoteBiotite + K-feldspar + magnetite ± quartz ± albite ±anhydriteChlorite + epidote + albite ± carbonate + sericite ±montmorillonite ± pyrite

Quartz ± chalcedony ± alunite ± barite ± pyriteAlunite + kaolinite/dickite + quartz/cristobalite ±pyrophyllite ± diaspore ± pyrite ± topaz ± andulusiteKaolinite/dickite + illite-smectite ± quartz ± pyriteChlorite + calcite + epidote + albite ± pyriteSericite + illite-smectite ± quartz ± calcite ± dolomite ±pyrite

Quartz + pyrite + hematiteQuartz ± muscovite ± carbonate + pyrite + pyrrhotiteTourmaline ± muscovite ± quartz ± pyrrhotiteAnkerite/siderite/calcite + quartz ± muscovite ± pyrrhotiteSericite + chlorite + quartz ± pyrrhotite ± pyrite ± albiteAlbite + chlorite + muscovite ± biotite

Si-metasomatismMg-metasomatismK-enrichmentNa-depletionCa, Mg, or Mn-metasomatism

K, Ca, Mg-metasomatismK, Ca, Mg-metasomatismK, Ca, Mg-metasomatism

Na, Ca, Mg-metasomatismK-metasomatism

K-metasomatism

Ca-Mg-metasomatism

Si-enrichmentCa, Mg, Na-depletion

K, Ca, Mg, Na-metasomatismCa, Mg-metasomatismK-metasomatism

Not used in VHMSliterature

ArgillicAdvanced argillicIntermediate argillic

PhyllicPotassic

Potassic

Propylitic

SilicicAdvanced argillic

Intermediate argillicPropyliticArgillic

Jasperiod

Tourmalinite

2.3 | ALTERATION FACIES -THE RECOMMENDED METHOD

We advocate a multi-faceted, descriptive approach to studyingaltered volcanic rocks. Different alteration facies can bedefined not only on the basis of their mineral assemblage andtexture, but also on distribution, intensity and composition(or compositional changes). This approach to describingand naming alteration facies is similar to the nomenclaturescheme adopted by McPhie et al. (1993) for volcanic faciesand the descriptive scheme for diagenetic calcite used by Folk(1965).

The four alteration variables are: mineral assemblage,texture, distribution and intensity (Fig. 2.3). However, becauseit is not always practical to provide information on all fourvariables, we suggest that the alteration mineral assemblageand at least one other variable be used. Ideally, descriptivenames for alteration facies follow the formula: intensity +distribution + texture + mineral assemblage.

The intensity variable (Section 2.5) provides informationon the degree of mineralogical, compositional and textural

modification (i.e. subtle, weak, moderate, strong or intense). Itis determined from petrographic descriptions in combinationwith compositional data (e.g. Na2O) and alteration indices.

The distribution variable (Sections 3.3 and 3.4) refers tothe mappable extent of the alteration facies and its relationshipto host facies or components, structures, mineralised rock,veins and other alteration assemblages (i.e. local or regional;footwall or hanging wall; stratabound, pipe or plume).

The texture variable (Sections 3.1 and 3.2) refers to thealteration texture that is superimposed on the rock, and istypically described in hand specimen and/or thin section. Itmay incorporate the shape, form, grainsize or fabric in thealtered rock (e.g. pervasive, selective or vein halo).

The alteration mineral assemblage (Section 2.4) is expressedas an abbreviated alteration mineral assemblage in which theminerals are listed in order of decreasing abundance (e.g. theassemblage feldspar > quartz > sericite becomes feldspar +quartz + sericite).

This approach produces alteration facies names such asweak, regional, selective, chlorite + sericite alteration facies orstrong, massive, footwall, quartz + sericite alteration facies.


FIGURE 2.3 | Descriptive names for alteration facies.

2.4 | ALTERATION MINERALASSEMBLAGE

Mineral assemblage refers to specific, and usually characteristic,observed mineral associations that may be in equilibrium ordisequilibrium. An equilibrium mineral assemblage is a groupof minerals formed at the same time, lacking any indicationof disequilibrium, such as replacement or veining textures,and hence interpreted to have formed due to the same process

and under the same fluid-rock conditions (Hemley and Jones,1964). Disequilibrium or metastable mineral assemblagesare common and caution must be exercised in equating co-existence with stable equilibrium (Meyer and Hemley, 1967;Rose and Burt, 1979).

In general, alteration is a process of re-equilibration.The pre-existing mineral constituents in a rock becomeunstable under changed physicochemical conditions (e.g.the addition of hydrothermal fluid) and progressivelyalter to a new stable mineral assemblage, with or without

24 | CHAPTER 2

metasomatic chemical changes. The alteration process may beonly partially completed and may result in a disequilibriumassemblage containing a mixture of the pre-existing and newalteration minerals. Indeed, disequilibrium assemblages aretypical of altered volcanic rocks. Common examples, at lowmetamorphic grade, are domainal devitrification of felsic glassand incipient sericitisation of feldspar crystals.

Subsequent overprinting alteration may complicate dis-equilibrium assemblages. Volcanic rocks commonly retainrelicts of primary minerals (especially as phenocrysts)and alteration minerals from several stages of diagenetic,metamorphic and/or hydrothermal alteration (e.g. Fig. 2.4).Equilibrium assemblages may be attained in zones of intensehydrothermal alteration or metamorphism, but primaryequilibrium assemblages are rarely preserved in ancientvolcanic rocks. This is true even in least-altered rocks.

When mapping altered rocks it is important to recognisedisequilibrium assemblages and correctly attribute mineralsto the various processes of formation. Equally important isan understanding of the effects and constraints that earlieralteration facies, at various scales, may impose on subsequentprocesses.

Tools for mineralogical determination

The first steps in the identification of alteration minerals areto apply the three essential field tools: the practised geologicaleye, the hand lens and the scriber. These are frequentlyadequate for useful descriptions of alteration mineralassemblages, mapping of altered zones and interpretation ofstyles or processes of alteration. Large-scale features must notbe overlooked; these provide the geological context that iscritical for interpretation of processes.

Simple chemical field tests, such as the use of dilute hydro-chloric acid for discriminating carbonates, and sodium cobaltnitrite for staining K-feldspar, can also be useful. However,when alteration minerals occur as fine-grained massesadditional instrumental techniques, such as microscopicpetrography, short wavelength infrared spectrometry, X-raydiffraction and micro-analyses, may be necessary to identifythem. In many situations, such as mineral exploration, thepractising geologist must rely largely on field skills, perhapsaugmented with limited laboratory work to substantiateand assist in developing an 'eye' for particular mineralassemblages.

Polished slabs

Stage 1: Hydration

Glassy plagioclase-phyriccoherent rhyolite with perliticfractures. Fracture surfacesare coated with clayminerals.

Stage 2: Diageneticalteration

Partly clay + zeolite-alteredplagioclase-phyric coherentrhyolite. The alterationfacies distribution iscontrolled by the perliticfracture pattern.

Stage 3: Hydrothermalalteration

Moderately sericite + quartz-altered plagioclase-phyriccoherent rhyolite.Pervasively developedsericite + quartz hasreplaced all glass andpreviously altered domains.Some clay-altered relictshave been altered to sericite.Plagioclase phenocrysts arepartly altered to sericite.

Stage 4: Hydrothermalalteration

Intense chlorite + pyrite-altered plagioclase-phyriccoherent rhyolite. Vein-halochlorite + pyrite associatedwith cross-cutting chlorite +carbonate veins hasoverprinted and destroyedearlier clay and sericite +quartz alterationassemblages and textures.

FIGURE 2.4 | Cartoons of the microscopic textural and mineralogical evolutionof an originally glassy plagioclase-phyric coherent rhyolite. Overprintinghydration, diagenesis and two stages of hydrothermal alteration are visible in thefinal rock.

The identification of primary and alteration minerals, textures,and overprinting relationships can often be facilitated by thecareful examination of polished slabs using a hand lens orsimple binocular microscope. Polished slabs can be made fromdrill core or hand specimens that have been sawn to producea relatively flat surface, which is subsequently ground smoothusing a diamond lap. At this stage many coarser minerals andtextures will be visible on the wet surface. The resolution offiner features can be improved by polishing the slab surface ona rotating metal lap with 220 to 400-grit zinc or iron powderand then finer powder (with water) on a glass plate. In theabsence of polishing equipment, it is sometimes beneficial to

buff the sawn surface with wet sandpaper. Steel wool can beused to clean tarnished sulfides. Polished slabs are the cheapestand most readily available tools for the field geologist.

Petrography

Examination of standard 75 x 25 mm thin sections orpolished thin sections with a polarising microscope is anexcellent and relatively inexpensive method of mineralidentification. It is the best way of resolving small-scalespatial relationships between minerals to assist determinationof alteration reactions, paragenesis and likely processes.


Petrography is most effective if carried out by the personwho maps and samples the rocks. This requires access topreparation facilities and a polarising microscope. This is nota practical solution for mineral explorers, but is still widelyapplicable in academia. The alternative is to send selectedsamples to a consultant petrographer with complete detailsof the geological context and the underlying objectives. Manyprofessional petrographers are unapologetic petrologists,principally interested in petrogenesis and not enthusiasticabout the obscuring effect of alteration. Therefore, it isimperative that the client informs the petrographer of theimportance of alteration mineral assemblages.

Short-wavelength infrared spectroscopy

The development of portable field instruments like thePIMA (portable infrared mineral analyser), has increased theuse of short-wavelength infrared (SWIR) spectroscopy inmineral exploration and related research (Thompson et al.,1999). The technique identifies phyllosilicates, hydroxylatedsilicates, carbonates and sulfates in most types of dry geologicalsamples and can also provide information on crystallinity andcompositional variations in some minerals, such as clays, whitemica and chlorite. These minerals, particularly phyllosilicates,are common constituents of alteration mineral assemblagesand may be difficult to discriminate by other field or opticalmethods.

Portable SWIR analysis has significant limitations inresolving complex mineral assemblages, analysing darksamples with significant opaque components and inidentifying aspectral anhydrous minerals, such as quartz.It is an empirical method and does not supersede precisedeterminative methods such as X-ray diffraction. Nevertheless,portable SWIR has practical advantages including rapid in-field analyses of up to 30 samples per hour and no samplepreparation other than drying. It has many applications inthe recognition and mapping of altered zones in a variety ofmineral deposit styles. Thompson et al. (1999) listed recentlypublished SWIR studies in epithermal, Archaean greenstone,VHMS, uranium, evaporite and regolith environments.

Electron microprobe

Micro-analysis of mineral grains by electron microprobe hasbecome the standard tool for studies of mineral chemistryover the past few decades. It has fine resolution, down to afew microns diameter, and provides quantitative analysesof elements with atomic numbers greater than four (Be) atconcentrations of greater than about 0.01 wt% (Berry etal., 1983). Major element data can be used to estimate themolecular formulae of unidentified minerals and investigatespatial variations in mineral composition. Non-destructiveanalyses are made on standard polished petrographic thinsections or polished grain mounts. However, the electronmicroprobe is an expensive laboratory instrument; itrequires a skilled operator and the sample throughput is low.Consequently, it is essentially a research tool. It is rarely appliedin alteration studies, mineral exploration or mapping, but ispotentially useful for the verification of mineral identificationand spatial compositional variations interpreted by othermeans.

X-ray diffraction

X-ray diffraction (XRD) is the definitive method for theidentification of all crystalline minerals, including opaqueminerals and structural polymorphs with similar chemicalcompositions. Modern powder diffractometers can providesemi-automated analysis and computerised semi-quantitativemineral identifications from a small amount of powderedsample (Berry et al., 1983). Like the electron microprobe,these machines are mainly used as research tools. Commercialquantitative XRD is not widely available and is relativelyexpensive, currently around $75-95 per sample (AMDEL)in Australia.

2.5 | ALTERATION INTENSITY

Alteration intensity is an indication of how completely arock has reacted to produce new minerals and textures,and is independent of the alteration process. The alterationintensity does not reflect the new mineral species, only theirabundance. It is closely linked to textural and compositionalchanges because it reflects the extent to which pre-existingtextures and minerals (relicts of the original volcanic facies)are preserved, and the degree of metasomatism (Rose andBurt, 1979). Alteration intensity can be estimated bothqualitatively and quantitatively.

Qualitative estimates of alteration intensity

Qualitative estimates of the alteration intensity summarisethe textural and mineralogical changes that occurred duringalteration. They are based on the abundances of new alterationminerals, the degree of destruction of pre-existing minerals,the pervasiveness of alteration textures, and/or the degree ofpreservation of pre-existing textures. Although these featurescan be estimated to some extent in hand specimen, they arecommonly estimated petrographically. Many geoscientistsapply terms such as least altered, weakly altered, moderatelyaltered, strongly altered and intensely altered to describealteration intensity; however, these terms are subjective andare rarely well defined.

Simmons and Christenson (1994) determined alterationintensity by measuring the percentage conversion of primary tosecondary minerals, such that a weakly altered rock contained0-33% alteration minerals; moderately altered 33-67%; andstrongly altered 67-100%. Alteration intensity can also bemeasured by independently estimating the addition of newminerals in the groundmass and the destruction of primaryphenocrysts such as plagioclase.

In contrast, Guilbert et al. (in Guilbert and Park, 1986)proposed that alteration intensity be described in terms of boththe growth of new alteration minerals and the destruction ofpre-existing textures. Their two-part alteration intensity scaleincorporates estimates of the susceptibility of minerals toalteration and the pervasiveness of alteration minerals. Mineralsusceptibility is the degree to which minerals in the rock arealtered, S1-S10 (vol%), whereas pervasiveness is the degree to

Despite these attempts to quantify alteration intensity,it is still applied subjectively by most geologists. For thisreason we prefer to avoid a numerical system and retain thedescriptive terms subtle, weak, moderate, strong and intense.The term least altered is reserved for rocks that are less alteredthan their counterparts in the same environment. Least-altered rocks may be weakly to moderately altered, especiallyin hydrothermal environments where all rocks are altered tosome degree.

Here we define subtle, weak, moderate, strong and intensealteration based on the extent of growth of new alterationminerals, the destruction of primary minerals and textures,and pervasiveness of alteration textures (Table 2.5). Typicallywith increasing intensity of alteration, primary minerals areprogressively replaced, new minerals are more pervasivelydistributed, primary textures are less consistently preserved,and new textures are developed (Fig. 2.5). For example,Gustafson and Hunt (1975) noted that with increasingintensity of hydrothermal K-silicate alteration at the ElSalvador porphyry deposit in Chile, there is an increasingdegree of replacement of plagioclase phenocrysts by K-bearing phases until the phenocrysts are obliterated. Withprogressively more intense alteration, the mafic phenocrystsare replaced, the groundmass becomes coarser grained withK-feldspar overgrowths, magnetite and hematite disappear,and the abundance of veins increases.

Estimates of alteration intensity that incorporate texturalchanges are biased towards texturally destructive alterationstyles such as feldspar-destructive hydrothermal alteration.It is important to recognise that under some circumstancesalteration, particularly carbonate alteration and some formsof silicification, can enhance some primary or pre-existingtextures (e.g. Fig. 2.6: Titley, 1982; Allen, 1988). Forexample, carbonate nodules preserve delicate shard texturesin the Hercules footwall, western Tasmania (Fig. 2.6A: Allen,1997), and shards are preserved in quartz nodules and quartz+ chlorite (± muscovite) altered zones in the Gossan Hillfootwall, Western Australia (Fig. 2.6B and C: Sharpe andGemmell, 2001). Although these alteration styles preservepre-existing textures, they may still be recognised as intenselyaltered because of the pervasiveness of the new mineralassemblage.

Colour contrasts related to overprinting alterationassemblages or different mineral habits within an assemblagecan enhance textures, such as clast margins, whereas anotheralteration assemblage with lower colour contrast and of equalintensity may preserve textures just as well but textures maybe less discernable. The pervasiveness of alteration texturesand the degree of preservation of pre-existing textures aredependent on the resilience of the pre-existing textures, theintensity and style of alteration (Doyle, 2001; Gifkins andAllen, 2001).

Quantitative estimates of alteration intensity

Alteration indices

Alteration indices are simple, multi-component or normalisedratios of lithogeochemical composition data. They are usuallycalculated from composition data expressed as weightpercentages (wt%) or parts per million (ppm), although insome cases molar proportions are used. They are geochemicalrepresentations of hydrothermal mineral assemblages designedto facilitate discrimination of alteration styles, quantificationof alteration intensity, and exploration vectors. Alterationindices have been widely applied in research and explorationforVHMS deposits (Ishikawaetal., 1976; Large etal., 2001a).They have also been used to a lesser extent in sediment hostedZn-Pb-Ag deposits (Large et al., 2001a) and Archaean lodeAu deposits (Eilu et al., 1997; Bierlein et al., 2000).

Simple ratio indices, especially of molar proportions, aregenerally easily related to mineralogical changes (Eilu et al.,1997). However, that is not the case for some more complexindices where changes in the index value could be due tochanges in one or more of three or four components, and thusrelated to several mineral phases. Stanley and Madeisky (1996)noted that some empirically determined alteration indices arenot universally effective outside the district where they wereinitially developed, tend to generate many false anomalies, ormay fail to identify significant altered zones, because losses ofone component may cancel out gains in another.

Alteration indices are formulated by placing proportionsof components that were gained during alteration in thenumerator and components that were lost in the denominator,thus producing the highest values in the most intensely alteredrocks. In developing new indices, it is therefore useful to firstapply mass transfer techniques to determine the componentsgained and lost.

Because alteration indices are ratios, they are less affectedby closure than composition data (closure is discussedin Section 4.1). They respond only to changes in theconcentrations of those components used in the index, butnot to all other components of the rock. Nevertheless, andcontrary to the opinion of Eilu et al. (1997), alteration indicesare not independent of closure because each componentof composition data is affected by closure. Hence, majorcomponents that dominate igneous rock compositions, suchas SiO2 and A12O3 (which is also relatively immobile), arerarely used in alteration indices for volcanic rocks.

Simple indices are ratios formulated from two componentsof analytical data. For example, the S/Na2O ratio of Large etal. (2001a). S/Na2O shows high contrast in VHMS alterationsystems, typically with values less than 0.1 in least-alteredrocks and values several orders of magnitude greater in sulfide-bearing, intense, proximal altered footwall zones (Fig. 2.7).

Multi-component and normalised indices have two or morecomponents added together in either or both the numeratorand denominator of the index. The alkali based K2O/Na2O + K2O + CaO index of Date et al. (1983) is a typicalexample. It has a common structure for alteration indices: thecomponents of the numerator are also in the denominator.This has a normalising effect of limiting the possible rangeof values from zero to one. The normalisation in someindices involves multiplication by a factor of one hundred to

26 [ CHAPTER 2

which alteration minerals permeate the entire rock, PI—P10(vol%).


TABLE 2.5 | Descriptive alteration intensity terms (subtle, weak, moderate, strong and intense) defined on the extent of growth of new alterationminerals, the destruction of primary minerals and textures, and pervasiveness of alteration textures.

Subtle Phenocrysts and free New minerals have Primary volcanic, devitrification Minor replacement/recyrstailisationcrystals of feldspar, coated the surfaces of and hydration textures are clearly (micro- or cryptocrystalline,quartz, and mafic existing phenocrysts, visible with little or no modification, overgrowths, poikilitic, microlitic,minerals (amphiboles, fractures and clasts, spherulitic, variolitic and perlitic) andpyroxenes etc.) were and infilled open space infill textures,unaffected by alteration, (fractures, vesicles,Plagioclase may have pore space, etc.). Glassbeen dusted with has been devitrified.sericite, carbonate orhematite.

Weak Feldspar has been Patchy or domainal and Good preservation of most Replacement, dissolution,partly replaced by albite, disseminated selective textures (original groundmass, recrystallisation, deformation andsericite, carbonate, alteration styles. matrix textures and phenocrysts). infill textures. Most common textureshematite and/or epidote. Alteration commonly Delicate textures such as shards, include: pseudomorphs, cleavageMafic minerals have nucleated on existing pumice clasts, perlite and the fine and rim texture, core and zonalbeen partly replaced minerals, clasts or fibrous textures in spherulites texture, core and rim texture, skeletalby Mg- and Fe-rich fractures and interstitial show some modification. - texture, overgrowths, micro- orminerals, such as in glomerocrysts. - cryptocrystalline, dissolution vugs,chlorite, epidote and J stylolites, poikilitic, foliations, fiamme,Fe-oxides. and infill textures.

Moderate Feldspar has been partly Patchy or domainal Most textures modified and/or Replacement, dissolution,to completely replaced and disseminated destroyed by alteration. Delicate recrystallisation, deformation, infillby feldspar, sericite, alteration styles. textures commonly destroyed or and pseudotextures. In particular:carbonate, epidote, Individual domains may substantially modified. Coarser pseudomorphs, partial pseudomorphs,quartz and/or magnetite, have been texturally groundmass textures (perlite, overgrowths, disseminated nodules,with outlines still destructive (i.e. chlorite spherulites, amygdales and flow spheriods, micro- or cryptocrystalline,visible. Mafic minerals alteration of pumice banding) and clasts partially to dissolution vugs, stylolites, fiamme,commonly completely clasts). Selective completely recrystallised but still porphyroblasts, poikiloblasts, poikilitic,pseudomorphed. Minor alteration of individual clearly visible in domains. hornfelsic and augen textures, andrecrystallisation or clasts, groups of clasts foliations and lineations.replacement of quartz. or minerals. Vein-halo

alteration.

Strong Feldspar has been Domainal selective to Primary volcanic, devitrification Replacement, dissolution,completely replaced pervasive. Vein-halo and hydration textures almost recrystallisation, deformation, infillby chlorite, sericite, alteration. completely destroyed (regardless and pseudo textures. Including:carbonate and/or of grainsize). Pervasive pseudomorphs, nodules, spheroids,opaques (although , replacement of groundmass, micro-or cryptocrystalline,outlines still partly matrix and phenocrysts. Sparse dissolution vugs, stylolites, fiamme,visible) and quartz relict fiamme, amygdales and clast porphyroblasts, poikiloblasts, poikilitic,partly replaced or outlines preserved. granoblastic, decussate, hornfelsicrecrystallised. and augen textures, and foliations and

lineations.

Intense No primary minerals Transgresses textural All original rock textures including Replacement, dissolution,remain. Sparse outlines facies and unit contacts phenocrysts have been destroyed, recrystallisation, deformation, infill andafter primary minerals and primary textures. Weak pseudomorphs or outlines pseudo textures. Including: nodules,may still be visible. Pervasive, typically of coarse phenocrysts may spheriods, micro- or cryptocrystalline,

homogenous, alteration be visible. Primary rock type rare stylolites, granoblastic, decussateon a local scale. Vein- indeterminate. and hornfelsic textures, and foliationshalo alteration. and lineations.

28 | CHAPTER 2


A. Bubble-wall shardsDelicate bubble-wall and platy shards (S) have been preservedwithin a carbonate nodule in the proximal, carbonate zonebeneath the Hercules VHMS deposit. The carbonate nodulecomprises quartz + calcite + chlorite-altered pumice breccia.Plane polarised light.Sample MR96-57, Cambrian Hercules Pumice Formation,Central Volcanic Complex, Mount Read Volcanics, Hercules

footwall, western Tasmania.

B. Pumice shardsDelicate tube pumice clasts (P) have beeen preserved in intenselyquartz + chlorite (± muscovite)-altered pumice breccia from thefootwall to the Gossan Hill VHMS deposit. The tube vesicles havebeen coated in thin films of chlorite and filled with quartz, andvesicle walls have been altered to quartz. Plane polarised light.Sample 138752, Archaean Golden Grove Formation, GossanHill footwall, Western Australia.

C. ShardsThis quartz nodule (Q) from the footwall, quartz + chlorite(± muscovite) zone contains delicate shard textures. Planepolarised light.Sample 138795, Archaean Golden Grove Formation, GossanHill footwall, Western Australia.

FIGURE 2.6 | Photographs of intensely altered pumice breccias with delicate primary textures.

FIGURE 2.5 | Pairs of hand-specimen and thin-section photographs of increasing intensity of alteration in rhyolitic feldspar-phyric pumice breccia in the Herculesfootwall, northern Central Volcanic Complex, western Tasmania. (A) Hand-specimen and (B) thin-section photographs of subtle, domainal, albite + sericite- andsericite + chlorite-altered pumice breccia (sample MR96-63) showing excellent preservation of volcanic textures. Plagioclase crystals are partly replaced by albite.In albite-rich domains, tube vesicles and clast margins are lined with sericite and albite + quartz altered. In contrast, pumice clasts and shards in the chlorite-richdomains are pervasively sericite + chlorite altered. The Al = 40 and CCPI = 26. (C) Hand-specimen and (D) thin-section photographs of weak, domainal, albite +sericite- and sericite + chlorite-altered pumice breccia (sample MR96-54). Volcanic textures are well preserved in the albite-rich domains and poorly preserved in thechlorite-rich domains. Plagioclase crystals (P) are sericite ± albite ± opaques altered and have albite overgrowths or nodules (alb), which locally preserve delicatevesicular textures. Elsewhere vesicles are coated in sericite and filled with albite. Pumice walls are albite + quartz altered and sericite ± chlorite + hematite fiammeand stylolites are abundant. The Al = 58 and CCPI = 37. (E) Hand-specimen and (F) thin-section photographs of moderate, pervasive, albite + sericite-altered pumicebreccia (sample MR96-48) with partly preserved pumice textures and plagioclase crystals. Sericite fiamme (F) and sericite + hematite stylolites are abundant.Nodules or overgrowths of albite occur around calcite and albite + hematite-altered plagioclase crystals (P). The Al = 70 and CCPI = 38. (G) Hand-specimen and (H)thin-section photographs of strong, pervasive, quartz + sericite + pyrite-altered pumice breccia (sample MR96-50). Primary volcanic textures are faint, with sparsesericite-altered pumice clasts or fiamme (F). Plagioclase crystals (P) are polycrystalline-quartz ± pyrite altered. The Al = 98 and CCPI = 64. (I) Hand-specimen and(J) thin-section photographs of intense, schistose, quartz + sericite + pyrite-altered pumice breccia (sample MR96-46). No relict plagioclase or volcanic textures arepreserved in thin section: in hand specimen irregular lenses of sericite resemble fiamme (F). This alteration facies is pervasive and strongly foliated. The Al = 99 andCCPI = 30.

30 | CHAPTER 2

FIGURE 2.7 | West-east 1700mN section through the K-lens of the Rosebery VHMS deposit, western Tasmania, showing geology and

contoured S/Na2O data.

produce a potential range from zero to one hundred, whichis convenient for quantification of alteration intensity. Theclassic example is the Alteration Index (AI) of Ishikawa et al.(1976):

AI = 100(MgO + K2O)

MgO K2O CaO + Na2O

Originally devised as a measure of intensity of sericite andchlorite alteration associated with the Kuroko-VHMSdeposits, it is useful in many types of plagioclase-destructivehydrothermal alteration systems.

In some cases where there are large differences inmagnitudes between components, some components aremultiplied by appropriate factors to adjust their effect in theindex. An example is the AI mark 4 index,

AI mark 4 = 100(FeO + lOMnO)FeO + lOMnO + MgO + (SiO2/10)

which quantifies alteration in siliciclastic dolomites (Large etal., 2000).

Molar proportion alteration indices are said to be more easilyrelated to the stoichiometry of alteration reactions and henceto alteration assemblages (e.g. Eilu et al., 1997). The extrastep in converting composition data to molar proportions ofoxides or elements is easily achieved in computer spreadsheetsbut it significantly complicates manual calculations. Someexamples of molar indices are the 3K/A1 sericitisation indexand the CO2/CaO carbonation index used in explorationfor lode Au deposits (Davies et al., 1990). The ACNKindex of Hodges and Manojlovic (1993) used the molecularproportions of Al2O3/(CaO + Na2O + K2O) to quantifyintensity of alteration related to metamorphosed massivesulfide deposits at Snow Lake, Manitoba.


The AI-CCPI alteration indices and box plot

The well-known Alteration Index (AI) was developed in theKuroko VHMS deposits, Japan, to represent the principalcomponents gained (MgO and K2O) during chlorite andsericite alteration, and those lost (Na2O and CaO) duringthe breakdown of Na-plagioclase and volcanic glass (Ishikawaet al., 1976). The AI has since been widely used in VHMSmineral exploration to provide quantitative estimates of theintensity of alteration. It typically increases to maximum valuesin the proximal hydrothermal zones beneath massive sulfidelenses (e.g. Saeki and Date, 1980). The AI ranges from 0 to100. High (> 60) values reflect high MgO and K2O contentsrelative to CaO and Na2O, and may be related to intensehydrothermal sericite and chlorite alteration. In contrast, low(< 30) AI values reflect high CaO or Na2O contents that maybe due to intense albite or calcite alteration more typical ofregional diagenetic alteration or metamorphism. For example,at Hellyer AI increases from 35 to 95 from the margin to thecentre of the alteration pipe directly below the ore deposit(Fig. 2.8A and B: Gemmell and Large, 1992; Large et al.,2001a).

There is a strong inverse relationship between AI and Naconcentration (e.g. Fig. 2.8A and D) because loss of Na, andsometimes loss of Ca, is the major chemical change involvedin the breakdown of plagioclase. In many studies Na depletionis used instead of AI as the principal measure of alterationintensity (Date et al., 1983).

The Ishikawa alteration index has two major limitations(Large et al., 2001a). Firstly, it does not take carbonatealteration into account, even though this type of alterationcan be significant in some VHMS alteration systems. WhereCa-carbonates are present they cause a decrease in AI, evenwhere plagioclase destruction is extreme, because CaO isin the denominator. Secondly, the AI effectively measuresplagioclase destruction but does not differentiate chlorite-from sericite-altered rocks. Variations in relative proportionsof chlorite and sericite or spatial relationships between chloriteand sericite zones may be important guides to explorationin some VHMS alteration systems. A geochemical index toquantify the variation would be an improvement on subjectivevisual estimates.

The chlorite-carbonate-pyrite index,

CCPI = 100(FeO +MgO)FeO+MgO+Na2O+K2O

was developed to reflect the prominence of chlorite, Fe-Mg carbonates, and pyrite, which are common minerals inthe proximal altered zones of many VHMS deposits (Largeet al., 2001a). High values of CCPI reflect high FeO andMgO contents, suggesting intense alteration to Fe- or Mg-rich minerals such as chlorite, Fe-Mg-bearing carbonates(dolomite, ankerite or siderite), pyrite, magnetite or hematite.However, the CCPI of least-altered rocks is dependent onprimary composition and magmatic fractionation. Maficrocks with high primary FeO and MgO contents typicallyhave CCPI values greater than 50, whereas more evolved felsicrocks have lower CCPI values between 10 and 50. Thus theCCPI is not well suited to the study of altered mafic rocks.

hangingwall Yvolcaniclastic unit \

FIGURE 2.8 | Alteration intensity in the altered footwall zones at the HellyerVHMS deposit, western Tasmania. (A) Schematic cross-section of the alteredfootwall zones and variations in alteration intensity in these zones as measuredby (B) Alteration Index (Al), (C) Chlorite-carbonate-pyrite index (CCPI), and(D) Na2O. Modified after Gemmell and Large (1992) and Large et al. (2001a).

Used in conjunction with the AI, particularly graphicallyon x-y bivariate plots with AI as the x-axis, the CCPI providesan effective means of discriminating sericite-, chlorite- andcarbonate-rich altered zones. Furthermore, the AI-CCPIbivariate plot, termed the Alteration box plot by Large et al.(2001a), discriminates these VHMS-related hydrothermalalteration assemblages from diagenetic albite- or albite + K-feldspar-bearing assemblages.

Feldspar, phyllosilicate, carbonate and several otheralteration mineral compositions plot around the margins ofthe Alteration box plot (Fig. 2.9). Albite plots at the lowerleft, K-feldspar and pure muscovite at the lower right, chloriteat the top right, and carbonates along the upper margin etcetera. Calcite plots at the top left corner (although CCPI isindeterminate for pure calcite, the merest trace of Fe or Mgwill result in CCPI = 100), magnesite at the top right andthe Ca-Mg carbonates spread between them according to AI

32 | CHAPTER 2

0 10 20 30 40 50 60 70 80 90 100

Al (Ishikawa Alteration Index)

0 10 20 30 40 50 60 70 80 90 100


FIGURE 2.9 | Al - CCPI Alteration box plot for least-altered samples from theMount Read Volcanics, western Tasmania (modified after Large et al., 2001a).The data are classified according to Ti/Zr ratios, where rhyolites have Ti/Zr <10,dacites 10-20 and andesites and basalts >20, and show the effect of magmaticfractionation on the CCPI.

determined by Mg/Ca ratios. Similarly, Mn-carbonates (exceptpure rhodochrosite, which is indeterminate in both indices)plot along the top margin of the Alteration box plot; theirpositions determined by the inevitable minor concentrationsof Ca, Mg and Fe. Siderite, pyrite and Fe-oxides have CCPIvalues of 100, but are indeterminate for Al and thus plot as aline, rather than a point, along the top.

Large et al. (2001a) found that least-altered rocks in theMount Read Volcanics have an Al range of 20 to 65 and aCCPI range of 15 to 85 (Fig. 2.9). A compilation of 1734geochemical analyses for unaltered volcanic rocks fromvarious modern volcanic arcs shows a slightly smaller Al rangeof 20 to 60 and a slightly greater CCPI range of 10 to 90(Fig. 2.10). Hence, least-altered volcanic rocks plot within arectangle near the middle (somewhat left of centre) of theAI-CCPI bivariate plot. This is the least-altered 'box' thatinspired the term box plot. The extent and position of theleast-altered box may vary for data from different districts,according to the diversity of primary compositions.

Fluid-dominated pervasive hydrothermal alterationtends to produce simple equilibrium assemblages of only afew phases. Therefore, intensely altered samples tend to plotoutside the least-altered box and towards the positions ofthe dominant alteration minerals. For example, unalteredcalc-alkaline rhyolites plot in a box towards the centre of theAlteration box plot; with increasing intensity of alteration,altered samples plot progressively further away from theunaltered box (Fig. 2.11). The relative direction of movementaway from the unaltered box is controlled by the alterationassemblage and hence by the alteration process (Large et al.,2001a).

Large et al. (2001a) defined 10 different mineralogicaltrends on the Alteration box plot. Six of these trends relate tocommon VHMS hydrothermal alteration styles and four areassociated mainly with diagenetic alteration (Fig. 2.12).

FIGURE 2.10 | AI-CCPI Alteration box plot for 1734 analyses of rocks frommodern volcanic arcs; most are assumed to be unaltered. Geochemical dataare from Aleutian, Andean, Indonesian and Scotian volcanic arcs, and werecompiled by A.J. Stolz (electronic communication, 1998). These data (classifiedby SiO2 content) show the effects of magmatic differentiation on CCPI, and to alesser degree on Al. Mafic rocks have high CCPI and low to moderate Al valuesbecause of their high Fe, Mg and high Ca contents, respectively. In contrast,felsic rocks have low CCPI because of their low Fe, Mg and high K contents, andhigh Al due to their low Ca and high K contents.

It is important to note that neither Al nor CCPI includesSiO2; thus the Alteration box plot does not provide a directmeasure of the intensity of quartz or silica alteration. Asoutlined in Section 7.2, silica-altered rocks are importantaround some VHMS deposits, including the silicic core zonein the Hellyer alteration pipe (Gemmell and Large, 1992), thestockwork zones of some Kuroko deposits (Shirozu, 1974),and in some Cyprus-type alteration pipes (Lydon, 1984).

The Alteration box plot is a powerful tool for relatinglithogeochemical data to mineral assemblages and alterationintensity in VHMS systems, particularly in felsic volcanicrocks. It has obvious applications in mineral exploration forrecognising favourable alteration styles, delineating alteredzones and providing vectors to ore within large altered systems.A similar dual index approach may be useful for other deposittypes, with different indices designed to highlight specificalteration assemblages.

Element concentrations and mineral abundances

Element concentrations can also be used as guides to alterationintensity. Many alteration studies have used Na depletion asa measure of hydrothermal alteration intensity (e.g. Franklinet al., 1975; Date et al., 1979; 1983; Hashiguchi et al., 1983;Ashley et al., 1988). Typically unaltered modern arc calc-alkaline rhyolites have Na2O values between 3 and 5 wt%(Barrett et al., 1993; Stolz et al., 1996). Rhyolites with greaterthan 5 wt% Na2O are normally albitised, whereas rhyoliteswith less than 3 wt% Na2O reflect feldspar-destructivealteration styles (e.g. sericite, chlorite, pyrite, K-feldspar andcarbonate). The lower the Na2O content the more intenselyhydrothermally altered the rock; thus, Na2O typically

DESCRIBING ALTERED VOLCANIC ROCKS I 33

0 10 20 30 40 50 60 70 80 90 100


FIGURE 2.11 | Al - CCPI Alteration box plot for rhyolites in the northern CentralVolcanic Complex, western Tasmania. Samples are of rhyolitic pumice brecciasfrom the Rosebery and Hercules footwalls. With increasing intensity of footwallchlorite + sericite ± pyrite alteration, Al and CCPI values increase and samplesplot in the upper right of the Alteration box plot.

decreases towards the centre of VHMS alteration systems(e.g. Fig. 2.13).

Alteration mineral abundances quantitatively estimatedfrom whole-rock composition data can also provide measuresof alteration intensity. For example, Large et al. (2001b) foundthat calculated mineral percentages closely approximated thepetrographic estimates of alteration mineral abundances insamples from Rosebery. By plotting the calculated mineralabundances down hole they showed that diagenetic minerals,such as albite, decrease in abundance and hydrothermalminerals, such as sericite, quartz, chlorite and Mn-carbonate,increase in abundance with proximity to ore (Fig. 2.14).Mineral abundances can be calculated as percentages fromthe whole-rock analyses by the least-squares method outlinedin Herrmann and Berry (2002). A free copy of the MINSQ(least-squares spreadsheet method for calculating mineralproportions from whole-rock major element analyses) isavailable to download from the University of Tasmania'sCentre for Ore Deposit Research website <www.codes.utas.edu.au>.

An integrated approach to alteration intensity

A combined compositional and descriptive approach toestimating alteration intensity can also be used. In fact,the Alteration box plot is most powerful when used incombination with petrographic and/or other instrumentalmineralogical studies, such as X-ray diffraction (XRD) orshort wavelength infra-red spectral analysis (e.g. PIMA). Usedin this way the box plot reveals trends in the data, from leastaltered to intensely altered, which can be related to alterationprocesses and hence exploration targets.

FIGURE 2.12 | Schematic AI-CCPI Alteration box plots showing the 10alteration trends recognised by Large et al. (2001a). These provide a tool forgraphically discriminating prospective from non-prospective altered zones and/orsystems. (A) The six trends marked by arrows on this box plot are typical ofhydrothermally altered rocks associated with VHMS deposits. Trend 1: sericitealteration at the margins of the hydrothermal alteration halo in felsic volcanicrocks. Trend 2: footwall sericite + chlorite ± pyrite alteration in felsic and maficvolcanic rocks. Trend 3: chlorite ± sericite ± pyrite alteration, typical of footwall,chlorite-dominated zones in either felsic or mafic volcanic rocks. Trend 4: chlorite+ carbonate alteration typically developed proximal to massive sulfide lensesin the footwall of either felsic or mafic host rocks. Trend 5: sericite + carbonatealteration in the proximal hanging wall to ore deposits or along strike in the hostrocks. Trend 6: K-feldspar + sericite, an uncommon trend developed locallywithin footwall felsic volcanic rocks. (B) The four trends marked by arrows onthis box plot are mainly attributed to diagenetic processes and are unrelated tomineralisation. Trend 7: albite + chlorite alteration, typical of low temperatureseawater-volcanic rock interaction. Trend 8: epidote + calcite ± albite alterationcommon in intermediate and mafic volcanic rocks. Trend 9: K-feldspar + albitealteration. Trend 10: paragonitic sericite + albite alteration.

34 | CHAPTER 2

FIGURE 2.13 | Contoured Na2O data forthe west-east 1700 mN section throughthe K-lens of the Rosebery VHMSdeposit, western Tasmania. Modified afterLarge etal. (2001b).

In the northern Central Volcanic Complex (Mount ReadVolcanics), detailed petrographic descriptions were combinedwith compositional data to assess the range of AI and CCPIvalues for subtly to intensely altered rhyolites (Table 2.6 andFig. 2.15). The least-altered rhyolites were subtly altered andhave comparable alteration indices to unaltered modern arcrhyolites (AI = 30-60 and CCPI = 10-40). Intensely alteredrhyolites have mid to high alteration indices (AI = 40-100 andCCPI = 28-100). AI ranges for subtly, weakly, moderately,strongly or intensely altered rhyolites, dacites, andesites andbasalts are similar. In contrast, the CCPI is influenced by the

primary composition. For this reason the Alteration box plotshould never be used independently as a method of classifyingthe alteration system, but should be integrated with theprimary geochemical and/or petrographic data.

Using a combination of alteration mineral assemblage andcomposition data also enables separation of rock samples intoleast-altered, diagenetically altered and hydrothermally alteredsamples (Gifkins and Allen, 2001; Large et al., 2001a). Thetrend from subtly to intensely hydrothermally altered rocksassociated with VHMS deposits is characterised by increasesin both CCPI and AI, and decreases in Na2O (Table 2.7).

TABLE 2.6 | Alteration indices for altered rhyolites in the northern Central

Volcanic Complex, western Tasmania. The broad range in Al and CCPI values

reflects different alteration styles. For example, strongly altered rhyolites with low

Al values probably reflect diagenetic alteration, whereas high Al values reflect

hydrothermal alteration.

TABLE 2.7 | Alteration indices, Na2O contents and approximate mass changesfor hydrothermally altered rhyolites from the footwalls to the Rosebery andHercules VHMS deposits, western Tasmania. AI and CCPI increase and Na2Odecreases with increasing intensity of alteration.

nyuruinermai aiieiauon.

Alteration intensity

Subtle

Weak

Moderate

Strong

Intense

AI

30-55

25-60

10-75

5-90

40-100

CCPI

10-32

15^5

10-55

28-90

28-100

Na2O (wt%)

3.5-5

2-5.5

1-6

0.5-4.5

0-3

Alteration

intensity

Subtle

Weak

Moderate

Strong

Intense

AI

30-55

40-60

40-75

70-90

90-100

CCPI

10-32

15-45

10-55

30-90

28-100

Na2O

(wt%)

3.5-5

2-A

1-2

0.5-1

0-0.5

Mass changes

(g/100g)

<1

<10

5-30

15-60

15-100


FIGURE 2.14 | Variations in calculated mineral abundances in samples from DDH 120R through K lens of the Rosebery VHMSdeposit, western Tasmania. Increasing intensity of hydrothermai alteration towards the ore lens corresponds with increasing proportionsof chlorite, Mn-carbonate and calcite, and decreasing concentrations of quartz and albite.

36 | CHAPTER 2

0 10 20 30 40 50 60 70 80 90 100

Al (Ishikawa Alteration Index)• intense A moderate • subtle• strong o weak

FIGURE 2.15 | Bivariant plots of rhyolite samples from the northern CentralVolcanic Complex, western Tasmania, which have been classified qualitatively.(A) AI-CCPI Alteration box plot for subtly, weakly, moderately, strongly andintensely altered rhyolites. With increasing intensity of alteration rhyolitesplot away from the subtle box in all directions depending on the alterationcomposition and hence on the processes. (B) Al versus Na2O for subtly, weakly,moderately, strongly and intensely altered rhyolites. With increasing intensity ofalteration rhyolites plot away from the subtle field.

2.6 | ALTERATION DATASHEETS

A practical way of integrating different alteration data,constructing descriptive names and defining alterationfacies is to use alteration data sheets. These visually displaya combination of different data types for a specific rock oralteration facies on one page. They present all the relevantdata for a particular rock sample or alteration facies together,and graphically illustrate relationships between texture,mineral assemblage, composition and alteration zonation,particularly with respect to an ore body. Alteration data sheetsact as 'flash cards', incorporating the distinctive physical and

chemical characteristics of the altered rock or alteration facies,providing quick reference to the data collected and facilitatinginterpretation. Data sheets are used in Chapters 5, 6 and 7 toillustrate the dominant alteration facies or zones associatedwith each of the case studies. The information that is includedon the data sheets may vary because the relevant or availabledata varies in different volcanic successions and in differentdeposits. Where appropriate, data sheets may incorporate:• sample number• location information• geographical or geological feature• formation or group• succession• coordinates• map, cross-section, or alteration zonation model showing

the location of the sample or alteration facies• volcanic facies characteristics• descriptive name for the volcanic facies (see McPhie et al.

(1993) for guidelines)• relict primary minerals• composition (e.g. rhyolitic, dacitic, andesitic or basaltic)

estimated from relict primary minerals and/or geochemicaldata

• lithofacies characteristics• relict textures• interpretation of the volcanic facies and application of

genetic nomenclature (e.g. volcanogenic sedimentarydeposit, resedimented mass-flow or turbidite deposit,syneruptive mass-flow deposit, autobreccia, hyaloclastite,peperite, pyroclastic-flow deposit, pyroclastic-fall depositor pyroclastic-surge deposit)

• alteration facies characteristics• descriptive name for the alteration facies• alteration mineral assemblage• alteration textures• distribution or zonation of alteration facies• alteration intensity• relative timing• interpretation of the alteration process (i.e. diagenetic,

metamorphic, hydration, intrusion-related hydrothermalalteration, proximal or regional hydrothermal alterationand mineralisation, syntectonic hydrothermal alteration)

• photographs of distinctive features of the alteration faciesin outcrop, drill core, hand specimen or thin section

• composition data (whole-rock, mineral-chemistry orisotope analyses)

• chemical characteristics such as important mass changes,alteration indices (e.g. Al and CCPI) and immobileelement ratios (e.g. Ti/Zr)

• Alteration box plot, with the sample highlighted• other significant compositional plots, such as Ti/Zr-SiO2

bivariant plot, mass change bar graph, SWIR spectra etcetera.

I 37

3 | COMMON ALTERATION TEXTURES ANDZONATION PATTERNS

This chapter describes common alteration textures,pseudotextures, and alteration distribution and zonationpatterns in submarine volcanic successions, which can beobserved at a variety of scales: map, outcrop, hand specimenand thin section. It also discusses the use of overprintingrelationships in determining the paragenetic sequence.

Alteration textures, patterns of distribution and zonation,and overprinting relationships are fundamental elements indescribing and interpreting alteration facies (e.g. Fig. 2.3).Alteration textures can aid determination of equilibriummineral assemblages, alteration intensity, and overprintingrelationships. Alteration facies distribution and zonationpatterns can be used to interpret patterns of fluid flow, changesin physicochemical conditions and development of alterationsystems. Superimposed alteration patterns and overprintingtextures are important for determining paragenesis involvingmultiple stages of alteration and hence for understandingevolution of the system over time.

3.1 | ALTERATION TEXTURES

Typically alteration encompasses mineralogical and texturalchanges. Textural changes are changes in the shape, form,grainsize and orientation of grains within the rock and can betexturally destructive, preserve relicts of pre-existing texturesor enhance textures (McPhie et al, 1993; Doyle, 2001).Alteration textures are those that are superimposed on the rockby the processes of alteration (i.e. by hydration, dissolution,diagenesis, hydrothermal alteration, metamorphism anddeformation).

Changes in texture during alteration may involve: theprecipitation of minerals along fluid pathways; creationor infilling of pore space; the dissolution and replacementof earlier minerals and glass by subsequent minerals; andrecrystallisation. There are five common types of alterationtextures that occur in volcanic facies (Tables 3.1 and 3.2):(1) replacement textures, (2) infill textures, (3) dissolutiontextures, (4) recrystallisation textures, and (5) deformationtextures. In addition, the combined effects of a number ofdifferent overprinting alteration facies can result in false orpseudotextures (De Rosen-Spence et al., 1980; Allen, 1988).

Furthermore, there are two types of textures that arecommon in and unique to volcanic rocks, which, althoughnot alteration textures, influence subsequent alteration,especially the development of pseudotextures. These are high-temperature devitrification textures (i.e. spherulites, varioles,lithophysae and micropoikilitic texture) and perlitic fractures(Figs 3.1 and 3.2). The formation and alteration of perlite isdescribed in detail in Section 5.2.

Replacement textures

Most alteration forms by replacement, because pre-existingmineral phases and glass become unstable during changedgeothermal conditions and are readily substituted by new,more stable minerals. Replacement is the process of practicallysimultaneous solution and deposition of a new mineral ofpartly or completely different composition either in a pre-existing mineral or an aggregate of minerals (Lindgren, 1933).Although mineral exchange is essentially simultaneous,replacement may occur in stages, where intermediateproducts form, at least temporarily, before the final alteration

TABLE 3.1 | Types of textural changes that occur during alteration.

Replacement(metasomatism)

Infill

Dissolution

Staticrecrystallisation

Dynamicrecrystallisation

Deformation

Existing minerals or glass are replaced by oneor more new mineral species

A mineral or minerals are precipitated fromsolution into open space

Existing minerals or glass are leachedand removed by solution with or withoutreplacement

Recrystallisation of existing minerals to newgrains, and/or a change in morphology of thesame mineral species or composition

Recrystallisation of existing minerals to newgrains and/or a change in morphology and/ororientation of the same mineral species orcomposition

Existing component or texture is rotated,milled, broken, compressed, modified,distorted or fractured

38 | CHAPTER 3

TABLE 3.2 | Common macroscopic and microscopic alteration textures in volcanic rocks.

minerals. For example, relict radiating fibrous textures locallypreserved in feldspar-altered pumice and perlite clasts in theMount Read Volcanics, western Tasmania, suggest that anintermediate phase between felsic glass and feldspar, possiblyfibrous zeolites, occurred (Fig. 5.11: Gifkins and Allen,2001).

Replacement can range from the conversion of specificmineral phases or domains to new minerals (selectivealteration, Fig. 3.3B, C and D), to complete replacement ofa rock to a completely new mineral assemblage (pervasivealteration, Fig. 3.3A). Where alteration occurs dominantlyby diffusion, it may affect a large volume of rock. Elsewhereit may occur along well-defined fluid pathways (vein-haloalteration, Fig. 3.3E) with its effects restricted to a scale ofmillimetres to metres (Titley, 1994). It is worth noting thatthe terms pervasive, selective or vein-halo depend on the scaleof observation. For example, vein-halo alteration can appearpervasive when viewed in thin section.

Pervasive

Pervasive alteration is extensive alteration that has completelychanged the rock composition and texture at scales that rangefrom millimetres to kilometres (Rose and Burt, 1979; Titley,1982). Pervasive alteration is distributed without regard forpre-existing textures and can result in disseminated, massivemicrocrystalline or cryptocrystalline microscopic textures(Fig. 3.4A).

Selective

Selective alteration converts only specific pre-existing phasesto new mineral phases (Titley et al., 1978; Titley, 1982). Theoriginal rock texture may be only slightly modified duringselective alteration because only certain components in thehost (e.g. minerals, volcanic glass or clasts: Fig. 3.3B, C andD) are preferentially replaced, and others are left relativelyunaltered (Rose and Burt, 1979).

Replacement Pervasive Pervasive, selective, massive, disseminated, microcrystalline,cryptocrystalline

Selective Disseminated Disseminated, pseudomorph, overgrowth, cleavage and rimtexture, core and zonal replacement texture, microcrystalline,cryptocrystalline, spheroid, nodule, concretion

Domainal

Vein halo Pervasive, selective, disseminated, pseudomorph, overgrowth,cleavage and rim texture, core and zonal replacement texture,microcrystalline, cryptocrystalline, spheroid, nodule, concretion

Infill Incomplete infill Crustiform, fibrous, prismatic, spherulitic

Massive infill Microcrystalline, prismatic

Layered or banded infill Crustiform, colloform, comb, botryoidal

Dissolution Stylolitic foliation Stylolites, solution seams

Corrosion vug Open pore space ± infill textures (prismatic, fibrous andmassive)

Static Pervasive (hornfels) Equigranular, granoblastic, granophyric, decussate

recrystallisation Seledive Porphyroblastic, idioblastic, xenoblastic, poikiloblastic,

intergrowths, overgrowths, reaction rims, polycrystalline grains

Dynamic Foliation Slaty cleavage Cleavage, mineral alignment, granoblastic, porphyroblastic,

recrystallisation poikiloblastic

Schistosity

Layering (gneissosity) Differential layering, microcrystalline, granoblastic, granophyric

Lineation Aligned, strained, bent, kinked, flattened, twinned and broken

grains (crystals or clasts), cleavage

Cataclasite No foliation, porphyroblastic, microcrystalline

Mylonite Foliation, granular

Deformation Foliation Cleavage, aligned, strained, bent, kinked, flattened, twinned andbroken grains (crystals or clasts), fiamme, eutaxitic

Lineation Aligned, strained, bent, kinked, flattened, twinned and brokengrains (crystals or clasts), cleavage

Augen structure Cleavage

COMMON ALTERATION TEXTURES AND ZONATION PATTERNS | 39

A. Spherulites and obsidian in rhyolitePink, isolated spherulites and densely microspheruliticflow bands are enclosed in black obsidian in this flow-banded rhyolite. Spherulites are radiating aggregatesor bundles of acicular and fibrous crystals. They varyin shape from spherical to bow-tie shaped sheafs andaxiolitic bundles, and are commonly composed offeldspar or intergrowths of alkali feldspar, plagioclase,cristobalite or tridymite and clinopyroxene (Lofgren,1971b). Spherulites are typically the product of high-temperature (above the glass-transition temperature)devitrification of silicic glass (Lofgren, 1971a).Sample NG4, recent Ngongotaha lava dome, Hendersonsquarry, Rotorua, New Zealand.

B. Lithophysae in rhyoliteThis red albite + quartz + hematite-altered, flow-bandedquartz + plagioclase-phyric rhyolite contains abundantspherulites and lithophysae. The lithophysae are filledwith layered quartz.Sample from the Lower Devonian Snowy River Volcanics,Flukes Knob area, Victoria.

C. Varioles in basaltDark spots in this basalt outcrop are varioles: radial orsheaf-like aggregates of plagioclase and pyroxene, olivineor iron oxides, and are similar to spherulites, but onlyoccur in mafic facies (cf. Fowler et al., 1987; Williamsetal., 1982).Shirakawa quarry, Miocene Green Tuff Belt, Odate,Japan.

D. Micropoikilitic texture in thin sectionThe groundmass of this rhyolite is densely micropoikilitic;comprising patches of optically continuous quartz,which enclose variably oriented laths of sericitised albite.Poikilitic and micropoikilitic texture (snowflake texture)comprise an optically continuous crystal enclosingnumerous randomly oriented inclusions of a differentcomposition (Anderson, 1969). The boundaries betweenthe micropoikilitic quartz domains in this sample arehighlighted by concentrations of sericite. Cross polarisedlight.Sample 133921, Cambrian Mount Black Formation,Central Volcanic Complex, Mount Read Volcanics, MountBlack, western Tasmania.

FIGURE 3.1 | Examples of high-temperature devitrification textures.

40 | CHAPTER 3

A. Altered macroperliteRelict macroperlitic factures in this coherent dacite areenhanced by dark grey sericite + chlorite-altered zonesalong and adjacent to the perlitic fractures. The arcuateshape of the fractures is preserved in some areas. Theperlite cores are pink albite + quartz + sericite altered.Cambrian Mount Black Formation, Central Volcanics

Complex, Mount Read Volcanics, Pieman Road, western

Tasmania.

B. Relict perlite in thin sectionThe formerly glassy groundmass of this rhyolite preservesperlitic fractures. Perlitic fractures are a network of finetypically concentric, arcuate fractures that enclose glassyor originally glassy cores. Here, the perlitic fractures arefilled with dark, mixed-layer smectite/chlorite and thegroundmass adjacent to the fractures is clinoptilolitealtered. The perlitic cores are partly glassy and partlysmectite altered. Plane polarised light.Sample J6-737 m, Miocene Nishikurosawa Formation,Hokuroku Basin, Green Tuff Belt, Odate, Japan.

C. Banded perliteThis finely flow-banded, plagioclase-phyric rhyolitecontains an intersecting fracture network of sub-parallellong fractures linked by short cross fractures (bandedperlite) superimposed on the flow-banded texture.Sample KB257, Siluro-Devonian rhyolite, Ural Volcanics,

Ural Ridges area, New South Wales.

D. Banded perlite in thin sectionIn thin section, concentrations of sericite ± hematitemark the relict perlitic fractures. The pale flow bandscomprise a fine-grained mosaic of feldspar + quartz,whereas the darker bands consist of sericite + feldspar+ quartz + chlorite. Disseminated fine-grained hematiteoccurs throughout the groundmass. Plane polarisedlight.Sample KB257, Siluro-Devonian rhyolite, Ural Volcanics,

Ural Ridges area, New South Wales.

FIGURE 3.2 | Examples of perlite.

Two textural types of selective alteration occur: disseminatedalteration (or selective-pervasive alteration), which refers tothe replacement of selective pre-existing phases throughoutthe host rock; and domainal alteration, which refers to thealteration of patches, pods, or groups of clasts within the hostrock (Fig. 3.3F, G and H). In addition, selective alteration mayresult in concentrically zoned alteration facies within clasts oralteration halos around clasts (Fig. 3.31, J, K and L). Selectivealteration can result in a patchy or mottled appearance (e.g.Allen, 1988).

Common microscopic selective replacement texturesare pseudomorphs, partial pseudomorphs (cleavage and rimtexture, core and zonal texture, core and rim texture andskeletal texture), overgrowths on pre-existing components,and spheroids or nodules (Fig. 3.4B to L: Dimroth andLichtblau, 1979; Craig and Vaughan, 1981; Ineson, 1989).Carbonate and zeolite nodules are common in submarinevolcaniclastic facies and can have a wide variety of grainsizesfrom 0.2 to greater than 20 mm (Fig. 3.41 to L: Franklin etal., 1975; Lees, 1987; Khin Zaw and Large, 1992; Hill andOrth, 1994; Allen, 1997).

Vein halo

Vein-halo alteration involves the replacement of either thewhole rock (pervasive alteration: Fig. 3.3E) or specific pre-existing phases (selective alteration) in restricted areas, such asthe halos around veins, intrusion contacts, or at stratigraphiccontacts. Alteration progresses in fronts, moving out fromfractures or contacts into the adjacent wall rock. Vein-haloalteration has also been termed infiltration metasomatism,vein-veinlet, reaction rim, vein-wall-rock, vein-envelope,veinlet-controlled and fracture-controlled alteration (e.g.Titley et al., 1978; Titley, 1982; Thompson and Thompson,1996; Doyle, 2001).

Infill textures

Infill or open space-filling textures result from the precipitationof new mineral phases from solution into open spaces orcavities such as pore spaces, vesicles, inter-clast space, vugsand fractures (Taylor, 1992). Infill textures are characterisedby well-developed crystal faces, zoned crystals and mineralbanding (Craig and Vaughan, 1981). Silicate, carbonates,oxides, sulfates and sulfides all occur as void fill in alteredvolcanic rocks.

Infill results from the precipitation of minerals fromsolution. The first mineral to be deposited forms a cruston the cavity walls and grows inwards, generally with thedevelopment of inward facing crystal faces. Common infilltextures include incomplete infill, massive infill, and layeredor banded infill (Fig. 3.6: Taylor, 1992). These textures caninclude fibrous, prismatic, spherulitic or equant crystal shapesand exist on a range of scales from micrometres to metres(Dimroth and Lichtblau, 1979; Taylor, 1992).


Incomplete infill

Incomplete or partial infilling of veins and cavities ordissolution of void fill can leave an open vug in the centre (e.g.Fig.3.5A). In many cases, the resulting infill texture containswell-formed crystals that project into this vug.

Massive infill

Massive infill textures result from the continuous depositionof a mineral or aggregate of minerals until the cavity is filled(e.g. Fig. 3.5B). Massive infill commonly contains well-formedcrystals, especially quartz, feldspar, fluorite, cassiterite, galena,sphalerite and chalcopyrite crystals. Massive, microcrystallineforms also exist (Taylor, 1992).

Layered infill

Layered or banded infill textures result from the deposition ofa succession of minerals inwards from the cavity or fracturewall (Bateman, 1951). Layered infill textures do not generallycontain well-formed crystals, such as comb texture (e.g. Fig.3.5H), but vary from thin layers of individual minerals tocrustiform bands or colloform textures.

Dissolution textures

Dissolution textures are common in altered volcanic rocks(Allen, 1990; Allen and Cas, 1990; Marsaglia and Tazaki,1992; Gifkins and Allen, 2001). They form from the corrosionor leaching of pre-existing phases (either glass or mineralphases), with or without minor replacement by new mineralphases (Fisher and Schmincke, 1984). For example, leachingof rhyolitic glass is commonly accompanied by crystallisationof muscovite or clay minerals that absorb leached ions fromsolution (Karkhanis et al., 1980).

Dissolution textures include corrosion vugs or dissolutionpits, stylolites and solution seams (e.g. Fig. 3.6: Pettijohn,1957).

Corrosion vugs

Dissolution or corrosion of volcanic glass or pre-existingminerals can create open cavities or oversized pores (Fig. 3.6Ato F) in which infill can occur synchronous with dissolution orafter dissolution (Hay, 1963; Sheppard et al., 1988). In somecases pseudomorphs of minerals or originally glassy particles,such as glass shards, form by dissolution and precipitation(Riech, 1979; Sheppard etal., 1988). Riech (1979) recognisedinfill textured zeolites and calcite within clinopyroxenes,and proposed that clinopyroxenes were corroded duringdiagenesis, creating an open void that was subsequently filledwith zeolites and calcite. Similarly, Hay (1963) recognisedpartial to complete dissolution of glass shards followed by theprecipitation of authigenic minerals, especially zeolites, in thenew cavities as well as in original pore space.

42 | CHAPTER 3

A. Pervasively altered rhyoliteIntense, pervasive, fine-grained K-feldspar + quartzalteration has completely replaced the groundmass andplagioclase phenocrysts in this rhyolite.Sample 143286, Central Volcanic Complex, Mount ReadVolcanics, Mount Darwin, western Tasmania.

B. Selectively altered phenocrystsSericite has selectively altered the coarse prismaticfeldspar phenocrysts (F) in this latite. The pale green-grey, fine-grained groundmass is moderately andpervasively phengite + chlorite + ankerite altered, andthe amygdales are quartz filled.Sample 144369, Ordovician Lake Cowal Volcanics, Junee-Narromine Volcanic Belt, Endeavour 42 prospect, NewSouth Wales.

C. Selectively altered pumice clastsLarge pumice clasts (P) in this sample of crystal- andpumice-rich volcaniclastic breccia have been selectivelyaltered to orange albite + quartz, whereas the finergrained matrix has been altered to green sericite +chlorite + albite. The domainal alteration style enhancesits clastic appearance.Sample 131993, Cambrian Mount Julia Member,Tyndall Group, Mount Read Volcanics, Comstock, westernTasmania.

D. Selectively altered matrixIn this andesitic volcaniclastic breccia, the matrix ismoderately and selectively epidote altered. In contrast,the plagioclase-phyric clasts (C) are weakly chlorite +sericite altered.Sample 144805, Ordovician Mingelo Volcanics, Junee-Narromine Volcanic Belt, Peak Hill, New South Wales.

FIGURE 3.3 | Examples of replacement textures in hand specimen.

COMMON ALTERATION TEXTURES AND ZONATION PATTERNS I 43

E. Vein haloRed albite altered zones are restricted to 5 mm halos orselvages adjacent to quartz + actinolite + pyrite veinletsin this feldspar porphyritic dacite.Sample TH386 271.1 m, Cambro-Ordovician TrooperCreek Formation, Seventy Mile Range Group, MountWindsor Subprovince, Thalanga, Queensland.

F. Banded, domainal alteration faciesDiffuse and discontinuous pink and green bands in thismassive crystal-rich volcaniclastic sandstone are definedby domains of albite + quartz ± chlorite, and chlorite+ sericite + magnetite alteration facies, respectively. Thebands are not obviously consistent with grainsize orcomponent variations; they alternate on a 2—10 cm scale,are laterally extensive (10—20 m) and are commonly, butnot exclusively, bedding parallel.Sample 131982, Cambrian Mount Julia Member, TyndallGroup, Mount Read Volcanics, Lyell Comstock, westernTasmania.

G. Patchy, domainal alteration faciesThe domainal, green epidote + quartz and grey albite+ quartz + hematite alteration facies are distributed inirregular patches with diffuse margins in this coherentplagioclase-phyric dacite.Sample M142, Cambrian Mount Black Formation,Central Volcanic Complex, Mount Read Volcanics, Tullah,western Tasmania.

H. Domainal alteration facies in pseudobrecciaDomainal red albite + quartz and dark green epidote +sericite + albite alteration facies in this sample of macro-perlite gives it a pseudo-polymictic and -clastic texture.However, the red apparent clasts have diffuse marginsand identical phenocryst populations to the apparentmatrix.Sample 147550, Cambrian Mount Black Formation,CentralVolcanics Complex, Mount Read Volcanics, PiemanRoad, western Tasmania.

FIGURE 3.3 | Examples of replacement textures in hand specimen, cont.

44 | CHAPTER 3

I. Zonation within clastsThe andesite and basalt clasts in this polymictic volcanicbreccia are concentrically zoned, with sericite + quartz+ calcite-altered rims, and chlorite-altered cores. Someof the larger clasts have an additional quartz + sericite+ chlorite-altered core zone. The matrix has beenmoderately and pervasively quartz + sericite + calcite ±chlorite altered.Cambrian Que-Hellyer Volcanics, Mount Charter Group,

western volcano-sedimentary sequences, Mount Read

Volcanics, Hellyer, western Tasmania.

J. Zonation within clastsClasts in this basaltic pebble conglomerate displayheterogenous alteration facies, and some clasts areinternally zoned. The basalt clast (B) has a fine-grained,pale green sericite-rich rim, and darker sericite + chlorite-altered core.Sample 134632, Cambrian Red Lead Formation correlate,Dundas, Kapi Creek, western Tasmania.

K. Zonation within pillowsThis metamorphosed, amphibolite-grade lava-pillow hasa typical triangular, draped shape and is concentricallyzoned. The central red zone is coarse-grained, scapolite-poor and albite + hematite + sericite ± epidote altered.The average grainsize decreases, and scapolite grainsizeand abundance increases, in consecutive zones towardsthe rim. Biotite + calcite + hornblende + microcline +scapolite + epidote + quartz comprise the inter-pillowmatrix.Proterozoic Corella Formation, Mary Kathleen Group,

Malbon River, northwest Queensland.

L. Altered halos around clastsOrange albite + quartz alteration facies is distributed ina halo around a massive, albite-altered dacite clast (C) inthis crystal- and lithic-rich volcaniclastic sandstone. Themore pervasive green-grey domain is sericite + chlorite +quartz + albite altered.Sample 132090, Cambrian Mount Julia Member, TyndallGroup, Mount Read Volcanics, Anthony Road, westernTasmania.

FIGURE 3.3 | Examples of replacement textures in hand specimen, cont.


A. Microcrystalline texture in thin sectionThe groundmass of this rhyolite is a microcrystallinemosaic of quartz + feldspar + sericite. Quartz phenocrysts(Q) have been recrystllised. Microcrystalline texture(aphanitic) is a fine-grained granular texture where theindividual crystals can be distinguished in thin section.In contrast, cryptocrystalline texture (phaneritic) iswhere the crystals are too minute to be distinguishedeven with the aid of a microscope (Williams et al.,1982). Cross polarised light.Sample 133318, Cambro-Ordovician Mount WindsorFormation, Seventy Mile Range Group, Thalanga,Queensland.

B. Pseudomorphs in thin sectionThe plagioclase phenocrysts in this sericite + quartz+ tourmaline-altered andesite were pseudomorphedby tourmaline, and subsequently almost completelyreplaced by blue-green chlorite. Pseudomorphs arecrystals or aggregates of crystals that preserve the shapeof a pre-existing mineral or particle (e.g. glass shard orpumice clast) (Spry, 1976). Plane polarised light.Sample 145199, Ordovician Forest Reefs Volcanics, MolongVolcanic Belt, Black Rock, New South Wales.

C. Pseudomorphs in thin sectionThis thin section of plagioclase + quartz + pyroxene-phyric rhyolite shows an illite pseudomorph afterpyroxene. Plagioclase phenocrysts have been altered toK-feldspar and the groundmass comprises a fine-grainedmosaic of K-feldspar + quartz + chlorite + smectite.Cross polarised light.Sample KB495, Siluro-Devonian Coan rhyolite, MountHope Volcanics, Coan Gonn Peak, New South Wales.

D. Cleavage and rim texture in thin sectionThe plagioclase crystals in this basalt have been selectivelyaltered by sericite along cleavage planes. Cleavage andrim textures occur by selective alteration of mineral grainboundaries and cleavages. It is common in plagioclase,in which montmorillonite, sericite or calcite form alongthe cleavage planes (Sales and Meyer, 1948). Planepolarised light.Sample SVD87a-104.9 m from the Cambrian SterlingValley Volcanics, Mount Read Volcanics, Sterling Valley,western Tasmania.

FIGURE 3.4 | Examples of replacement textures in altered volcanic rocks.

46 I CHAPTER 3

E. Core and zonal texture in thin sectionZones within plagioclase phenocrysts in this subtly,smectite + calcite-altered diorite have been selectivelyaltered to sericite. These incomplete pseudomorphs,termed core and zonal texture, are particularly commonin zoned feldspar, amphibole and mica crystals wherethe cores, or one or more zones in zoned minerals,are altered (Barker, 1990). In plagioclase crystals, likethose pictured here, the calcic zones are typically alteredto calcite or sericite (Sales and Meyer, 1948). Planepolarised light.Sample 152958, Pliocene-Pleistocene Luise Volcano, LihirIsland, New Ireland Province, Ladolam epithermal Aumine, Papua New Guinea.

F. Core and zonal texture in thin sectionIn this example of core and zonal texture, the core zonesof plagioclase phenocrysts have been altered to sericite.The groundmass of this plagioclase + clinopyroxene-phyric basalt was subtly and pervasively smectite +calcite-altered. Plane polarised light.Sample 152830, Pliocene-Pleistocene Luise Volcano, LihirIsland, New Ireland Province, Ladolam epithermal Aumine, Papua New Guinea.

G. Overgrowth texture in thin sectionA discontinuous K-feldspar overgrowth encloses ahematite-altered plagioclase phenocryst (P) in thisstrongly and pervasively albite + quartz + sericite-alteredpumice breccia. K-feldspar nucleated on the plagioclasephenocryst, spread outwards filling vesicles, and replacedvesicle walls in the pumice clasts. Overgrowth texturesare mineral rims that may be composed of one or morecrystals of similar or different minerals. Cross polarisedlight.Sample 133814 from the Cambrian Hercules PumiceFormation, Central Volcanic Complex, Mount ReadVolcanics, Hercules footwall, western Tasmania.

H. Altered nodules in pumice brecciaBlue-green celadonite nodules have overprinted pumiceclasts in this polymictic volcanic breccia. These nodulesare composed of fine-grained aggregates of celadonite ±opal CT ± quartz, preserve uncompacted tube and roundvesicle pumice textures, and are surrounded by pervasivesmectite + mordenite + calcite alteration facies.Sample FK2, Miocene Byobu-iwa Member, TokiwaFormation, South Fossa Magna, Green Tuff Belt, FujikawaRiver, Japan.

FIGURE 3.4 | Examples of replacement textures in altered volcanic rocks, cont.


I. Carbonate spheroidsLarge dolomite spheroids are enclosed in the stronglychlorite + quartz + dolomite-altered matrix of thisformerly plagioclase-phyric andesite. Nodules andspheroids are spherical domains of alteration, whichmay comprise radiating aggregates of fibrous crystals,fine internally concentric structures, or mosaics ofanhedral grains, with or without cores (Allen, 1997; HillandOrth, 1995).Sample 135756, Cambrian Que-Hellyer Volcanics,Mount Charter Group, western volcano-sedimentarysequences, Mount Read Volcanics, Hellyer footwall, westernTasmania.

J. Carbonate spheroids in thin sectionIn thin section, the dolomite spheroids displayconcentric zones and a coarse, radiating, fibrous texture.This compositional zonation in the spheroids probablyindicates multiple stages of carbonate alteration (cf. Hilland Orth, 1995). Plane polarised light.Sample 135756, Cambrian Que-Hellyer Volcanics,Mount Charter Group, western volcano-sedimentarysequences, Mount Read Volcanics, Hellyer footwall, westernTasmania.

K. Carbonate spheroidsCarbonate spheroids are concentrated in individual bedsin this strongly chlorite + carbonate + pyrite-alteredlaminated volcaniclastic sandstone. The larger spheroids,which are up to 2 mm in diameter, have coalesced.Sample 138601, Archaean Mb5 Golden Grove Formation,Luke Creek Group, Murchison Volcanics, Golden Grove,Western Australia.

L. Carbonate spheroids in thin sectionThin section examination shows these carbonatespheroids are supported in a fine-grained quartz +sericite + carbonate matrix. Plane polarised light.Sample 138601, Archaean Mb5 Golden Grove Formation,Luke Creek Group, Murchison Volcanics, Golden Grove,Western Australia.

FIGURE 3.4 | Examples of replacement textures in altered volcanic rocks, cont.

48 | CHAPTER 3

A. Incomplete infill in fracturesIncomplete filling of fractures in this altered diorite leftsub-planar vugs. The pyrite fill has botryoidal surfaces,representing rounded shapes of either spheruliticradiating aggregates of fibrous crystals or fine-concentricinternal structures (cf. Jensen and Bateman, 1981). Thismassive plagioclase-phyric diorite has been pervasivelyK-feldspar + pyrite (>quartz + illite) altered, anddissolution of primary mafic minerals produced a fine,spongy, porous texture.Sample 152959, Pliocene-Pleistocene Luise Volcano, LihirLsland, New Ireland Province, Ladolam epithermal Aumine, Papua New Guinea.

B. Massive infillPale green epidote altered halos surround massivechlorite-filled amygdales (A) in this basalt sample.Sample 144753, Ordovician, Junee-Narromine VolcanicBelt, Boundary Prospect, Lake Cowal, New South Wales.

C. Layered infill stringer veinThis banded vein consists of successive layers, from thevein wall to centre, of quartz, quartz and intergrownchalcopyrite and pyrite, and dolomite. A thin dolomitevein has overprinted the stringer vein at an obliqueangle. These veins are hosted in strongly and pervasivelysericite + chlorite + albite + pyrite-altered andesite.Cambrian Que-Hellyer Volcanics, western volcano-sedimentary sequences, Mount Read Volcanics, Hellyerfootwall, western Tasmania.

D. Layered infill in amygdalesAmygdales (A) in this basalt clast, from a basalt-mudstone peperite, contain concentric layers of quartzand calcite, which have grown inwards from the vesiclewalls. The basalt groundmass has been pervasivelysericite + chlorite + calcite altered. The clast grainsizedecreases towards the clast rim, to the left of the field ofview in this photograph.Sample 76836, Cambrian Que-Hellyer Volcanics, westernvolcano-sedimentary sequences, Mount Read Volcanics,Hellyer, western Tasmania.

FIGURE 3.5 | Examples of infill textures in altered volcanic rocks.


E. Layered infill texture in vesiclesVesicles (V) in this plagioclase-phyric pumice clast havebeen filled with roughly concentric layers of tan-colouredmordenite, dark smectite and clear clinoptilolite. Thezeolites occur in clusters or aggregates of fine, radiatingfibres. The originally glassy vesicle walls (W) have beenreplaced by mordenite + K-feldspar ± smectite. Planepolarised light.Sample OH8-387 m, Miocene Onnagawa Formation,Hokuroku Basin, Green Tuff Belt, Odate, Japan.

F. Layered infill texture in amygdalesAmygdales in this subtly altered perlitic rhyolite havebeen filled with bands of fine-grained montmorilloniteand unknown radiating fibrous minerals. Pale polarisedlight.Sample J6-737 m, Miocene Nishikurosawa Formation,Hokuroku Basin, Green Tuff Belt, Odate, Japan.

G. Layered infill in amygdalesThe amygdales in this palagonite-altered trachytic basaltclast from a crystal- and lithic-rich pumice breccia arefilled with layers of montmorillonite and fibrous zeolites.Plane polarised light.Sample OH8-387 m, Miocene Onnagawa Formation,Hokuroku Basin, Green Tuff Belt, Odate, Japan.

H. Comb textureThis example of comb texture shows layers of prominentsparry quartz + amethyst ± carbonate crystals projectinginwards from the vein or cavity wall.Sample T5> Cretaceous, andesite, Fresnillo epithermal

district, Mexico.

FIGURE 3.5 | Examples of infill textures in altered volcanic rocks, cont.

50 | CHAPTER 3

A. Dissolution vugsThis hand specimen of polymictic breccia has a spongyor vuggy porous texture due to the dissolution of primarymafic igneous minerals and glass. It has been intenselyand pervasively adularia + illite + pyrite altered with illitereplacing plagioclase crystals, secondary K-feldspar inthe altered matrix, and disseminated pyrite.Sample 152726, Pliocene-Pleistocene Luise Volcano, LihirIsland, New Ireland Province, Ladolam epithermal Aumine, Papua New Guinea.

B. Dissolution vugs in thin sectionIn thin section, irregularly shaped, empty, corrosionor dissolution vugs (V) are conspicuous in the matrixand clasts. Some vugs cut across clast margins. Planepolarised light.Sample 152726, Pliocene-Pleistocene Luise Volcano, LihirIsland, New Lreland Province, Ladolam epithermal Aumine, Papua New Guinea.

C. Filled dissolution vug in thin sectionCorrosion vugs, created by the dissolution of volcanicglass or pre-existing minerals, are commonly filledby subsequent mineral precipitation from solution.Successive layers of montmorillonite and zeolite havefilled an irregular vug (V) in this thin section. The vugoccurs in the matrix and in a basalt clast, crossing theclast-matrix contact. Plane polarised light.Sample OH8-387 m, Miocene Onnagawa Formation,Hokuroku Basin, Green Tuff Belt, Odate, Japan.

D. Vuggy quartzThe prominent features in this quartz-rich sample are thecorrosion vugs, which were generated by the dissolutionof pumice clasts and crystals from this pumice and lithictuff.Miocene rhyodacitic pumice and lithic tuff, Pierina Au-Agdeposit, Peru.

FIGURE 3.6 | Examples of dissolution textures in altered volcanic rocks.


E. Kaolinite + dickite-altered andesiteLarge (up to 4 mm), blocky feldspar phenocrysts havebeen kaolinite altered in this sample of massive, coherentandesite.Miocene andesitic lava, Pierina Au-Ag deposit, Peru.

F. Vuggy quartzIn this sample of vuggy quartz, which is equivalent tothe previous kaolinite + dickite-altered andesite, thefeldspar phenocrysts have been dissolved resulting inblocky vugs. The groundmass is composed dominantlyof quartz.Miocene andesitic lava, Pierina Au-Ag deposit, Peru.

G. Stylolite in thin sectionStylolites (S2) in this rhyolitic pumice breccia haveconcentrated fine-grained opaques and sericite ±chlorite. The stylolites define the compaction foliationand are crenulated by the dominant regional cleavage{S2) defined by alignment of sericite in the subtly albite+ quartz + sericite-altered matrix.Sample 147422, Cambrian Kershaw Pumice Formation,Central Volcanics Complex, Mount Read Volcanics,Rosebery, western Tasmania.

E. Solution seams in thin sectionThese analcime-filled solution seams occur in a smectite-rich fiamme, extending from the damme terminationsinto the shard- and crystal-rich matrix of a crystal-richpumice breccia. They are interpreted to have formedby dissolution and precipitation under the influence oflithostatic load during diagenesis. Plane polarised light.Sample FK7, Miocene Wadaira Tuff Member, TokiwaFormation, South Fossa Magna, Green Tuff Belt, Wadaira,Japan.

FIGURE 3.6 | Examples of dissolution textures in altered volcanic rocks, cont.

52 | CHAPTER 3

Vuggy silica (quartz) alteration facies is characterised byfine-grained, microcrystalline quartz and abundant open vugsor pores, which may be partly infilled (e.g. Fig. 3.6D, E andF). It is common in high-sulfidation epithermal systems andresults from the extensive leaching of all phases, except SiO2

and TiO2, from volcanic rocks by hot acid solutions (Whiteand Hedenquist, 1990).

Stylolites

Stylolites are common in altered volcaniclastic rocks (Allen,1990; Allen and Cas, 1990; Marsaglia and Tazaki, 1992;Gifkins and Allen, 2001). They are surfaces of dissolutionassociated with strain (pressure solution). They are roughlyplanar surfaces that exhibit mutual column and socket inter-digitation and may branch. Stylolites result from mechanicalcompaction and removal of elements by diffusion andprecipitation (Merino et al., 1983). They indicate volumeloss and may form parallel or sub-parallel to bedding duringburial, or at high angles to bedding during folding. Stylolitesoften contain a residue of insoluble material and mineralsprecipitated from solution. RecrystaUisation, dissolution,grain growth, grain orientation, pressure twinning, fracturingand residual accumulation of minerals along stylolites arecommon (Amstutz and Park, 1967).

Irregular, anastomosing, bedding-parallel stylolites havebeen recognised in originally glassy volcanic facies, especiallypumice breccias, in the Mount Read Volcanics (Allen, 1990;Allen and Cas, 1990; Gifkins and Allen, 2001). These areseams that concentrate fine-grained opaques and sericite atthe margins of originally glassy clasts, along tube vesicle wallsin pumice clasts and in the matrix (Fig. 3.6D: Gifkins, 2001).These stylolites are interpreted as diagenetic compaction anddissolution fabrics that formed by the dissolution of solublecomponents, particularly glass, and by the precipitation ofclays and Fe-oxides as a result of pressure during burial (Allen,1990; Allen and Cas, 1990; Gifkins, 2001).

Solution seams

Solution seams are non-sutured, discontinuous mineral-filled seams that may form during diagenesis as a result ofstress-related dissolution of soluble components and re-precipitation (Merino et al., 1983). Analcime-filled solutionseams in pumice-rich rocks from the Green Tuff Belt (Japan)are anastomosing and roughly parallel to bedding. They occurin the fine-grained matrix and within fiamme and partlycompacted pumice clasts (Fig. 3.6E: Gifkins et al., in press).

Static recrystaUisation textures

RecrystaUisation is the transformation of a mineral or glassto a new grainsize, morphology or orientation of the samemineral species or minerals of the same composition (i.e.neomorphism, Folk, 1965). Pre-existing minerals recrystalliseto a new grainsize in an attempt to assume a more stable formby minimising the ratio of the surface area to the volumeduring changed physical conditions (Yardley, 1989).

RecrystaUisation textures are produced by changes inthe size, shape and arrangement of minerals in a rock. Withincreasing temperature, recrystaUisation generally involves thechange from fine to coarse grainsize (aggrading), except forstatic recrystaUisation to hornfels and dynamic recrystaUisationwhere large, strained grains are replaced by a mosaic of tiny,unstrained crystals (Folk, 1965). Minerals may be directedor randomly orientated (non-directed: Spry, 1976). Directedtextures occur where recrystaUisation is accompanied by stress(dynamic recrystaUisation). Non-directed textures occur wherethe pressure is equal in all directions (static recrystaUisation).

Common macroscopic and microscopic recrystaUisationtextures include mineral overgrowths, porphyroblasts,poikoblasts, and hornfelsic, granoblastic, granophyric anddecussate textures (Fig. 3.7).

Dynamic recrystaUisation textures

Directed fabrics or textures are common to regionalmetamorphic rocks where recrystaUisation is accompaniedby stress (dynamic recrystaUisation). These textures includesubgrains (granoblastic, porphyroblastic and poikiloblastictextures), foliations, layering and lineations.

In common usage the term foliation is non-genetic anddescribes any planar, spaced or pervasive fabric in the rock.Foliations may form during diagenesis, metamorphismor tectonic deformation. Planar foliations are due to thepreferred orientation of minerals, particularly micas, alignedperpendicular to the maximum compression direction(Yardley, 1989). Planar foliations are subdivided on the basisof grainsize and overall appearance of the altered rock. Theseinclude slaty cleavage, schistosity and gnessic layers.

In fault zones or zones of intense ductile shear, twocharacteristic textures occur: cataclastic and mylonitictextures. Cataclase refers to a fine-grained fault gouge brecciawith an unfoliated matrix (Sibson, 1977). Ideally, cataclasis ismechanical fragmentation without recrystaUisation, howeverthis rarely occurs in nature (cf. Sibson, 1977). Mylonite is aterm used for strongly foliated fine-grained rocks in whichthe grainsize has been reduced by recrystaUisation (Bell andEtheridge, 1973).

Deformation textures

Deformation textures are stress activated and develop inresponse to overburden pressures or regional tectonic stress.Because volcanic deposits typically have high initial porositiesthey are easily modified by mechanical compaction duringburial, tectonic deformation or, in the case of pyroclasticdeposits, welding (Peterson, 1979; Allen, 1988; Branneyand Sparks, 1990). Textural modification associated withcompaction is essentially a result of increased pressure causingthe re-arrangement and deformation of grains and reductionof intergranular pore space (Deelman, 1975).

Deformation textures result from the rotation, brittlefracturing, flattening and distortion of existing grains orfabrics, especially clasts that were previously altered to softminerals (McBride, 1978; Galloway, 1979; Craig andVaughan, 1981; Branney and Sparks, 1990).


A. PorphyroblastsThe large, strongly altered, cordierite porphyroblasts givethis rhyolite a distinctive coarsely spotty texture. Thistexture inspired the terms dalmatianite, which was appliedby early workers in the Noranda Camp, and the spottedfades, which was applied by Riverin and Hodgson (1980),for the cordierite-altered Amulet rhyolite and Millenbachandesite. The porphyroblasts are enclosed in a groundmassof chloritised biotite + sericite + quartz. Porphyroblasts aremetamorphic crystals that are surrounded by a much finergrained matrix of other minerals (Spry, 1976). These largeminerals have formed at the expense of the matrix and arethe metamorphic equivalent of phenocrysts.Archaean Amulet Rhyolite Formation, Noranda, Abitibigreenstone belt, Amulet Upper A deposit, Canada.

B. Porphyroblasts in gneissThis biotite + garnet gneiss is characterised by spotty5 mm diameter garnet porphyroblasts in a medium- tofine-grained quartz + feldspar + biotite groundmass. Thegarnet porphyroblasts commonly have a biotite rim. Agneiss is a rock with coarsely differentiated layeringdenned by the segregation of minerals of differentcomposition (typically dark and light minerals) inmedium- to coarse-grained, granular rocks. Layeringforms parallel to the tectonic foliation and in this casedeviates around the garnet porphyroblasts. The precursoris interpreted to have been a felsic volcaniclastic rock.Sample 154061, Proterozoic Potosi gneiss, Harp prospect,Broken Hill Block, New South Wales.

C. Porphyroblasts in thin sectionThis sample of garnet hornfels, from the contact zonebetween rhyolite and a diorite intrusion, containseuhedral garnet porphyroblasts in the biotite + muscovite+ quartz groundmass. Porphyroblasts, like these, withwell-developed crystal shapes are idioblastic or euhedral,whereas those with poorly developed crystal shape arexenoblastic or anhedral (Yardley, 1989). Plane polarisedlight.Sample 140868, Cambro-Ordovician Mount WindsorFormation, east Thalanga, Mount Charter, Queensland.

D. Poikiloblast in thin sectionThis thin section of amphibolite displays an amphibolepoikiloblast with quartz and biotite inclusion trails.Poikiloblasts are porphyroblasts that contain numerousinclusions that may or may not show a preferredorientation (Barker, 1990). Poikiloblastic texture isanalogous to poikilitic or micropoikilitic texture. Usuallythe inclusions are minerals that also occur in the matrix(Yardley, 1989). In this sample, the biotite inclusiontrails display snowball rotation indicating syntectonicgrowth of the amphibole. Plane polarised light.Sample 3215, Proterozoic Corella Formation, MaryKathleen Group, Malbon River, northwest Queensland.

FIGURE 3.7 | Examples of static recrystallisation textures in altered volcanic rocks.

54 | CHAPTER 3

Common deformation textures include intergranulartextures and fabrics such as foliations, lineations, and augen-structure, and intragranular textures such as strained, bent,kinked, flattened, twinned and broken grains (crystals orclasts), as well as irregular grain contacts (e.g. Fig. 3.8A,B and C: Deelman, 1976; Spry, 1976). Deformation canmodify pre-existing textures such as volcanic, hydration anddevitrification textures (e.g. Fig. 3.8D to H). Fiamme andeutaxitic deformation textures are unique to volcanic facies(e.g. Fig. 3.9: Ross and Smith, 1960; Allen and Cas, 1990;McPhie et al., 1993; Gifkins et al., in press). Fiamme andeutaxitic textures are characteristic of, but are not restrictedto, welded ignimbrites (e.g. Fig. 3.9A and B: Ross and Smith,1960; Smith, I960), welded pyroclastic fall deposits (e.g.Sparks and Wright, 1979), welded autobreccia (e.g. Sparkset al., 1993) and pyroclastic deposits that have undergonesecondary welding as a result of contact with hot lava orintrusions (e.g. Ross and Smith, I960; Christiansen andLipman, 1966; Schmincke, 1967; McPhie and Hunns, 1995).Similar fiamme and eutaxitic texture also occur in non-weldedaltered pumice-rich rocks (e.g. Fig. 3.9C and D: Fiske, 1969;Allen, 1988; Branney and Sparks, 1990; Gifkins et al., inpress) and felsic lavas (e.g. Pichler, 1981; Allen, 1988).

The terms fiamme and eutaxitic texture are used hereinto describe the rock texture and not to imply any particularorigin. Fiamme are flame-like, glassy or devitrified lenses,which define a pre-tectonic foliation (cf. McPhie et al., 1993).Fiamme may have a wide variety of sizes (0.5 mm to 1 m),length to height ratios (up to 40:1), shapes (e.g. flame-like,bow tie, irregular branching and blocky) and internal textures(aphyric, porphyritic, vesicular or stylolitic) (Gifkins et al., inpress). Eutaxitic texture is the pre-tectonic foliation definedby the parallel alignment of fiamme (cf. Fritsch and Reiss,1868; Ross and Smith, I960; Smith, I960). Eutaxitic texturetypically imparts a blotchy or streaky appearance to the rockdue to the colour contrast between the darker fiamme andpaler matrix (e.g. Fig. 3.9A and C).

3.2 I PSEUDOTEXTURES

The incomplete destruction of primary textures and thecombined effects of a number of different overprintingalteration styles (polyphase alteration) can result in significanttextural modification and the development of false textures orpseudotextures (De Rosen-Spence et al., 1980; Allen, 1988;McPhie et al., 1993). Pseudotextures are alteration texturesthat modify or obscure primary volcanic textures and oftenlead to incorrect interpretation of the primary volcanic facies.Allen (1988) described examples of altered silicic lavas andautobreccias from Benambra, New South Wales, that have theremarkably deceptive appearance of welded and non-weldedpyroclastic facies and thinly bedded tuffaceous rocks.

Pseudotextures can be subdivided into pseudoclastictextures (pseudobreccia, pseudogranular, false thin-beddedvolcaniclastic) or false pyroclastic textures (false shards,false pyroclastic or eutaxitic: Fig. 3.10). However, strongpervasive alteration can also produce false massive texturesthat resemble either massive volcaniclastic or coherent facies(Allen, 1988; McPhie et al., 1993; Doyle and Huston, 1999;

Doyle, 2001). Polyphase and patchy alteration of monomicticvolcaniclastic facies can also result in false clast-supported andfalse polymictic textures.

Pseudoclastic textures

The most common pseudoclastic textures are pseudobrecciaand false pyroclastic texture (also referred to as false eutaxitictexture). Other pseudoclastic textures include false thin-bedded volcaniclastic and pseudogranular textures.

Pseudobreccias have the appearance of breccias, but formas a result of alteration of coherent facies (Carozzi, 1960; Allen,1988). In outcrop they resemble coarse-grained, monomicticor polymictic, clast- to matrix-supported breccias comprisingangular to sub-rounded clasts in a fine-grained matrix (Fig.3.10A, BandC).

False pyroclastic textures occur in both coherent faciesand in situ hyaloclastite. In outcrop and hand specimen theymay have a eutaxitic texture and contain abundant fiamme(e.g. Fig. 3.10D). In thin section they appear to containsplintery and arcuate fragments, which may closely resemblepyroclastic glass shards (false shards: Fig. 3.10E).

Both pseudobreccia and false pyroclastic texture developas a result of two-phase alteration of fractured (perlitic orquench fractured) coherent or autoclastic facies and domain-controlled alteration of nodular devitrification textures incoherent facies (e.g. Fig. 3.11: Allen, 1988).

Networks of intersecting quench and/or perlitic fracturesmay control polyphase alteration in the fractured glassy partsof coherent lavas and intrusions because they are permeablepathways for fluid flow. Initially glass immediately adjacent tothe perlitic fractures is altered, then, as the fractures are filled,replacement fronts migrate away from the perlitic fracturestowards the core. This may either obscure the continuity ofthe perlitic fractures or, if alteration is incomplete, enhancethe perlitic fractures. False shard textures develop either dueto the preservation of less altered, relatively siliceous sliversbetween two or more fractures, or as altered segments ofthe fractures themselves (e.g. Fig. 3.10E: Allen, 1988). Theshape of false shards is a function of the shape of the fracturenetwork. For example, cuspate false shards are produced fromclassical perlite, whereas those resembling flattened or weldedshards result from banded perlite (e.g. Fig. 3.2D). False clastsdevelop where altered perlitic glass is partly overprinted bya subsequent alteration phase, thereby preserving isolatedrelicts of the earlier phase. Alternatively, the earlier phasemay be incomplete, leaving isolated kernels of glass that aresubsequently altered to a different mineral assemblage.

Pseudobreccia may also result from domainal or selectivealteration of nodular devitrification textures: spherulitesand lithophysae (e.g. Fig. 3.10F and G). Spherulites andlithophysae are typically recrystallised to quartzo-feldspathiccompositions, whereas the interstitial originally glassydomains are altered to phyllosilicate-rich assemblages (Allen,1988). Consequently, the originally glassy and crystallinedomains differ in alteration mineralogy and colour, and thespherulites appear as rounded siliceous clasts in a fine-grainedphyllosilicate-rich matrix.

Pseudogranular or sandy textures resemble well-sortedsandstones. These result from the recrystallisation of densely


A. Augen schistThis sample of quartz-augen schist comprises largelenticular quartz-rich domains (Q) enclosed in a stronglyfoliated, sericite + quartz ± chlorite-altered matrix. Augentexture is common in deformed, strongly porphyriticcoherent and crystal-bearing volcaniclastic rocks. Thisaugen schist probably resulted from the superposition ofa strong regional cleavage on an altered pumice breccia.The cleavage anastomoses around competent silicifiedpumice clasts.

Sample 040617, Cambrian western volcano-sedimentarysequences, Mount Read Volcanics, Rosebery hanging wall,western Tasmania.

B. Broken crystals in andesiteThis deformed andesite contains broken plagioclasephenocrysts in a strongly foliated, sericite + chlorite+ magnetite ± epidote-altered groundmass. Broken orfractured grains may result from mechanical pressureduring tectonic deformation (McBride, 1978). Typically,feldspar crystals have been fractured along theircleavage planes, whereas quartz crystals have developedconchoidal fractures (Taylor, 1950; Sippel, 1968). Crosspolarised light.

Sample 144387, Ordovician Lake Cowal Volcanic

Complex, Junee-Narromine Volcanic Belt, Lake Cowal,

Gateway Prospect, New South Wales.

C. Deformed grainsIn this amphibolite grade volcaniclastic siltstone, thequartz grains are deformed polycrystalline grains withundulose extinction, and elongated parallel to theregional cleavage. Cross polarised light.Sample GA9, Early Proterozoic Supra crustal succession,

Bergslagen mining district, Garpenberg, Sweden.

D. Deformed clasts and pillowsPillow fragments and clasts in this basaltic hyaloclastitewere deformed and stretched parallel to the regionalfoliation. The clast shapes are irregular and difficult torecognise as pillow or hyaloclastite fragments. Despitethis, many clasts preserve an internal zonation.Amphibolite, Proterozoic Corella Formation, Mary

Kathleen Group, Malbon River, northwest Queensland.

FIGURE 3.8 | Examples of deformation textures, deformed clasts and pre-existing textures in altered volcanic rocks.

56 | CHAPTER 3

E. Deformed clastsLens-shaped siliceous clasts (Q in this volcaniclasticbreccia have been rotated and stretched into the strongtectonic cleavage. The fine-grained matrix has beenfoliated and chlorite + sericite + quartz altered.Sample 133520, Cambro-Ordovician Trooper CreekFormation, Seventy Mile Range Group, Mount WindsorSubprovince, central Thalanga, Queensland.

F. Folded pumice clastThis sample of rhyolite-, pumice- and crystal-rich brecciacontains a folded tube pumice clast with an axial planarcleavage (Sj) defined by aligned sericite. The pumiceclast has been albite + quartz + sericite altered. Planepolarised light.Sample KB304B, Siluro-Devonian Ural Volcanics, UralRidges area, New South Wales.

G. Deformed relict perliteRelict perlitic fractures in this jigsaw-fit andesiticbreccia are elongate and flattened, especially adjacentto competent phenocrysts. The groundmass has beensericite + chlorite + calcite + albite altered and the perliticfractures chlorite filled. Plane polarised light.Sample 76902, Cambrian Que-Hellyer Volcanics, westernvolcano-sedimentary sequences, Mount Read Volcanics,Hellyer, western Tasmania.

H. Deformed grainsIn this sample of volcaniclastic sandstone, strongly'deformed feldspar grains and clasts have been rotatedparallel to the strong cleavage. Plane polarised light.Sample 133520, Cambro-Ordovician Trooper CreekFormation, Seventy Mile Range Group, Mount WindsorSubprovince, central Thalanga, Queensland.

FIGURE 3.8 | Examples of deformation textures, deformed clasts and pre-existing textures in altered volcanic rocks, cont.


A. Fiamme and eutaxitic texture in weldedignimbriteDark flame-like obsidian lenses or fiamme (F) are alignedin this sample of subaerial welded rhyolitic ignimbrite.Fiamme are commonly interpreted as flattened pumiceclasts. The fiamme in this sample are interpreted to resultfrom the plastic deformation, flattening and sinteringtogether of hot glassy pumice clasts during welding(cf. Smith, I960). The bedding-parallel alignment offlattened, elongate fiamme and glass shards defines theeutaxitic texture.Sample OW7, Pleistocene Owahoroa ignimbrite, WhitiangaGroup, Coromandel Volcanic Zone, Owharoa Falls, NewZealand.

B. Fiamme in thin sectionIn thin section, the former pumice clasts, fiamme (F),lack uncompacted vesicles, have feathery terminationsand are enclosed in domains of cuspate and platy shards(5), and quartz, feldspar and biotite crystal fragments.Although some shards have preserved bubble-wallshapes, others were plastically deformed and compacted,especially adjacent to crystals. Plane polarised light.Sample OW11, Pleistocene Owahoroa ignimbrite,Whitianga Group, Coromandel Volcanic Zone, OwharoaFalls, New Zealand.

C. Fiamme and eutaxitic texture in non-weldedpumice brecciaDark, plagioclase-phyric, wispy, chlorite-rich fiamme areenclosed in pale domains of quartz + chlorite + pyrite-altered pumice clasts, shards and crystal fragments inthis non-welded rhyolitic pumice breccia. The bedding-parallel alignment of fiamme defines the eutaxitic texture.Alteration and compaction of pumice clasts duringdiagenesis formed these apparent welding textures.Sample 133809, Cambrian Hercules Pumice Formation,Central Volcanic Complex, Mount Read Volcanics, Herculesfootwall, western Fasmania.

D. Fiamme in thin sectionIn thin section, the chlorite fiamme (F) have featheryterminations and lack internal textures other than sparseplagioclase phenocrysts and hematite-rich stylolites. Thepale quartz + chlorite + pyrite-altered domains containuncompacted tube pumice clasts (P). Plagioclase crystalsare dusted with hematite and sericite. Plane polarisedlight.Sample 133811, Cambrian Hercules Pumice Formation,Central Volcanic Complex, Mount Read Volcanics, Herculesfootwall, western Fasmania.

FIGURE 3.9 | Examples of fiamme and eutaxitic texture.

58 I CHAPTER 3

A. Pseudobreccia in perlitic rhyoliteSericite + chlorite-altered perlitic fractures and pinkalbite-altered perlitic cores result in the pseudoclastictexture in this plagioclase-phyric coherent rhyolite.Sample BBP248-504.7 m from the Cambrian CentralVolcanic Complex, Mount Read Voleanics, Boco, westernTasmania.

B. Pseudobreccia in macroperlitic dacitePolyphase alteration of macroperlite in this plagioclase-phyric dacite has resulted in dark chlorite + epidote-richdomains along and adjacent to the perlitic fractures.This overprinted and enclosed earlier, pale grey albite+ sericite-altered perlitic cores. The chlorite + epidote-rich domains resemble the matrix in a matrix-supportedbreccia. Locally, the arcuate perlitic fractures are welldefined.Cambrian Sterling Valley Volcanics, Mount Read Volcanics,

Sterling Saddle, western Tasmania.

C. False clastsIn this coherent andesite, plagioclase phenocrysts havebeen extensively replaced by epidote, pyroxenes bychlorite and the groundmass domains by green epidote+ chlorite and orange albite. The domainal distributionof the alteration facies gives the andesite a patchypseudoclastic texture. The false clasts have both sharpand diffuse margins, which are transgressed locally byaltered plagioclase phenocrysts.Sample 145147, Ordovician Forest Reefs Volcanics, Molong

Volcanic Belt, Cooramilla, New South Wales.

D. False pyroclastic textureThe most conspicuous feature of this sample is the wispydark green chlorite-rich lenses that resemble fiamme.However, these lenses are aligned in the tectonic cleavageand occur in an evenly porphyritic rhyolite. The lensesare interpreted to result from domainal chlorite andsericite + quartz + biotite alteration of the groundmassin a coherent quartz + plagioclase-phyric rhyolite.Sample 140727, Cambro-Ordovician Mount WindsorFormation, MountWindsorSubprovince, centralThalanga,Queensland.

E. False shards in thin sectionPolyphase chlorite + biotite and K-feldspar alterationof perlitic fractures, and their subsequent deformationhave resulted in irregular arcuate and platy shapes, whichresemble shards (arrows) in this quartz + plagioclase+ pyroxene-phyric coherent rhyolite. Plane polarisedlight.Sample KB499, Siluro-Devonian Coan rhyolite, MountHope Volcanics, Mount Hope area, New South Wales.

FfGURE 3.10 | Examples of pseudotextures in altered volcanic rocks.


F. Pseudoclastic texture in devitrified rhyoliteThis rhyolite contains silicified nodules that are composedof coalesced spherulites in a fine-grained sericite-alteredgroundmass. These nodular devitrification textures givea clastic appearance to the hand specimen and outcrop.Late Devonian Bunga Beds, Boyd Volcanic Complex,Bengunnu Point, New South Wales.

G. False clasts in thin sectionThe clastic texture in this rhyolite comes from theuneven distribution of strongly chlorite + hematite-altered spherulites in a strongly foliated, chlorite +sericite + feldspar + hematite-altered groundmass. Thefoliation has wrapped around the altered spherulites,which have preserved fibrous textures and quartz cores.Plane polarised light.Sample KB536D, Siluro-Devonian Mount Hope Volcanics,Boolahbone tank, Mount Hope area, New South Wales.

I. Pseudogranular texture in thin sectionRecrystallised micropoikilitic textures in the groundmassof this aphyric rhyolite resemble sand-sized, roundedgrains in a well-sorted quartzo-feldspathic sandstone.The margins of the micropoikilitic domains are markedby concentrations of sericite, which enhance the granulartexture. Plane polarised light.Sample 147448, Cambrian Kershaw Pumice Formation,Central Volcanics Complex, Mount Read Volcanics,

western Tasmania.

H. Pseudogranular textureThis altered dacite has a fine granular texture in handspecimen and resembles massive sandstone. However, inthin section it has a densely microspherulitic groundmassin which the recrystallised quartz + albite spherulitesare separated by cuspate sericite-rich domains. Thefine-grained, densely packed spherulites give the handspecimen its sandy texture.Sample MR25, Cambrian Kershaw Pumice Formation,Central Volcanics Complex, Mount Read Volcanics, MountRead summit, western Tasmania.

FIGURE 3.10 | Examples of pseudotextures in altered volcanic rocks, cont.

60 | CHAPTER 3

J. False thin-bedded volcaniclastic textureFlow banding in this microspherulitic plagioclase-phyricdacite is defined by alternating pale albite + quartz anddarker albite + sericite + quartz layers. The planar,repetitive, thin flow banding resembles thin bedding inclastic facies such as tuffaceous siltstones.Sample 76772, Cambrian Que-Hellyer Volcanics, westernvolcano-sedimentary sequences, Mount Read Volcanics,Hellyer, western Tasmania.

K. False thin-bedded volcaniclastic texture in thinsectionIn thin section, this finely flow-banded rhyoliteresembles a thin-bedded volcaniclastic facies withfractured plagioclase crystals. However, axiolitic andbow-tie shaped spherulitic textures are locally preservedin the groundmass. Cross polarised light.Sample KB132A, Siluro-Devonian Ural Volcanics, Ural

• area, New South Wales.

L. False volcaniclastic texture in thin sectionFractured and broken plagioclase crystals, and re-crystallised spherulites in the groundmass of this flow-banded rhyolite contribute to its pseudoclastic texture.Cross polarised light.Sample 133837 Cambrian Mount Black Formation,Central Volcanic Complex, Mount Read Volcanics, MountRead summit, western Tasmania.

M. False polymictic, matrix-supported textureThe dark chlorite-rimmed clasts in this plagioclase-phyric basaltic breccia appear to be supported in acompositionally different, pale calcite + chlorite-alteredmatrix. However, the matrix comprises jigsaw-fit, blockyand splintery clasts of perlitic basalt that are identicalto the darker clasts. The chlorite-rimmed clasts appearsubrounded, because chlorite alteration of the clastmargins and adjacent matrix has obscured the blockyand splintery shapes.Sample 76833, Cambrian Que-Hellyer Volcanics, westernvolcano-sedimentary sequences, Mount Read Volcanics,Hellyer, western Tasmania.



N. False polymictic textureOverprinting domainal albite + hematite and epidotealteration facies in this plagiocla.se + hornblende-phyricdacite gives the sample a heterogenous appearance. Theabundance of pink, albitised plagioclase phenocrysts in thered and green domains is equivalent, although they appearmore abundant in the epidote-altered domains due to thecolour contrast between the phenocrysts and epidote-altered groundmass. Colour differences between the twoalteration facies and more prominent phenocrysts in theepidote-altered domains obscure the massive, coherenttexture and uniform composition of this sample.Sample 147557, Cambrian Mount Black Formation,Central Volcanic Complex, Mount Read Volcanics, PiemanRoad, western Tasmania.

O. False matrix-supported textureClasts in this monomictic dacite breccia displayjigsaw-fit texture. However, feldspar + quartz + sericitealteration facies has replaced the groundmass adjacentto the quench fractures between clasts, enhancing thematrix, and imparting an apparent matrix-supportedtexture. The feldspar + quartz + sericite-altered matrixhas been more resistant to weathering than the sericite +chlorite-altered clasts and forms ridges on the outcrop.The larger clasts are perlitic, plagioclase + hornblendeporphyritic with planar and curviplanar margins typicalof clasts produced by quench fragmentation.Cambrian, Mount Black Formation, CentralVolcanic Complex,

Mount Read Volcanics, Tullah lakeside, western Tasmania.

P. False matrix-supported textureIn this in situ andesitic hyaloclastite, blocky chlorite-altered plagioclase + pyroxene-phyric clasts appear tobe supported in a pyrite + quartz + sericite-rich matrix.However, thin section inspection reveals relict clastswith jigsaw-fit textures preserved in the false matrixdomains.Sample 144710, Ordovician Lake Cowal Volcanics, Junee-Narromine Volcanic Belt, Boundary Prospect, New SouthWales.

Q. False coherent textureThis albite + quartz + sericite- and chlorite + epidote-altered pumice and rhyolite breccia resembles a coherent,feldspar porphyritic facies. The albite-altered plagioclasecrystals, pseudophenocrysts, are evenly distributed ina fine-grained, sericite-rich false groundmass. On theleft side of the photograph, pervasive, massive albite +quartz alteration facies obscures the plagioclase crystalsin a pseudo-aphyric texture.Sample 147402, Cambrian Kershaw Pumice Formation,CentralVolcanic Complex, Mount Read Volcanics, Rosebery,western Tasmania.


62 | CHAPTER 3

FIGURE 3.11 | Sketches summarising the relationship between primary volcanic, high-temperature devitrification and hydration textures, and pseudotextures.(A) Pseudotextures in classical perlite (after Allen, 1988). False shards may be produced either (i) as siliceous segments between phyllosilicate-altered perliticfractures or (ii) phyllosilicate-altered sections of the perlitic fractures, (iv) Alternatively, pervasive phyllosilicate alteration can obscure the perlitic fractures resulting in amassive texture.

(B) Pseudotextures in flow-banded facies comprising alternating crystalline (spherulitic) and glassy flow bands (modified from Doyle, 2001). (i) Glassy domains arealtered to phyllosilicate-rich assemblages and spherulitic bands to quartzo-feidspathic assemblages. Consequently the originally glassy and crystalline domainsdiffer in alteration colour and mineralogy. The spherulites superficially appear as rounded siliceous dasts in a fine-grained phyllosilicate-rich matrix, resulting in apseudobreccia texture, (v) In parallel flow-banded facies this can result in a false thin-bedded voicaniciastic texture. The crystalline nature of the quartzo-feldspathic-altered bands enhances the granular appearance of the false beds.

(C) Pseudotextures from in situ quench fragmented porphyritic glass (modified from Doyle, 2001). Alteration progresses as fronts away from the fractures and intothe non-fractured glass, (i) Where phyllosilicate alteration is incomplete the remaining kernels of glass are subsequently altered to quartzo-feidspathic assemblages.The further the alteration progresses away from the fractures, the more matrix-supported the resulting pseudobreccia texture, (ii) Pseudoclastic texture also developsas a result of the complete replacement by the first alteration phase (phyllosilicate alteration) and only partial replacement by the second alteration phase (quartzo-feidspathic alteration), thereby preserving isolated relics of the earlier phase, (iii) False polymictic texture develops in quench fragmented porphyritic glass aspolyphase alteration results in colour variations between the false dasts and matrix. Varying intensities of alteration enhance the polymictic appearance. Phenocrystsare more prominent in dark phyllosilicate domains than in paler siliceous domains, resulting in an apparent variation in crystal content consistent with a polymicticrock, (iv) Pseudomassive texture develops from pervasive phyllosilicate alteration.

(D) Pseudotextures from porphyritic autobreccia or hyaloclastite (modified from Doyle, 2001). (i) and (ii) alteration commences along fractures and in the matrix ofautobreccia but gradually progresses into the dasts, greatly enhancing the clastic, matrix-supported appearance. The result is commonly false matrix-supportedtexture, (iii) Polyphase alteration results in pseudoclasts and pseudomatrix that comprise different alteration mineral assemblages and colours, and therefore appearto have different primary compositions (false polymictic texture). Phenocrysts are more prominent in dark phyllosilicate-altered domains, (iv) Pseudomassive textureoccurs where alteration has been extensive and pervasive.

(E) Pseudotextures from porphyritic pumice breccia (modified after Gifkins, 2001). False eutaxitic texture results from phyllosilicate alteration of originally glassypumice dasts or shards within the pumice breccia. The phyllosilicate-altered pumice dasts and shards are flattened by compaction (i) or tectonic deformation

(ii) resulting in a texture that resembles eutaxitic texture in welded ignimbrites. (iii) False polymictic texture as a result of colour variations and apparent variationsin phenocryst content due to polyphase alteration, (iv) False coherent textures result from the complete and pervasive phyllosilicate alteration and subsequentcompaction of originally glassy shards and pumice dasts producing a massive textured rock in which the original dasts are indistinguishable.

spherulitic or micropoikilitic groundmass of coherent felsicfades (e.g. Fig. 3.10HandI).

False thin-bedded volcaniclastic texture resembles thinlybedded and tectonically folded volcaniclastic rocks, and canoccur in altered, planar flow-banded and flow-folded lavasand intrusions (e.g. Fig. 3.10J and K). The false thin-beddedvolcaniclastic texture is due to the planar and uniform characterof flow banding and the apparent textural and compositionaldifferences between flow bands (Allen, 1988). This apparentcompositional difference between flow bands results fromdomainal alteration: originally glassy bands were altered todark phyllosilicate assemblages and contain well-preservedphenocrysts, whereas the originally cryptocrystalline bandswere altered to pale quartzo-feldspathic assemblages andcontain altered polycrystalline phenocrysts that are hardlyrecognisable. Phenocrysts may also be extensively fracturedand dismembered during deformation, thereby contributingto the pseudoclastic and finer grained appearance (e.g. Fig.3.10L). Discrepancies in colour and the apparent relativeproportions of phenocrysts enhance the compositionalcontrast between adjacent bands.

False polymictic texture

Domainal phyllosilicate alteration may impart a heterogenousappearance due to colour variation and to variable preservationof phenocrysts. Pseudobreccias and pseudoclastic faciesmay appear polymictic due to patchy or mottled alterationdistribution resulting in colour variations and variablepreservation of phenocrysts. This is particularly common inpseudobreccias because the pseudoclasts and pseudomatrixcontain different alteration minerals, and hence colours:therefore they appear to have different primary compositions(e.g. Fig. 3.10M and N). In addition, phenocrysts (especiallyfeldspar or quartz) are more prominent in dark phyllosilicatedomains than in paler siliceous domains, resulting in anapparent variation in crystal content between differentalteration domains consistent with different compositions(e.g. Fig. 3.10N).

False matrix-supported texture

Alteration along fractures in coherent facies and in the matrixof autoclastic facies (hyaloclastite and autobreccia) results incolour variation between the clasts and fracture selvages ormatrix, and greatly enhances the clastic appearance. Wherealteration has progressed into the clasts from the clast margins,the clasts appear to be isolated within a matrix (e.g. Fig. 3.10Oand P). The result is an apparently matrix-supported brecciathat is difficult to recognise as hyaloclastite or autobreccia(Allen, 1988).

False coherent textures

False coherent textures, such as pseudomassive texture,are less common than pseudoclastic textures. They resultfrom the complete and pervasive alteration of an originallyvolcaniclastic facies. For example, pervasive phyllosilicate


alteration of pumice breccia produces a massive texturedrock in which the original clasts are indistinguishable (Fig.3.10Q).

3.3 I ALTERATION DISTRIBUTION

Alteration distribution refers to the mappable extent of thealteration facies and its relationship to host rocks, structures,mineralised rock and other alteration mineral assemblages or

zones.Although alteration distribution, zonation and textural

relationships are easily observed in thin section and handspecimen they are more difficult to assess on a macroscopicscale. This is due to the typically sparse exposure in ancientvolcanic successions, structural complications, commonlypatchy mode of occurrence of alteration facies in volcanicrocks, and the considerable amount of detailed work requiredto determine the distribution and zonation. Closely spaceddrill holes may adequately delineate the alteration faciesdistribution and zonation at prospect scales. However,alteration zones may be superimposed and the original zonationpatterns modified. It is important to recognise disequilibriumalteration mineral assemblages and overprinting textureswhen determining zonation patterns, so as to account for anysuperposition caused by subsequent alteration styles.

Is the alteration facies limited in extent or widespread?

Typically the macroscopic alteration distribution is describedas either regional or local in extent. Regional alteration stylesare widespread, affecting extensive (hundreds of metres totens of kilometres) expanses of rock. Local alteration stylesare limited in extent and can refer to alteration on a scaleof centimetres, such as wall-rock alteration associated witha fault or vein, to hundreds of metres, such as the extensivehydrothermal alteration associated with the Rosebery VHMSdeposit (western Tasmania) that extends up to 100 mstratigraphically beneath the ore lenses, 10-20 m into thehanging wall and 500 m along strike (Large et al., 2001b).

Is there a relationship between the distribution of the

alteration facies and stratigraphy?

Alteration facies can be distributed within individual volcanicunits or formations or cut across stratigraphic contacts.Alteration facies that are confined to individual units orformations and are either concordant or discordant within theunit are described as stratabound or semiconformable (Guilbertand Park, 1986; Large, 1992). In contrast, alteration pipes andalteration plumes refer to the distribution pattern or shapeof alteration systems or facies that transgress stratigraphiccontacts (pipes, e.g. Sangster, 1972; Gemmell and Large,1992; Doyle and Huston, 1999, and plumes, e.g. Jack, 1989;Gemmell and Fulton, 2001).

64 | CHAPTER 3

Is the alteration fades associated with other alteration

fades, structures, mineralised rock or particular units?

Locally distributed alteration fades are commonly associatedwith structures, such as veins or fractures, ore bodies orparticular volcanic units: forming halos, envelopes orselvages. They may also occur around isolated clasts withinclastic facies. In zoned alteration systems or halos, they arecommonly spatially and temporally associated with otheralteration facies.

3.4 I ALTERATION ZONATION PATTERNS

Alteration zonation is a regular pattern in the spatial distributionof mineral species, mineral assemblages, major or traceelements or textures, and reflects mineralogical and chemicalchanges that relate to fluid-rock ratios and temperaturegradients (e.g. Meyer and Hemley, 1967). In general, the haloof altered rock is divided mesoscopically into altered zones.These are imprecisely defined as the areal extent of differentalteration mineral assemblages or facies. Boundaries betweenhydrothermally altered zones associated with VHMS depositsare commonly sharp, and reflect changes in the abundance ofparticular alteration mineral species such as chlorite, sericiteor quartz. However, in some cases alteration facies overlap,resulting in diffuse boundaries between altered zones (e.g.Titley, 1982; Doyle and Huston, 1999).

In reality, altered zones are identified and named usingindex minerals - the dominant alteration minerals - and zoneboundaries - isograds - are drawn at the first appearance ofan index mineral characteristic of the next zone. Thus theclinoptilolite + mordenite zone is placed at the first appearanceof clinoptilolite or mordenite (e.g. Iijima, 1978). In some casesan altered zone may contain more than one alteration facies.For example, at Thalanga the tremolite + chlorite ± carbonatezone includes intense pervasive stratabound chlorite + tremolite(no carbonate), intense pervasive stratabound chlorite +calcite, and intense stratabound tremolite + dolomite + calcite(no chlorite) facies. The expression of zoning may be limitedby exposure and modified by structural and compositionalhomogeneity, faulting and/or intrusions.

There are three scales of zonation that are generally appliedto the relative distribution of metals (e.g. Kutina et al., 1965)

TABLE 3.3 I Scales of alteration zonation.

Regional zonation

District zonation

Local zonation

Sample-scalezonation

Alteration facies or minerals occur in zonesthroughout a region

Altered zones are associated with a clusterof ore bodies, fractures or intrusions

Altered zones are associated with anindividual ore body, fracture or intrusion

Altered zones are associated with small-scale features such as clasts or minerals inthe rock

and that can also be applied to the distribution of alterationfacies or minerals (Table 3.3): (1) regional zoning, (2) districtzoning, and (3) local or ore body zoning. In addition, alteredzones may occur within or surrounding individual clastsin clastic facies. Thus the dimensions of altered zones mayvary from a few centimetres to several tens of kilometres(cf. Bohlke et al., 1980; Galley, 1993). These variations indimension are a function of the size of the alteration system(regional versus local systems), and changes in physical andchemical conditions such as the porosity and composition ofthe host rock, fluid-rock ratio and composition of the fluid(Rose and Burt, 1979).

Different alteration processes result in different zonationpatterns. Regional alteration processes such as diagenesisand metamorphism, produce thick, flat-lying, verticallyzoned systems. Local hydro thermal systems and contactmetamorphism result in footwall and hanging wall alteredzones or concentrically zoned altered halos or contact aureoles.Zonation related to faults or fractures is generally parallel tothe structure, commonly cross cutting stratigraphic contacts.

Regional diagenetic zones

Diagenesis develops in response to increasing temperaturewith depth during burial; as a result it forms a sequenceof flat-lying altered zones (Fig. 3.12: Iijima, 1974, 1978;Fisher and Schmincke, 1984; Utada, 1991). These zonesare characterised by mineral assemblages, which reflect thereaction of glass, primary minerals and diagenetic mineralswith interstitial pore water and seawater at a particulartemperature — anywhere between 0° and 250°C (Hay, 1978;Iijima, 1978; Utada, 1991).

Sequences of diagenetic zones are typically between 500 mand 6 km thick (Hay, 1978) and individual altered zones varyfrom a few metres to several kilometres in thickness (e.g. Hay,1978; Iijima, 1978; Vierecketal., 1982; Passagliaetal., 1995).The thickness of diagenetically altered zones is dependent onthe geothermal gradient, rate of burial, and the porosity andpermeability of the volcanic succession.

The Miocene Hokuroku Basin, part of the Green TuffBelt in northern Honshu, Japan, is an excellent example ofdiagenetic zonation (Fig. 3.12). The Hokuroku Basin containsa 3 to 6 km thick submarine volcanic succession dominatedby rhyolitic to dacitic and minor basaltic coherent and clasticvolcanic facies (Horikoshi, 1969; Iijima, 1978; Tanimura etal., 1983; Urabe, 1987; Utada, 1991). Diagenetic alterationin the Hokuroku Basin has produced a series of flat-lyingdiagenetic zones, which grade vertically from fresh glass atthe top, to smectites, zeolites and albite at depth (Fig. 3.12:Iijima, 1978).

Regional metamorphic zones

Regional metamorphic zones develop due to regionalchanges in temperature and pressure. Unless metamorphismis related solely to orogenic deformation, progressive burialresults in dehydration and increasing metamorphic gradewith depth. This produces a vertical sequence of flat-lying,regional metamorphic zones (Fig. 3.13B). It is generally

FIGURE 3.12 | Schematic cross-section of vertically developed diagenetic

zones in the central Hokuroku Basin, Japan (after lijima, 1974; Hay, 1978; lijima,

1978).


assumed that higher-grade metamorphic rocks formerlyhad mineral assemblages typical of lower-grade zones andwere altered progressively as metamorphism proceeded.However, variations in geothermal gradient, variable ratesof sedimentation and folding and faulting can modify theprogression of metamorphism and the resulting metamorphiczones, in particular, folding and faulting may affect theposition of mineralogically determined isograds by promotingreactions or causing local increases in temperature or heat flow(Coombs et al., 1959).

Coombs (1954) first recognised vertical metamorphiczoning of low-temperature, high-pressure metamorphicmineral assemblages in a 10 km section of the upliftedPermian to Jurassic Wakatipu Metamorphic Belt, in theNew Zealand geosyncline (Fig. 3.13A). The WakatipuMetamorphic Belt comprises submarine emplaced rhyoliticto basaltic volcaniclastic and volcanogenic sedimentary rocks,which were progressively metamorphosed (Coombs, 1954).There is a continuous gradational sequence of regionalmetamorphic zones characterised by zeolite, prehnite +pumpellyite, pumpellyite + actinolite and greenschist facies (abiotite zone and chlorite zone) at depth (Fig. 3.15B: Coombs,1954; Coombs etal., 1959; Landis and Coombs, 1967). Thesezones have an even thickness over a wide area (>10 km by300 km) suggesting that the geothermal gradient, estimatedat l4-25°C/km, was consistent throughout the WakatipuMetamorphic Belt (Landis and Coombs, 1967).

FIGURE 3.13 | Alteration zonation patterns associated with regionalmetamorphism. (A) Regional metamorphic zones in the Wakatipu metamorphicbelt, South Island, New Zealand, (after Coombs et al.,1959, and Landisand Coombs,1967). (B) Schematic cross-section reconstruction of regionalmetamorphic zones and isotherms during peak metamorphism of the Wakatipumetamorphic belt, New Zealand (after Landis and Coombs, 1967).

66 | CHAPTER 3

Regional, deep, semi-conformable altered zones

Deep, semi-conformable altered zones are interpreted to bethe products of hydrothermai alteration within regional,subseafloor hydrothermai systems (Galley, 1993). Theseregional hydrothermai systems involve the large-scaleconvection of modified seawater driven by the emplacementof synvolcanic intrusions into the subsurface (Spooner andFyfe, 1973; Norton, 1984; Cathles, 1993; Galley, 1993).The upper contacts of sub-surface intrusions are typicallysub-parallel to the volcanic strata and hence the overlyingisotherms are also semi-conformable (Galley, 1993). Theresult is a series of vertically stacked, sub-horizontal alteredzones, which are characterised by mineral assemblages thatreflect reactions between host volcanic facies and modifiedseawater at temperatures transitional with diagenesis andgreenschist facies metamorphism (Galley, 1993; Skirrowand Franklin, 1994). Sequences of deep, semi-conformablealtered zones may be up to 20 km wide and between 1 and2 km thick (Gibson et al., 1983; Galley, 1993; Skirrow andFranklin, 1994).

One of the first comprehensive descriptions of deep, semi-conformable altered zones is by Gibson et al. (1983). Theseauthors documented a series of vertically stacked altered zonesin the Noranda district of the Archaean Abitibi belt, Canada

(Fig. 3.14). The Amulet Rhyolite Formation comprisescoherent and clastic basalts and andesites with minorrhyolitic lava domes. The volcanic rocks are variably alteredand the distribution of the alteration facies can be relatedto stratigraphic depth and primary host rock composition.Within the andesites there are four regionally extensivealteration facies, from top to bottom: chlorite + actinolite +albite + epidote + quartz; silicified; mottled epidote + quartz;and chlorite alteration facies (Fig. 3.14B). The upper chlorite+ actinolite + albite + epidote + quartz alteration facies isthe result of low-grade regional metamorphism, whereas theother three alteration facies are interpreted to be the productsof regional deep, semi-conformable alteration (Gibson et al.,1983).

Local contact metamorphic or hydrothermallyaltered halos

Halos or contact aureoles associated with granitoids, thicksynvolcanic sills, cryptodomes and dykes reflect changesin the composition and temperature of the magmatic orhydrothermai fluid away from the intrusion, and its interactionwith local pore water (Rose and Burt, 1979; Einsele et al.,1980; Yardley, 1989). The result is a progressive sequenceof roughly concentric altered zones or metamorphic zones

FIGURE 3.14 | Diagram showing alteration zonation patterns associated with regional hydrothermai alteration.(A) A schematic cross-section through a vertically stacked sequence of deep, semi-conformable altered zones in the AmuletRhyolite Formation, Noranda district, Abitibi belt, Canada (after Gibson, 1989, in Galley, 1993). (B) A reconstructed sectionof deep, semi-conformable altered zones in the Amulet Rhyolite Formation at Turcotte Lake (after Gibson et al., 1983).


that correspond with decreasing temperature away from theintrusion contact (e.g. Fig. 3.15). For example, submarinebasaltic lavas and breccias of the Upper Triassic KarmutsenSubgroup, Vancouver Island, Canada, have undergone low-pressure, high-temperature contact metamorphism related tothe shallow emplacement of the Jurassic Coast Range Batholith(Carson, 1973; Kuniyoshi and Liou, 1976). Intrusion of theCoast Range Batholith resulted in two locally developedcontact metamorphic zones: hornblende + plagioclase zoneand epidote + actinolite zone, which are superimposed ona prehnite + pumpellyite facies regional metamorphic zone(Fig. 3.15A: Kuniyoshi and Liou, 1976). The hornblende+ plagioclase zone is approximately 2600 m wide and theepidote + actinolite zone less than 900 m wide (Kuniyoshiand Liou, 1976).

Local hydrothermally altered halos around oredeposits

Local halos of hydrothermally altered rock aroundmineral deposits result from the interaction between thehydrothermal fluid/s responsible for mineralisation and thesurrounding country rock. In the case of VHMS deposits,hydrothermal fluid temperatures range up to 35O°C (Large,1977, 1992; Urabe et al., 1983; Gemmell and Large, 1992).The migration of hydrothermal fluids and mixing with coldseawater at the margins of the hydrothermal system producesthree-dimensional altered zones composed of successivelylower temperature mineral assemblages away from the site ofmineralisation (Date et al., 1983; Urabe et al., 1983; Large,1992; Schardt et al., 2001). The shape, dimensions anddistribution of fluid circulation (and thus the altered zones)are usually closely related to initial patterns of permeability,porosity and compositional contrasts in the volcanic succession(e.g. Yamagishi and Dimroth, 1985; Large, 1992; Doyle,2001). Accordingly, the altered halos exhibit a wide varietyof shapes and zonation around VHMS deposits, typicallywith intense proximal footwall and weak hanging wall zones(Large, 1992). In contrast, altered halos around porphyrydeposits may have a more concentric zonation, comprisinga potassic core zone, an outer propylitic zone and, in someexamples, minor phyllic zones (Gustafson and Hunt, 1975).Later hydrothermal alteration and mineralisation zones mayoverprint these initial zones, such as the widespread phyllicand argillic zones at the El Salvador deposit, Chile and BatuHijau in Indonesia (Gustafson and Hunt, 1975; McMillanand Panteleyev, 1998; Garwin, 2002).

Hydrothermally altered zones associated with VHMSdeposits can extend laterally up to 500 m and persist at least500 m stratigraphically (Iijima, 1974; Utada et al., 1974;Eastoe et al., 1987; Gemmell and Large, 1992; Gemmelland Fulton, 2001). Footwall zones occur either as diffusestratiform zones (e.g. Rosebery, Fig. 3.16C: Green et al., 1981;Large et al., 2001b) or well-defined and zoned alterationpipes (e.g. Hellyer, Fig. 3.16A: Millenbach, Fig. 3.16D:Franklin et al., 1981; Morton and Franklin, 1987; McArthur,1989; Gemmell and Large, 1992). Hanging wall alteration isnormally of lower intensity, and developed above the thickestpart of the ore body either as a diffuse stratabound zone (e.g.Woodlawn and Scuddles) or an alteration plume (e.g. Hellyer,

FIGURE 3.15 | Diagram showing alteration zonation patterns associated withcontact metamorphism. (A) Contact metamorphic zones in the northeasternVancouver Island, Canada (from Kuniyoshi and Liou, 1976). (B) Schematiccross-section of contact-metamorphic zonation.

Fig. 3.16A: Large and Both, 1980; Jack, 1989; Large, 1992;Gemmell and Fulton, 2001).

In the Hokuroku district, local hydrothermally alteredhalos associated with the Kuroko deposits comprise up tofour altered zones in a roughly concentric distribution aroundthe ore deposits (Fig. 3.16B: Urabe et al., 1983; Utada et al.,1988; Utada, 1991). From core to margin the altered zonesare: sericite + quartz ± pyrite zone; chlorite + sericite + quartzzone; mixed-layer clay (illite-montmorillonite) zone; kaoliniteor smectite-montmorillonite zone (Date et al., 1983; Urabeet al., 1983). These altered zones vary in thickness fromapproximately 16 m to 100 m, transgress stratigraphicboundaries, and grade into and interfinger with regionaldiagenetic zones such as the clinoptilolite + mordenite zonein Figure 3.16B (Utada, 1970, 1991; Iijima, 1974).

In contrast, the mineral assemblages in altered zonesassociated with Australian VHMS deposits reflect higher-grade regional metamorphic overprints. Some have proximalchlorite zones and halos of quartz + sericite, sericite andcarbonate zones (Fig. 3.16A and C: Allen, 1988; Large, 1992;Doyle, 2001).

Vein and fracture altered halos

Altered halos around fractures or veins are typically limitedin width (from millimetres to tens of metres) occurring

68 | CHAPTER 3

FIGURE 3.16 | Schematic cross-sections showing hydrothermal alteration zonationpatterns associated with VHMS deposits. (A) Hydrothermally altered zones developed inthe footwall alteration pipe and hanging wall alteration plume to the Hellyer VHMS deposit(after Gemmell and Fulton, 2001). (B) Hydrothermally altered halo developed aroundthe Uwamuki group deposits, Kosaka VHMS mine, Japan (after Urabe et al., 1983).(C) Hydrothermally altered halo developed around the K-lens VHMS ore lens at Rosebery(after Large et al., 2001b). (D) Hydrothermally altered zones developed in the footwallalteration pipe and hanging wall alteration plume to an ore lens in the Millenbach VHMSdeposit, Canada (after Riverin and Hodgson, 1980).

either as narrow selvages around individual veins or as widepervasive fracture-controlled altered zones. Decreasing fluidtemperatures and changes in fluid chemistry (especially pH)away from fractures or faults results in alteration zonationparallel to the fracture or vein surface.

3.5 | OVERPRINTING RELATIONSHIPSAND TIMING OF ALTERATION

The aim in determining overprinting relationships is toestablish a time sequence for mineral growth, which is referredto as a paragenetic sequence (Kutina et al., 1965). Normallyparagenesis refers to the growth of minerals from oldest to


youngest. The paragenetic sequence is typically depictedusing a horizontal bar chart (e.g. Fig. 3.17) or a space-timediagram, such the schematic diagram by Wilson et al. (2003)that shows alteration and vein stages relative to the intrusivehistory of the Ridgeway Complex at the Ridgeway Au-Cuporphyry deposit in New South Wales.

Unfortunately, interpreting timing relationships betweendifferent alteration facies can be difficult and confusingas there are few unambiguous overprinting textures. Thesuperposition of many different alteration phases on thesame rock can obscure or complicate alteration textures. Inaddition, while one mineral is being deposited under certainconditions at one place, other minerals are forming elsewhereunder different conditions. Thus the deposition of onemineral may overlap with another in both time and space.

FIGURE 3.17 | Interpretation of the relative timing of alteration mineral assemblages in the northern Central Volcanic Complex, western

Tasmania (modified after Gifkins and Allen, 2001). S1 is the bedding-parallel, stylolitic foliation interpreted as a diagenetic compaction

and dissolution fabric (Allen, 1990). S2 is the regional tectonic cleavage related to Devonian folding. Assemblages in brackets refer to

pre-metamorphic equivalents. For example, early regional sericite is probably a metamorphosed equivalent of early clays. The feldspar

+ quartz + sericite alteration mineral assemblage may include the growth of early zeolites and the replacement of zeolites and glass by

K-feldspar and albite during diagenesis.

70 | CHAPTER 3

Method

The relative timing of alteration mineral assemblages isdetermined by documenting overprinting relationshipsamong the alteration mineral assemblages, mineral deposits,successive volcanic units or intrusions, the compactionfoliation and regional tectonic cleavages. This requiresdetailed examination of textural and mineralogical featuresthat are visible in the field, hand specimen (including drillcore) and thin section. Although thin sections can be useful indisplaying the inter-relationships between different minerals,they may be too small to show larger cross-cutting features(e.g. veins) and textural features in coarse-grained rocks.Paragenetic determination relies on a representative suite ofsamples that includes all rock types and alteration facies inthe alteration system. Because overprinting relationships arenot always straightforward the inspection of a large numberof samples is required for a systematic understanding of theoverprinting textures.

Overprinting textures

Overprinting texture refers to any texture or geometry thatcan be used to infer that one mineral or group of mineralshas formed later than another mineral. Table 3.4 describesthe three main types of overprinting textures. There is a widerange of overprinting textures, some of which are describedin Sections 3.1 and 3.2; however, the reader is referredto Ramdohr (1980), Craig and Vaughan (1981), Ineson(1989) and Taylor (1996) for more detailed discussions ofoverprinting textures and their significance.

The possible relationships among alteration mineralassemblages, mineralised rocks, successive volcanic units orintrusions, the compaction foliation and regional tectoniccleavages and their implication for the timing of alterationare described in Table 3.5.

Recent work (Allen, 1997; Gifkins, 2001; Gifkins andAllen, 2001) in pumice breccias around the Rosebery VHMSdeposit has unravelled a complex sequence of hydrothermaland non-hydrothermal alteration mineral assemblages.This work was based on overprinting relationships betweenalteration mineral assemblages, compaction foliation (S:) andtectonic cleavage (S2: Fig. 3.17). Alteration mineral assemblagesthat infill primary porosity, preserve delicate uncompacted

TABLE 3.4 | Types of overprinting textures.

Mineral superposition Where one mineral or a group of mineralscan be seen to have nucleated upon apre-existing mineral or grain; this appliesto replacement, infill and recrystallisationtextures

Where fractures fragment pre-existingminerals or grains, provide space forsubsequent infill and focus wall-rockalteration

Foliation superposition Where foliations or lineations rotate, modify,distort or fracture pre-existing minerals orcomponents

Fracture superposition

pumice textures and/or are overprinted by Sj and S2 areinterpreted to be early, prior to compaction and completelithification. These include thin films of sericite, chlorite +sericite + hematite and calcite (or their pre-metamorphicequivalents, smectites, calcite and palagonite) that coated alloriginal surfaces — including vesicle walls, plagioclase crystals,shards and fractures — and pre-date the infilling of vesicles bysubsequent minerals. Feldspar + quartz + sericite, chlorite +sericite, chlorite + sericite + hematite, chlorite, sericite andquartz + sericite alteration facies have filled and preservedvesicles, and replaced glass, indicating that they (or their pre-metamorphic equivalents such as zeolites and clays) formedprior to compaction and deformation.

Alteration features that define the bedding-parallelcompaction foliation, such as chlorite + sericite + hematiteand chlorite + sericite fiamme, are crenulated by the regionalcleavage (S2) and are interpreted to be pre- to syn-S,.

Alteration features that overprint the early alterationmineral assemblages and Sp but which are strongly foliatedby S2, are broadly syntectonic or post-Sj pre-S2. A chlorite+ epidote alteration assemblage has overprinted the earlysericite films and chlorite + sericite fiamme but is stronglyfoliated by S2 and interpreted to be post-Sj and pre- to syn-S2.Chlorite + calcite + magnetite has replaced and post-dated thechlorite + sericite fiamme and chlorite + sericite + hematitestylolites that define Sv and which are commonly transposedinto the S2 cleavage and associated with syn-S2 chlorite veins.Chlorite ± pyrite is associated with shear zones parallel to S2,and undeformed post-S2 brittle fractures and faults. Sericite+ calcite + actinolite ± epidote recrystallised earlier alterationmineral assemblages and are aligned parallel to S2.

Alteration mineral assemblages that are unaffected or onlyvery weakly affected by the S2 foliation are interpreted to bepost-deformation. These alteration mineral assemblages arerelated to chlorite + quartz + calcite vein infill and associatedvein wall-rock replacement and altered halos associated withDevonian granites, such as quartz ± calcite ± tourmaline +magnetite veins.

Early Mn-carbonate, chlorite, sericite and quartz + sericitealteration facies are spatially associated with ore at Rosebery(Fig. 3.16C), whereas the other alteration facies are regionallyextensive, spatially associated with synvolcanic or Devonianintrusions, veins, faults or shear zones. The overprintingrelationships, combined with spatial associations anddistributions of the alteration facies, suggest that diageneticalteration began shortly after eruption in the Cambrian andcontinued until the transition to regional metamorphism.Hydrothermal alteration associated with the VHMS depositsat Rosebery and Hercules commenced prior to burialcompaction, but was synchronous with diagenesis and theintrusion of synvolcanic sills. Peak regional metamorphismwas synchronous with Devonian deformation and isoverprinted by contact metamorphic assemblages associatedwith the emplacement of Devonian granites.

The interpretation of textures to determine a parageneticsequence can result in misleading results unless the subsequenteffects of metamorphism and deformation upon the texturesand minerals are appreciated. For example, in the Que RiverVHMS deposit galena commonly occurs in the S2 cleavage.As a result, conventional interpretation may conclude thatit formed syn-S2 (Devonian). This interpretation would

suggest that galena post-dates the other sulfides (sphaleriteand pyrite), which are overprinted by S2. However, geologicalevidence (Large et al., 1989) combined with lead isotope data(Gulson and Porritt, 1987) indicates that galena depositionoccurred on the seafloor at the same time as the other sulfidesin the Cambrian. In this case, the textural evidence constrains


the timing of recrystallisation of the galena to syn-S2, whereasthe geological and isotopic evidence suggests syn-depositionalformation of galena. Misinterpretations of this type arecommon for soft and easily recrystallised minerals suchsericite, chlorite, clay minerals, galena and chalcopyrite.

as

TABLE 3.51 Overprinting relationships in altered volcanic rocks and their implications for the relative timing of alteration (modified afterAllen and Large, 1996)

Alteration facies to primary volcanic textureAlteration facies truncated by clast marginsClasts with different alteration facies in the same rockAlteration facies infills primary porosityAlteration facies cross cuts primary porosity or clast marginsRock contains relict high-temperature devitrification textures

Alteration facies to successive volcanic units or intrusionsAlteration facies truncated by younger less-altered rocksAlteration facies overprints intercalated intrusionsAlteration facies overprints younger rocks

Between alteration facies and diagenetic compaction texturesAlteration facies protects primary texture from compactionAlteration facies defines the compaction foliation (e.g. fiamme)Alteration facies overprints the compaction foliation

Hydrothermal assemblages and diagenetic faciesHydrothermal facies overprinted by early diagenetic faciesHydrothermal facies overprinted by late diagenetic faciesHydrothermal facies overprints late diagenetic facies

Alteration facies to early bedding-parallel stylolitic dissolution foliation (S1)Alteration facies overprinted by stylolitic foliationAlteration facies overprints stylolitic foliation

Alteration facies to tectonic foliations and lineations (>S2)Alteration textures deformed and cut by tectonic foliationAlteration textures less deformed than primary volcanic texturesUndeformed alteration textures

Alteration facies to metamorphic assemblages and texturesAlteration facies overprinted by metamorphic assemblagesAlteration facies overprints metamorphic assemblages

Megascopic alteration facies distributionRegional distributionStratabound in formerly permeable faciesLocalised in fractured domains of coherent volcanic rocksRestricted to faults and shear zones

Overprinting relationships between all alteration assemblages

Pre-fragmentationPre-frag mentationPre-lithificationPost-lithificationPost-devitrification

Older syn-volcanic/intrusionSyn- to post-intrusion and overlying unitsSyn- to post-younger rocks

Pre-compactionPre- to syn-compactionSyn- to post-compaction

Pre- to early-diageneticSyn-diageneticPost-diagenetic

Pre- S1

Post-S1

Pre-cleavageSyn-cleavagePost-cleavage

Pre-metamorphicPost-metamorphic

Diagenetic or metamorphicPre- to syn-lithificationPost-lithification and post-fracturingSyn- to post-faulting

Provides a paragenetic sequence fordifferent alteration assemblages

72

I 73

4 | GEOCHEMISTRY OF ALTERED ROCKS

This chapter reviews three principal types of geochemicaldata that are used to characterise, quantify and interpret theprocesses of alteration: (1) whole-rock lithogeochemistryincluding major, trace and immobile trace elements; (2)mineral chemistry or composition of individual mineral phasesin alteration mineral assemblages; and (3) stable isotope ratiosof whole rocks and specific alteration minerals.

Geochemistry is the study of the abundance anddistribution of chemical elements in the earth (Mason, 1966).It has particular applications for the interpretation of alteredrocks. Geochemical data can aid recognition of the type andcomposition of the precursor rocks before alteration. Thechemical compositions of metasomatised or altered rocksand mineral phases reflect compositional changes duringalteration and provide clues about the alteration processes.On a utilitarian level, geochemical data are routinely usedin mineral exploration to identify broad chemical halosand gradients or vectors toward mineralised zones, andto discriminate between prospective and non-prospectivetargets.

4.1 | LITHOGEOCHEMISTRY

Literally, lithogeochemistry is the determination and study ofthe abundances and distribution of elements in stones. In thischapter, we understand it to mean the whole-rock chemistryof the 10 or 12 major elements in rocks (usually expressedas oxides) and various groups of trace elements, such as rareearth elements (REE) and immobile elements. Galley's (1995)review of lithogeochemical applications in massive-sulfideexploration also included mineral and isotope chemistry butthose methods are considered here in separate Sections (4.2and 4.3).

There are three main applications of lithogeochemistry inmineral exploration (Eilu et al., 1997):(1) identification or discrimination of prospective and non-

prospective areas and lithological units(2) recognition of large alteration and geochemical halos to

increase the size of exploration targets(3) definition of exploration vectors based on compositional

gradients around ore deposits.

Several aspects of lithogeochemistry that attempt to 'seethrough' the effects of alteration to reveal the compositionsof the protoliths, identify the process or type of alterationand quantify the chemical changes associated with alterationhave particular relevance to studies of altered rocks. Chemo-stratigraphy uses immobile elements to identify lithologiccorrelations, magmatic affinities and geotectonic settingsin otherwise unrecognisable altered rocks. Mass-transfertechniques, also based on immobile elements, are used to inferthe compositions of unaltered precursors and quantitativelyestimate the major element changes that occurred. Rare-earth-element geochemistry can facilitate recognition of rockalteration processes (e.g. Eu anomalies in exhalites). Alterationindices may assist in the discrimination of alteration styles orfacies, and the quantification of alteration intensity (Section2.4).

If composition data are available, lithogeochemicalmethods certainly contribute to the interpretation andevaluation of altered rocks. However, they are not quicksolutions to all problems and may, if used in isolation, leadto false conclusions. In every case it is important to considerother geological information: field relationships, distributionor zonation of alteration mineral assemblages, macro- andmicro-scale textures, mineralogy et cetera.

Before describing the lithogeochemical methods, weprovide explanations of some common lithogeochemicalterms (Table 4.1), brief descriptions of geochemical samplingand analytical techniques, and draw particular attention tothe phenomenon of closure in composition data.

Sampling and analytical methods

How do we acquire lithogeochemical data? As in many otherfields, the quality of interpretation rests on the quality ofthe data. Rocks are particularly difficult to analyse becauseof their wide compositional range and the chemical diversityof the elements of interest. An inappropriate analytical tech-nique can lead to an expensive list of useless numbers andto the discredit of the unwary geologist who attempts to usethem. This section offers advice on effective methods (theyare never inexpensive) and some pitfalls to be avoided.

74 | CHAPTER-

TABLE 4.1 | Explanation of lithogeochemical terms.

Composition data

Mass

Units of mass change

Absolute mass change (Aa)

Relative mass change (Ar)

Net mass change

Proportion

Data expressed as part of a whole (e.g. in weight percent, wt%, or parts per million, ppm).

A quantity of material, whole rock or its components, measured in weight units (g).

Usually expressed in grams per hundred grams (g/100 g) to avoid confusion with composition data in weightpercent, although they are effectively the same.

Mass changes that are proportional to the whole initial mass of the rock. Usually expressed in grams per hundredgrams (g/100 g). It may refer to individual components or the sum of all components (net mass change).

The absolute mass change in a component expressed as a percent proportion of the initial composition of thatcomponent in a rock. Relative mass changes distort the perception of chemical processes. For example, theaddition of a small quantity (say 0.5 g/100 g) of MnO initially present at 0.5% would produce a 100% relative gainin MnO. In comparison, addition of the same quantity of SiO2 (0.5 g/100 g) to the same rock initially containing50% SiO2 would represent only 1% relative gain. This gives the false impression that MnO was a vastlymore important element of mass change than SiO2. Absolute mass changes are preferable because they arequantitative rather than proportional, and accurately reflect the quantities of materials added to or subtracted fromthe system.

The sum of positive and negative rock component mass changes relative to the initial mass of the rock. Usuallyexpressed in grams per hundred grams (g/100 g).

The amount of a component expressed as a proportion of the whole rock. Usually given in units of weightpercent (wt%) for major elements and oxides, and in parts per million (ppm) for trace elements. It is analogous to"concentration" in chemical solutions.

Sampling strategy

The number and size of samples to be used in a lithogeochemicalinvestigation depends on many factors. These include theavailable budget, degree of exposure, geological complexity,compositional homogeneity of rock units or altered zones,the elements of interest and the volume of rock that can bephysically removed from the outcrop or drill-core shed. Therecan be no universally applicable strategy but generally, moreis best. Obviously, a good understanding of the geologicalcontext is fundamental to any subsequent lithogeochemicalinterpretation. Smash and grab sampling may lead to in-adequate and often meaningless interpretations.

Field methods should avoid sampling rocks potentiallymodified by weathering or other superficial secondaryprocesses, unless these processes are under investigation. Thesampler must also recognise potential overprinted alterationfacies. For example, a sample containing a dense network ofmetamorphic quartz veins is of little value if the purpose isto study synvolcanic pervasive alteration phases; it would bebetter to select a sample without veins. Handling, transportand subsequent storage of lithogeochemical samples shouldnot allow contamination or cross contamination. Whensampling drill cores or cuttings stored at a mine site, be awarethat mines and concentrators are dusty places, and expect tosee the signature of the ore in your trace-element data if thesamples are not clean.

In all but very coarse-grained rocks, samples of around1 kg should be adequate (Potts, 1987). In fine-grained orglassy volcanic rocks, samples in the range 200 to 500 g aregenerally sufficient for major element analyses. However, forlow-level trace elements (<1 ppm) the 'nugget effect' may besignificant if elements of interest are concentrated in sparselydisseminated grains (e.g. zircon). Potts (1987) provided illus-trative tables of the sample weights required for analyticalprecision, at various concentrations and grainsizes. These

represent the quantities of sample that actually undergo analy-sis and should not be confused with the amount of pulverisedsample from which smaller portions are taken for analysis.

Sample preparation

All whole-rock analytical techniques (except neutron activa-tion analysis) require samples to be crushed and ground tofine powder for direct analysis, acid digestion or fusion. Thereare opportunities here for sample contaminations from wearon the grinding machinery, cross contamination betweensamples in a batch, and contamination from previous batchesof samples put through the mill. Reputable laboratories willuse routine procedures to minimise those problems. However,if low-level trace elements are important, it may help to informthe laboratory personnel and request extra care.

Jenner (1998) cautioned against the use of new diamondsaw blades and recycled coolant for trimming slabs. Hesuggested that if sawn rock slabs must be analysed, they shouldbe washed in acetone and distilled water. If the rock samplesinclude overprinting phases, such as veins and weatheredrinds, these are best cut out before crushing. The practiceof crushing to -10 mm and attempting to pick out therock chips without extraneous materials does not eliminatecontamination.

Jaw crushers are infamous contaminators. Crevicesbetween their moving parts harbour dust that causes crosscontamination, and the wear of plates directly contributes steelparticles. Magnetic separation of steel particles is inadvisablebecause it may lead to selective removal of magnetic minerals.A better alternative is to crush rock samples down to ^5 mmfragments in a hydraulic press fitted with tungsten carbideplates. The crushed rock, including any fine dust, can then beconed and quartered down to about 80 or 100 g for pulverisingto a final grainsize of less than 75 (Am in a swing mill.

Materials used in swing mills include tungsten carbide,chrome steel, alumina ceramic, and agate. All of these arepotential contaminators of some elements. Chrome steel isthe least expensive, but seriously compromises Cr and low-level Fe analyses. Tungsten carbide is popular and durable butwill contribute Co and W to the sample, and possibly also Ti,Ta and Nb (Potts, 1987). Agate mills are fragile, expensive andslow but are the least likely to contaminate samples destinedfor low-level trace element determination.

Precision, accuracy and reference materials

Precision is a measure of the repeatability of data within ananalytical session, and reproducibility in different analyticalsessions over longer intervals. Analysing several duplicatesamples within batches and subsequent batches provides anestimate of the precision of a technique. Precision is reportedas the co-efficient of variation (CV), which is equivalent torelative standard deviation (RSD) expressed as a percentageof the mean.

where Cr' is the true or accepted proportion of element i inthe reference material, and Ca

] is the apparent or analysedproportion of element i in the reference material.

Reputable analytical laboratories will routinely run refer-ence materials to calibrate their instruments and maintain anacceptable level of accuracy. It is helpful if the client, uponsubmittal, gives the analyst an indication of the rock types,mineral assemblages and compositional range of the samples,to enable an appropriate selection of reference materials.However, laboratories do not routinely report their analysesof reference materials for comparison with accepted values.Even if they do, the potential for fudging the results providesthe user with a less than satisfactory assurance of accuracy.The solution is for the user to acquire some appropriategeochemical reference materials and include a sample or two,preferably disguised, in each analytical batch. This can involveconsiderable expense. Potts (1987) suggested that a minimumof 10 reference materials should be submitted with each batchfor a full assessment of accuracy.

In many cases, the user may not be greatly concernedwith the 'correct' values provided that the data are relativelyconsistent or precise within and between batches. Submitting

GEOCHEMISTRY OF ALTERERD ROCKS | 75

representative duplicate samples allows the user to monitorlong-term precision and calibration drift. It permits greaterconfidence in the analytical data, which justifies the additionalexpense. Keeping the reference materials and duplicatesanonymous can be problematic, even if the sample numbersare disguised. A few small packets of powdered or crushed rockstand out prominently in a batch of drill-core samples andanalysts have no difficulty in recognising them as standards.They may analyse them with special care, develop their owncomparative data over successive batches, and be tempted toadjust any outliers. However, the recognition that the clientis prepared to carry out independent quality control generallyhas a positive effect on analytical practice and discouragescomplacency at the laboratory.

Limit of detection

Limit of detection of an element is commonly understood to be'the lowest concentration that can be confidently measured bya particular method on an average sample' (e.g. Anonymous,1997). However, as pointed out in detail by Potts (1987), thelevels of confidence are frequently not stated. That obscuresthe reality that the quoted detection limits are often belowthe levels at which reliably quantitative measurements arepossible.

Potts (1987) proposed three new terms for better definitionof the much abused 'detection limit':• Lower limit of detection (LLD) for a signal level of three

standard deviations higher than the mean background(mean + 3s). This is the lowest level at which an elementcan be recognised but not quantitatively estimated.

• Limit of determination (LOD) representing a level sixstandard deviations above the background signal (mean+ 6s). It is the lowest level at which the signal can bequantitatively measured for a confident analysis.

• Limit of quantitation (LOQ) is set at 10 standard deviationshigher than background (mean + 10s) to provide extraconfidence in the analysis in cases where there are legal,commercial or statutory implications placed on theinterpretation of detection limits.According to Jenner (1998) the limit of detection

commonly quoted by analysts is equivalent to Pott's (1987)lower limit of detection (LLD; mean + 3s). Jenner (1998)gave the example that if LLD for an element is 0.3 ppm,then data reported between 0.3 and 1 ppm may or may notbe significantly different, but the element is recognisablypresent. Data above 1 ppm may be considered quantitative,or in other words, data at 1.1 and 1.5 ppm can be confidentlyregarded as different.

The consequence is that one cannot place much relianceon data reported at close to the lower limit of detection. Asafe rule of thumb is to treat with circumspection any data ofless than one order of magnitude above the quoted detectionlimit. Therefore, select an analytical method that will providequantitative data at an order of magnitude lower than thethreshold of interest.

where s> is the standard deviation of element i in the samplesanalysed, and %J is the sample mean.

Accuracy is a measure of how close analytical data lie to the'true' values. It may be evaluated by including in each analyticalbatch some international geochemical reference materials, orin-house standards for which the true values are relativelywell determined. It is important that such reference materialshave compositions near the compositional range of the rocksbeing analysed (i.e. a standard rhyolite is inappropriate whenanalysing a batch of basalts). For each element accuracy isexpressed as a positive or negative percentage difference,relative to the true or accepted values of the standard.

76 | CHAPTER 4

Analytical techniques

The most popular methods for analysis of whole-rock samplesare X-ray fluorescence spectrometry (XRF), inductively coupledplasma atomic emission spectrometry (ICP-AES) and inductivelycoupled plasma mass spectrometry (ICP-MS). These methodsoffer good precision for a large number of major and trace elementsover wide concentration ranges. XRF remains the preferredmethod for major elements (Robinson, 2001). However, somecommercial laboratories have recendy converted to ICP-AES,presumably for its rapid throughput and ability to measure mostmajor and trace elements. ICP-MS is used only for trace elementdeterminations. Neutron activation analysis (NAA) provideslow limits of detection for REE, some platinum-group elementsand some high-field-strength elements but it requires access toa nuclear reactor, produces radioactive waste and has slow turn-around of weeks or months.

The geochemical laboratory at the University ofTasmania's School of Earth Sciences uses a combination ofXRF on flux-fused and powdered samples (for major elementsand moderate-abundance trace elements, respectively) andsolution or laser-ablation ICP-MS for low-abundance traceelements (Robinson, 2001).

To obtain the appropriate quality of data, it is importantto involve the analysts in the selection of analytical methods.Inform them of the approximate compositional range in thesamples and the geochemical objectives of the analyses.

X-ray fluorescence (XRF) can be used to analyse up toabout 60 elements with atomic numbers greater than 10(Na upwards) at concentrations from 100% down to a fewparts per million (Rollinson, 1993). Detection limits fortrace elements in the range 0.5 to 2 ppm are achievable, andprecision for major elements approaches less than 1% RSD(Robinson, 2001). However, routine procedures used incommercial laboratories generally result in higher detectionlimits of 2-10 ppm.

Optimum detection limits and precision for traceelements, at concentrations below about 0.2%, are obtainedby analysis of 6—10 g undiluted rock powder pressed into apill. The rock powder must be ground to a grainsize of lessthan 75 [im to ensure homogeneity in the sample.

Major element concentrations are determined on glassdiscs made by fusing a small amount of powdered rockdiluted with lithium metaborate and tetraborate fluxes. Thefusion produces a homogenous glass, which enables analysisof the light major elements and minimises X-ray absorptionand enhancement matrix effects. The composition of therock influences the type and dilution factor of the flux tobe used, especially in sulfide-, base-metal- and carbonate-bearing samples. Accordingly, at the geochemical laboratoryin the University of Tasmania's School of Earth Sciences,approximate proportions of S, Fe, Ca, Ba, Cu, Pb and Znare first determined on the powder pills, along with the traceelements, to enable appropriate selection of fusion fluxes.

In sulfide-bearing samples, fusion should be a two-stageprocess with LiNO3 in the flux to oxidise sulfur. Initially, thecarefully weighed sample-flux mixture is heated and held for10 minutes at 700°C to ensure that the sulfur is retained assulfate, and not evolved. It is then heated to 1000°C for afurther 10 minutes to complete the fusion and the melt is castinto a 32-mm diameter glass disc.

In the past decade, inductively coupled plasma (ICP)analysis has revolutionised geochemical analysis, particularlyfor trace elements. There are two separate methods of analysisknown as ICP-AES and ICP-MS. Both use inductivelycoupled high-temperature argon plasma to generate atomicand ionic emissions in the sample. In ICP-AES (atomicemission spectrometry) the spectrum of atomic emissions ismeasured by an array of photomultipliers. This method issometimes called ICP-OES (optical emission spectrometry).ICP-MS uses a mass spectrometer to measure ionic particlesin plasma-sample gases. ICP-AES provides low limits ofdetection typically 2—10 ppm for trace elements and 10—100 ppm for major elements. ICP-MS enables determinationsof heavier trace elements at extreme detection limits, up tofour orders of magnitude lower than ICP-AES (e.g. REE andHFSE detection limits in the range of 0.1 to 2 ppb).

ICP methods are able to measure most elements at lowdetection limits with high precision using linear calibrationover eight orders of magnitude (Robinson, 2001). Up to 50elements can be analysed simultaneously in a few minutes onsamples of less than 100 mg.

There are a number of disadvantages to ICP related tothe requirement that the rock sample must be dissolved indilute solution before introduction to the plasma. Ensuringcomplete dissolution of rocks, including refractory phasessuch as zircon and REE minerals, and maintaining them insolution without contamination, is a difficult task (Yu et al.,2001). Samples are typically digested by strong acid cocktailsin sealed teflon vials for one or two days, dried by evaporationand then the residue is re-dissolved in dilute nitric acid readyfor analysis. In some cases, samples are fused with lithium-borate fluxes or sodium peroxide before acid dissolution. Itis advisable to test for complete solution by comparing theICP-MS data with XRF determinations of some relativelyimmobile elements, such as Zr, Nb, Y and REE (Robinson,2001). The likely loss of some volatile elements during fusionand evaporation may render the method unsuitable for Hg,Tl, Sb, Se and As.

H2O and CO2

A combination of XRF and ICP-AES or ICP-MS can provideaccurate analyses of most major elements and a surfeit oftrace elements. However, they cannot determine H2O andCO2 as XRF is limited to elements of atomic numbers greaterthan 10, and ICP, which measures the samples in solution,obviously cannot determine water. Loss on ignition (LOI) is apoor substitute for H2O and CO2 analyses. For geochemicalanalyses of altered rocks that contain significant hydrousminerals or carbonates, it is preferable to separately determineH2O and CO2.

Hydrogen and total carbon in rock samples can be analysedby the 'C-H-N elemental analyser', an instrument designed forroutine determinations of C, H and N in organic compounds(Potts, 1987). It is a relatively straightforward and inexpensivetechnique requiring only about 25 mg of finely ground rockpowder. On the, generally safe, assumption that alteredvolcanic rocks do not contain significant organic substances,the total H and C determinations can be recalculated to H2Oand CO2 for inclusion in major element composition data.

Interference

Some instrumental analytical methods are susceptible toinaccuracies caused by peak overlap, particularly in mineralisedrocks with exotic-metal contents. For instance, in XRF analyseson powdered samples, high concentrations of Ba and Pbinterfere with determination of Ti and Zr, respectively. Theseproblems may be minimised by instrument configuration or,in extreme cases, another method of analysis. Therefore, itis important that the analyst is advised which elements arecritically important for the geochemical interpretation andthe approximate compositional ranges of the rocks submitted,so that the appropriate analytical methods and procedures areapplied. For example, if the purpose of analyses is to inferthe volcanic precursors of Pb + Zn + Ba-bearing mineralisedrocks by immobile element chemostratigraphy, it wouldhelp the analyst to know that accurate Ti and Zr analyses arepriorities, and that the mineral assemblage includes Pb- andZn-sulfides, and barite.

Reporting data

Major elements in rocks are conventionally reported in weightpercent (wt%), mostly as oxides and in order of decreasingcation valency: SiO2, TiO2, A12O3, FeO (or Fe2O3 total),MnO, MgO, CaO, Na2O, K2O, P2O5, H2O+, CO2, S andTotal. The percentages are relative to the dry sample; analysedafter driving off moisture (H2O~) by heating for a few hoursat 105°C. H2O+ in the analysis represents structural water incrystals or glass. All other elements are typically regarded astrace elements although some may be present at greater thanthe conventional 0.1% cut off. Trace elements are usuallyreported in alphabetic order in parts per million (ppm),which is equivalent to the SI unit |ig/g preferred by analysts.It is often useful to re-order the trace elements into groupsaccording to their geochemical characteristics (e.g. immobiletrace elements, rare earths, etc.) to facilitate plotting. It isimportant to note that geochemical analyses are usuallyreported to one decimal place more than can be quoted withconfidence (Potts, 1987).

Loss on ignition

Hydrogen and carbon, mainly as H 2 O and carbonate, aresignificant components of some silicate rocks, particularlyaltered rocks. However, since XRF or ICP methods cannotdetermine these light elements, it has become common toreport loss on ignition (LOI) as a proxy. LOI is determinedby igniting a weighed sample at 1000-1050°C for at least12 hours and then weighing the residue. Expressed in weightpercent (as for the major elements), it is often assumed thatLOI represents the combined proportions of H2O+ and CO2

in the rock. However, for several reasons pointed out by Potts(1987), the determination of LOI has little geochemicalvalue:• The maximum temperature may not be high enough to

dehydrate some minerals, such as talc, topaz, cordieriteand epidote.

• Oxidation of ferrous iron (e.g. in silicates, magnetite and


sulfides) can produce a weight gain, partly offsetting theLOI. Even if FeO and Fe are separately determined, itis not reasonable to assume that they will be completelyoxidised in the LOI process.

• LOI may include a contribution from volatilisation ofalkali metals, sulfur oxides and fluorine; the loss may beonly partial and not predictable. For example, in samplescontaining both sulfide and carbonate, some SO2 mayreact with CaO and remain in the ignited sample.The variables affecting LOI and the wide range of

compositions in altered and mineralised rocks may leadto unpredictable results that are not readily interpretable.Consequently, LOI is not a reliable substitute for H2O

+ andCO2 analyses, and it may significantly underestimate theevolved volatiles.

Totals

The sum of the elements in a major-element analysis isfrequently taken as an indication of analytical reliability.Considering the shortcomings of LOI determinations, it isunreasonable to expect that their inclusion in a major elementanalyses will provide totals close to 100%. Nor should oneexpect that the error values on determinations of individualelements should cancel each other to produce totals of 100%.Further ambiguity is due to the usual practice of analysingand reporting total Fe as Fe2O3, irrespective of its actualoxidation state.

Sums of XRF major element oxides, sulfur and LOI datawill exceed 100%. This is because the sulfur is measuredtwice: in the fused disc XRF and in the LOI determinations.However, simple subtraction of XRF determined sulfur doesnot solve the problem because the sulfur may not be entirelyevolved in the LOI process. Reasonable totals are obtained ifH2O+ and CO2 are separately determined by another methodand summed with XRF major element oxides and sulfur.

Low totals do not necessarily indicate erroneous analyses;it may be that some significant elements were not determined.For example, boron constitutes about 10% of tourmalines andsome micas contain significant proportions of Li, Fl, Rb andCe (Deer et al., 1966), and mineralised rocks may containsignificant proportions of base metals.

Potts (1987) cautioned that, although it is nice to findtotals near 100%, it is not a satisfactory test of the qualityof analytical data. He considered that the only acceptableway of checking accuracy of modern instrumental analyticaltechniques, such as XRF and ICP, was to analyse appropriategeochemical reference materials. By applying this qualitycontrol practice to all of the methods used to produce a set ofmajor-element data, the user can have reasonable faith in therelative accuracy of the individual elements analysed. Likewise,confidence in the accuracy of individual determinationspermits acceptance of odd totals as accumulated errors, orindications of incomplete analyses.

Recalculating to volatile free

It is common practice, particularly in petrological literature,to recalculate major-element analyses to 100% 'anhydrous'

78 | CHAPTER 4

or 'volatile free'. The recalculation involves multiplying theproportion of each major-element oxide in an analysis, exceptLOI, by a factor derived from the formula 100/2 Q, where 2C| is the weight percent sum of all major oxides (not includingLOI) in the analysis.

Thus the 10 major oxides are commonly adjusted to sumto exactly 100%. LOI is reported, but not included in thetotal (e.g. Crawford et al., 1992; Stolz, 1995). The objectof recalculation is to remove apparent variations in theproportions of major oxides that may be due to differencesin LOI values. This may be a valid approach for studies ofpetrogenesis and fractionation in unaltered or weakly alteredrocks, where alteration was limited to hydration or minorcalcite vesicle fillings (Crawford et al., 1992). In these cases,additions of H2O+ and CO2 may be the only significantmetasomatic changes, although CaO may be added if it isnot otherwise derived from the decalcification of plagioclase.It is not unreasonable to 'subtract' the estimated additionsdue to hydration and carbonation to determine equivalentanhydrous magmatic compositions.

In more altered rocks, however, many of the other majorelements, particularly Si, Fe, Mg, Ca, Na and S, are also likelyto be involved in significant mass changes. Thus recalculatingto volatile free would distort the pattern of mass changes,artificially increasing the changes of other elements relative tothe constituents of LOI. In cases where the 10 major oxides(SiO2 to P2O5) sum to low totals, volatile free recalculationfollowed by immobile-element-based mass change estimatescould result in actual small net mass losses of some elementsappearing as mass gains. Similarly, it may upwardly distortthe estimates of net mass change, with implications forinterpretation of volume changes.

Barrett and MacLean (1994a) recommended using volatilefree recalculated major element data for mass change estimates.They neglected to mention whether the same recalculationfactor should also, for consistency, be applied to all the traceelements. Failure to similarly adjust the trace elements couldlead to small inconsistencies in immobile element ratios(e.g. Ti/Zr, which is usually calculated from major elementTiO2 data). It could also positively skew subsequent masschange estimates based on immobile elements. The volatilefree recalculation has a proportionally greater effect on datafor altered rocks because they generally have higher LOIthan their unaltered precursors. Upward recalculation of alltrace elements, as well as majors, can avoid those problems.However, comprehensive recalculation is unnecessary as bothchemostratigraphic methods and mass change estimates arebased on ratios of elements and hence are unaffected byrecalculation.

The method used in some studies (e.g. Gemmell andLarge, 1992), of recalculating the 10 major oxides but notother major elements, such as sulfur, is also inconsistent. Ithas the effect of under-estimating mass changes in sulfurrelative to the other oxides.

In dealing with altered rocks, it is preferable to obtain asnear as practical 'complete' major element analyses. The majorelement suite should include sulfur, H2O+ and CO2 and anyother elements likely to exceed 0.1% (e.g. base metals). Ensureadequate quality controls and then trust in the accuracy of the

data. They do not require fudging or normalisation at the riskof introducing new errors and misinterpretations. Otherwise,where analyses of altered rocks are limited to the usual 10major oxides and LOI, it is best to treat LOI as a (somewhatfuzzy) component in its own right, and not to recalculate toartificial 100% totals. This particularly applies to data that areto be used in mass transfer calculations.

Closure in composition data

Closure, also known as the constant sum effect, affects allanalytical data expressed as proportions of a whole; that is,as composition data (or as 'concentrations' in Stanley andMadeisky, 1996). The total of all elements in a rock analysismust sum to 100%, or one million ppm, and so forth,according to the units of measurement. Chemical masstransfers that change the total mass of a system by addingor removing some elements (e.g. hydrothermal alteration)will affect the proportions of all elements, even those notinvolved in the mass transfers (Fig. 4.1). Consequently, theapparent differences in composition data between unalteredand altered rocks do not accurately reflect the real materialchanges, except in systems where there has been no net masschanges

Closure particularly obscures mass changes in the majorelements such as Si, Al and Fe, which exist in high proportionsin primary unaltered rocks. For example, intense silicificationof felsic volcanic rocks may not be apparent in compositiondata (e.g. footwall alteration zone of the Thalanga deposit,Herrmann and Hill, 2001). Closure is less of a problem forelements of low initial proportions. For example, in felsicvolcanic rocks, Ca and Na concentrations average only a fewweight percent and their depletion associated with plagioclase-destructive alteration is usually evident in the compositiondata.

Trace elements, by definition, are present at lowproportions in background rocks. They tend to providehigh-contrast anomalies in mineralised and altered rocks,commonly orders of magnitude greater than backgroundlevels, and are therefore practically unaffected by closure.

FIGURE 4.1 | Schematic illustration of the effect of closure in compositiondata (after Eilu et al., 1997). Alteration leading to a net mass gain in the systemresults in a lower proportion (or concentration) of element i, although the actualmass of element i is not changed (AM, = 0). Mp = initial mass and MA = alteredmass of total system.

GEOCHEMISTRY OF ALTERERD ROCKS I 79

Rollinson (1993) presented a detailed discussion of theproblem of closure in geochemistry and suggested somesolutions. It is not a trivial matter in the study of altered rocks.The techniques for estimating alteration-related mass transfersovercome the problem of closure by quantifying the amountsof elements or oxides gained by or lost from altered rocks.

Chemostratigraphy

Major element compositions are routinely used to classifyvolcanic rocks in terms of petrogenesis and tectonic setting(e.g. Pearce and Cann, 1973). However, the same method isnot applicable to altered rocks because many of the majorelements, especially Si, Fe, Mg, Ca, Na and K, are relativelymobile during alteration. Consequently, compositionalchanges related to alteration may considerably outweightheir primary variations. Fortunately, several elements arechemically immobile during most types of alteration and thesecan be reliably used to classify and correlate altered volcanicrocks. In this context, immobile means elements that areneither added to, nor taken from, the rock during alteration.Immobile elements may be involved in phase changes andperhaps be mobile at millimetre scale, but their mass in thealtered rock remains unchanged. Although the proportions(concentrations) of immobile elements may change, due tonet mass changes in the size of the system, their inter-elementratios remain the same.

Immobile elements

The high-field-strength elements Ti, Zr, Nb and Y arerelatively immobile during hydrothermal, diagenetic andweathering alteration, and during regional metamorphism upto mid-amphibolite facies. Ratios of these immobile elementsare the basis of tectono-magmatic discrimination diagramsdeveloped in the 1970s (e.g. Pearce and Cann, 1973; Floydand Winchester, 1978).

Many studies of VHMS deposits have shown that Al, Ti,Zr, Nb, Y, heavy REE (Lu, Yb), Hf, Ta and Th, and in somecases P, Sc, V and Cr, remain essentially immobile duringalteration. Their immobility has been documented even in themost intense hydrothermally altered zones directly beneathdeposits (e.g. MacLean and Kranidiotis, 1987; Skirrow andFranklin, 1994; Barrett and MacLean, 1994a). Barrett andMacLean (1994a) recognised some mobility of the light REEin proximal, intense chlorite altered zones beneath somedeposits and suggested that they may be useful as explorationvectors. Y and Nb show considerable scatter in some datasetsand this may be partly attributable to primary variations(Ewart, 1979) or slight chemical mobility in some systems.Analytical precision may also be a factor, particularly for Nb,which typically occurs at low concentrations not much aboveXRF detection limits.

In practice, Ti and Zr are the most reliably immobileelements. They can be inexpensively and accurately analysedby XRF on pressed powder pellets and they exist at easilydetectable levels in most volcanic rocks, unlike the heavyREE, Sc, Nb, Ta, Hf and Th, which generally exist at lessthan 20 ppm.

Incompatible elements

Incompatible elements are those that tend to be excluded fromthe lattices of minerals crystallising from magmas and areinstead partitioned into the melt phase. Hence, incompatibleelements exist at highest proportions in the most evolvedfelsic rocks.

The high-field-strength elements (HFSE) Zr, Y and Nbare generally incompatible, except in some calc-alkaline suites.They have similar low magmatic liquid-solid distribution co-efficients and so tend to retain similar inter-element ratiosthroughout a single magmatic fractionation series, and on x-ybivariate plots form linear trends that project from the zeroorigin. Subsequent alteration involving net mass gain or losscan change the proportion of incompatible elements in thewhole rock but their primary ratios are preserved. As a resultalteration trends coincide with the primary fractionationtrends.

The gradients of these trends vary according to magmaticaffinity (MacLean and Barrett, 1993; Barrett and MacLean,1994a). Samples from different magmatic suites thus produceseparate linear trends of magmatic enrichment, which regressto the origins on incompatible-incompatible immobileelement plots (e.g. Fig. 4.2).

Incompatible-incompatible element ratios and bivariateplots have chemostratigraphic applications even inhydro thermally altered samples in which the major elementsmay not be reliable discriminants. Incompatible-incompatibleelement ratios are used to identify magmatic affinities,favourable volcanic suites and terranes.

Compatible elements

Compatible elements have high magmatic distribution co-efficients (>1) and are preferentially taken up by mineral

FIGURE 4.2 | Schematic Y-Zr (incompatible-incompatible) plot used fordetermination of magmatic affinities in altered volcanic rocks (modified afterMacLean and Barrett, 1993).

80 | CHAPTER 4

phases crystallising from magma. Consequently, compatibleelements are depleted from the melt phase (this is oppositeto the enrichment in incompatible elements). Thus, therelative proportion of compatible and incompatible elementsin residual melts changes as fractionation proceeds. Batchesof magma that are successively tapped off and emplaced aseruptive or intrusive units will have successively smallercompatible-incompatible element ratios.

Aluminium, Ti, Cr, Sc and V are generally compatibleduring crystallisation and immobile during alteration(Barrett and MacLean, 1994a). Bivariate plots of immobilecompatible-incompatible element data for least-alteredsamples from a particular magmatic affinity should showsmooth fractionation trends, generally with negative slopes(Fig. 4.3).

Subsequent net mass gains and losses imposed by alterationwill produce, on compatible-incompatible immobile elementplots, separate alteration lines for each chemically distinctrock unit (Fig. 4.3). This property is particularly useful forthe discrimination and correlation of initially homogenousbut subsequently altered volcanic units that may be otherwiseunidentifiable.

Chemostratigraphic fingerprinting is commonly donewith Ti/Zr ratios and TiO2—Zr bivariate plots. However,in tholeiitic suites Ti enrichment parallels the typical Fe-enrichment trend at the mafic end of the fractionationseries (MacLean and Barrett, 1993). In other words, Ti isinitially incompatible in tholeiitic fractionation series upto about the composition of basaltic-andesite. On a TiO2—Zr plot, the mafic end of the tholeiitic series has a positivefractionation trend, essentially similar to the alteration trendssuperimposed by subsequent metasomatic net mass changes(Fig. 4.3). Therefore, TiO2-Zr is not a reliable discriminantof altered mafic tholeiites. A reasonable substitute is A12O3—Zr, which has a near linear, slightly negative trend in tholeiites(MacLean and Barrett, 1993). Titanium becomes compatiblein tholeiitic magmas more evolved than basaltic-andesite andthe TiO2-Zr fractionation curve then has a negative slope.Two examples of chemostratigraphic discrimination andcorrelation in the Mount Read Volcanics are presented inFigures 4.4 and 4.5.

FIGURE 4.3 | Schematic TiO2-Zr (compatible-incompatible) plot showing thenegative curvilinear fractionation trend typical of co-genetic calc-alkaline volcanicsuites (after Barrett and MacLean, 1994a). Mafic tholeiites may show a positivetrend up to the composition of basaltic-andesite, because of TiO2 incompatibilityin the early stage of magmatic differentiation.

It is worth emphasising that immobile element ratios mustbe used with discretion in chemostratigraphic discriminationand correlation. This method is most effective in rockswith primary compositional homogeneity, such as coherentlavas and sills, and possibly in some massive syneruptivevolcaniclastic units such as pumice breccias. Processesof magma generation, crystallisation and fractionationdetermine the immobile element ratios of magmas. Therefore,immobile element ratios are likely to be uniform within singlecoherent eruptive or intrusive emplacement units. However,volcaniclastic debris may be subject to unhomogenisingprocesses. Both lateral and vertical compositional variationsmay occur in volcaniclastic units as a result of:• mechanical sorting of components of different densities,

such as clasts, pumice, scoria, crystals (e.g. Ti-oxides andzircon) and glass shards, during eruption and transport

• winnowing of glass shards from turbidity currents• mixing of debris from a variety of volcanic sources in

variable proportionsor

• incorporation of extraneous clasts into the base ofvolcaniclastic mass flows.Gifkins (2001) showed that thick, graded, rhyolitic

pumice breccia units in the Central Volcanic Complex,western Tasmania, have Ti/Zr ratios that vary from -5 nearthe crystal- and lithic-rich bases, to -9 in the fine-grained,shard-rich tops of the units. This is consistent with anincreased concentration of zircon crystals towards the base ofthe units but may also reflect the abundance of felsic clastsin the basal portions. In contrast, graded rhyolitic volcanicbreccia units in the Rosebery hanging wall, western Tasmania,display the opposite trend. Large et al. (2001b) suggested thatdecreasing Ti/Zr ratios towards the top of these units weredue to physical fractionation of Zr-poor crystal and lithiccomponents from Zr-bearing originally glassy pumice andshards during emplacement.

Testing immobility

It is preferable that element immobility in any system isestablished, rather than assumed, before proceeding withchemostratigraphic interpretations of altered rocks. This isalso an important preparatory step for methods of estimatingmass transfers of mobile elements.

The simplest test is to plot potentially immobile elementson x-y bivariate diagrams with the origins at zero. If possible,the tests should use data from unaltered and variably alteredsamples of a single, originally compositionally uniform volcanicunit, such as a coherent lava or sill. If the selected elementsare immobile, the data points for a single-precursor systemshould align in a highly correlated linear trend, which projectsto the origin of the plot and through the data points of theleast-altered samples. These linear trends, or alteration lines,are due to net mass gains and losses of the mobile elementsin the altered rock samples (Fig. 4.6). Typically there is somedata scatter due to analytical errors and slight inhomogeneitiesin the primary rock. However, if both elements are immobile,calculated linear correlation co-efficients (r) for alterationlines should exceed -0.85 (Barrett and MacLean, 1994a).In contrast, elements that were mobile during alteration are


0 10 20 30 40 50 0 25 50 75 100Ti/Zr Alteration Index

100(K2O+MgO)(K20+Mg0+Ca0+Na20)

FIGURE 4.4 | Graphic log and down-hole Ti/Zr and Al plots of drill hole NC4, near Lake Newton, western Tasmania. The Ti/Zr

data clearly delineates units of different primary volcanic compositions despite effects of strong hydrothermal alteration in rocks

intersected in the middle and lower part of the hole. The high Ti/Zr ratios at -200-230 m led to recognition of an altered mafic

volcanic breccia unit, which had previously been interpreted as a zone of chlorite-altered felsic volcaniclastic rocks. Another

altered mafic unit occurs below 530 m. Quartz and feldspar crystal-rich sandstones in the upper 100 m have a range of Ti/Z

ratios, which suggests that they do not have a unique provenance.

readily identifiable by their erratic distribution or near totalremoval (Finlow-Bates and Stumpfl, 1981).

Mass transfer techniques

Mass transfer techniques aim to quantify the amounts ofindividual elements added to and subtracted from the rockduring alteration in order to overcome the distortions ofclosure that are inherent in composition data.

As noted by Barrett and MacLean (1994b), significantmass change anomalies may not be apparent in untreated

compositional data, due to closure. The results of masstransfer calculations are usually easy to relate to mineralassemblages and may reveal clues about the composition,source and temperature of hydrothermal fluids (e.g. Barrettand MacLean, 1994b). Mass change data have been usedto infer hydrothermal water-rock ratios (e.g. MacLean andHoy, 1991). They are also used in the interpretation ofwhole-rock 618O and REE data (e.g. MacLean and Barrett,1993). Thus, they may enable discrimination of favourablealteration systems and altered zones within systems. Whenplotted spatially mass transfer data can be used as quantitativeexploration vectors (Section 8.2).

82 I CHAPTER 4

FIGURE 4.5 | Chemostratigraphic correlation diagram of the volcano-sedimentary succession that hosts the Rosebery massive sulfide deposit, western Tasmania.A thin unit of dacitic pumice breccia, between the massive sulfide lens and the feldspar + quartz + biotite porphyry sill intersected in hole 120R, is texturallyindistinguishable from the footwall rhyolitic pumice breccias (Ti/Zr = 7-9) but has a distinctive Ti/Zr ratio of between 12 and 14.

FIGURE 4.6 | TiO2-Zr plot of data from a tholeiitic volcanicsuite (after MacLean and Barrett, 1993). The data points forleast-altered basalt, andesite and rhyolite samples definethe magmatic 'fractionation curve'. Two linear arrays of datarepresent variably altered samples of originally homogenousunits of andesite and rhyolite. These highly correlatedalteration lines intersect the least-altered data points onthe fractionation curve and project towards the origin. Datapoints on the alteration lines below the fractionation curverepresent net mass gains, those above the curve representnet mass losses. The positive slope at the basaltic end of thefractionation curve is due to incompatibility of TiO2 in the earlystages of differentiation of tholeiitic magmas. This may leadto confusion of fractionation and alteration trends in mafictholeiites.

There are several approaches to estimating mass transfers.These include the mathematically complex graphicalmethod of Gresens (1967); subsequently simplified by Grant(1986) and Huston (1993) to produce the isocon method;the immobile element techniques of MacLean and Barrett(1993), and the Pearce element ratio analysis of Stanley andMadeisky (1996). These techniques all depend on recognitionof immobile elements. They also, with the exception of thePearce element ratio analysis, depend on the identification ofprecursor-rock compositions.

The determination of unaltered precursor compositions isoften problematic in the altered, lithologically and structurallycomplex rocks that host ore deposits. The safest approachis to consider the geological context, established from fieldrelationships and rock textures, in combination with variousimmobile element tests (e.g. MacLean and Barrett, 1993;Stanley and Madeisky, 1996). This ensures that precursors areappropriately matched to the altered rocks under investigation.The isocon method does not include a procedure for selectingprecursors and it commonly produces erroneous resultsbecause incorrect geological assumptions are applied. For thisreason, we prefer the more rigorous MacLean and Barrett(1993) method, which is also simplest to calculate.

Another potential difficulty arises from primary compo-sitional variations in volcanic rocks related to magmaticfractionation or volcaniclastic mixing processes. Huston(1993) recognised this as a deficiency of the isocon method.He suggested examining standard deviations of data fromleast-altered samples to screen their suitability, estimatingerrors for the consequent mass changes and recognising thatsmall apparent mass changes may be artefacts of primaryinhomogeneities, not due to alteration. The Pearce elementratio analysis method (Stanley and Madeisky, 1996) and themultiple-precursor variant of the MacLean and Barrett (1993)method attempt to overcome the limitations of primaryvariability. Of these two, we prefer the MacLean and Barrettapproach, again for its relative ease of calculation.


The MacLean and Barrett multiple-precursor methodinvolves sorting samples into affinity groups, calculatingprimary variability trends (fractionation curves) and usingimmobile elements to synthesise primary compositions foraltered samples that may fall between compositions of theavailable least-altered samples. This method relies on theassumptions that co-genetic groupings can be recognisedand separated, and compositional variations are smoothlinear or curvilinear trends for every major element. Thisis generally true for magmatic fractionation, and thereforeapplicable to volcanic suites, but the variations may be erraticif volcaniclastic, sedimentary or multiple alteration processeswere involved (Eilu et al., 1997). Thus, due caution mustbe exercised when using this method in mixed provenancevolcano-sedimentary successions. Filtering sample setsthrough geological field evidence and petrographic texturalobservations are obvious fundamental precautions. Samplesets from restricted prospect-scale areas are less likely to havecomplex primary compositional variations than district-scalesample sets. Peculiar aberrations or significant mass changesin elements otherwise expected to be immobile, such as Al,should arouse suspicion that primary fractionation trendsmay have been incorrectly modelled.

Single precursor mass transfer technique

The MacLean and Barrett method of estimating mass transfersin single-precursor systems incorporates testing of immobilityas a fundamental step. The analytical data for potentiallyimmobile elements, from an initially homogenous butvariably altered lithological unit, are plotted on x-y bivariatediagrams and linear regressions are calculated. The existenceof highly correlated (r >0.85) trends that pass through theorigin enables selection of the optimal element to be usedas the immobile monitor in the mass change calculations.This graphical process also highlights any outliers or samples

FIGURE 4.7 | Diagram representing the calculation ofmass transfers for an altered rock with an initial mass of100 units (after MacLean and Barrett, 1993). The proportionof immobile element in the altered rock decreases (isdiluted) because of net mass gains in mobile element(s).The ratio of immobile elements Z°/Za represents the factor,which multiplied by the proportion of immobile element inthe altered rock, would restore it to a mass unit, rather thana proportion of the whole. The same factor applied to allother elements produces the 'reconstructed composition' inmass units. The mass transfers in each element are thencalculated by subtracting its primary composition from thereconstructed composition. In this example, a large massgain in X combined with a small mass loss in Y produces alarge net mass gain.

84 | CHAPTER 4

from different precursors, which should be eliminated ortreated separately. It identifies any primary compositionalinhomogeneities in the rocks, which may require treatmentby the multiple-precursor method (MacLean and Barrett,1993) or Pearce element ratio analysis (Stanley and Madeisky,1996). These more complex methods are not explained here;the reader is referred to the relevant references for details.

The single-precursor mass transfer method proceedsby calculating the ratios of the proportions of an immobileelement in the altered and unaltered samples. Each of themobile element proportions is then multiplied by that ratioto obtain a reconstructed composition (Fig. 4.7). The masschange of each element is found by subtracting its percent

proportion in the precursor from that in the reconstructedcomposition.

The steps below and the flow chart in Figure 4.8 outlinethe procedure for estimating mass changes in single-precursorsystems. Figures 4.9, 4.10 and 4.11 present a worked examplebased on compositions of Thalanga footwall rhyolites,Queensland.(1) Acquire and tabulate the lithogeochemical analyses (e.g.

Figure 4.9). The first part of this section (4.1) presentssome guidance on sampling and analytical techniques.Barrett and MacLean (1994a) recommended using majorelement composition data that are recalculated to 'volatile-free' totals of 100%. However, it is unreasonable to expect

Comments related toThalanga data in Figures 4.9,4.10 and 4.11.

Values below detection replacedby arbitrary small values; e.g.0.5 x detection limit or zero(Figures 4.9 and 4.10).

Data should preferably be frommappable single-precursorunits. This example includes 34rhyolites from diverse volcanicunits. However, the remarkableuniformity of im m obile- el em entratios suggests they wereessentially co-magmatic andcan be treated as a single-precursor system.

No obvious outliers. Yttrium(Fig. 4.10) typically showsconsiderable scatter, which maybe a result of primary variation.

AI?O3 andZr both have three offour correlation co-efficients>0.85 (Fig. 4.10). Either couldbe used as the immobilemonitor component. Zirconiumwas selected because ofmarginally higher averagecorrelation (0.86 vs 0.83).

Appropriate $ symbols in theformula enable it to be filledacross and down in thespreadsheet and maintain cellreferences to Z°, Za and C°.

Large positive and negative netmass changes are mainly due toSiO2 gains and losses, oraddition of Fe and S in pyriticsamples (Fig. 4.11), Note thatbecause of closure, the SiO2

changes are not obvious in theanalytical data.

FIGURE 4.8 | Flow chart for mass-change calculations by the single-precursor method (after MacLean and Barrett, 1993).

elements of an analysis to total exactly 100%, and theuse of recalculated data may produce positive distortionsin subsequent mass transfer estimates. In dealing withaltered rocks, it is preferable to obtain accurate analyses ofall the major elements (including S, CO, and H2O+) andinclude them in the mass transfer calculations. If S, CO2

and H2O+ data are not available, and losses on ignition(LOI) are significant (>2%), then LOI could be includedas a separate, albeit loosely denned, component of masschange.

(2) Test for immobility. Plot analytical data for potentiallyimmobile elements (e.g. Al, Ti, Zr, Nb, Y, etc.) on x-ybivariate diagrams with the origins at zero (e.g. Fig. 4.10).If possible, use data from variably altered and unalteredsamples of a single uniform rock unit.

(3) Inspect for outliers and, with geological considerations,decide whether to cull them, or treat the data as a multiple-precursor system. Readers are referred to MacLean andBarrett (1993) for details of the multiple-precursormethod.

(4) The immobile element data for single-precursor systemsshould plot on highly correlated linear trends (r >0.85)that pass through the origin and through the data pointsfor least-altered samples. Evaluate the bivariate diagrams


and correlation factors to determine which elementconsistently occurs in highly correlated trends and is mostsuitable as an immobile monitor (i.e. was uniform in theprimary rock and least mobile during alteration).

(5) Select a composition for the precursor. This could be froma single unaltered sample or an average of several unalteredsamples.

(6) Calculate the absolute mass change for each componentusing the formula:

Aa = [Z7 Z a . Ca ] - C°

where Aa is absolute mass change expressed in g/100 gCa = wt% proportion of component in altered rockC° = wt% proportion of component in precursorZa = proportion of immobile element in altered rockZ° = proportion of immobile element in precursor.

The mass changes may be calculated from compositionsof individual altered samples or from average compositionsof sample groups representing certain mineral assemblagesor altered zones.

(7) For visual comparison, plot the absolute mass changes forindividual elements on a bar graph (e.g. Fig. 4.11).

1

2345

6789101 11213141516171819202122232425262728293031323334353637

38

39404142

43444546474849

A B C

Whole-rock major and trace element composition data

Sample no.

* 140802* 140727

* 140808* 140724

* 140902TH394-142TH144B-34TH41A-575TH005-256TH238-236TH038-191TH085-125TH148-159TH085A-422TH270-278TH270-145TH085-312TH018-266TH038-054TH085-159TH085A-348TH238-194TH270-220TH085A-335TH085-204TH085-215TH085A-241TH085-188TH270-313TH061-086TH41A-713TH061-157TH270-381TH085A-384Precursorcomposition

Mass Changes

140727140808140724

140902

Alteration facies SiO,

east-altered footwall 76.40-noderate, foliated sericite + chlorite 70.90

strong, pervasive quartz + sericite + pyrite ± chlorite 75.70ntense, pervasive quartz + pyrite 67.00

ntense, microcrystalline quartz + K-feldspar 84.20east-altered footwall 77.39east-altered footwall 77.25

76.7177.8074.6277.8175.2582.0269.9764.3772.7072.7366.8275.3775.2966.9374.4163.3878.0471.2267.0871.5974.5472.5674.3975.8778.4855.3649.30

average of 140802, TH394-142 and TH144B-34 77.01

{Formula in cell C451 ={$Q$38/$O5"C5)- '• .1 C$38] is ?lled across and down. ' units h\ Placement of $ symbols is critical to \ DL 0.05i maintain cell references for Zo, Za and Co. \

\

V_ siOj

moderate, foliated sericite + chlorite ' >• -15strong, pervasive quartz + sericite + pyrite ± chlorite 7ntense, pervasive quartz + pyrite 43

ntense, microcrystalline quartz + K-fe!dspar 72units

D

TiO2

0.110.10

0.070.05

0.060.090.100.070.070.090.070.070.080.110.120.090.080.080.080.090.100.090.050.060.050.070.070.060.100.110.100.080.100.10

0.10

0.01

TiO2

0000

g/100q

E

AIA

11.9014.50

11.406.60

7.4011.8511.9010.188.83

13.4010.5411.138.84

13.0617.5612.2212.3311.2610.7012.7914.5712.187.999.488.28

10.4411.038.40

12.5513.7512.1711.6013.9013.47

11.88

0.05

AIA

1101

9/100g

F

FeA

1.641.69

5.1414.34

0.601.132.090.991.901.401.512.320.682.804.732.395.482.464.471.965.701.42

13.464.459.188.626.916.954.931.561.700.88

14.3619.14

1.62

0.01

FeA

04

24-1

g/100g

G

MnO

0.040.04

0.080.02

0.000.020.030.080.060.020.090.060.020.030.070.090.010.130.020.090.080.050.130.040.090.130.060.050.070.030.050.010.140.30

0.03

0/

0.01

MnO

0000

q/100q

H

MgO

0.673.94

2.381.67

0.080.760.684.332.143.543.131.901.051.984.354.241.109.882.912.514.704.165.881.763.976.422.532.773.881.191.120.614.907.79

0.70

0.01

MgO

322-1

q/100q

I

CaO

1.420.12

0.000.05

0.110.460.901.753.080.120.253.030.240.280.010.270.000.010.000.080.000.070.000.020.000.010.080.000.000.340.080.120.110.19

0.93

0.01

CaO

-1-1-1-1

q/100q

J

Na2O

2.271.05

0.210.06

0.344.792.771.570.841.820.900.590.530.380.240.220.170.120.120.100.090.080.000.160.150.120.110.090.030.950.420.370.180.08

3.28

0.05

Na.O

-2-3•3

-3q/100g

K

K2O

4.043.92

2.391.99

5.871.823.602.803.062.233.403.685.333.694.154.733.755.082.434.412.913.570.912.451.511.592.822.142.506.336.367.602.400.93

3.15

0.01

K2O

0-107

q/100q

L

PA

0.020.01

0.010.01

0.010.010.020.020.050.010.010.020.020.050.020.010.010.020.010.010.010.010.010.010.010.010.020.010.010.010.020.020.020.02

0.02

0.01

PA

0000

g/100g

M

S

0.010.00

0.6510.61

0.480.000.240.011.330.000.061.390.302.031.080.843.470.531.610.810.860.447.641.574.972.893.594.360.440.000.660.135.925.12

0.08

0/

0.01

s

01

191

q/100q

N

COj

0.000.10

0.000.00

0.150.15

0.990.00

0.26

0.00

0.18

0.10

0/

0.1

co2

0000

q/100q

O

Total

98.5296.37

98.03102.40

99.1598.4799.7498.51

100.1597.2597.7799.4499.1194.3896.7097.8099.1396.3997.7298.1495.9596.7499.4598.0499.4397.3898.8199.3797.0798.6698.55

100.0897.3996.44

98.91

P

LOI

0.602.67

2.758.24

0.490.490.771.382.122.492.081.751.033.394.232.564.072.653.542.504.112.797.602.955.335.024.734.503.171.241.710.767.227.83

0.62

Q | R

Zr Nb

146 13162 19

128 1479 7

80 9137 14141 12113 12114 13140 13119 13116 12114 11162 17195 20130 14129 15132 15119 14137 16176 16138 1683 9

105 1184 9

116 12120 1492 11

140 16158 18157 17122 13154 19145 16

141 13

ppm ppm1 1

S T

Y

4054

3644203433305730414228494636354527484639193017293729395247376344

35

ppm1

^Formula in cell 045 is Tiled •1 down [ =sum(C45:N45)i

Net

-159

8476

q/100q

FIGURE 4.9 | Example of layout and calculation of mass changes by the single-precursor method for 34 rhyolites from the Thalanga footwall, Queensland. Data from

Paulick (1999) and Paulick et al. (2001). Values below limit of detection replaced by zero. Sample numbers with * are featured in the Thalanga data sheets (Section 7.7).

86 | CHAPTER 4

FIGURE 4.10 | Bivariate plots of potentially immobile elements: AI2O3, TiO2, Zr, Nb and Y (Thalanga data from Figure 4.9). (A) TiO2 versus Zr, (B) AI2O3 versus Zr,(C) Nb versus Zr, (D) Y versus Zr, (E) TiO2 versus Nb, (F) AI2O3 versus Nb, (G) Y versus Nb, (H) TiO2 versus Y, (I) AI2O3 versus Y, (J) TiO2 versus AI2O3, (K) Na2Oversus Zr, and (L) sulfur versus Zr. These plots facilitate recognition of compositional outliers, which should be excluded from single-precursor calculations, andselection of the least-mobile element to serve as the immobile monitor in mass-change calculations. In this example, Zr and AI2O3 both show consistent highlycorrelated trends that project to the zero origin. Note the considerable scatters, and hence poorer linear correlations, in plots involving Y and the major elements sulfurand Na2O. These are consistent with primary compositional variations or significant chemical mobility, particularly for sulfur and Na2O.

FIGURE 4.11 | Bar graph showing estimated absolute mass changes of major elements in four samples representing major alteration fades in the Thalanga footwall,Queensland (data from Figure 4.9).

GEOCHEMISTRY OF ALTERERD ROCKS I 87

Rare-earth-element geochemistry related toalterationIn the past few decades, trace elements have become basic toolsor pathfinders in ore deposit exploration and in petrogeneticinterpretation. Immobile trace elements have valuable appli-cations in studies of altered rocks. The rare earth elements(REE) have some special properties during alteration, whichmay be useful in interpretation and should be understood inorder to avoid false conclusions.

Mobility of light REE and effects of net mass change

Rare earth elements (with the exception of Eu) are generallyincompatible during igneous fractionation. Heavy REE (Luand Yb) are essentially immobile, whereas the light REE maybe variably mobile during alteration (MacLean and Barrett,1993). Lanthanum is the most likely to be affected and themobility of the other REE decreases towards the heavy REE(Barrett and MacLean, 1994a). These incremental changes inthe lighter REE modify the slopes of chondrite-normalisedREE patterns, which may confuse petrogenetic interpretation.Therefore, immobility of REE needs to be established beforethey are used to infer magmatic affinity. Plotting geochemicaldata for each rare earth element against a reliably immobileelement, such as Zr, is a means of testing for scatter andmobility. If the elements were immobile, REE and Zr data

FIGURE 4.12 | Modifications in REE patterns due to hydrothermal alterationillustrated by REE profiles of variably altered rhyolites from the Ansil andDelbridge deposits, Canada (after Barrett and MacLean, 1994b). (A) Partialleaching of the mobile light REE in the Ansil footwall is reflected in profiles withdifferent slopes converging towards the least-mobile heavy REE. (B) Immobilityof all REE is evident in the sub-parallel profiles for rhyolites from the Delbridgefootwall. However, net mass changes in mobile major elements have causedchanges in the proportions of the REE, producing vertical shifts in the profiles.

from altered single-precursor systems or co-genetic volcanicsuites should produce highly correlated linear trends onbivariate plots.

In systems where REE were chemically mobile, bothpositive and negative shifts in light REE concentrations havebeen recorded. Significant mobility of REE appear to occurin proximal altered zones associated with VHMS deposits(MacLean and Barrett, 1993). The greater mobility of thelight REE may produce a fan shaped array in chondrite-normalised plots with the REE profiles converging towardsthe heavy, least-mobile REE (e.g. Fig. 4.12A)

Even in cases of REE immobility, net mass changesassociated with alteration may produce significant verticalshifts in chondrite-normalised REE patterns. The slopes ofthe REE patterns are retained, but the y-axis magnitudes aremodified (downward by net mass gain and upward by netmass loss; Fig. 4.12B).

Europium anomalies in seafloor sediments

Recent studies of sediments and hydrothermal precipitates inmodern and ancient massive sulfide environments have foundthem relatively enriched in light REE, particularly in Eu(Barrett et al, 1990; Peter and Goodfellow, 1996; Shikazono,1999). The explanation is that Eu exists in a divalent state infelsic magmas and hence is compatible in feldspars, unlike theother trivalent REE, which remain incompatible (Rollinson,1993). The divalent Eu+2 in feldspars may be liberated bysubseafloor hydrothermal alteration to sericite or chlorite,whereas the incompatible REE that are concentrated inalteration resistant phases, remain relatively immobile. Theliberated Eu+2 is transported by reduced acidic hydrothermalfluids and may ultimately be precipitated by oxidation atthe seafloor. Therefore, in felsic volcanic successions, alteredVHMS-footwall zones tend to be depleted in Eu. In contrast,positive Eu anomalies exist in seafloor sediments and jaspers,and probably indicate proximity to hydrothermal vents(Barrett et al., 1990). Both phenomena have significance formassive sulfide exploration in modern and ancient submarinevolcanic environments. The recognition of positive Euanomalies in stratiform jasper lenses recently contributedto the discovery of a small satellite massive sulfide depositat Thalanga in the Mount Windsor Volcanics (Miller et al.,2001).

4.2 I MINERAL CHEMISTRY

Principles

Minerals, by definition, are natural homogenous solids ofdefinite chemical composition and definite atomic structure(Dana, 1957). However, many minerals do not have simplydefined chemical formulas. Their compositions may liebetween limits defined by two or more end-member formulas,effectively forming solid solutions. Mineral crystal structurescan accommodate various impurities where atoms and ions

88 | CHAPTER 4

of suitable size and charge can substitute for others in thelattice, occupy interstices in the lattice or be omitted from aproportion of lattice sites.

The considerable variety of linked tetrahedral crystalstructures in silicate minerals permits a wide range of chemicalsubstitutions and interstitial solid solutions. A frequentlycited example is the olivine series in which Mg2+ and Fe2+

ions, having similar charge and size, substitute for eachother between the end member compositions of forsterite(Mg2SiO4) and fayalite (Fe2SiO4).

The sheet-like structures of phyllosilicates allow a greatrange of ionic substitutions and interstitial contaminants.For example, muscovite, with the ideal formula ofK2Al4[Si6Al2O20] (OH)4, commonly contains the isomorphoussubstitutions of Na, Rb, Cs, Ca and Ba for K; Mg, Fe2+, Fe3+,Mn, Li, Cr, Ti, and V for octahedral Al, and F for OH andtetrahedral cation proportions from Si6Al2 to Si7Al] (Deeret al., 1966). Layered clays (particularly smectites) alsoaccommodate many cationic substitutions and exchanges aswell as physical mixtures and inter-layered structures of morethan one clay mineral (e.g. smectite-illite).

The causes of such variations are both physical andchemical. Temperature, pressure, fluid pH, JO2, and cationsolubility may all affect the stability and composition ofthe minerals formed. Chemical and temperature gradientsin hydrothermal alteration systems commonly producespatial zonation of alteration mineral assemblages. Similarly,they may produce gradational and zonal variations in somealteration mineral compositions. For example, the Mncontent of metasomatic pyroxenes associated with Zn skarnsgenerally increases systematically along the fluid pathway, andcan be used to identify proximal and distal skarns, and alteredzones (Meinert, 1993).

The standard methods for determining mineral com-positions and crystal structure are electron microprobe andX-ray diffraction analyses. Although Galley (1995) noted thatthese methods are more readily available than previously, theimpracticality of mineral chemistry as an exploration tool, andthe complexity and cost of these laboratory techniques meansthat they are rarely used other than in academic research.The development of field portable short wavelength infra red(SWIR) spectrometers and spectral interpretation softwareduring the last decade, has allowed mineral chemistry to bepractically integrated into exploration programs for a varietyof deposit types (Thompson et al., 1999). SWIR spectralanalysis can determine compositional and crystal-structuralvariations in white micas, smectites, clays, chlorites, biotitesand carbonates (Pontual et al., 1997). These minerals areprominent in altered volcanic rocks and are also prone tosignificant compositional variations. The technique reliablyestimates variations in white mica composition but appearsto be less effective at analysing chlorites in typical mixedphyllosilicate assemblages (Herrmann et al., 2001). Carbonateshave relatively weak SWIR absorptions, which tend to beobscured in mixed assemblages containing phyllosilicates andare thus less amenable to spectral analysis.

Although there are many possible applications in mineraldeposit exploration, we do not know of any cases wheresystematic investigations of mineral chemistry have led toan ore discovery. As pointed out by Simmons and Browne(2000), the extent to which patterns of mineral distribution

and chemical variations can be used in exploration dependslargely on whether they are related to a single phase ofhydrothermal activity that produced equilibrium mineralassemblages.

Applications

The three main applications of mineral chemistry in alterationstudies are:(1) interpretation of the processes and physicochemical

conditions of alteration(2) discrimination or identification of metasomatic alteration

and mineralisation styles from mineral compositions(3) determination of spatial variations and exploration vectors

to ore.

Interpretation of processes

The compositional and crystal-structure variations in someminerals are diagnostic of particular alteration processesbecause of physicochemical influences on mineral stability andcomposition. Thus, mineral chemistry may be used to inferthe geological environment in which an alteration mineralassemblage formed. However, these kinds of compositionalvariations may not always be universally applicable; they mayrequire orientation testing to determine their usefulness indifferent districts or sites.

For example, Dill et al. (1997) found that the kaolin-alunite deposits in felsic volcanic rocks of western Peru couldbe classified into hypogene (hydrothermal) and supergene(weathering) types on the basis of the chemical variationsin kaolinite. Hydrothermal kaolinite tended to be rich inBa, Sr and sulfur, whereas weathering-related kaolin claysconcentrated Cr, Ti, Nb and REE. This approach has directapplications to mineral exploration because hypogene kaolinitealteration in Peru is associated with high-sulfidation epithermalAu-Ag deposits. Similarly, Yang et al. (1999) alluded to low-and high-crystallinity forms of kaolinite in the Comstockdistrict of Nevada, which they respectively attributed to low-temperature weathering and higher temperature hydrothermalalteration processes. They suggested that spectral recognitionof distinctive kaolinite could be used in satellite-borne remotesensing to detect prospective hydrothermal altered zones.

Discrimination of hydrothermal alteration styles

Mineral compositions can be used to identify or fingerprinthydrothermal-alteration mineral assemblages, and possibleassociations with mineralised rock. This application is usefulat early stages of mineral exploration to discriminate betweeneconomically favourable and less favourable alteration andmineralisation styles. For example, Zn-skarn assemblages arecommonly dominated by pyroxene with varying amountsof garnet, amphibole, bustamite, chlorite and carbonate,which may all be Mn enriched. Manganese-rich pyroxene(johannsenite) is virtually diagnostic of distal Zn skarns(Meinert, 1983) and could be used as an index mineral inexploration for this type of deposit.

Another example that is more relevant to submarinevolcanic successions, is the recognition of a class of pyritic-alteration systems in the Mount Read Volcanics that havesome characteristics of high-sulfidation epithermal deposits.These include Western Tharsis and Lyell-Comstock, whichcontain sub-economic Cu + Au resources (Huston andKamprad, 2000; Corbett, 2001), and Basin Lake and Chester,which appear to be barren (Boda, 1991; Green and Taheri,1992; Williams, 2000; Williams and Davidson, 2004). Whitemicas in the central zones of these systems have distinctive,non-phengitic sodic compositions (Herrmann et al., 2001).The recognition of sodic white mica, along with low 634Svalues in pyrite and the presence of pyrophyllite, enables thediscrimination of this type of alteration system from thoseassociated with economic Zn-rich polymetallic VHMSdeposits (i.e. Rosebery and Hellyer). The altered footwallzones of Zn-rich polymetallic VHMS deposits containnormal potassic to slightly phengitic white micas. The factthat the compositional variations in white mica can be simplydetermined by SWIR spectral analysis makes this a practicalmethod for selecting and ranking exploration targets.

Mineral chemistry exploration vectors

Many documented studies have shown spatial variations inmineral chemistry in altered zones around VHMS deposits.Chlorite has received the most attention, particularly in thelast two decades, and there are few deposits for which nodata are available. There has also been significant interest incarbonate and, to a lesser extent, white mica compositions.

Chlorite compositions are typically Mg-rich in theproximal altered zones of VHMS deposits. They commonlyshow systematic distal trends to more Fe-rich compositions.These trends are typically recognisable over several hundredmetres, both laterally and stratigraphically into the footwall,away from the ore. Some examples include the Seneca andCorbet deposits in Canada (Urabe et al., 1983), the Arcticdeposit in Alaska (Schmidt, 1988), and the Thalanga depositin north Queensland (Paulick et al., 2001). However, thereare many cases where the opposite trend exists and Fe-richchlorites occur in proximal altered zones. The Aznacollar andMasa Valverde deposits are two examples in the Iberian pyritebelt (Sanchez-Espana et al., 2000). At the Home deposit,Canada (MacLean and Hoy, 1991), chlorites in proximalchlorite-rich zones are more Fe-rich than in the enclosingsericite + chlorite zone (Fig. 4.13). Similarly, at MattagamiLake (Abitibi belt, Canada) there is a general trend of Feenrichment in chlorites upwards towards the ore position andoutwards from the core of the altered footwall zone (Costaet al., 1983). In northern Turkey, the dacite-hosted depositsof the eastern Black Sea province have altered footwall zonesof Mg chlorite and sericite (Cagatay, 1993). In contrast, thewestern Black Sea ophiolite-hosted pyritic Cu deposits ofthe Kure district are associated with Fe-rich chlorites andtrends of Fe enrichment toward ore. Some deposits exhibitinconsistent patterns of variations in chlorite composition.McLeod (1987) found that Mg chlorites around the MountChalmers deposit (Queensland) have a stratigraphic upwardstrend of Fe enrichment in the footwall and a sharp reversal toMg enrichment in the mineralised zone. Two recent regional-


scale studies by Hannington et al. (2003a, 2003b) showedcontrasting compositional trends in chlorites associatedwith VHMS deposits in the Noranda district, Canada, andKristineberg deposits of the Skellefte district, Sweden. Inthe Noranda district, the moderately Fe-rich compositionsof chlorites (Fe/Fe+Mg 0.4-0.9) associated with sulfidedeposits and surrounding district-scale hydrothermallyaltered zones, contribute to discrimination of prospective andnon-prospective volcanic centres. However, chlorites in theKristineberg district are distinctly Mg rich (Fe/Fe+Mg 0.05-0.5) and show little variation between proximal and distalalteration facies. These studies also found the chloritesassociated with sulfide deposits contained significant Mn (upto 1% MnO) and Zn (up to 0.5% ZnO) suggesting that thesecould be used as proximity indicators in exploration.

McLeod and Stan ton (1984) investigated several easternAustralian VHMS deposits and showed that chlorites insphalerite-rich ores are relatively Mg rich compared tothose in chalcopyrite-rich ores. Furthermore, the Mg/Feratios of chlorites are related to Mg/Fe ratios of co-existingphyllosilicates and the Fe content of co-existing sphalerite.Therefore, zonal compositional variations in chlorite maybe reflected in other alteration mineral species, such aswhite mica, which may be more easily measured by SWIR.Importantly, McLeod and Stan ton (1984) concluded that thecompositions of chlorites and other phyllosilicates had notbeen significantly modified by subsequent greenschist faciesmetamorphism.

Variations in chlorite composition have also been used,with some success, as empirical and thermodynamicallycalculated geothermometers to estimate temperature gradientsin hydrothermal systems above 200°C (e.g. Cathelineauand Nieva, 1985; Walshe, 1986). They are sensitive tore-equilibration and therefore not reliable indicators ofhydrothermal temperatures in subsequently metamorphosedterrains (Green and Taheri, 1992).

FIGURE 4.13 | Aliv-Mg-Fe cation plot showing trend to Fe-rich chlorite withproximity to the Cu-Au VHMS deposit at the Home Mine, Quebec, Canada (afterMacLean and Hoy, 1991). Where Ab = albite, Ep = epidote, Mt = magnetite andSer = sericite.

90 | CHAPTER 4

White micas, commonly referred to as sericite, arenearly ubiquitous in massive sulfide-related hydrothermalalteration systems and they can vary considerably from theideal muscovite formula of K2Al4[Si6Al2O20](OH)4 (Deer etal., 1966). The term phengite refers to white micas in whichFe, Mg and some other cations substitute for Al in octahedralsites and the charge balances are maintained by increased Si/Alratios in tetrahedral sites. Phengitic micas form solid solutionsbetween the end members of muscovite and celadonite:K2(Mg,Fe2+)2(Al,Fe3+)2Si8O20](OH)4. Barium-rich phengiticmicas, in which Ba substitutes for K in inter-layer sites, alsoexist in some sediment-hosted sulfide and VHMS deposits(e.g. Schmidt, 1988; Jiang et al., 1996; Leistel et al., 1998).At low to moderate temperatures, there may be limited Nasubstitution for K, with Na/Na+K ratios up to about 0.2.The sodic muscovites generally have low phengite contents.White micas formed at temperature below 300°C may have asignificant proportion of vacancies in inter-layer sites normallyoccupied by K, as well as phengite-like Fe-Mg substitutionin octahedral sites. These are commonly termed illites; theyform complex solid solutions between three end-members:muscovite, celadonite and pyrophyllite. Yang's (1998) reviewprovides a more detailed description of variations in whitemica compositions.

Although white mica composition has been examinedin a number of massive sulfide-related hydrothermalalteration systems, few studies were systematic enough toevaluate its usefulness as an exploration tool. White micasin the proximal altered zones of the weakly metamorphosedHellyer deposit, western Tasmania, are more phengitic thanthe normal muscovites in distal altered zones (Yang, 1998).At the nearby but slightly more metamorphosed Que Riverdeposit, Offler and Whitford (1992) found considerablesmall-scale compositional variations in mica, even withinsingle samples, due to a complex alteration history.Although the metamorphic phases preserve hydrothermalalteration compositional trends, no convincing vectors wererecognised, possibly because of structural complications.Around the Arctic deposit, Alaska, white micas span almostthe entire compositional range between muscovite andceladonite (Schmidt, 1988). Metamorphic micas outside thehydrothermal altered zones are highly phengitic. Micas in avariety of proximal alteration mineral assemblages are variablyphengitic; some contain up to 0.4 cations of Ba per formulaunit and the least-phengitic types are significantly sodic. It isnot clear whether these variations are systematic enough to beused as broad exploration vectors.

Altered zones in the Iberian pyrite belt also have a confusingvariety of mica composition patterns. Plimer and de Carvalho(1982) found that white micas in altered footwall zonesaround the Salgadinho Cu deposit are phengitic, and appearto show increase in Fe/Fe+Mg ratios towards the mineralisedzone. In contrast, in the Rio Tinto deposit the proximal alteredzones contain muscovite and the distal altered zones (up to2500 m from the deposit) contain micas of more phengiticcomposition (Leistel et al., 1998). The Masa Valverde depositis associated with Ba-rich muscovites and some ore bodies inthe Aljustrel district have extensive halos of sodic white mica(Leistel et al., 1998; Carvalho and Barriga, 2000).

With the possible exception of Ba substitution, most of thecompositional variations in white micas are semi-quantifiable

by SWIR spectral analysis. This has been demonstratedby several recent studies in the Mount Read Volcanics(Herrmann et al., 2001). Figure 4.14 provides an example ofvariations in wavelengths of Al-OH bond-related absorptionfeatures in SWIR spectra of white mica in samples taken atintervals from a single drill hole. These wavelength variationsare directly related to white mica compositional variations.Some alteration systems, particularly those associated withdisseminated Cu-Au deposits and/or kaolinite ± pyrophylliteassemblages, exhibit compositional gradients in white micacompositions that are measurable over a few hundred metres.The background white mica compositions are commonlyvariably phengitic and tend to non-phengitic muscoviteor sodic-muscovite in proximal altered zones (e.g. Hustonand Kamprad, 2000; Herrmann et al., 2001). Figure 4.15illustrates variations in wavelengths of Al-OH absorptionfeatures, related to white mica composition, spatially aroundthe Western Tharsis deposit.

Carbonates are a third group of minerals that canaccommodate compositional variations and are common insome VHMS altered zones. Documentation of carbonatecompositional trends and zonal distributions is fairly sparse.However, it seems that massive-sulfide-related carbonates aretypically Fe-, Mg- or Mn-bearing phases, and backgrounddiagenetic or metamorphic carbonates are commonly calcic.Documented examples include the Hokuroku district inJapan (Shikazono et al., 1998), the Rosebery deposit inwestern Tasmania (Large et al., 2001b), and the South Baydeposit in northwest Ontario (Urabe et al., 1983).

FIGURE 4.14 | Stack of selected SWIR hull quotient spectra of core samplesfrom a diamond-drill hole through the altered zone at the Chester deposit,western Tasmania. Annotations on the left side are depths in metres down thehole. The spectral features are almost entirely attributable to white mica in thealteration mineral assemblages. Note the distinct variation in wavelengths ofthe Al-OH absorption features at around 2200 nm. These indicate that the holeintersected mineral assemblages containing normal potassic muscovite in theupper part, sodic white mica from about 100 to 250 m and muscovite to slightlyphengitic white mica in the lower part.


FIGURE 4.15 | Cross-section of the Western Tharsis deposit (westernTasmania) showing zonation of wavelengths of AI-OH absorption features inSWIR spectra. The background of 2200-2210 nm, corresponding to slightlyphengitic white mica, decreases over a few hundred metres to 2194-2198 nm,attributable to non-phengitic, slightly sodic white mica in the proximal alteredzone associated with disseminated pyrite and chalcopyrite.

Inevitably, there are exceptions, such as the deposits ofthe northern Iberian pyrite belt, which have calcite, ankeriteand dolomite in proximal alteration mineral assemblages(Sanchez-Espana et al., 2000). The Mattabi deposit in theSturgeon Lake area, Canada, is underlain by a funnel shapedsiderite-rich altered zone grading outwards to dolomite,which is widespread on a district scale in the footwall andhanging-wall volcanic rocks.

Hydrothermal carbonates in the Rosebery-Hercules area,western Tasmania, are conspicuously Mn rich, (Khin Zaw andLarge, 1992; Large et al., 2001b). Large et al. (2001b) showedthat Mn-siderite and ankerite carbonates in the footwall ofthe Rosebery deposit increase in Mn content towards ore (Fig.4.16). Magnesium contents of carbonates in altered footwallzones of the South Bay deposit, Canada, increase steadilytowards ore over distances of tens to hundreds of metres(Urabeetal., 1983).

Chlorite, white mica and carbonate all have potential asmineral exploration vectors, at least on a prospect or depositscale. However, the considerable diversity of compositionaltrends in the published data indicate that exploration vectorsneed to be empirically established on a district or depositspecific basis, and are not universally applicable.

FIGURE 4,16 | Downhole plot of drill hole 120R illustrating the distribution ofMn-rich carbonates (kutnahorite and manganosiderite-rhodochrosite) in proximityto K-lens of the Rosebery Pb-Zn VHMS deposit, western Tasmania. Magnesium-carbonates occur in the footwall and in a thin unit of altered pumice brecciaimmediately above the ore lens; carbonates more than 50 m above ore in thehanging wall sequence are Ca rich.

92 I CHAPTER 4

4.3 I STABLE ISOTOPES TABLE 4.2 | Natural abundances of H, C, 0 and S isotopes, and standards incommon use (data from Rollinson, 1993).

Theoretical background

Isotope geochemistry is a diverse and rather specialisedscience. This section aims to provide a bare outline ofaspects that have particular relevance to interpretation ofaltered volcanic rocks. It includes only a brief introductionto theoretical principles, necessary to grasp the applications.We recommend that interested readers supplement this byreferring to other textbooks — Rollinson (1993) provides anexcellent working basis.

Isotopes are distinct atomic forms of elements thathave the same number of protons but different numbersof neutrons in their nuclei. Of the 92 naturally occurringelements, 60 consist of more than one isotope and manyof them have two or more stable isotopes. That means thatthey are non-radioactive, and do not change naturally, ordecay, into other radiogenic elements by emissions of sub-atomic particles from their nuclei. Some natural radiogenicisotopes have important geological uses in geochronology,petrogenesis and metallogenesis, because their rates of decayare constant and measurable. Stable isotopes also have manygeological applications, mainly based on their propertiesof isotopic fractionation. The stable isotopes of the lightelements H, C, O and sulfur have received the most attentionfrom geochemists because they are naturally abundant in thehydrosphere and in crustal rocks, not least in altered rocks.

Informal isotopic notation uses the chemical symbol of theelement preceded by the mass number of the isotope writtenas a superscript. Thus 17O denotes the oxygen isotope with17 nucleons, comprising eight protons and nine neutrons. Asingle isotope, which usually has equal numbers of protonsand neutrons, typically dominates the isotopic compositionof each element (Table 4.2). Therefore isotopic ratios are verysmall numbers (e.g. for the average abundances of oxygenisotopes, 18O /16O = 0.002). To avoid direct comparison ofthese unconvincingly small ratios, stable isotopic proportionsare expressed in parts per thousand (i.e. per mil, %o) relativeto a standard material (i.e. delta form). For example:

Stable isotopes undergo fractionation (or selectivepartitioning into different phases) according to thermodynamicproperties that are related to their differing atomic weightsand consequent ionic bond strengths (Faure, 1986; Rollinson,1993). Fractionation may occur by several physicochemicalprocesses of which the most geologically important areisotopic exchange reactions between phases. The degree offractionation is controlled by physical and chemical factors,which vary according to the elements and fractionationprocesses involved. Thus, O-isotopic fractionation is largelydependent on temperature, whereas S-isotopic fractionationis influenced by temperature, pH, / O 2 , and the activities ofsulfur and other cations involved with sulfate.

Isotopic applications in alteration studies

Isotopic studies of alteration mineral assemblages associatedwith mineralised zones may help to estimate alterationtemperatures and water-rock ratios, interpret fluid origins,discriminate between alteration styles and identify alteredhalos around ore deposits. However, it is worth repeatingOhmoto's (1986) cautionary advice to integrate isotopicstudies with geologic, mineralogic and geochemical data. Hestated: 'there is more than one process, which may producethe same isotopic characteristics (in an ore deposit) andthe same geological process may produce entirely differentisotopic characteristics in different conditions. Therefore,isotopic data alone cannot provide a unique answer to anygeological problem, especially when the data are limited toisotopes of one element.'

Diagenetic and hydro thermal alteration of volcanic rocksinvariably involves hydration reactions, between mineralsand water, and so the amount and isotopic compositionof water are important variables. Apart from geologicvariability, sample preparation, isotopic analytical methodsand calibration of fractionation factors also introducesignificant uncertainties. Experimentally, empirically andthermodynamically determined isotopic fractionation factorsprovide a confusing diversity of choice for use in isotopiccalculations. The Laboratoire de geochimie isotopique atUniversite Laval, Quebec, has a comprehensive compilationof fractionation factors from many published sources andis accessible at <www.ggl.ulaval.ca/personnel/beaudoin/labo>(Beaudoin and Therrien, 1999).

Geothermometers

The temperature dependency of isotope fractionationsbetween mineral pairs forms the basis of isotope geo-

1H 99.9844 Std mean ocean water (SMOW), Vienna-SMOW (V-SMOW) or PDB belemnite.

2D 0.015612C 98.89 PDB belemnite13C 1.11

16O 99.7630 Std mean ocean water (SMOW), Vienna-SMOW (V-SMOW) or PDB belemnite.

17O 0.037518O 0.199532S 95.02 Troilite in Canon Diablo meteorite (CDT)

33S 0.7534S 4.2136S 0.02


thermometry. Provided that the paired minerals formed inequilibrium, that their fractionation factors are known andare significantly different, their original isotopic compositionshave been retained and can be separately determined, then acombination of equations can be solved for temperature offormation.

This approach is useful with O isotopes, because O iscommon to, and abundant in, silicates and other alterationphases, such as carbonates and sulfates. It is also applicableto S-isotopic compositions of minerals in complex sulfideand ore assemblages. Hydrogen isotopes are not generallyreliable as geothermometers because they are readily modifiedby subsequent fluid interactions. Furthermore, the mineralfractionation factors are relatively insensitive to temperatureand are not well calibrated (Ohmoto, 1986).

It is usually difficult to physically separate fine-grainedminerals for isotopic analysis and to petrographically demon-strate equilibrium between the analysed phases. However,close agreement between several temperature estimates of twoor more pairs of minerals in a single assemblage (e.g. quartz +magnetite, muscovite + chlorite and calcite + chlorite) wouldinspire reasonable confidence in their isotopic equilibriumand the calculated temperature.

In fluid-dominated hydrothermal systems, mineral-waterO-isotope fractionation factors can be used to estimate relativetemperatures. Although it is difficult to reliably measurethe isotopic composition of the water from fluid inclusions(Nesbitt, 1996), an assumed value can provide approximateor relative temperature estimates. This approach is used inthe determination of ocean palaeo-temperatures from 618Ovalues of the carbonate shells of marine organisms (Rollinson,1993) and also has applications in mineral exploration (e.g.Miller et al., 2001).

Fluid origins

Natural waters have a broad range of H- and O-isotopic compositions because of fractionation effectsin the hydrosphere, lithosphere and mantle (Fig. 4.17).Consequently, it may be possible to infer the source or sourcesof alteration fluids, and something about their evolution,from their isotopic signature.

Fluid inclusions in hydrothermal minerals may permitdirect measurement, but commonly the fluid compositions arecalculated from isotopic compositions of alteration mineralswith known fractionation characteristics, that are assumed tohave been in equilibrium with the hydrothermal fluid. Isotopiccomposition of a single hydrothermal mineral may constrainthe fluid composition if independent temperature estimates,such as fluid inclusion data, are available. Otherwise, isotopiccompositions of mineral pairs in equilibrium can be used (asoutlined above) to deduce temperature, which can then beapplied in the mineral-water fractionation relationship toestimate fluid-isotopic composition.

Water-rock ratios

Knowledge of water-rock ratios may help to determinethe processes of alteration and interpret the hydrology of

FIGURE 4.17 | 6D- and S18O-isotopic compositions of natural waters (from

Taylor, 1979; Ohmoto, 1986). SMOW is standard mean ocean water with 6D-

and618O values of 0%o.

hydro thermal alteration systems. Water-rock ratios can beestimated from whole-rock O-isotope data or from inferencesof mass transfers and solubilities (e.g. Ohmoto et al., 1983).

Unaltered mafic volcanic rocks have initial 618O valuesin the range 6 to 7.5%o, slightly higher than the mantlevalue of 5.7%o, and unaltered felsic volcanic rocks havevalues up to about 10%o, (Hoefs, 1973). The 618O values ofhydrothermally altered volcanic rocks will differ from initialvalues, depending on the temperature and mineral assemblage(which affect fractionation), the initial isotopic compositionof the water and the quantity of water that reacted with a givenamount of rock. The water-rock ratio is usually expressed inatomic proportions of oxygen.

Taylor (1979) presented the following equations expressingthese relationships in closed and open hydrothermal systems:

closed systems

where the superscripts and subscripts i, f, w and r, respectivelyrefer to initial, final, water and rock.

These equations can be plotted as curves of the typeillustrated in Figure 4.18, which relate 818O / to water-rockratios. Thus, a measured final 518Or can be used to estimatethe amount of water involved in hydrothermal or diageneticalteration, under assumed (or otherwise determined) valuesfor temperature, whole-rock fractionation factors and theinitial isotopic compositions of fluid and rock.

However, as discussed in some detail by Ohmoto (1986)and noted by Green and Taheri (1992), natural geologicsystems are not likely to be simple isothermic, closed or opensystems. Rates of isotopic re-equilibration vary according totemperature, and the isotopic compositions of both rock and

94 | CHAPTER 4

30

FIGURE 4.18 | Curves illustrating relationships between water-rock ratioand final-rock 618O at various equilibration temperatures under parameters of:618Owatef = 0 %o, S18 Cy = 7 %0, fractionation factor Aw' = (2.68 x K W ) - 3.53(plagioclase). Solid and dashed lines represent open and closed systems,respectively. Note that re-equilibration with a small amount of water at lowtemperature can produce a large increase in rock 618O.

fluid change incrementally along the flow path. The final rock5I8O reflects an integrated history of fluid-rock reaction andoffers only broad constraints on temperature and water-rockratio. Green and Taheri (1992) suggested that conditionsof diagenesis might approximate a closed system, whereassubmarine hydrothermal convection is more analogous to anopen system. Water-rock ratios calculated under assumptionsof either closed or open systems are likely to represent theminimum values because of kinetic and incremental factorsaffecting rates of re-equilibration. Natural open systems mayrequire water-rock ratios one or two orders of magnitudegreater to achieve equivalent shifts in the isotopic compositionof the rock (Ohmoto, 1986).

Nevertheless, consideration of water-rock ratios isimportant in evaluation of whole-rock 618O data. This willbe further explained in the following section on isotopicexploration vectors.

Oxygen-isotope exploration vectors

Early isotopic studies (e.g. O'Neil and Silberman, 1974;Taylor, 1974) discovered the link between terrestrial epi-thermal Au-Ag deposits, meteoric-hydrothermal convectionand very broad halos of low 618O in volcanic host rocks. Theseextensive isotopic halos had obvious potential as semi-regionalexploration vectors and stimulated further investigations intovolcanic successions hosting other deposit types.

Among them was the landmark study by Green et al.(1983) on whole-rock O-isotope geochemistry in the hostrocks to VHMS deposits in the Hokuroku district, Japan. Theyfound concentric zonation of whole-rock 618O values aroundthe cluster of Fukuzawa ore bodies ranging from 6.7 + 1.3%o

FIGURE 4.19 | Cross-section illustrating the distribution of whole-rock 518O values (black contours), and altered footwall zones around the Fukuzawa deposits,

Hokuroku district, Japan (modified after Green et al., 1983).

in the proximal sericite + chlorite zone and 11.1 + 2.5%oin the surrounding 1—3 km-wide montmorillonite zone,to 16.9 ± 2.7%o in the outer zeolite zone (Fig. 4.19). The618O anomaly is significantly broader and less variable thanelemental geochemical halos; it extends up to 1 km laterallybeyond the Na2O depletion anomaly and at least 400 m intothe hanging wall above the mineralised zone. The wide extentof the whole-rock 518O anomaly is advantageous for regionalexploration. It has particular application in deformed terrainswhere the original mineral assemblages of hydrothermallyaltered zones has been obscured by subsequent metamorphism,because the hydrothermal whole-rock 618O patterns may stillbe preserved. This is because regional metamorphism typicallyinvolves low water-rock ratios.

The observed whole-rock 618O values are consistent withisotopic exchange between the host rocks and large amountsof seawater (0%o, w/r >1) at different temperatures. Isotopicmodelling, using a plagioclase fractionation factor as anaverage felsic volcanic rock value, showed that high 618Ovalues in the diagenetic zeolite zone could be produced byinteraction with fluid of virtually any source (magmatic, seaor meteoric) at low temperatures and relatively small water-rock ratios. This may be due to the high fractionation factorsbetween silicates and water at temperatures below 100°C.However, low 818O values of the proximal sericite + chloritezone are more consistent with conditions of equilibrium with


either meteoric waters at unrealistically low water-rock ratios(-8%o, <0.2) or seawater at 200-300°C and moderate to largewater-rock ratios. The submarine volcanic environment andimplications of hydrothermal mass transfers favour the latterinterpretation.

Cathles (1983) carried out detailed thermal, geochemicaland isotopic analysis of a hypothetical, but geologicallyrealistic, submarine intrusion-heated convective hydrothermalsystem. His model produced O-isotopic results that wereconsistent with 518O data observed by Green et al. (1983)in the Hokuroku district. It predicts that rocks in the shallowsubstrate become isotopically heavier by reaction withdown-welling seawater at low temperatures. Lower isotopicfractionation, due to increased temperatures at depths greaterthan about 2 km below the seafloor, produces a zone of lowrock 518O. As the convective system evolves, and dependingon permeability and rate of isotopic exchange, the deep-low 818O zone migrates up through the shallow-high 818Oanomaly, to produce a low 518O isotopic anomaly around thevent site (Fig. 4.20).

Chapter 8 summarises some other deposit- and district-scale isotopic studies, which illustrate the possible complexitiesin submarine volcanic successions, but indicate significantpotential for whole-rock O-isotope geochemistry in targetingmineral exploration: potential that has not been widelyapplied outside academic studies.

FIGURE 4.20 | Modelled distribution of changes in whole-rock 618O values due to hydrothermal alteration generated by the convection of fluidaround a subseafloor intrusion (after Cathles, 1983). The intrusion is 1 km wide and 3.25 km deep, and emplaced with its top 1.75 km below theseafloor. Re-equilibration with down-welling, low-temperature seawater produces a shallow zone of higher 818O. Increasing temperatures at depthcreate a sub-horizontal zone of low 818O, which propagates up to the seafloor resulting in the characteristic low 818O surrounded by a halo ofpositive anomalies. Note that the contours represent shifts from the initial rock 818O values, not the actual rock 818O values.

96

I 97

5 | SEAFLOOR- AND BURIAL-RELATEDALTERATION

This chapter discusses the alteration processes and theirproducts (textures, minerals and zones) that occur immediatelyafter deposition and during burial of volcanic facies insubmarine environments. It encompasses the relatively low-temperature processes of hydration, diagenesis and earJyburial metamorphism.

Provided sufficient time, burial-related alteration ultimatelyresults in the lithification of clastic facies in the succession.Oxidation, hydration, dissolution, dehydration, ion exchange,and hydrolysis reactions result in the breakdown of volcanicglass, precipitation of authigenic minerals in pore space, andreplacement of glass and magmatic minerals by new minerals.Alteration mineral assemblages may change over time dueto changing physical and chemical conditions during burial,and may progress to low-pressure, high-temperature regionalmetamorphic assemblages at depth (Coombs et al., 1959).

The recognition and description of burial-related alterationstyles in submarine volcanic successions has implicationsfor exploration and ore genesis studies, because of dramaticchanges in porosity and permeability, which result fromcementation, compaction and dissolution during diagenesis.These changes influence subsequent fluid pathways and thesites of hydrothermal venting and mineralisation. Althoughit has been frequently assumed that compositional changesassociated with diagenetic alteration are limited, they mayinvolve mass changes of up to 9% (Gifkins and Allen, 2001).Diagenetic and burial metamorphic mineral assemblages andthe thickness of altered zones can also be used to determine abasin's thermal history (e.g. Utada, 1991).

5.1 | ALTERATION RELATED TO SEA-FLOOR PROCESSES AND BURIAL

Distinctive weathering and burial-related alteration processesoccur in submarine volcanic successions because of rapidaccumulation rates, and the presence of abundant glass andseawater. Silicate glasses are more susceptible than mineralsto alteration, because they lack well-developed crystalstructures and thus will readily devitrify, dissolve or alter tominerals. Glass fragments are especially prone to alterationbecause of their reactivity and large surface area to volume

ratio. At elevated temperatures, volcanic glass readily altersin the presence of alkaline fluids, but the rate of alteration isreduced under dry conditions or in the presence of pure water(Lofgren, 1970). For example, hydration and devitrificationrates of feJsic voicaniclastic Facies increase one to Eve orders

of magnitude in the presence of seawater (Lofgren, 1970,1971b).

Another important aspect of burial-related alteration involcanic and igneous rocks is that anhydrous primary igneousminerals that have crystallised at high temperatures (e.g.olivine and pyroxene) become unstable and alter to hydrousminerals at lower temperatures. The extent of these retrogradereactions depends on the availability of water and the rockpermeability.

The effects of diagenesis and burial metamorphism onthick, proximal volcanic successions are relatively poorlyunderstood and documented, and detailed studies arealmost exclusively limited to well-sorted, fine-grained felsicvoicaniclastic facies.

The Ocean Drilling Program (ODP) in fore-arc andback-arc basins in the western Pacific region has providedimportant information on the behaviour of volcaniccomponents during early low-temperature alteration andlithification, and the factors controlling the intensity anddepth of diagenetic alteration in Miocene to Recent felsic tointermediate sandstones (e.g. Hein and Scholl, 1978; Taylorand Surdam, 1981; Klein and Lee, 1984; Hay and Guldman,1987; Marsaglia and Tazaki, 1992; Tazaki and Fyfe, 1992;Torres et al., 1995). Studies in mafic volcanic successions havegenerally been limited to seafloor alteration (e.g. Bonatti,1965; Hay and Iijima, 1968a; Honnorez, 1978; Zhou andFyfe, 1989).

Limited work in uplifted and eroded ancient submarinesuccessions provides data on diagenetic and burialmetamorphic minerals, textures and zones that formed atdepths greater than 1 km (e.g. in New Zealand, Coombs,1954; Coombs et al., 1959; in Canada, Kuniyoshi and Liou,1976; Starkey and Frost, 1990; in Australia, Smith, 1969;Smith et al., 1982; Gifkins and Allen, 2001; Gifkins et al., inpress; and in Japan, Hay and Iijima, 1968a; Seki et al., 1969;Iijima and Utada, 1972; Utada, 1991).

Active geothermal regions provide direct measurements oftemperatures, alteration mineral assemblages and pore water

98 | CHAPTER 5

chemistry at relatively shallow depths, less than 2 km (e.g.Coombs et al., 1959; White and Sigvaldason, 1962; Vierecket al., 1982). In addition, experimental work on the alterationof natural and synthetic glasses by modified seawater providesestimates of alteration mineralogy, temperature ranges formineral species, fluid-rock ratios, elemental variations in glass,and variations in fluid chemistry over time. Basalt-seawaterexperiments were performed by: Hajash (1975, 1977),Keene et al. (1976), Seyfried and Bischoff (1977), Mottl andSeyfried (1977), Seyfried et al. (1978), Hajash and Archer(1980), Seyfried and Mottl (1982), and Ghiara et al. (1993).Rhyolite-seawater experiments were conducted by: Ellis andMahon (1964), Sakai et al. (1978), Hajash and Chandler(1981), Shiraki et al. (1987) and Shiraki and Iiyama (1990).

Physical conditions

Early studies assumed that burial-related alteration mineralassemblages and zonation patterns in submarine volcanicsuccessions were controlled by pressure and temperatureconditions. However, it is now believed that the compositionand pressure of intergranular fluids and the composition ofthe primary facies are more important (e.g. Miyashiro andShido, 1970; Surdam, 1973). Differences in the mineralassemblage, intensity, stratigraphic position and sequence ofburial-related altered zones may be explained by variations in:primary rock composition, pore-fluid composition, pore-fluidpressure, geothermal gradient and hence temperature, burialhistory and sediment accumulation rate, interaction time orage, fluid-rock ratio, porosity and permeability, and tectonicsetting (Hay, 1966; Surdam, 1973; Furnes, 1975; Boles andCoombs, 1977; Ratterman and Surdam, 1981; Lee and Klein,1986; Marsaglia and Tazaki, 1992; Ghiara et al., 1993).

Temperatures reached during diagenesis and burialmetamorphism are directly related to the geothermal gradientand in submarine settings these range from 0°C at the seafloorto 250°C at a depth of 2-10 km (Alt and Honnorez, 1984;Morrow and Mcllreath, 1990; Alt, 1995b; Torres et al.,1995). In modern volcanic successions, measured geothermalgradients average 40°C/km, although some are as high as200°C/km (Palmasson et al., 1979; Viereck et al., 1982).High geothermal gradients, associated with magmatism andregions of lithospheric extension such as back-arc basins andrifts, can enhance diagenetic reactions by increasing reactionrates (Boles, 1977; Surdam and Boles, 1979; Torres et al.,1995).

The geothermal gradient may have varied in different partsof a geosyncline or basin; it was likely to have been lowestwhere the sediment was thickest and where sedimentationoccurred most rapidly (Coombs et al., 1959). Taylor et al.(1990) proposed that examples of minimal diagenesis in somebasins may be explained by rapid sediment accumulationrates that did not allow sufficient time for diageneticreactions to occur at depth or for the development of pore-fluid gradients. In addition, magmatism provides heat tothe geothermal system, locally increasing the geothermalgradient and compressing isograds near volcanic centres orlarge intrusions (e.g. Schiffman et al., 1984; Neuhoff et al.,1997). Coeval volcanism, plutonism and rapid burial mayestablish short-lived elevated geothermal gradients in many

submarine volcanic successions. The result is low-pressure,high-temperature diagenesis and metamorphism, and thesuppression of some facies or zones (e.g. pumpellyite-actinolite facies, Patuki ophiolite sequence, New Zealand,Sivell, 1984).

Definitions

The term spilite refers to an altered basalt or dolerite,commonly porphyritic and vesicular, in which Ca-plagioclasehas been albitised and is accompanied by chlorite, calcite,epidote, prehnite or other low-temperature hydrous mineralstypical of greenschist facies (e.g. Cann, 1969; Jolly and Smith,1972; Grapes, 1976). Spilites are interpreted to result fromseawater-basalt interaction during diagenesis on or nearthe seafloor (Coombs, 1974; Turner, 1980). Similarly, theterm keratophyre, although originally restricted to lavas, hasbeen applied to all felsic rocks that contain albite or albite-oligoclase, chlorite, epidote and calcite.

5.2 | HYDRATION

Hydration of glass is typically the first stage of alteration ofvolcanic facies in submarine settings and occurs during low-temperature (<50°C) seafloor weathering and the early stagesof diagenesis. Hydrated glasses (e.g. perlite or palagonite) arevery susceptible to alteration (Lipman, 1965). Hydrationfacilitates subsequent reactions as it increases the alkalinityof the pore fluid, which assists glass dissolution, promotescrystallisation, and may produce perlitic fractures, whichfurther increase porosity and permeability (Lofgren, 1970;Friedman and Long, 1984; Noh and Boles, 1989; Casey andBunker, 1990).

Hydration involves the diffusion of water into solid glass;typically accompanied by a volume change (e.g. reaction R5.1from Noh and Boles, 1989). As water is rapidly absorbed onto glass surfaces, hydration initially affects the outer surfacesof glassy clasts, lavas or shallow intrusions, margins alongfractures in glassy facies, pillow margins, and densely weldedpyroclastic deposits. This is followed by the slow diffusionof water into the glass as hydration proceeds inwards alonghydration fronts defined by strain birefringence, and changesin glass colour and refractive indices (Ross and Smith, 1955;Friedman et al., 1966; Lofgren, 1971a). The rate of diffusionis dependent on composition and temperature and, hence,the extent of alteration is dependent on the time that glasshas been in contact with water (O'Keefe, 1984). Most glasseswill not undergo hydration to great thicknesses unless parallelreactions relax the glass structure allowing water penetration(Casey and Bunker, 1990).

dacitic glass + nH2O -^ perlitic glass + Na+ + (OH)- (R5.1)

Hydration increases the H2O content of glass, reorganisesthe glass structure and may form palagonite or silica gels.Changes in the glass structure may include volume changes,and the formation of evenly spaced tiny bubbles and perliticfractures. Boundaries between glass and hydrated glass are

typically sharp (Peacock, 1926; Lofgren, 1971a; Fisher andSchmincke, 1984).

Palagonite

Palagonite is a dull, resinous, yellow-orange to brown wax-like substance formed from hydrous altered sideromelane(basaltic) glass (Fig. 5.1). It is a mineraloid mixture of relicthydrated glass, nontronite, montmorillonite and other sheetsilicates (Hay and Iijima, 1968b; Honnorez, 1969; Jakobssonand Moore, 1986). Eggleton and Keller (1982) describedpalagonite as a transitional alteration phase between volcanic

SEAFLOOR-AND BURIAL-RELATED ALTERATION | 99

glass and smectite; however, the end product may not alwaysbe smectite.

There are two main varieties: gel-palagonite and fibro-palagonite (Peacock, 1926). Gel-palagonite is isotropic, darkbrown and commonly banded, forming directly adjacent tounaltered glass (Peacock, 1926; Zhou and Fyfe, 1989). Fibro-palagonite is orange-yellow, transparent and birefringent(Zhou and Fyfe, 1989).

Palagonite is widespread in submarine basaltic facies andcommon around the edges of glassy grains in basaltic tuffs, inpillow rinds, along fractures in glass, and in originally glassyvesicle walls (Moore, 1966; Baragar et al., 1977; Friedmanand Long, 1984). Partly altered basaltic pillows typically

A. Gel-palagonite in pillow basaltThe sideromelane groundmass of this plagioclase +augite-phyric basalt is altered to yellow-brown palagoniteadjacent to the vesicle (V). The gel-palagonite exhibitsbanding parallel to the vesicle wall and perpendicularcontraction cracks. Plane polarised light.Sample 153254, Miocene Waitakere Group, Muriwai,Northland region, New Zealand.

B. Palagonite-altered basalt clast rindThe basalt clast in this polymictic conglomerate has athin palagonitised rind. The plagioclase-phyric clast isconcentrically zoned with an unaltered sideromelanecore (C), yellow-brown gel-palagonite altered zone (P)and a brown fibro-palagonite rim (R). The conglomeratematrix includes palagonitised basaltic shards and crystalfragments. Plane polarised light.Sample 131562, Tertiary Macquarie Plains volcanics,Bushy Park, Tasmania.

C. Banded palagoniteThe palagonite-altered rind on this basalt clast displaysfine concentric banding. Plane polarised light.Sample 131562, Tertiary Macquarie Plains volcanics,Bushy Park, Tasmania.

FIGURE 5.1 | Photomicrographs of palagonite.

1 0 0 | CHAPTER 5

have glassy cores successively surrounded by concentric zonesof gel-palagonite and fibro-palagonite (+ smectite), whichare enhanced by bands of fine Fe- and Ti-oxides (Fig. 5.2:Dimroth and Lichtblau, 1979; Zhou and Fyfe, 1989).

Palagonites have variable compositions with 10-20 wt%H2O (Brey and Schmincke, 1980; Eggleton and Keller, 1982;Pichler et al., 1999). Compared with sideromelane, Fe2+ isoxidised, K2O, FeO, TiO2 and Cl may be locally gained, andNa2O, Al2O3, SiO2 and CaO lost (Baragar et al., 1977, 1979;Jakobsson and Moore, 1986; Zhou and Fyfe, 1989). However,whole-rock compositions are not significantly changed, exceptfor H2O. Palagonitisation is typically accompanied by thegrowth of authigenic minerals in open pore spaces (Fig. 5.2)and these commonly account for the elements lost from theglass (e.g. Baragar et al., 1979; Jakobsson and Moore, 1986).

Genesis of palagonite

Zhou and Fyfe (1989) and others have proposed a two-stage solution-precipitation mechanism for palagonitisationof sideromelane based on physical characteristics, chemicalchanges and the presence of etch or dissolution pits at alterationfronts. The first stage is Ti constant: glass is dissolved andgel-palagonite formed. There is a dramatic reduction in theglass volume due to the loss of greater than 60% of the SiO2,A12O3, MgO, CaO and Na2O. The second stage is volumeconstant: gel-palagonite is replaced by fibro-palagonite, andzeolites begin to fill adjacent fractures and vesicles. CaOand Na2O are lost, and K2O and SiO2, Al2O3 and MgO aregained from solution. Titanium and Fe3+ are localised intonearby fracture-filling clay and oxide minerals.

The rate of palagonitisation is temperature dependent anddoubles with every 12°C increase in temperature (Jakobssonand Moore, 1986). Palagonitisation proceeds rapidly attemperatures above 50°C and up to 150°C (Jakobsson,1972, 1978). Jakobsson and Moore (1986) noted thatpalagonitisation of glass varied from less than 40% at 60°C,through 90% at 100°C and was complete at temperaturesabove 120°C. They also found that both gel- and fibro-palagonite occurred below 87°C, but only fibro-palagoniteoccurred above this temperature.

The thickness of palagonite rinds is time and temperaturedependent. Palagonite rinds in pillow basalts systematically

increase in thickness with time and doubles for every 8°Ctemperature increase (Moore, 1966; Jakobsson and Moore,1986).

PerlitePerlite is a textural term referring to networks of fine fracturesor cracks that range from concentric arcuate fracturesenclosing cores of glass (classical perlite; e.g. Fig. 5.3A and B)to long sub-parallel fractures linked by short cross fractures(banded or ladder perlite) (Fig. 3.2C and D: Ross and Smith,1955; Friedman et al., 1966; Allen, 1988). Perlitic fracturesare a common feature of glassy rock fragments, felsic lavas andsynvolcanic sills, and also occur in the glassy rinds of mafic tointermediate lavas.

Felsic perlites typically contain 2-6.5 wt% H2O comparedwith non-hydrated obsidian, which contains a few tenths ofone percent (Ross and Smith, 1955; Noh and Boles, 1989).In addition to gains in H2O, perlites typically gain K2O, andlose Na2O and to a lesser degree CaO and SiO2 (Lipmanet al., 1969; Fisher and Schmincke, 1984; Noh and Boles,1989). Iron is oxidised, volatile components Cl2 and F2 maybe lost, and 8O18 isotope values modified by interactionwith external fluids (Lipman, 1965; Jezek and Noble, 1978;Cerling et al., 1985). These compositional changes are mostintense along the perlitic fractures (Jezek and Noble, 1978;Fisher and Schmincke, 1984).

Genesis of perlite

A debate continues over the origin of perlite and theimportance of hydration (Ross and Smith, 1955; Friedmanand Smith, 1958; Friedman et al., 1966) versus coolingcontraction (Marshall, 1961; Yamagishi and Goto, 1992).The formation of perlite is favoured by hydration of rapidlycooled glass (i.e. glass with a high degree of under cooling)either during cooling or later at low temperatures (Friedmanetal, 1966; Noh and Boles, 1989; Drysdale, 1991). However,it is also possible that perlitic fractures form in response tostrain inherited from rapid cooling contraction, during theconversion of melts to glass, and associated volume changes(Ross and Smith, 1955; Friedman et al., 1966; Davis andMcPhie, 1996).

FIGURE. 5.2 | Sequence of palagonite alteration and zeolite cementation stages in phonolitic glass fragments (after Brey and Schmincke, 1980, in Fisher andSchmincke, 1984). (A) Glassy shards, perhaps with montmorillonite rim cements. (B) Hydration and development of perlitic fractures accompanied by partialdissolution and alteration of glass shards to gel-palagonite. (C) Complete dissolution and alteration of hydrated glass shards to gel-palagonite, accompanied by theprecipitation of zeolites on to glass surfaces. (D) Alteration of gel-palagonite to fibro-palagonite and precipitation of zeolites into open spaces.

SEAFLOOR-AND BURIAL-RELATED ALTERATION | 1 0 1

A. Perlite in thin sectionThe glassy groundmass of this quartz latite exhibitsclassical perlitic fractures comprising intersectingand overlapping arcuate cracks. The perlitic fracturesenclose cores of unaltered and locally oxidised glass.Arcuate glassy false shard textures occur where perliticfractures intersect (arrow). Amygdales have been filledwith zeolites. Plane polarised light.Sample ET7-4, Wereldsend Formation, Pilchard Gorge,Etendeka, Namibia.

B. Perlite in partly altered rhyoliteWell-developed perlitic fractures are abundant in thispartly glassy rhyolite. The perlitic fractures have beenlined with fine-grained, dark green to brown smectites,enhancing the fracture pattern. Perlite cores have beenpartly altered to smectites and zeolites. Amygdales havebeen filled with cristobalite. Plane polarised light.Sample 147582, Miocene Nishikurosawa Formation,Hokuroku Basin, Green Tuff Belt, Odate, Japan.

C. Relict perlite and amygdales in altered rhyoliteIn this diagenetically altered rhyolite, relict perliticfractures are conspicuous where glass adjacent to thefractures has been altered to dark green mixed layersmectite-chlorite. Elsewhere in the pervasively zeolitealtered domains the perlitic fractures have been obscured.The amygdales have been filled with layers of cristobaliteand fibrous chlorite. Plane polarised light.Sample J6-735 m, Miocene Nishikurosawa Formation,Hokuroku Basin, Green Tuff Belt, Odate, Japan.

D. Relict perlite in altered basaltIn this hydro thermally altered jigsaw-fit basaltic breccia,perlitic fractures are only weakly discernable due tomultiple overprinting alteration facies. The pervasivesericite + quartz + pyrite and nodular carbonatealteration facies obscure the perlitic fracture pattern.Plane polarised light.Sample 76833, Cambrian Que-Hellyer Volcanics, westernvolcanosedimentary sequences, Mount Read Volcanics,western Tasmania.

FIGURE 5.3 | Photomicrographs of fresh and altered perlite.

1 0 2 | CHAPTER 5

Alteration of perlite

Perlite commonly undergoes subsequent alteration todiagenetic mineral assemblages that include smectite, Fe-oxides, zeolites, K-rich gel-like glass, low-cristobalite, K-feldspar, chlorite, sericite and carbonate (Noh and Boles,1989). Alteration begins by dissolution of hydrated glassand crystallisation of smectite, carbonate or Fe-oxides alongperlitic fractures (e.g. Noh and Boles, 1989). This commonlyaccentuates the fracture pattern (e.g. Fig. 5.3B). As alterationprogresses, glass dissolution with continued precipitationadvances inwards and the perlitic fractures become diffuse andindistinct (e.g. Fig. 5.3C, D and Allen, 1988). Dissolution ofremaining glassy cores is succeeded by formation of zeolites,such as clinoptilolite or mordenite, or gel-like glass, whichare ultimately replaced by K-feldspar (e.g. Noh and Boles,1989).

5.3 | DIAGENESIS (GLASS TO ZEOLITEFACIES)

Diagenesis encompasses the low-temperature and low-pressure alteration processes that occur during progressiveburial of sediments and rocks. It can be defined as theprocesses (excluding weathering) that change their characterand composition, between the moment of deposition, andthe onset of metamorphism (Larsen and Chilingar, 1979).

Submarine diagenesis involves low-temperature processes,ranging from bottom water temperatures up to crystallisationof unequivocally metamorphic minerals such as laumontite,wairakite, chlorite and pumpellyite (Winkler, 1979; Bohlkeet al., 1980). It is impossible to define a unique pressure andtemperature range that would characterise the transitionbetween diagenesis and metamorphism, because of the greatlycontrasting degrees of mineral stability that characterisedifferent rock types and the wide range of conditions underwhich the common diagenetic minerals crystallise. Generally,diagenesis in submarine settings occurs at pressures of 0.1 to10 MPa (1 bar to 1 kbar) and temperatures ranging from 0 to250°C (Alt and Honnorez, 1984; Morrow and Mcllreath,1990; Alt, 1995b). Temperatures and pore water salinitiesincrease, and seawater-rock ratios decrease with burial depth(Hanor, 1979; Alt-Epping and Smith, 1997).

Submarine diagenesis encompasses compaction,dissolution and leaching of components, precipitation ofnew minerals, and recrystallisation in response to changesin pressure, temperature and chemical conditions in thesubseafloor. New minerals directly replace glass, form mineralovergrowths, fill primary and secondary pore spaces, andform cements, all of which dramatically reduce the porosityand permeability and promote lithification.

With increasing diagenesis, porosity and permeabilitytypically decrease. However, reversals in this trend canoccur during fracturing or if a major component of the rockbecomes under saturated and secondary porosity is formed bydissolution. This can occur where deeply buried sediments areinfiltrated by fresh or brackish ground water, or can be dueto the release of water of crystallisation from clay minerals(Morrow and Mcllreath, 1990).

The process of dissolution involves corrosion or leachingof pre-existing phases (either glass or mineral phases), with orwithout minor replacement by new minerals (Morrow andMcllreath, 1990). It is a complex process involving manydistinct reaction steps and pathways. It can modify glass andmost primary igneous minerals. Dissolution may ultimatelylead to the formation of secondary porosity (e.g. dissolutionvugs), replacement of glass and minerals, and developmentof solution seams or stylolites (Amstutz and Park, 1967;Marsaglia and Tazaki, 1992).

Despite changes in mineral assemblage, many pre-existingtextures (primary volcanic, high-temperature devitrificationand hydration textures) are preserved and sometimes enhancedduring diagenesis. Figure 5.4 shows some examples of texturesin unaltered volcanic rocks, and their diagenetically alteredand in some cases metamorphosed equivalents.

Submarine diagenesis may involve multiple stages orepisodes of diagenesis (Bohlke et al., 1980; Morrow andMcllreath, 1990). Diagenesis of most ancient sedimentarysuccessions involved repeated exposure to diagenetic realmsas they underwent cycles of subsidence and uplift. Generally,however, the imprint of the first stages of diagenesis ispreserved because of the large initial porosity reduction andlithification (Morrow and Mcllreath, 1990).

Diagenetic minerals

There are three main types of minerals typical of seafloorweathering and diagenesis in volcanic successions: layeredsilicates, zeolites and carbonates. Figure 5.5 provides estimatesof their formation temperatures.

Layered silicates

The layered silicates include clay minerals, mixed-layeredminerals, micas, chlorite and prehnite. The common clayminerals in volcanic facies can be divided in to two groups:(1) smectites (e.g. montmorillonite, nontronite andsaponite),and (2) illite group clay minerals (e.g. celadonite, glauconiteand illite).

Smectites are swelling clay minerals that readily exchangeCa and Na cations. They typically result from the alterationof volcanic grains under alkaline conditions where Mg andCa ions are available (Deer et al., 1966). Smectites form rimson glass surfaces, replace both felsic and mafic glass, andpseudomorph glass shards and olivine crystals (Sheppard andGude, 1968; Schmincke and von Rad, 1976; Viereck et al.,1982). Smectites initially forms blebs and web-like arrays onglass surfaces, and become better crystallised as diagenesisproceeds (Hein and Scholl, 1978). The term bentonite refersto felsic tuff that is composed of almost pure smectite (Garyetal., 1974).

In contrast, the illite group are K- and Al-rich mineralsthat typically form in neutral to alkaline conditions from thebreakdown of feldspars and micas (Deer et al., 1966). Theytypically occur as vesicle fill and pseudomorphs of felsic glassshards and pumice (Schmincke and von Rad, 1976; Iijima,1978). Celadonite and glauconite are less common thanillite.

SEAFLOOR-AND BURIAL-RELATED ALTERATION I 1 0 3

A. Flow bandingThis devitrified flow-banded plagioclase-phyric rhyolitecontains alternating dark and light flow bands. The darkbands are dominantly obsidian, whereas the pale bandscontain fine spherulites and lithophysae.Sample NG1, < 140 ka Ngongotaha lava dome, Hendersonsquarry, Rotorua, New Zealand.

B. This diagenetically altered and metamorphosed flow-banded plagioclase-phyric rhyolite contains alternatingorange albite + quartz and grey sericite-rich bands. Inthin section, the orange bands contain relict spherulites,whereas the grey bands are microcrystalline.Sample 147481, Cambrian Central Volcanic Complex,Mount Read Volcanics, Mount Block, western Tasmania.

C. SpherulitesIn thin section, fresh spherulites consist of radial crystalfibres; typically feldspar intergrown with cristobalite,tridymite or clinopyroxene. Many of these spherulitesenclose plagioclase phenocrysts and are separated bysmall cuspate lenses of dark brown obsidian. Planepolarised light.Sample NG4, <14O ka Ngongotaha lava dome, Hendersonsquarry, Rotorua, New Zealand.

D. Recrystallised spherulites in this greenschist faciesrhyolite are composed of albite, quartz and sericite.Fine sericite trails preserve a radial pattern within thespherulites. The boundaries between the spherulites aremarked by concentrations of sericite. Plane polarisedlight.Sample 147528, Cambrian Central Volcanic Complex,Mount Read Volcanics, Mount Black, western Tasmania.

E. Tube pumice clastsThis unaltered, semi-consolidated, dacitic pumicebreccia contains glassy tube pumice clasts and plagioclasecrystals in a matrix of fine glass shards. The pumice \clast pictured here displays a fine fibrous texture, whichmay be preserved during subsequent alteration. Planepolarised light.Sample from the -1 Ma trachydacitic pumice breccias,Efate Pumice Tormation, Vanuatu.

FIGURE 5.4 | Photographs of unaltered and diagenetically altered volcanic textures.

1 0 4 | CHAPTER 5

F. Pumice clasts in this diagenetically-altered andmetamorphosed rhyolitic pumice breccia preserve thefine tube vesicle structure. The originally glassy vesiclewalls have been altered to albite + quartz + hematite, thevesicles have been lined with sericite and filled with albite.Plagioclase crystals in this sample have been completelyreplaced by albite and hematite. Plane polarised light.Sample 133815, Cambrian Hercules Pumice Formation,Central Volcanic Complex, Mount Read Volcanics, Hercules

footwall, western Tasmania.

G. Many tube pumice clasts locally preserve roundvesicles adjacent to phenocrysts. In this diageneticallyaltered pumice breccia, round and tube vesicles adjacentto a cluster of plagioclase phenocrysts have been filledwith mordenite. As a result, the vesicles have retainedtheir shapes during burial compaction. Plane polarisedlight.Sample OH8-369 m, Miocene Onnagawa Formation,Hokuroku Basin, Green Tuff Belt, Odate, Japan.

H. Similarly, this pumice breccia, which has beendiagenetically altered and metamorphosed to greenschistfacies, contains round and tube vesicles adjacent tohematite-altered plagioclase phenocrysts. The vesicles(V) have beeen filled with sericite and albite. Planepolarised light.Sample 147499, Cambrian Kershaw Pumice Formation,Central Volcanic Complex, Mount Read Volcanics, eastHercules, western Tasmania.

I. Palagonitised rinds on clastsThe rim of this basalt clast has been altered to orange-brown palagonite. Palagonite has also formed rimsaround the vesicles in the clast. Plane polarised light.Sample 131562, Tertiary Macquarie Plains volcanics,Bushy Park, Tasmania.

J. The sericite + albite + hematite-altered rim (R) on thisbasalt clast may be the metamorphosed equivalent of apalagonite-altered rind. Plane polarised light.Sample 147572, Cambrian Sterling Valley Volcanics,Central Volcanic Complex, Mount Read Volcanics, SterlingValley, western Tasmania.

FIGURE 5.4 | Photographs of unaltered and diagenetically altered volcanic textures, cont.


Carbonates

FIGURE 5.5 | Temperature estimates for the growth of common diagenetic andburial metamorphic minerals, and palagonite (data from Thompson, 1971; Seki,1972; Merino, 1975; Grapes, 1976; Kastnerand Gieskes, 1976; Seyfried andBischoff, 1979; Bohlke et al., 1980; Munha et al., 1980; Boles, 1982; Viereck etal., 1982; Jakobsson and Moore, 1986; Bish and Aronson, 1993; Ogihara, 1996;Ylagan et al., 1996; Bodon and Cooke, 1998).

Zeolites

Zeolites are hydrous Al-silicates containing Na and Ca(Table 5.1). The most common zeolites in marine settingsare clinoptilolite, mordenite, phillipsite and analcime(Marsaglia and Tazaki, 1992). A variety of fibro-radiated andbladed zeolites fill pore spaces, cement volcaniclastic particlesand replace glass in altered volcanic facies (Miyashiro andShido, 1970; Schmincke and von Rad, 1976). Most zeolitesprecipitate in open space on to smectite or chlorite films oroccur as overgrowths on detrital grains such as plagioclasecrystal fragments (e.g. Schmincke and von Rad, 1976). Otherscrystallise directly from glass via dissolution reactions withsmectite (e.g. Noh and Boles, 1989) and may pseudomorphglass shards (e.g. Walton, 1975).

Diagenetic carbonates are dominantly calcite and dolomite.They typically fill originally open spaces such as vesicles, occuras cements in volcaniclastic facies (e.g. Hay, 1977), as spheroidsor nodules, and as euhedral crystals replacing palagonite (e.g.Dimroth and Lichtblau, 1979), rock fragments, olivine andplagioclase crystals.

Other diagenetic minerals

Other diagenetic minerals include silica phases (e.g. low-cristobalite, opal CT, chert and quartz), Fe-oxides (e.g.hematite), Ti-rich minerals (e.g. leucoxene), anhydrite, pyrite,epidote and feldspars (albite and K-feldspar). These mainlyreplace glass, primary crystal phases and earlier alterationminerals. Silica phases and feldspars also occur as overgrowthson primary plagioclase and quartz crystals (e.g. Noh andBoles, 1989; Tsolis-Katagas and Katagas, 1989).

Diagenetic zones

Diagenetic mineral assemblages commonly show a thickvertical zonation (e.g. Fig. 5.6 and Section 5.5). Diageneticzones have been described by a number of authors in modernand ancient submarine felsic to intermediate volcanicsuccessions (e.g. Iijima, 1974; Walton, 1975; Iijima, 1978;Ratterman and Surdam, 1981; Sheppard et al., 1988;Williams et al., 1989; Utada, 1991; Passaglia et al., 1995;Ogihara, 1996). Sequences of diagenetic zones are between500 m and 6 km thick, with individual altered zones varyingfrom a few metres to several kilometres in thickness. Thisvertical zonation corresponds to progressive mineral reactionsthat occur in response to changes in pore water chemistry andtemperature with depth of burial, and is very similar to burialmetamorphism (Coombs, 1954). Some altered zones may beabsent or combined.

Diagenetic zones in felsic volcanic successions

Diagenetic zones in felsic volcanic successions can be groupedinto four main zones (Table 5.2): (I) partially altered zones,(II) alkali-rich zeolite zones, (III) late-stage zeolite + calcitezones, and (IV) albite zones. At depth Zone IV may pass in toa prehnite + pumpellyite zone, which represents the transitionto greenschist facies metamorphic zones (Iijima, 1974, 1978;Utada, 1991).

Partially altered zones are characterised by silica andclay minerals, they lack zeolites, contain unaltered and partlyaltered glass, and unaltered primary minerals such as plagioclase(Iijima, 1974, 1978). Alteration mineral assemblages aredominated by smectites (commonly montmorillonite) + low-cristobalite or opal-CT (Iijima, 1974, 1978; Walton, 1975;Sheppard et al., 1988; Passaglia et al., 1995). Primary porespaces, such as vesicles, have typically been partially filled withlow-cristobalite, glassy clasts have been coated in thin films ofsmectite, and some originally glassy shards and pumice clastsaltered to smectite. Coherent facies were relatively unaltered.

106 | CHAPTER 5

TABLE 5.1 | Common zeolites and their occurrences in submarine volcanic facies. Zeolite formulas are from Deer et al. (1966).

A Na-rich late stage zeolite, which replaces earlier alkali zeolites in both coherent and clastic

volcanic facies of rhyolitic to basaltic composition (e.g. lijima, 1974; Ratterman and Surdam, 1981;

Torres et al., 1995)

Analcime Na[AISi2O6].H2O

Chabazite Ca[AI2Si4012].6H20 -

thomsonite NaCa2[(AI,Si)5O10]2.6H2O

Clinoptilolite (Na,K)4CaAI6Si30O72.H2O

Heulandite (Ca,Na2)[AI2Si7O18].6H2O

Laumontite Ca[AI2Si4012].4H20

Mordenite (Ca, Na2,K2)[AI2Si10O24].7H2O

Phillipsite (1/2Ca,Na,K)3[AI3Si5O16].6H2O

Wairakite CaAI2Si4012.H20

Restricted to mafic facies, typically replacing palagonite (e.g. Brey and Schmincke, 1980; Dimroth

and Lichtblau, 1979)

Occurs as a cement and replaces glass in felsic volcanic facies (e.g. Noh and Boles, 1989;

Ratterman and Surdam, 1981; Torres et al., 1995)

Occurs as cements in felsic volcaniclastic facies (e.g. Ratterman and Surdam, 1981)

A calcic zeolite, which occurs at depth in originally glassy felsic volcanic facies

Only derived from felsic volcanic facies and commonly coexists with smectite and silica phases

(i.e. opal, quartz, tridymite and cristobalite) (Noh and Boles, 1989; Ratterman and Surdam,

1981; Sheppard et al., 1988; Sheppard and Gude, 1968; Torres et al , 1995; Tsolis-Katagas and

Katagas, 1989; Utada, 1970)

Occurs mainly in basaltic lavas and less commonly in volcaniclastic facies where it replaces

basaltic glass and palagonite (Taylor and Surdam, 1981), it commonly contains inclusions of Fe-

oxyhydroxides and smectites (Brey and Schmincke, 1980)

A common alteration product in basaltic facies at depth in modern geothermal systems (e.g. Boles,

1977; Hay, 1977)

Table 5.2 | Common diagenetic zones and their alteration mineral assemblages for thick submarine volcanic successions.

Zone I: partially altered zone

unaltered glass + smectite (montmorillonite) + low-cristobalite/opal-CT

Zone II: alkali-rich zeolite zone

(a) clinoptilolite + smectite (montmorillonite) + low-cristobalite/opal-CT

(b) Ca-clinoptilolite + mordenite + smectite + K-feldspar ± quartz

Zone III: late stage zeolite + calcite zone

(a) analcime + heulandite + clacite + phyllosilicate minerals (smectite,chlorite, mixed layer minerals) + K-feldspar ± quartz ± pyrite

(b) analcime + laumontite + clacite + phyllosilicate minerals (illite,chlorite, smectite) + K-feldspar + quartz

Zone IV: albite zone

albite + phyllosilicate minerals (prehenite, pumpellyite, chlorite,sericite) + quartz ± K-feldspar ± laumontite ± calcite

Zone I: partially altered zone

unaltered glass + palagonite + smectite + illite + low-cristobalite/adularia + Fe/Mn/Ti oxides + unaltered glass

Zone II: calcic-zeolite zone

phillipsite/chabzite + phyllosilicate minerals (chlorite, smectite,sericite) + Fe/Mn/Ti-oxides ± K-feidspar

Zone III: late-stage zeolite zone

analcime± natrolite (± heulandite ± laumontite) + chlorite + K-feldspar + Fe/Mn/Ti-oxides ± calcite

Zone IV: epidote zone

epidote + chlorite + albite + calcite + sphene ± prehnite

SEAFLOOR-AND BURIAL-RELATED ALTERATION I 107

FIGURE. 5.6 | East-west schematic cross-sections showing the depth distribution of regional diagenetic zones and localhydrothermal zones associated with the Kuroko deposits in the Green Tuff Belt, Japan (after lijima, 1974,1978). (A) OdateBasin, Hokoroku district. (B) Odate to Hanawa Basin, Hokoroku district. (C) Diagenetic zones in the Neogene and Palaeogeneformations of Hokkaido. Traces of cross-sections A and B are shown on the regional map of the Hokuroku Basin (Fig. 5.15).

1 0 8 | CHAPTER 5

These early clay-rich zones are associated with minorinitial gains in K2O and Al2O3, and losses in Na2O and, tolesser degrees, CaO and SiO2 (Noh and Boles, 1989)

Alkali-rich zeolite zones are commonly characterised byassemblages of clinoptilolite + mordenite + smectite (typicallymontmorillonite, saponite or mixed-layer illite/smectite) +low-cristobalite ± quartz ± opal ± K-feldspar. Opal-CT andlow-cristobalite occur in the upper parts of these zones, whereasquartz and K-feldspar occur in the lower parts (Walton, 1975;Sheppard et al., 1988). Plagioclase is rarely altered.

In general, clinoptilolite and mordenite has filled porespaces such as primary vesicles and dissolution voids. Glassyclasts have been coated in montmorillonite or silica rims andcompletely altered to zeolites (e.g. mordenite), low-cristobalite,quartz, clay minerals and K-feldspar (Iijima and Utada, 1971;Iijima, 1974, 1978). The glassy groundmass of lavas and sills,and the cores of blocky clasts may have been partly altered.Pumice-rich facies in these zones contain dark green, variablyflattened phyllosilicate-rich fiamme (e.g. saponite fiammein pumice breccia in the Hokuroku Basin, Iijima, 1974).Tuffaceous mudstones may be rich in montmorillonite, lowcristobalite and quartz.

Alteration to alkali-zeolites and phyllosilicate mineralsresulted in whole-rock gains of MgO and Fe2O3, and lossesof SiO2, Na2O, K2O, and variable changes in CaO (Noh andBoles, 1989; Tsolis-Katagas and Katagas, 1989; Passaglia etal., 1995).

Late-stage zeolite + calcite zones are characterised bymineral assemblages containing analcime and calcite (Iijima,1974). In some successions, two late-stage zeolite + calcitezones have been defined (e.g. Iijima, 1978): (a) analcime +heulandite + calcite ± phyllosilicate minerals ± K-feldspar± quartz ± pyrite, and (b) analcime + laumontite + calcite± chlorite ± illite ± sericite ± K-feldspar. The phyllosilicateminerals are typically smectites, chlorite and mixed-layerminerals such as illite/smectite, saponite/chlorite or swellingchlorite. These mineral assemblages may also contain relictclinoptilolite and/or mordenite (Iijima, 1974; Walton, 1975;Sheppard et al., 1988).

Plagioclase phenocrysts have remained unaltered orhave been analcime ± calcite altered. Analcime has replacedmordenite- or clinoptilolite-altered felsic glass fragmentsand pumice clasts. Saponite, smectite, chlorite and mixed-layer mineral fiamme are typically common in pumice-richrocks. Calcite may occur as euhedral crystals, concretions orveinlets.

Albite zones are commonly characterised by albite+ laumontite ± calcite + prehnite ± chlorite ± sericite ±pumpellyite ± quartz ± K-feldspar (Iijima and Utada, 1971;Iijima, 1974). Plagioclase phenocrysts have been extensivelyalbitised, and albite + laumontite have replaced plagioclasecrystals, originally glassy shards and pumice clasts, and filledpore spaces.

Mass gains in CaO, SiO2, Na2O, Sr and Ba in these zonesare consistent with seafloor albitisation (Boles and Coombs,1977; Boles, 1982).

Diagenetic zones in mafic volcanic successions

Diagenetic mineral assemblages in mafic volcanic successionscontain palagonite, several species of calcic zeolites, Fe/Ti/Mn-oxides and abundant clay minerals of the smectite-chloriteseries typically distributed in four zones (Table 5.2: Baragar etal., 1979; Zhou and Fyfe, 1989; Utada, 1991).

Partially altered zones contain some fresh basaltic glassand have alteration mineral assemblages of palagonite ± Fe/Mn/Ti-oxides (e.g. maghemite or magnetite) + clay minerals(smectites, illites and mixed-layer minerals) ± low-cristobalite.

Compositional changes include major gains of H2O, veryminor gains of K2O, FeO, TiO2 and Cl and losses of Na2O,A12O3, SiO2 and CaO (Baragar et al., 1977, 1979; Jakobssonand Moore, 1986; Zhou and Fyfe, 1989).

Calcic-zeolite zones are characterised by phillipsiteor chabazite ± chlorite + smectite + Fe/Mn/Ti-oxides + K-feldspar.

Whole-rock gains in MgO in these zones are consistentwith the formation of smectite, chlorite and other Mg-silicates during diagenesis (cf. Hajash and Chandler, 1981;Shiraki and Iiyama, 1990).

Late-stage zeolite zones are characterised by analcime ±laumontite ± natrolite ± chabazite + heulandite ± mesolite +chlorite + Fe/Mn/Ti-oxides + K-feldspar.

Epidote zones are characterised by epidote + chlorite +albite + sphene ± calcite ± prehnite.

Genesis of diagenetic minerals and zones

In submarine volcanic facies, dissolution, cementationand lithification begin shortly after deposition (<1 Ma);however, major diagenetic changes develop through a seriesof recognisable stages over tens of millions of years (Marsagliaand Tazaki, 1992). The paragenesis from glass to smectites toalkali zeolites may be explained by a sequence of hydrationand dissolution reactions in most glass bearing rocks (Fig.5.7). Later reactions involve transitions from less stable tomore stable mineral assemblages, such as clinoptilolite toNa-clinoptilolite + mordenite or K-rich gel-like glass to K-feldspar (Noh and Boles, 1989).

There are four stages of clastic diagenesis after initialhydration and oxidation (e.g. Fig. 5.8): (1) formation ofclay mineral rims on glassy surfaces, (2) partial to completedissolution of glass and compaction, (3) precipitation ofauthigenic minerals, especially zeolites and calcite, in openpore spaces, and (4) alteration and replacement of mineralphases (Hay, 1963; Fisher and Schmincke, 1984; Pichler etal., 1999). Stages two and three may overlap.

Stage 1: coating surfaces

The initial stage of diagenesis in volcaniclastic facies ischaracterised by the precipitation of thin rim cements, whichcoat all originally glassy surfaces and some crystal surfaces(Fig. 5.9A and B). Rim cements help to preserve shard andclast outlines during subsequent replacement (e.g. Walton,1975). Rim cements may be accompanied by dissolution ofintermediate to mafic glass, alteration of rhyolitic glass to clay

minerals and minor precipitation of calcite or clinoptilolitecements (Fig. 5.9C and D: Marsaglia and Tazaki, 1992;Torres et al., 1995).

In felsic volcaniclastic facies, the outer walls of most glassshards and vesicles in pumice clasts are lined with thin films ofsmectite, calcite, opal or rarely chlorite (e.g. Henneberger andBrowne, 1988; Sheppard et al., 1988; Noh and Boles, 1989;Tsolis-Katagas and Katagas, 1989; Marsaglia and Tazaki,1992; Torres et al., 1995). Only rarely are fine-grained, glassyfragments such as shards and pumice completely replaced bysmectite. Smectite rims probably precipitate from alkalinefluids during the dissolution of hydrated glass surfaces. Thismay follow the reactions:


and high activities of Mg (Hay, 1978; Hajash and Chandler,1981).

The initial stage of diagenesis in basalts involves thepalagonitisation of basaltic glass. Palagonitisation initiates onglass surfaces, along fractures and around vesicles in a similarway to the thin smectite coating on rhyolitic glass shards (e.g.Zhou and Fyfe, 1989). Associated with palagonitisation is theprecipitation of Fe- and Ti-oxides, which coat all surfaces,thermal contraction cracks, vesicles, and the palagonitisationfront (e.g. Dimroth and Lichtblau, 1979).

Stage 2: dissolution of glass and compaction

perlitic glass + 3.88K+ + 0.65H+ + 15.4H2O -»smectite + 9.5gel-like glass + 4.03Na+ + 0.25Ca2+ +10.55H4SiO4

(Noh and Boles, 1989) (R5.2)

or

rhyolitic glass + Mg2+ + H2ONa-Ca montmorillonite + SiO2 + Na+ + K+ + Fe2+

(R5.3)

The formation of smectite and other Mg-silicates duringrhyolite-seawater interaction does not require significant gainsin alkalis, but is favoured by high ratios of H/Na and K/Ca,

Large-scale dissolution of glass and crystals is typicallyaccompanied and followed by the precipitation of authigenicmineral cements and lithification after a few million years(Marsaglia and Tazaki, 1992; Torres et al., 1995). Thedissolution of glass fragments, olivine and amphiboles occursrapidly at shallow burial depths prior to extensive cementationand lithification (Smith, 1991; Marsaglia and Tazaki, 1992).With increasing depths of burial, dissolution of feldsparmicrolites and the glassy groundmasses of coherent faciesoccurs (Marsaglia and Tazaki, 1992). In contrast, plagioclasecrystal fragments and phenocrysts undergo only minordissolution early in the diagenetic history.

Elements leached from the glass during dissolutionreactions are consumed by the formation of new minerals.

FIGURE. 5.7 | Flow diagrams showing thesuccessive development of alteration mineralassemblages in volcanic glass during diagenesis.(A) Alteration of silicic glass to day minerals,zeolites and silicates (after Hay, 1978; lijima, 1978;Utada, 1991). (B) Alteration of basaltic glass topalagonite, clay minerals, zeolites and oxides(Honnorez, 1978; lijima, 1978; after Brey andSchmincke, 1980; Viereck et al., 1982; Fisher andSchmincke, 1984).

FIGURE 5.8 | Schematic model for the microscopic textural evolution and reduction in porosity in non-welded pumice breccias during diagenesis (after Gifkins,2001). (A) Stage 1: thin films of smectite (green lines) coat original surfaces, such as vesicle walls, crystals, shards and lithic clasts. (B) Stage 2: primary porosityis filled, and originally glassy shards and vesicle walls replaced or partly replaced by zeolites (bronze), clay minerals (green) or carbonates. Zeolite or K-feldsparovergrowths (orange) may develop on plagioclase crystals. (C) Stage 3: glass is dissolved, altered to clay minerals and compacted, producing phyllosilicate-richfiamme. Clays, zeolites and Fe-oxides precipitate synchronously with the dissolution of glass, forming stylolites. After compaction, more stable diagenetic ormetamorphic minerals replace any remaining glass and less stable minerals (i.e. Stage 4).

For example, Na released by the hydration and dissolution ofrhyolitic glass (reactions 5.2 and 5.3) may be consumed bythe precipitation of mordenite in vesicles (e.g. Fig.5.10E) ordissolution vugs (Sheppard et al., 1988).

Dissolution may be accompanied by compaction thatreduces the pore space geometry by rotating grains, deformingsoft grains and crushing grains. This promotes lithification involcaniclastic facies by pressure welding clasts so that theirmargins interpenetrate (Taylor et al., 1990; Marsaglia andTazaki, 1992). Generally, lithification of volcaniclastic faciesis associated with the formation of diagenetic clay mineralor carbonate rims, which act as cohesive binders: cements.However, Marsaglia and Tazaki (1992) in their study ofmodern partially altered sandstones at ODP Site 788 (Japan)suggested that lithification could be related to a combinationof compaction and brown glass dissolution. Sandstones ina transitional zone, between the unlithified and cementedzones, appeared to be lithified as a result of compaction and/or pressure welding with minor cementation by phillipsiteand smectite/chlorite rim cements. The rim cements formedwhere sufficient glass had dissolved to produce favourableconditions for smectite or zeolite precipitation.

Compaction may also bend and flatten clasts, particularlypumice clasts. Gifkins etal. (in press) suggested that compactionduring burial of clay-altered pumice clasts flattened the clastsand modified tube vesicle structures resulting in bedding-parallel phyllosilicate lenses (i.e. fiamme, Fig. 5.9G). Theyalso proposed that stylolites in pumice breccias in the MountRead Volcanics (western Tasmania) and the Green Tuff Belt(Japan) resulted from the dissolution of soluble components,particularly glass, and the precipitation of clays and Fe-oxidesduring compaction (Fig. 5.9H).

Stage 3: filling pore space and cementation

Precipitation of low-cristobalite and zeolites as pore-fillcements follows the early rim cements in both felsic andmafic volcanic facies (e.g. Klein and Lee, 1984; Zhou andFyfe, 1989). Zeolites also fill vesicles and dissolution voids inglass, and directly replace glass, forming shard pseudomorphsor altering the glassy cores of perlite (Dimroth and Lichtblau,1979; Noh and Boles, 1989; Tsolis-Katagas and Katagas,1989; Passaglia et al., 1995). Many of the zeolites fillingvesicles have fibrous radiating textures (e.g. mordenite,Fig. 5.10A and B), which may be partly preserved duringsubsequent metamorphic recrystallisation (e.g. Fig. 5.IOCand D). Dimroth and Lichtblau (1979) described fibro-radial textures defined by a dusting of fine oxides in Archaeanbasaltic hyaloclastite of the Noranda District (Canada), whichsuggest the former presence of fibro-palagonite or zeolites.

The formation of alkali-rich zeolites as pore-fill cementsinvolves hydration and dissolution of glass by saline, alkalinesolutions (Ratterman and Surdam, 1981; Noh and Boles,1989). For example, the formation of clinoptilolite fromhydrated felsic glass in reaction R5-4, which consumes Caand Si released during reaction R5.1:

perlitic glass + 0.1Ca2+ + 0.1H4SiO4 + 0.1 H+ + H2O -»clinoptilolite + 0.1K+ + 0.2Na+ (R5.4)

The formation of rim and pore-fill cements results in alithified rock. Cements reduce the porosity, strengthen thegrain framework and reduce the amount of compaction.The initial 35-40% porosity of a well-sorted sandstone canbe reduced to 15-20% by early clay mineral, carbonate orzeolite cements (Helmold and van de Kamp, 1984). In theclinoptilolite + mordenite zone, Henneberger and Browne(1988) found the porosity of pumice breccias was reducedby half, from 34-47% to 20-50%. Alteration in the quartz +adularia zone further reduced porosity to 4—23%.

SEAFLOOR- AND BURIAL-RELATED ALTERATION I 1 1 1

A. Clay rim cement in pumice brecciaSmectite films on all glass and crystal surfaces recordthe initial stage of diagenesis in this partially alteredrhyolitic pumice breccia. Green-brown smectite hascoated bubble-wall shards, plagioclase and quartz crystalfragments, and lined vesicles. Some originally glassyshards have been completely replaced by smectite;however, larger clasts are still glassy (G).Sample J6-295 m, Miocene Onnagawa Formation,Hokuroku Basin, Green Tuff Belt, Odate, Japan.

B. Clay-lined vesiclesRound and elongate vesicles (V) in this pumice clast arecoated in irregular, fine-grained, pale brown smectitefilms. Small vesicles have been completely filled wihsmectite, larger vesicles are unfilled, and the originallyglassy vesicle walls have been altered to mordenite.Sample FK5B, Miocene Tokiwa Formation, South FossaMagna, Green Tuff Belt, Odawara, Japan.

C. Pore-filling cementsIn this pumice breccia sample, calcite cement binds theunaltered glassy and partly calcite-altered tube pumiceclasts. Plane polarised light.Sample Y2A, Quaternary Yali pumice breccia, Yali Island,eastern Aegean, Greece.

D. In crossed nicols, the glassy pumice clasts are isotropicand the calcite cement, calcite-filled tube vesicles, andaltered shards and pumice clasts are evident.

FIGURE 5.9 | Examples of textures that record the different steps in the evolution of pumice clasts during diagenesis.

1 1 2 I CHAPTER 5

E. Zeolite-filled vesicles in pumiceThe smectite-lined vesicles (V) in this pumice clast havebeen infilled with layered fibrous zeolites: mordeniteand clinoptilolite. Originally glassy shards have beenaltered to smectite and vesicle walls to mordenite +smectite. Fine-grained nodules of analcime overprintedthe mordenite and smectite altered tube pumice clast(P). Plagioclase and quartz crystals are unaltered.Sample OH8-537 m, Miocene Onnagawa Formation,Hokuroku Basin, Green Tuff Belt, Odate, Japan.

F. Clay-altered pumice clastOther pumice ciasts may be completely altered to clayminerals, like this dark green uncompacted smectite-altered pumice. Shards and fine-grained ciasts in thematrix have been altered to smectites (montmorilloniteand saponite) + mordenite.Sample OH8-387 m, Miocene Onnagawa Formation,Hokuroku Basin, Green Tuff Belt, Odate, Japan.

G. Clay-altered and compacted pumiceDuring burial, lithostatic pressure may lead to theflattening of soft clay-altered pumice ciasts. The mixedlayer smectite-chlorite fiamme (F) in this pumiceand lithic breccia roughly define a bedding-parallelcompaction fabric. The fiamme have a fibrous internaltexture, wispy terminations and flame-like shapes.Some fiamme are also plagioclase porphyritic. They areinterpreted to be diagenetically altered and compactedpumice ciasts.Sample 147583, Miocene Nishikurosawa Formation,Hokuroku Basin, Green Tuff Belt, Odate, Japan.

H. Dissolution fabrics in pumice brecciaThe dissolution of glass commonly accompaniescompaction during diagenesis. Solution seams andstylolites, like the one pictured here, are interpretedto record the dissolution of soluble components. Thisstylolite is an anastomosing sutured structure thatconcentrates clay minerals and oxides.Sample FK7, Miocene Wadaira Tuff Member, TokiwaFormation, South Fossa Magna, Green Tuff Belt, Wadaira,Japan.

FIGURE 5.9 | Examples of textures that record the different steps in the evolution of pumice ciasts during diagenesis, cont.


A. Fibrous zeolites in vesicles B. In crossed nicols, radial extinction patterns accentuate theVesicles adjacent to this plagioclase crystal in a pumice clast fibrous nature of the vesicle-filling zeolites.have been lined with smectite and filled with fibrous radiatingmordenite. Plane polarised light.Sample 147580, Miocene Onnagawa Formation, HokurokuBasin, Green Tuff Belt, Odate, Japan.

The vesicles (V) in this altered pumice clast are faintly visible albite-filled vesicles mimic pre-existing fibrous textures.in plane polarised light because they are lined with sericite.Sample 133815, Cambrian Hercules Pumice Formation, CentralVolcanic Complex, Mount Read Volcanics, Hercules footwall,western Tasmania.

E. Fibrous feldspar in perlite coresIn plane polarised light, perlitic fractures are conspicuous inthe groundmass of this altered plagioclase-phyric rhyolite.Sample 147541, Cambrian Kershaw Pumice Formation, CentralVolcanic Complex, Mount Read Volcanics, Murchison Highway,western Tasmania.

F. In crossed nicols, overlapping arcuate perlitic fracturesare defined by concentrations of sericite and radial fibroustextures are preserved in the extinction pattern of the albite +quartz + sericite-altered perlite cores (C).

FIGURE 5.10 | Photomicrographs of relict fibrous textures in vesicles and originally glassy domains in diagenetically altered fades.

1 1 4 I CHAPTER 5

Stage 4: alteration and replacement of mineral phases

Later reactions involve the dissolution and replacementof earlier diagenetic phases, remaining gel-like glass, andmagmatic minerals such as plagioclase (Torres et al., 1995).With time and/or increasing temperature, zeolite assemblagesare especially susceptible to replacement because zeolitecrystallisation is controlled by temperature, fluid pressure,and rock and fluid composition. Transitions from unstablezeolites (e.g. phillipsite, clinoptilolite and heulandite) to morestable phases (e.g. mordenite, analcime, laumontite and K-feldspar) are common. These changes reflect dehydrationreactions (Ratterman and Surdam, 1981; Noh and Boles,1989; Smith, 1991):

2.75clinoptilolite + 0.75Na+ + 3.0H4SiO4 -»Na-clinoptilolite + 2.25mordenite + 0.13K+ +0.27Mg2+ + 0.08H+ + 4.46H2O (R5.5)

Clinoptilolite and mordenite may also react with solutionto form analcime, adularia, quartz and calcite (Iijima, 1974).

mordenite + 2Na+ + CO32" —»

analcime + 6SiO2 + CaCO3 + 5H2O

clinoptilolite + Ca2+ + 2HCO3" —» analcime +adularia + 5SiO2 + CaCO3 + CO2 8H2O

(R5.6)

(R5.7)

Increasing diagenesis may favour the formation of calciczeolites, particularly in mafic facies, because of increased Ca/Naactivity ratios due to albitisation of calcic plagioclase (Utada,1991). In addition, chlorite and epidote may crystallise.Chlorite and epidote have been observed as direct alterationproducts of dacitic to basaltic glasses, and clays at relativelyshallow depths (420 m) in modern geothermal regions(White and Sigvaldason, 1962). The development of epidotedepends on the availability of Fe3+ and is probably controlledby the earlier formation of palagonite. Where palagonite isabsent the reaction of basaltic glass to form chlorite releasesCa, which may be consumed by the precipitation of epidote(Baragar et al., 1979). Chlorite and epidote are also typicalof low-temperature metamorphism and the growth of theseminerals may bridge the boundary between diagenesis andmetamorphism where glass and diagenetic clays are replacedby phyllosilicates.

K-feldspar is common in diagenetically altered originallyglassy volcanic facies. Munhaetal. (1980) suggested that below150°C, Na in glass might be exchanged for K in seawater,resulting in precipitation of K-feldspar and K-smectite.However, K-feldspar has not been reported as a direct productof glass alteration, thus some intermediate phases appear tobe required. Iijima and Hay (1968), Surdam and Sheppard(1978) and Hay and Guldman (1987) recognised that K-feldspar replaced earlier mordenite, analcime, clinoptiloliteand phillipsite. In contrast, Noh and Boles (1989) proposedthat K-feldspar crystallised from silicic glass via a series ofhydration and dissolution reactions, which included anintermediate phase of K-rich gel-like glass.

dacitic glass + nH2O —> perlitic glass + Na+ + OH~ (R5.8)

This is probably followed by either reaction R5.9 orR5.10, which fixes Mg from seawater.

12.5perlitic glass + 3.88K+ + 0.65H+ +15.4H2O -»smectite + 9.5gel-like glass + 4.03Na+ + 0.25Ca2+

10.55H4SiO4

2.79perlitic glass + 0.2Mg2+ + 0.2Fe2+ + 0.32H+ +5.27H2O —» 1.27smectite + gel-like glass +1.0Na+ + 0.12Ca2+ + 3.61H4SiO4

+

(R5.9)

(R5.10)

The K liberated during zeolite forming reactions (e.g.reaction R5.4) is then fixed in K-feldspar formation.

gel-like glass + 0.5K+ + 0.2H+ +0.2H2O1.3K-feldspar + 0.1 Na+ + 0.1 Ca2

0.1 Mg2+ + 0.1Fe2++ 0.8H4SiO4 +

(R5.ll)

Albite is common as a replacement product of plagioclaseand K-feldspar in rhyolitic to basaltic rocks (Munha etal., 1980; Boles, 1982; Torres et al., 1995). Albitisation ofplagioclase occurs by dissolution and replacement (Boles,1982; Shiraki and Iiyama, 1990; Torres et al., 1995). In manycases, albitisation appears to have progressed preferentiallyalong plagioclase grain fractures and cleavage traces,suggesting that fluid films can infiltrate the crystal alonglattice defects and planes of weakness, promoting albitisation(Boles, 1982). Microlites in the glassy groundmass of slowlycooled lavas and sills may serve as nuclei for the crystallisationof albite (Dimroth and Lichtblau, 1979). The replacement ofplagioclase and K-feldspar by albite could reflect the exchangeof K in the rock with Na in seawater at greater depths andtemperatures (105-120°C) (Iijima and Utada, 1972; Merino,1975; Munha et al., 1980; Boles, 1982). The Na and Sirequired for albite crystallisation are supplied by diffusionfrom seawater and earlier diagenetic reactions (Boles andCoombs, 1977; Boles, 1982).

2SiO2 + 0.5H2O + H+ + Na+ + Ca-plagioclasealbite + 0.5Al,Si,Os(OH)4 + Ca2+ (R5.12)

Albite is typically riddled with minute inclusions of clays,sericite and calcite, which form as by-products and consumeCa and Al released during this reaction.


5.4 | REGIONAL METAMORPHISM(ZEOLITE TO AMPHIBOLITEFACIES)

Transition from diagenesis to regionalmetamorphism

Diagenesis progresses gradually to regional metamorphismwith increasing temperature, pressure and, commonly, depthof burial. Many sedimentologists consider that the boundarybetween diagenesis and metamorphism occurs when a rock hasless than 5% interconnected pore space, whereas metamorphicpetrologists tend to define it by mineral assemblages that arenot stable in sedimentary environments (e.g. Coombs, 1954;Blatt et al., 1972; Winkler, 1979; Turner, 1980; Morrowand Mcllreath, 1990; Bevins and Robinson, 1992). Thesedefinitions do not necessarily coincide. Diagenesis andlow-grade metamorphism can both have temperatures andpressures in the 200-300°C and less than 1 kbar range. Infact, no clear distinction exists between the processes, pressureand temperature conditions, fabrics (textures) and mineralassemblages of diagenesis and low-grade metamorphism.

Diagenesis involves hydration of glass, compaction,dissolution, cementation and minor recrystallisation: meta-somatic processes involving minor chemical exchangebetween the host facies and trapped fluid at low temperatures(up to 250°C). In contrast, metamorphism involves mainlyisochemical processes with substantial recrystallisation,but chemical changes that are limited to dehydration anddecarbonation (Fyfe et al., 1958). During progressive burialthere is a stage when metasomatic reactions occur but thetemperatures are generally considered too high for diagenesis;this transitional stage is referred to as burial metamorphism(Coombs, 1954).

The transition from diagenesis to burial metamorphismhas been studied in andesitic to rhyolitic rocks in New Zealand(e.g. Coombs, 1954; Coombs et al., 1959; Boles, 1974; Bolesand Coombs, 1975, 1977), Chile (e.g. Levi, 1970), the USA(e.g. Dickinson, 1962; Sheppard and Gude, 1973) and Japan(e.g. Utada, 1970; Iijima and Utada, 1972; Iijima, 1978), andin basaltic rocks in Canada (e.g. Surdam, 1973; Kuniyoshiand Liou, 1976), Australia (e.g. Smith, 1969; Hellman etal., 1977; Smith et al., 1982) and Iceland (e.g. Viereck et al.,1982).

Burial metamorphismBurial metamorphism was defined by Coombs (1954) to coverprogressive mineral changes that can be directly correlatedwith increases in temperature and burial depth in thicksedimentary or volcanic successions. It is a form of regionalmetamorphism that affects thick sedimentary or volcanicsuccessions in subsiding basins, where the basal parts attainlow-grade metamorphic conditions without the deformationor folding typical of regional metamorphism.

Burial metamorphism, like diagenesis, rarely attainsequilibrium mineral assemblages, and penetrative deformationfabrics are absent. Alteration minerals common to burialmetamorphism in submarine volcanic successions are:zeolites (heulandite, stilbite, laumontite, analcime), prehnite,

pumpellyite, epidote, albite, K-feldspar, phyllosilicate minerals(smectites, chlorite, celadonite, sericite), calcite, siderite,quartz, apatite, sphene, pyrite and Fe-oxides (Surdam, 1973;Boles and Coombs, 1977).

Burial metamorphic facies

A metamorphic facies is defined by a group of metamorphicmineral assemblages occurring in spatially associated rocktypes of diverse chemical composition, which are interpretedto have formed during restricted temperature and pressureconditions (Fig. 5.11).

Low-grade facies typical of burial metamorphism involcanic successions are characterised by hydrous minerals andcarbonates, whereas high-grades facies are typically coarsergrained and contain anhydrous and CO2-poor minerals.The low-grade facies are (Table 5.3): zeolite, prehnite +pumpellyite, pumpellyite + actinolite, lawsonite + albite +chlorite, blueschist, and greenschist facies (Turner, 1980).

Although burial metamorphic facies are widespread,they are commonly heterogenous with patchy and domainaltextures (e.g. epidote metadomains of Smith, 1968, 1974,1977; Smith and Smith, 1976). This domainal style ofalteration may reflect mobilisation and local redistributionof elements on a scale of centimetres to metres (e.g. theloss of Fetotal, Mg, Na and K and gain of Ca from epidotemetadomains balance losses and gains from the enclosingalbite domains, Smith, 1977), rather than the significantaddition of elements.

Burial metamorphic zones

Burial metamorphism is a progressive process that producesa sequence of regionally extensive metamorphic zones (e.g.Figs 3.13 and 5.12). Metamorphic zones are mappable groupsof rocks that have similar metamorphic grade. Adjacentmetamorphic zones, like altered zones, are separated by linesof equal grade (isograds), which are delineated by the firstappearance of an index mineral or minerals within the samerock type or composition.

FIGURE 5.11 | Pressure and temperature diagram showing the fields for

regional metamorphic facies (after Turner, 1980; Bevins and Robinson, 1992).

Boundaries between the fields are gradational. Abbreviations used are PA =

pumpellyite + actinolite and PP = prehnite + pumpellyite.

116 I CHAPTER 5

Table 5.3 | Common burial and regional metamorphic facies and mineral assemblages for submarine volcanic successions (from Coombs, 1954; Humphrisand Thompson, 1978; Turner, 1980; Sivell, 1984; Yardley, 1989).

Zeolite Zeolites (laumonite, analcime, heulandite, stilbite, natrolite, mesolite, wairakite) + mixed-layer clays +

quartz + calcite ± muscovite

Prehnite + pumpellyite Prehnite + pumpellyite ± chlorite + albite + quartz ± epidote ± calcite ± sphene ± rare garnet

Pumpellyite + actinolite Pumpellyite + actinolite + albite + chlorite + sphene + quartz + muscovite + calcite ± lawsonite

Lawsonite + albite + chlorite Lawsonite + albite + chlorite + quartz ± pumpellyite ± epidote ± actinolite ± sphene + muscovite ± calcite

Blueschist Glaucophanic amphibole + albite + actinolite/phengite + quartz ± epidote ± chlorite ± sphene ±

pumpellyite ± stilpnomelane ± calcite

Greenschist Actinolite + epidote + albite ± chlorite ± calcite ± tremolite ± talc + quartz ± sphene ± magnetite ± biotite

Amphibolite Hornblende + plagioclase ± epidote ± garnet ± biotite ± quartz + muscovite + calcite ± sphene ±

magnetite

Drilling in modern geothermal regions has revealedpatterns of low-grade metamorphic zones that can be relateddirectly to temperature and fluid composition at shallowdepths (e.g. Fig. 5.13: Coombs et al., 1959; White andSigvaldason, 1962; Vierecketal., 1982). It is generally assumedthat higher-grade rocks formerly had mineral assemblagestypical of lower grade zones, which were progressively alteredas metamorphism proceeded. Near isograds, index mineralsfrom the lower zones locally overgrow lower grade minerals.For example, in basaltic lavas and tuffs in east Greenland nearthe boundary between two zeolite facies burial metamorphiczones, the analcime and mesolite + scolecite zones, vesiclesare lined with analcime and filled with mesolite (Neuhoff etal., 1997).

The generalised sequence of burial metamorphic facieswith increasing depth is (e.g. Fig. 3.13): zeolite, prehnite +pumpellyite, pumpellyite + actinolite, lawsonite + albite +chlorite, blueschist, and greenschist facies (Coombs, 1954;Coombs et al., 1959; Turner, 1980). Departures from thispattern are common.

Zeolite facies

Coombs (1954) and Coombs et al. (1959) proposed thatregionally extensive zeolite facies in Triassic volcaniclastic rocksof New Zealand bridged the transition between diagenesis andconventional metamorphism. The zeolite facies embraces co-existing assemblages of Ca + Al- and Na + Al-rich zeolites andquartz (Coombs et al., 1959). No single mineral assemblage

is diagnostic of the zeolite facies as mineral assemblages aresensitive to primary rock composition, fluid composition,burial history and geothermal gradient. Typically the zeolitefacies contains heulandite, laumontite, analcime, quartz,albite and smectites + prehnite and pumpellyite (Turner,1980; Vierecketal., 1982).

Fyfe et al. (1958) and Coombs (1954), in studies ofTriassic submarine volcano-sedimentary rocks of TarinagaturaHills (New Zealand), divided the zeolite facies into threemineral assemblages that correlate with increasing depth: (1)heulandite + analcime + quartz ± (montmorillonite + celadonite+ sphene), (2) laumontite + albite + quartz ± chlorite, and(3) quartz + albite + adularia. However, at Hokonui Hills,50 km east of Tarinagatura Hills, Boles and Coombs (1975)found no correlation between mineral assemblage and depthof the zeolite facies rocks. In contrast, Neuhoff et al. (1997)divided the zeolite facies in the Tertiary flood basalts of eastGreenland into five diagenetic and burial metamorphic zonesbased on the index minerals: (1) chabazite + thomsonite, (2)analcime, (3) mesolite + scolecite, (4) heulandite + stilbiteand (5) laumonite (Fig. 5.12).

Genesis

The development of the burial metamorphic zones is complex.In many cases, it is progressive through a number of stagesthat may overlap with diagenesis. However, the transitionto burial metamorphism generally involves dehydration anddecarbonation accompanied by the release of silica, such as in

FIGURE 5.12 | Cross-section in the Borggraven region, east Greenland, showing the regional extent and vertical distribution of burial-

metamorphic zeolite zones (after Neuhoff et al., 1997). The thin discontinuous lines represent the dips of selected lavas in this section.

SEAFLOOR- AND BURIAL-RELATED ALTERATION | 1 1 7

FIGURE 5.13 | Schematic cross-section showing the low-temperature altered

zones and average temperature measurements for the boundaries in the

Wairakei geothermal field, Taupo Volcanic Zone, New Zealand (modified from

Coombs etal . , 1959, after Steiner, 1953; Banwelletal., 1957).

reactions R5.13 and R5.15, below (Coombs, 1954; Boles andCoombs, 1977).

Rocks in the Tarinagatura Hills record two stages ofzeolite facies alteration: an upper heulandite + analcimediagenetic stage and a lower laumonite + albite + quartz burialmetamorphic stage (Coombs, 1954). In the upper part of thesuccession, volcanic glass has been replaced by heulanditeand, less commonly, analcime (Fig. 5.14). These zeolites co-exist with newly crystallised quartz and fine-grained smectite.In the lower part, analcime has been replaced by albite andheulandite by laumonite + quartz. Smectite, chlorite, sericiteand mixed-layer minerals occur and prehnite and pumpellyiteappear as accessory minerals. These stages can be summarisedby the reactions:

heulandite —> laumontite + quartz + 2H2O (R5.13)

analcime + quartz —> albite + H2O (R5.14)

laumontite + calcite —* prehnite + quartz +

H2O + CO2 (R5.15)

Some successions lack textural evidence for shallowdiagenetic and zeolite zones as precursors to higher-grademineral assemblages, suggesting that metamorphism to highergrades is not always progressive (e.g. Neuhoff et al., 1997) orthat early textures are destroyed.

FIGURE 5.14 | Down-hole mineral distributions at Tarinagatura Hills, New

Zealand (after Coombs, 1954).

1 1 8 I CHAPTER 5

5.5 | DIAGENETIC ALTERATION IN THEHOKUROKU BASIN

The Middle Miocene Hokuroku Basin, part of the GreenTuff Belt in northern Honshu, Japan, is the most frequentlycited example of diagenetic zones in a submarine volcanicsuccession that hosts VHMS deposits. Iijima (1974, 1978),Iijima and Utada (1971) and Utada (1991) described fourflat-lying, vertically stacked, zeolite-dominated altered zonesthat grade into clay-rich hydrothermal zones proximal to theKuroko VHMS deposits (Fig. 5.6A and B).

Geological setting

The Hokuroku Basin is a 30 x 30 km submarine basincontaining a 3 to 6 km thick bimodal volcanic succession ofcalc-alkaline rhyolites and tholeiitic basalts with some locallyabundant andesites (Figs 5.6 and 5.15: Dudas et al., 1983;Urabe, 1987). Here, the Nishikurosawa and OnnagawaFormations dominate the stratigraphy.

The lower Nishikurosawa or Hotakizawa Formation is upto 650 m thick and includes intercalated basaltic lavas andbreccias, rhyolitic lavas, and laminated mudstone (Tanimuraet al., 1983). The upper Nishikurosawa Formation is athick (<400 m) succession of rhyolitic lavas, domes andinterbedded pumice-rich facies and mudstone (Fig. 5.16:Ishikawa, 1983; Urabe, 1987; Yamagishi, 1987). The upperNishikurosawa Formation hosts the Kuroko VHMS depositsand is conformably overlain by the Onnagawa Formation(Nakajima, 1988).

The Onnagawa Formation comprises a sequence ofpumice-rich breccia, sandstone, siltstone and black mudstonewith abundant felsic synvolcanic intrusions and local basalticlavas (Fig. 5.16: Ohmoto and Takahashi, 1983; Tanimuraet al., 1983; Nakajima, 1988). The lower pumice breccia isextensive and has been correlated across the Hokuroku Basin(Urabe, 1987).

The volcanic succession is relatively undeformed buthas undergone regional diagenesis and local hydrothermalalteration and mineralisation (Utada, 1970; Tanimura et al.,1983). Generally the stratigraphy has a gentle dip with open,N-S-trending folds (Tanimura et al., 1983).

Noquchi depression

LEGENDFunakawa Formation

Onnagawa Formation

Nishikurosawa Formation

Daijima Formation

Rhyolite

Andesite

Dolerite

Quartz diorite

Other rocks

Drill holes in this study

Major mine

Township

Major fault

Railway

FIGURE 5.15 | Geology of the Hokuroku Basin showing the major lithostratigraphic units and inset the distribution of the Green Tuff Belt in Japan (after Sato, 1974;Tanimura et al., 1983). The trace of Ijima's (1974) cross-section A (in Fig. 5.6) is represented by solid line AA', B by the dashed line BB' and our cross-section (Fig.5.18) by the dotted line BC.


FIGURE 5.16 | Graphic log of drill core OH-8 in the Odate Basin, western Hokuroku Basin (Japan) showing lithology, bedforms, and diagenetic and hydrothermalzones in pumice-rich facies of the Upper Nishikurosawa and Onnagawa Formations.

Alteration facies and zonesThe regional diagenetic zones are, from top to base (Figs5.6 and 5.17): (I) partially altered zone, (II) clinoptilolite +mordenite zone (e.g. data sheets HK1 and 2), (III) analcimezone (e.g. data sheets HK3 and 4), and (IV) laumontite +albitezone (Iijima, 1974, 1978; Hay, 1978; Iijima, 1978).

The partially altered zone is distributed beneath theQuaternary gravels and overlies the clinoptilolite + mordenitezone in the eastern Odate Basin, western Hokuroku Basin(Fig. 5. 6A). This zone has a maximum thickness of 60 mand is characterised by well preserved volcanic textures,the absence of zeolites and the presence of unaltered glassypumice clasts and plagioclase crystals (Iijima, 1974). Rocks inthis zone are pale grey. Glass shards and parts of pumice clasts

have been altered to montmorillonite and vesicles are emptyor have been partly filled with low-cristobalite.

The clinoptilolite + mordenite zone is 160—250 m thickand widely distributed in the shallower part of the Odate Basinin the upper Onnagawa Formation. Regionally it overliesthe analcime zone, but in the east of the Odate Basin it isrepeated below the analcime + heulandite zone (Figs 5.6B and5.17). In the upper part of this zone, some glass shards andpumice clasts are unaltered (e.g. data sheet HK1). Typicallythe surfaces of shards, pumice clasts and vesicles have beencoated in a thin film of smectite or low-cristobalite. Vesicleswere filled sequentially with mordenite and clinoptilolite.Originally glassy shards and clasts have been altered to smectite(montmorillonite, saponite or mixed-layer illite/smectite) ±mordenite. In the lower part of this zone, dark green saponite

120 I CHAPTER 5

FIGURE 5.17 | Schematic cross-section of the westernHokuroku Basin (Japan) showing lithology and alteredzones. The locations of the six data sheets are markedon the section.

or mixed-layer smectite/chlorite and pale green mordenitedomains are common in the pumice-rich rocks (e.g. datasheet HK2). Plagioclase crystals are typically unaltered.

The analcime zone is at least 7 km wide and 150-200 mthick, and occurs in the Upper Nishikurosawa and lowerOnnagawa Formations (below the Ml mudstone). It isapproximately equivalent to the ore position, grading intothe hydrothermal montmorillonite zone and overlying thesericite + chlorite zone (e.g. data sheets HK5 and 6) associatedwith the Kuroko VHMS deposits (Figs 5.6 and 5.17). In theeast of the Odate Basin, an analcime + calcite zone occurswithin the clinoptilolite + mordenite zone (Fig. 5.6A: Iijima,1974). However, it is different from the regional analcimezone, containing disseminated pyrite and calcite concretionsand veinlets, and is considered to result from hydrothermalalteration related to Kuroko mineralisation (Yoshida andUtada, 1968; Iijima, 1974).

The regional analcime zone is characterised by analcime,which has completely replaced some shards, and roundedcrystals of analcime that have replaced clinoptilolite- andmordenite-altered shards and pumice clasts (e.g. data sheetsHK 3 and 4). The internal structure of analcime-altereduncompacted tube pumice clasts is not as well preservedas those in the clinoptilolite + mordenite zone. Plagioclase

crystals are unaltered or have been partly analcime and calcitealtered. Dark green saponite, saponite + chlorite and chloritefiamme are common in this zone. They are interpreted asaltered and compacted pumice clasts (Gifkins et al., in press).Analcime-filled solution seams are common in the lower partof this zone.

The albite + laumontite zone occurs beneath the sericite+ chlorite zone at depth in the Odate Basin (Fig. 5.6). It ischaracterised by albite + laumontite ± calcite ± chlorite +sericite (Iijima and Utada, 1971). Laumontite has replacedplagioclase crystals, originally glassy shards and pumice clastsand filled pore space (Iijima, 1974). Albite has replacedplagioclase crystals.

Genesis of altered zones

The four zeolite zones in the Hokuroku Basin are interpretedto have formed during submarine diagenesis of the mainlyglassy felsic volcanic succession (Iijima and Utada, 1971;Iijima, 1974, 1978; Utada, 1991). The alteration patternis interpreted to have formed beneath the seafloor, whilesedimentation continued and the rocks were progressivelyburied. Ptygmatic folds in near vertical calcite veinlets in


Diagenetic zonesZONE!

I | Partially altered zone lacking zeolitesZONE III I Clinoptilolite + mordenite zoneZONE III

— Analcime + calcite zone

Hydrothermal zones^B Montmorillonite and transitional zonesBi Sericite + chlorite zone with plagioclaseI | Sericite + chlorite zone lacking

plagioclase

f.-' .'•] GravelI ?• I Dacitic lavaI T l Basaltic lava/sill— Massive sulfide orei i Felsic volcanic facies

FIGURE 5.18 | Model for the development and

deformation of the altered zones in the Hokuroku

Basin from the Middle Miocene to Recent

(after lijima, 1974). (A) Late Nishikurosawa

stage (Middle Miocene). (B) Onnagawa stage

with deposition of the M1 mudstone. (C) Late

Funakawa stage (Late Miocene). (D) Recent.

Zone III are considered to result from compaction that tookplace when the volcanic facies were most deeply buried.Thus the diagenetic zones probably formed before the endof deposition of the upper Miocene Funakawa Formation,which overlies the Onnagawa Formation (lijima, 1974).The development of the diagenetic zones can be described infour stages (Fig. 5.18): late Nishikurosawa stage, Onnagawastage, late Funakawa stage and recent (lijima, 1974). Duringthese stages, mineralogical changes progressed with depth assuccessive reactions between volcanic glass and interstitialmodified seawater occurred, originally forming zeolites andthen albite as the temperature and pressure increased (Utada,1991). These mineralogical changes were accompanied bytextural and compositional changes.

During the late Nishikurosawa stage (Fig. 5.18A) in themiddle Miocene, felsic glass was hydrated and began to alterto montmorillonite and low-cristobalite (lijima, 1974). Thisresulted in a shallow, regionally extensive partly altered zone.Submarine hydrothermal activity associated with VHMSmineralisation also commenced (lijima, 1974).

The clinoptilolite + mordenite zone formed during thelate Nishikurosawa and early Onnagawa stage (Fig. 5.18Aand B), when partly altered facies were buried to a depth

of approximately 100 m (lijima, 1974). This increased thealkalinity of the fluid resulting in the dissolution of felsic glassand the precipitation of alkali clinoptilolite and mordenitein vesicles, interstitial voids and dissolution cavities.Simultaneous reactions altered glassy clasts and tube pumicewalls to saponite. Reaction rates were possibly accelerated as aresult of hydrothermal fluids circulating within the successionincreasing the geothermal gradient and concentrations of Naand K in the fluid. By the end of the Onnagawa stage thehydrothermal sericite + chlorite and montmorillonite zonesassociated with the ore deposits had formed (lijima, 1974).

During the late Funakawa stage (Fig. 5.18C), clinoptiloliteand mordenite reacted with solution to form analcime,adularia, quartz and calcite (Ogihara, 1996). As a result, theanalcime + heulandite zone was superimposed on the deeperpart of the clinoptilolite + mordenite zone (lijima, 1974). TheFunakawa stage corresponds to the time of deepest burial andcompaction of altered pumice clasts to form fiamme (lijima,1974). The diagenetic analcime + heulandite zone formedcontemporaneously with hydrothermal zones surroundingthe Kuroko VHMS deposits (lijima, 1978; Ohmoto, 1978).

Since the Funakawa stage, the altered zones have beenmildly deformed and eroded (Fig. 5.18D: lijima, 1974).

1 2 2 | CHAPTER 5

Subtle, patchy mordenite + smectite-chlorite alteration facies HK1

Sample No.

Alteration facies

Alteration zone

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

J6-294

subtle, patchy mordenite + smectite-chlorite

clinoptilolite + mordenite zone

Yoneshiro River

Onnagawa Formation

Green Tuff Belt

pumice breccia

plagioclase + quartz

tube pumice clasts, bubble-wall shards,crystal fragments, non-vesicular volcanicclasts

rhyolite

graded bed

syneruptive, mass-flow emplaced pumicebreccia

partly glassy, mordenite + saponite +montmorillonite + smectite-chlorite +K-feldspar + pyrite

saponite films in vesicles, mordenite ±saponite filled vesicles and pore space,smectite-chlorite fiamme, disseminatedpyrite

patchy

excellent

subtle

early

Weak, pervasive mordenite + smectite alteration facies


HK2

Sample No.

Alteration facies

Alteration zone

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

OH8-369

weak, pervasive mordenite + smectite

clinoptilolite + mordenite zone

Odate city

Onnagawa Formation

Green Tuff Belt

pumice + lithic breccia


tube pumice clasts, bubble-wall shards,crystal fragments, non-vesicular volcanicclasts

rhyolite

graded bed


mordenite + smectite-chlorite + K-feldspar+ calcite + pyrite + glauconite

smectite films in vesicles, mordenite ±smectite filled vesicles and pore space,mordenite-altered glass shards andvesicle walls, smectite-chlorite fiamme,disseminated pyrite, microcrystalline lithicclasts

pervasive

excellent

weak

early

diagenetic

1 2 4 | CHAPTER 5

Subtle, pervasi¥e smectite-chlorite + mordenite + analcime alteration facies HK3

Sample No.

Alteration facies

Alteration zone

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

OH8-511

subtle, pervasive smectite-chlorite +mordenite + analcime

analcime zone

Odate city

Onnagawa Formation

Green Tuff Belt

pumice breccia


tube pumice clasts, crystal fragments

rhyolite

massive


smectite-chlorite + mordenite + analcime +sericite + pyrite

analcime solution seams, smectite-chloritefiamme, mordenite filled vesicles, analcimereplacing mordenite + smectite-alteredpumice clasts

pervasive

good

subtle

early

diagenetic

Weak, pervasive analcime + mordenite alteration facies


HK4

Sample No.

Alteration facies

Alteration zone

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

OH8-537

weak, pervasive analcime + mordenite

analcime zone

Odate city

Onnagawa Formation

Green Tuff Belt

pumice breccia

plagioclase

tube pumice clasts, bubble-wall shards,crystal fragments

rhyolite

graded bed


analcime + mordenite + clinoptilolite +smectite-chlorite + pyrite + sericite

mordenite and clinoptilolite filledvesicles, mordenite and analcime alteredvesicle walls, analcime overgrowths onplagioclase crystal fragments

pervasive

moderate

weak

early

diagenetic

1 2 6 | CHAPTER 5

Strong, pervasive quartz + sericite alteration fades HK5

Sample No.

Alteration facies

Alteration zone

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

OH8-794

strong, pervasive quartz + sericite

sericite + chlorite zone

Odate city


Green Tuff Belt


plagioclase

volcanic clasts

rhyolite

graded bed


quartz + K-feldspar + sericite + chlorite+pyrite

pseudomorphs after plagioclase crystalsand clasts in pervasive crystalline matrix,quartz-filled dissolution vugs

pervasive

poor

strong

early

hydrothermal

moderate, pervasive sericite + chlorite alteration facies


HK6

Sample No.

Alteration facies

Alteration zone

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

HO20-485

moderate, pervasive sericite + chlorite

sericite + chlorite zone

near Fukazawa deposit


Green Tuff Belt


nil

clasts?

rhyolite

graded bed

synemptive, mass-flow emplaced pumicebreccia

chlorite + sericite + pyrite + montmorillonite

dissolution vugs after crystals, sericite +chlorite fiamme, disseminated pyrite

pervasive

poor

moderate

early

hydrothermal

1 2 8 | CHAPTER 5

5.6 | DIAGENETIC ALTERATION IN THEMOUNT READ VOLCANICS

Gifkins (2001) recognised two regionally developedCambrian diagenetic zones (albite zone and epidote zone)within the northern Central Volcanic Complex in theMount Read Volcanics. Although the rocks currently havemineral assemblages consistent with greenschist faciesmetamorphism, diagenetic alteration facies were identifiedbased on combinations of mineral assemblages, overprintingrelationships, textures, distribution, alteration intensity andwhole-rock geochemistry. The original diagenetic mineralassemblages were inferred from local relict textures and bycomparison with younger diagenetically altered volcanicsuccessions (Gifkins and Allen, 2001).

Locally, the diagenetic alteration facies merge withhydrothermal alteration facies in the margins of the Roseberyand Hercules VHMS systems (Allen, 1997). Diagenetic andhydrothermal alteration facies are interpreted to have hadsimilar timing (e.g. Fig. 3.20), and the hydrothermal systemmay have contributed heat and fluid to intensify the diageneticsystem (Gifkins, 2001).

Geological setting

For a detailed description of the geology of the Mount ReadVolcanics refer to Section 1.5. The northern Central VolcanicComplex is exposed in an approximately 30 by 6 km arealocated north and west of the Henty fault, and east of theRosebery fault (Fig. 1.5). It includes three compositionallyand texturally different formations (Fig. 5.19): the SterlingValley Volcanics, the Mount Black Formation, and theKershaw (or Hercules) Pumice Formation (Gifkins, 2001).The Sterling Valley Volcanics (>1.5 km thick) are composedof dacitic to basaltic lavas and sills, and polymictic maficvolcaniclastic facies interpreted as resedimented syneruptivehyaloclastite, autobreccia, pillow lava and scoria. The MountBlack Formation is a laterally extensive (>20 km), thicksuccession (>1.6 km) of mainly feldspar-phyric massive, flow-banded and autobrecciated lavas, domes, cryptodomes andsynvolcanic sills (e.g. data sheets CVC1 CVC5 and CVC6).The Kershaw Pumice Formation, which conformablyoverlies the Mount Black Formation, is a laterally extensive(>16 km), relatively thick (>800 m) succession dominatedby non-welded pumice breccia (e.g. data sheets CVC3 andCVC4), pumice-rich sandstone and shard-rich siltstone, withlesser proportions of pumice-lithic clast-rich breccia andsandstone, and massive, flow-banded and brecciated rhyoliticand dacitic lavas and intrusions (e.g. data sheet CVC2). Theupper part of the Kershaw Pumice Formation and the base ofthe overlying White Spur Formation host the Rosebery andHercules VHMS deposits.

Abundant spherulites, lithophysae, micropoikilitic textureand relict perlite indicate that volcanic rocks in the northernCentral Volcanic Complex were initially partly crystalline andpartly glassy (Gifkins and Allen, 2001).

Alteration facies and zonation

Regionally distributed diagenetic albite and epidote zonesformed before, or were synchronous with, stylolitic S,compaction foliation. The alteration intensity is generallyweak with volcanic textures and albite-altered plagioclasecrystals preserved. Locally the distribution and intensity of thediagenetic alteration facies is patchy, reflecting the complexityof the original volcanic facies (Fig. 5.20).

The albite zone is characterised by pervasive albite + quartz+ sericite (e.g. data sheets CVC2 and CVC5), domainal albite+ quartz + sericite with sericite + hematite ± chlorite (e.g. datasheet CVC3) and pervasive sericite (e.g. data sheet CVC4)alteration facies. It is thick (>2 km) and encompasses theKershaw Pumice Formation and most of the Mount BlackFormation (Fig. 5.21). The albite + quartz + sericite-rich faciesare associated with minor increases in SiO2, CaO, Na2O andtotal mass, and decreases in K2O and A12O3 consistent withseafloor albitisation (cf. Boles and Coombs, 1977; Boles,1982). The sericite + hematite + chlorite alteration facies isassociated with minor increases in K2O and A12O3 consistentwith the conversion of silicic glass to clay minerals (cf. Nohand Boles, 1989; Passaglia et al., 1995). The abundance ofhematite may reflect the oxidation of Fe3+ during alteration ofglass to clays (e.g. Klein and Lee, 1984).

The epidote zone is characterised by pervasive albite +quartz + sericite, pervasive chlorite + sericite, pervasive chlorite+ epidote and domainal chlorite + epidote with albite + quartz+ sericite (e.g. data sheet CVC6) alteration facies. The epidotezone is less extensive than the albite zone and is restricted to theMount Black Formation and Sterling Valley Volcanics at thestratigraphic base of the northern Central Volcanic Complex,adjacent to the Henty fault (Fig. 5.19). In the epidote zone,chlorite + sericite and chlorite + epidote altered felsic rockshave gained MgO, consistent with the formation of smectite,chlorite and other Mg-silicates during diagenesis (cf. Hajashand Chandler, 1981; Shiraki and Iiyama, 1990).

Genesis of alteration facies

The epidote zone occurs in the core of the regional anticlinein the Sterling Valley, suggesting that it is associated with thedeepest stratigraphic level in the northern Central VolcanicComplex: the lower Mount Black Formation and SterlingValley Volcanics (Gifkins, 2001). The change from the albitezone to the epidote zone with stratigraphic depth is consistentwith diagenetic alteration zonation (cf. Iijima, 1974, 1978).Thick (>1 km) diagenetic zones with high-temperaturemineral assemblages (albite + quartz + sericite and chlorite+ epidote) suggest that they developed in response to a high-grade diagenetic alteration system that involved an elevatedgeothermal gradient (cf. Utada, 1991).

Albite + quartz + sericite, sericite + hematite ± chlorite,and sericite alteration facies are the metamorphosedequivalents of diagenetic alteration facies that coated originalsurfaces, filled primary porosity and replaced glass in thenorthern Central Volcanic Complex prior to or synchronouswith diagenetic compaction. Thin films of sericite, carbonateand hematite replaced clays that had coated original glassysurfaces at the onset of diagenesis. Albite + quartz + sericite,


FIGURE 5.19 | Geology of the northern Mount Read Volcanics in western Tasmania, showing the major lithostratigraphic units and altered

zones in the northern Central Volcanic Complex (after Gifkins, 2001). Locations of the six data sheets are marked on the map.

FIGURE 5.21 | Schematic cross-section of the northern Central Volcanic

Complex stratigraphy and altered zones, western Tasmania (after Gifkins, 2001).

FIGURE 5.20 | Detailed cross-section in the Rosebery hanging wall (western Tasmania) showing the complex distribution of volcanic and alteration fades (after

Gifkins and Allen, 2001).


Stage 1: Onset of regional synvolcanic diagenesisThin films of sericite, hematite and calcite coat original surfaces, includingvesicle walls, plagioclase crystals, shards and fractures. These films arethe metamorphic equivalents of low-temperature smectite, palagonite andcalcite rim cements. This early stage probably involved interaction withseawater trapped in the volcanic succession. Modified seawater may havebeen expelled from the succession in response to overburden pressure,and migrated towards the seafloor as diffuse unfocused flow.

Stage 2: Diagenesis and synchronous hydrothermalalteration and mineralisationZeolite or clay mineral cements began to fill primary pore spaces, vesiclesand perlitic fractures. Subsequently, these were extensively replacedby K-feldspar or albite and chlorite. Zeolitisation probably occurred attemperatures between 40 and 100°C. Locally, hydrothermal fluids alteredthe succession. Hydrothermal fluid flow was unfocussed and in placesponded beneath the coherent facies of sills and lavas. The Rosebery andHercules VHMS deposits and their altered halos are interpreted to haveformed during this stage at temperatures greater than 300°C.

Stage 3: Continuing diagenetic alteration andcompaction synchronous with deposition of the WhiteSpur FormationDissolution and alteration of glass to clays, sericite and chlorite occurredsynchronous with compaction. Replacement of earlier zeolites by K-feldspar occurred below 150°C, albitisation of plagioclase phenocrysts andalbite replacement of K-feldspar occurred at temperatures between 100and 190°C. Large volumes of fluid were probably displaced as a result ofcompaction under the weight of the accumulating White Spur Formation.Rapid and variable sedimentation rates may have over-pressured thepore fluid, promoting lateral fluid flow along permeable layers. Weakhanging wall alteration developed during continued hydrothermal alterationassociated with the formation of the Rosebery deposit.

Stage 4: Transition between diagenesis and regionalmetamorphismMore stable, higher-temperature mineral assemblages replaced remainingglass, phenocrysts and early alteration minerals. Chlorite + epidotealteration facies developed at depth in both mafic and felsic volcanicfacies: probably at high (>200°C) temperatures.

Stage 5: Devonian metamorphism and deformationGreenschist facies mineral assemblages and tectonic fabrics overprinteddiagenetic and hydrothermal alteration facies. Deformation modified pre-existing volcanic and alteration textures and produced folds, faults andshear zones. The distribution of syn-S2 alteration facies suggests thatmetamorphic fluid migration was restricted to regional structures suchas faults and shear zones. Mineral assemblages in intermediate andmafic rocks in the Mount Read Volcanics indicate that the peak regionalmetamorphic temperature was between 370 and 450°C.

FIGURE 5.22 | Model for the post-depositional evolution of the northern Central Volcanic Complex, western Tasmania (after Gifkins, 2001). Schematic cross-

sections are not to scale.

1 3 2 | CHAPTER 5

chlorite + sericite and sericite + hematite + chlorite replacedzeolites and clays that filled pore space and altered glass, priorto and synchronous with diagenetic compaction. In pumice-rich facies, a bedding-parallel stylolitic foliation reflectsthe dissolution of glass during compaction and fiamme areinterpreted as diagenetically altered and flattened pumiceclasts (Gifkins et al., in press). Diagenetic alteration involvedsignificant mineralogical and textural changes but only minorchanges in composition consistent with the interaction ofrhyolitic and basaltic glass with seawater during burial.

The chlorite + epidote alteration mineral assemblage maybe transitional between diagenesis and burial metamorphism.It developed after lithification and compaction but pre-datedregional deformation associated with peak metamorphism.The chlorite + epidote facies replaced earlier clay or chlorite+ sericite-rich facies and filled any remaining pore space.Negligible absolute and total mass changes associated withchlorite + epidote alteration suggest that it grew in responseto increasing temperature with increasing depth of burial latein the diagenetic history (Gifkins, 2001).

Mineral assemblages in these diagenetic zones reflect thereaction of glass with interstitial fluid at elevated temperatures.The albite zone probably formed at temperatures between100 and 190°C (cf. Iijima and Utada, 1971; Thompson,1971; Merino, 1975; Munha et al., 1980; Boles, 1982).The epidote zone is characterised by chlorite + epidote,chlorite + sericite and albite + quartz + sericite indicatingformation at temperatures of at least 200°C (cf. Seki, 1972;Kristmannsdottir, 1976).

The regional diagenetic alteration and metamorphism ofthe northern Central Volcanic Complex can be described infive successive stages (Fig. 5.22): (1) the onset of diagenesis;(2) formation of diagenetic cements, and synchronoushydrothermal alteration and mineralisation; (3) diageneticalteration and compaction synchronous with emplacementof the White Spur Formation; (4) replacement of earlydiagenetic minerals and remaining glass by more stable mineralassemblages; and (5) regional Devonian metamorphism anddeformation.

SEAFLOOR-AND BURIAL-RELATED ALTERATION | 133

Subtle, pervasive albite + quartz + chlorite alteration faciesLeast-altered rhyoiite

CVC1

Sample no.

Alteration facies

Alteration zone

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

133921

subtle, pervasive albite + quartz + chlorite

albite zone

Mount Black

Mount Black Formation


massive, plagioclase-phyric rhyolite

plagioclase

porphyritic, micropoikilitic

rhyolite

massive

coherent facies

albite + quartz + sericite + chlorite +hematitealbite + quartz ± sericite ± chloritepseudomorphs of plagioclase,micropoikilitic albite + quartz, interstitialchlorite, disseminated hematite

pervasive

excellent

subtle

pre-S2

diagenetic

Hand specimen photograph

GeochemistrySiO2 74.58 K2O 4.34 Cu 2 Al 56TiO2 0.27 P2O5 0.03 Pb 3 CCPI 22AI2O3 13.85 S <0.01 Zn 17 Ti/Zr 5.98Fe2O3 2.08 Total 100.32 Th 22MnO 0.01 Zr 270MgO 0.38 Rb 136 Nb 17CaO 0.13 Sr 96 Y 41Na2O 3.54 Ba 988

Photomicrograph (ppl)

1 3 4 | CHAPTER 5

Weak, pervasive albite + quartz * sericite alteration facies CVC2

Sample no.

Alteration facies

Alteration zone

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

147407

weak, pervasive albite + quartz + sericite

albite zone

120R-438.5 m

Kershaw Pumice Formation


jigsaw fit, monomictic, plagioclase-phyricrhyolite breccia

plagiociase

porphyritic, perlitic fractures, jigsaw fitclasts

rhyolite

massive

in situ hyaloclastite

albite + quartz + sericite > chlorite + pyrite> calcite

albite ± calcite pseudomorphs ofplagiociase, microcrystaiiine groundmass,calcite veins, chlorite filled perlitic fractures

pervasive

moderate

weak

pre-S2

diagenetic

Hand specimen photograph Photomicrograph (xn)

GeochemistrySiO2 74.01 K2O 1.82 Cu 2 Al 26TiO2 0.23 P2O5 0.03 Pb 4 CCPI 29AI2O3 12.21 S 0.01 Zn 19 Ti/Zr 5.31Fe2O3 2.17 Total 100.61 Th 12MnO 0.09 Zr 258MgO 0.49 Rb 76 Nb 16CaO 2.41 Sr 113 Y 36Na2O 4.07 Ba 513


Moderate, domainal albite + quartz + sericite with sencite + hematite ± chloritealteration facies

CVC3

Sample no.

Alteration facies

Alteration zone

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

147410

moderate, domainal albite + quartz +sencitealbite zone

120R-524.5m



graded, plagioclase-phyric pumice breccia

plagioclase

tube pumice clasts, fiamme, plagioclasecrystal fragments, blocky rhyolite clastsrhyolite

normally graded

syn-eruptive, mass-flow-emplaced pumicebrecciaalbite + quartz + sericite + chlorite +hematite + calcitesericite fiamme, hematite styioiites, albiteveins, recrystallised albite + quartz +sericite pumice clasts and matrix, albite +sericite + calcite altered plagioclasedomainal

poor

moderate

pre-S2

diagenetic

GeochemistrySiO2

TiO2

AI2O3

Fe2O3

MnOMgOCaO

Na2O

76.080.19

10.661.67

0.090.472.39

4.71

K2O

P 2 O 5

STotal

RbSr

Ba

0.880.030.01

99.93

34144

280

CuPbZnTh

ZrNbY

14

2710

2101337

AlCCPiTi/Zr

1626

5.43


1 3 6 | CHAPTER 5

Weak, pervasive sericite alteration facies CVC4

Sample no.

Alteration facies

Alteration zone

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

147552

weak, pervasive sericite

albite zone

Pieman Road



massive, plagioclase-phyric pumicebrecciaplagioclase

tube pumice ciasts, bubble wall shards,plagioclase crystal fragments, fiammerhyolitenormally graded

syn-eruptive, mass-flow-emplaced pumicebrecciasericite + albite + calcite + chlorite +hematitesericite fiamme, hematite stylolites,disseminated calcite rhombs, albite +sericite altered pumice ciasts and shards

pervasive

good

weak

pre-S2

diagenetic


GeochemistryGeochemistrySiO2 70.91 K2O 3.16 Cu 4 Al 48TiO2 0.31 P2O5 0.07 Pb 2 CCPI 36AI2O3 14.08 S 0.01 Zn 48 Ti/Zr 7.41Fe2O3 2.78 Total 99.75 ThMnO 0.07 Zr 251MgO 0.77 Rb 124 Nb 13CaO 1.66 Sr 87 Y 28Na2O 2.68 Ba 786

Photomicrograph (ppl)


I

Subtle, pervasive aibite + quartz + chlorite alteration faciesLeast-altered dacite

CVC5

Sample no.

Alteration facies

Alteration zone

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

147435

subtle, pervasive aibite + quartz + chlorite

aibite zone

MBD4-18.4m



massive, plagioclase + hornblende-phyricdacite

plagioclase, hornblende

porphyritic, glomeroporphyritic clusters,micropoikilitic

dacite

massive

coherent facies

aibite + quartz + chlorite + epidote

aibite + quartz micropoikilitic groundmasswith interstitial chlorite + epidote, aibitepseudomorphs of plagioclase, epidote +chlorite altered hornblende

pervasive

excellent

subtle

pre-S2

diagenetic

GeochemistrySiO2 67.53 K2O 3.95 Cu 4 Al 47TiO2 0.52 P2O5 0.13 Pb 4 CCPI 41AI2O3 14.51 S 0.01 Zn 51 Ti/Zr 14.48Fe2O3 4.37 Total 99.51 Th 15MnO 0.06 Zr 216MgO 1.3 Rb 102 Nb 12CaO 2.38 Sr 242 Y 34Na,0 3.56 Ba 958

Hand specimen photograph Photomicrograph (ppl)

1 3 8 | CHAPTER 5

Moderate, domainal chlorite + epidote alteration facies CVC6

Sample no.

Alteration facies

Alteration zone

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

147557

moderate, domainal chlorite + epidote

epidote zone

Pieman Road



jigsaw fit, monomictic plagioclase +homblende-phyric dacite brecciaplagioclase, hornblende

glomeroporphyritic, perlitic fractures,jigsaw-fit clastsdacite

massive


albite + quartz + chlorite + epidote

microcrystalline groundmass with domainalalbite + quartz and chlorite + epidotefacies, plagioclase phenocrysts albite orchlorite ± epidote altered, hornblendealtered to chlorite + epidotedomainal

good

weak

pre-S2

diagenetic

GeochemistryGeochemistrySiO2 67.93 K2O 2.05 Cu 10 Al 31TiO2 0.59 P2O5 0.13 Pb 3 CCPI 44AI2O3 14.32 S 0.01 Zn 28 Ti/Zr 17.68Fe2O3 4.57 Total 99.65 ThMnO 0.06 Zr 201MgO 1.33 Rb 41 Nb 12CaO 2.67 Sr 151 Y 31Na2O 4.97 Ba 826


139

6 | SYNVOLCANIC INTRUSION-RELATEDALTERATION

The spatial and genetic associations between intrusionsand altered zones are widely appreciated in porphyry andepithermal districts (Lowell and Guilbert, 1970; Titley, 1982;Henley and Brown, 1985). Similar relationships also existin VHMS districts, where synsedimentary or synvolcanicintrusions are commonly altered and surrounded by halos ofaltered rocks. In some VHMS districts (e.g. Snow Lake andSturgeon Lake, Canada), there are spatial associations betweensynvolcanic intrusions and broad-scale, semi-conformablealtered zones and clusters of VHMS deposits in the overlyingsuccessions (Spooner and Fyfe, 1973; Campbell et al., 1981;Gibson and Watkinson, 1990; Galley, 1993; Hannington etal., 2003a). It has been suggested that synvolcanic intrusionswere heat sources (Spooner and Fyfe, 1973; Ohmoto andRye, 1974; Solomon, 1976; Cathles, 1977; Franklin et al.,1981; Polya et al., 1986; Galley, 1993; Large et al., 1996),and perhaps also volatile and metal sources (Urabe and Sato,1978; Stanton, 1990; Yang and Scott, 1996; Hannington etal., 1999) for subseafloor hydrothermal systems that formedaltered zones and VHMS deposits.

Synvolcanic intrusive sills, cryptodomes, dykes and sub-volcanic plutons are volumetrically important in submarinevolcanic successions (Polya et al., 1986; McPhie and Allen,

1992; Doyle and Huston, 1999; Galley, 2003). They may becomposite intrusions of variable volumes up to 1000 km3,typically emplaced at depths up to 4 km below the seafloor(Nielsen et al., 1981; Galley, 2003; Whalen et al., 2004).Intrusions and intrusion-related altered zones that significantlypost-date volcanism are also common in ancient submarinevolcanic successions; however, they are not the focus of thischapter.

Alteration can occur within intrusions (deuteric andlocal hydrothermal alteration), locally in the immediatehost rocks (contact alteration) or regionally in the hostsuccession (regional hydrothermal alteration) (Fig. 6.1). Thischapter describes the role of intrusions in generating regionalhydrothermal systems, regional hydrothermally altered zones,altered zones within intrusions and contact altered zonesaround both small-volume, near-seafloor and larger, deeperintrusions in submarine volcanic successions. The final sectionpresents a case study of contact altered zones associated withthe Darwin Granite in the southern Mount Read Volcanics,western Tasmania. The recognition of altered zones relatedto synvolcanic intrusions can provide insights into fluid-flowand thermal histories of VHMS districts, and thereby assistmineral exploration.

FIGURE 6.1 | A cartoon of the variety of altered zones associated with synvolcanic intrusions. (A) A deuteric altered zone withinthe top of a large volume intrusion. (B) A fracture-controlled hydrothermally altered zone at the margins of an intrusion and in thesurrounding host rocks. (C) Contact-altered zones around synvolcanic sills emplaced into unconsolidated sediment immediatelybelow the seafloor. (D) Concentric contact-altered zones around a large volume intrusion emplaced at depth. (E) Regionalhydrothermally altered zones related to emplacement of a subvolcanic pluton. (F) Afootwall alteration pipe beneath a VHMS deposit.

140 I CHAPTER 6

6.1 | THE ROLE OF INTRUSIONS INGENERATING HYDROTHERMALSYSTEMS

The most active hydrothermal systems are those related tomagma-induced thermal anomalies (Alt, 1999; Butterfleld,2000). The magma chamber provides heat to overlyingstrata and active volcanism contributes heat from its eruptiveproducts, intrusions and feeder dykes. The transfer of heat andmass away from the intrusion may occur by either conductiononly, or conduction and infiltration. Conduction generallyinvolves only minor diffusion of elements, although Weaver etal. (1990) suggested that at near solidus temperatures vapour-phase expulsion may produce local mineral and chemicalvariations (loss of Na, halogens and REE) in volcanic glass.In contrast, conduction accompanied by infiltration andcirculation of hot fluid can remove heat from the magmaticsystem much faster than conduction alone, and effectivelytransport elements considerable distances, up to hundreds ofkilometres, through the succession.

Thermal metamorphism related to the shallowemplacementof synvolcanic intrusions in dry successions typically resultsin limited alteration with little or no mass transfer. In rarecases, magmatic fluids exsolved from the crystallising magmahydrothermally alter dry host facies. Vapour-phase expulsionof some elemental species as complexes (e.g. fluoride,chloride, hydroxide, sulfide and carbon dioxide) may resultin minor losses as glassy rocks devitrify, and glassy clasts maybe welded by elevated temperatures in the contact zones (e.g.Christiansen and Lipman, 1966).

The effects of intrusions emplaced into water-saturatedsuccessions are very different because water mobilises heatand soluble elements. Trapped seawater in submarine volcanicsuccessions is heated by intrusions, initiating convection andmetasomatic alteration in the overlying succession. Thus,almost all intrusion-related alteration in submarine volcanicsuccessions involves some degree of metasomatism bymagmatic fluid, modified seawater, or both.

Subseafloor regional hydrothermal systems

Studies of the petrology, geochemistry and oxygen isotopesof hydrothermally altered volcanic and plutonic rocksfrom ophiolite complexes provide insight into subseafloorhydrothermal systems, fluid generation and circulation, and

the role of intrusions (e.g. Lydon and Jamieson, 1984; Altet al., 1986; Gillis and Robinson, 1990; Bettison-Varga etal., 1992; Kelley et al., 1992). The convection cell model forhydrothermal systems and the formation of VHMS depositsis based on observations from VHMS deposits and the upperpart of the Cretaceous Troodos Massif in Cyprus, wherehydrothermal convection was driven by emplacement of late,high-level gabbro stocks into the fractured and permeablecrust (e.g. Spooner et al., 1974; Lydon and Jamieson, 1984;Bettison-Varga et al., 1992). This model involves the circulationof seawater in approximately 10 km diameter cells to depthsof 3-5 km within the crust (Fig. 6.2). Initially, increasedtemperatures in the host succession drive dehydration anddecarbonation reactions, and fluids migrate away from theintrusion. Buoyant heated connate seawater rises through thepermeable volcanic succession, drawing down cold seawater,which is heated as it descends. In this way, magma drivesconvective circulation of seawater between the seafloor andthe intrusion (Norton, 1984; de-Ronde et al., 1994; Galley,2003). Fluid flow is focused along joints, fractures and faultsformed during extension or in response to intrusive pressures(Bettison-Varga et al., 1992). Alternatively, the multi-tieredconvection model involves a high-temperature (450-700°C)cell, which circulates recycled modified seawater in plutonicrocks at depth, overlain by a low-temperature (350-400°C)cell (Gregory and Taylor, 1981; Norton et al., 1984; Alabasterand Pearce, 1985; Kelley et al., 1992).

Submarine hydrothermal systems comprise three parts:a down-flow or recharge zone; a high-temperature reactionzone; and an up-flow or discharge zone (Fig. 6.3: Spoonerand Fyfe, 1973; Alt, 1999). The locations of the recharge anddischarge zones are commonly controlled by faults (Schardt etal., in press). Seawater percolates down through the rechargezone, and is slowly heated and chemically modified by low-temperature reactions (White, 1970; Gibson et al., 1983;Galley, 1993; Alt, 1999). The reaction zone is a porousreservoir near the heat source where heated seawater reactswith the host rocks, exchanging some elements (Norton,1984; de-Ronde et al., 1994; von Damm, 1995; Butterfield,2000; Schardt etal., in press). Hot buoyant hydrothermal fluid(modified seawater) ascends rapidly to the seafloor throughthe discharge zone, which is characterised by cooling of thefluid, alteration of the host rock, and mineral precipitation(Skirrow and Franklin, 1994; Schardt et al., in press). Therising hydrothermal fluid cools by adiabatic decompression,conductive heat loss, and mixing with cold seawater in theshallow subsurface (Mottl, 1983; Butterfield, 2000). In well-

FIGURE 6.2 | Simple convection cell model for the genesis ofthe Cyprus VHMS deposits (modified after Heaton and Sheppard,1977, and Spooner, 1977, in Lydon and Jamieson, 1984).

established hydrothermal systems, the discharge zone may befocused, intensely altered and veined. Surface discharge ontothe seafloor may produce high-temperature (150—350°C)features such as black smokers (Goodfellow and Franklin,1993; Rona et al., 1993). The temperature of the dischargingfluids on the seafloor initially increases, and then graduallydecreases to ambient temperatures, in a time scale of 100 to10,000 years (Ohmoto, 1996).

Regional hydrothermal systems are interpreted to berelated to large volume intrusions, as the volume of circulatingfluid in a hydrothermal system theoretically cannot be greaterthan the volume of the intrusion (Cathles, 1981). However,small-volume near-seafloor intrusions, which are unlikely togenerate significant hydrothermal systems, may be relatedto larger plutons or stocks at depth that were capable ofgenerating hydrothermal convection (e.g. Bettison-Varga etal., 1992).

6.2 | REGIONAL ALTERED ZONESASSOCIATED WITH INTRUSIONS

The products of regional-scale hydrothermal alterationsystems in ancient submarine volcanic successions arerecorded by cross-cutting recharge and discharge zones, andbroad, regional-scale, semi-conformable altered zones orreaction zones (Galley, 1993).

Recharge zones

Very little is known about altered zones associated withrecharge. They are rarely recognised except in studies ofmodern crustal alteration beneath mid-ocean ridges (e.g.Mottl, 1983; Saccocia et al., 1994; Alt, 1999), hydrothermal Discharge ZOneS

SYNVOLCANIC INTRUSION-RELATED ALTERATION | 1 4 1

alteration studies in ophiolites (e.g. Schiffman et al., 1987;Schiffman and Smith, 1988), and studies of O- and S-isotopecompositions in ancient hydrothermally altered systems (e.g.Cathles, 1993; Davidson and Kitto, 1997). Rocks in modernrecharge zones are pervasively altered at low to moderatetemperatures. At less than 150°C, oxidation, the fixation ofalkalis (mainly Ca and Na), and Mg-metasomatism producessericite, hematite and clays (Alt, 1999). At higher temperatures(150-350°C) anhydrite precipitates, alkalis are leached andMg is consumed by chlorite in the rock (Alt, 1999).

Schiffman and Smith (1988) proposed that the distributionof epidosites in the Troodos ophiolite represent areas ofhigh-temperature alteration involving high fluid-rock ratios.Epidosites are granoblastic, fine- to medium-grained rocks,with little or no relict igneous textures, composed of epidote,quart and chlorite. They are inferred to record reaction zonesin which circulating modified seawater reacted with hostrocks to form metal-rich hydrothermal fluids, and appeardiagnostic of the up-welling and deep recharge parts of thehydrothermal system beneath VHMS deposits. Co-incidentwhole-rock O-isotope patterns support their formation inproximal recharge zones and up-flow conduits beneath VHMSdeposits. Regionally extensive, depth-dependent 618O profilesin the sheeted dyke complex reflect oxygen exchange duringprograde regional hydrothermal alteration involving diffusedown-welling of cold seawater (Fig. 6.4). However, surfaces ofequal whole-rock 618O are not horizontal but nearly verticalin the central epidosite zone. This suggests up-flow of hotmodified seawater within the epidosite zone.

The spatial association between gabbro intrusions andthe epidosite zones in the sheeted dyke complex indicates agenetic link between the emplacement of these intrusions andfocused high-temperature hydrothermal up-flow (Richardsonet al., 1987; Bettison-Varga et al., 1992).

Discordant footwall alteration pipes and feldspar-destructivezones that directly underlie VHMS deposits are widelyinterpreted as discharge zones through which metal-bearinghydrothermal fluid ascended to the seafloor (Sangster,1972; Large, 1977; Lydon, 1984; Galley, 1993; Skirrow andFranklin, 1994; Brauhart et al., 1998). They are characterised

FIGURE 6.3 | Model of an active geothermal system illustrating the recharge,reaction or reservoir and discharge zones. Seawater is drawn down in broadrecharge zones or along faults and reacts at increasing temperatures. High-temperature reactions (>350°C) occur in the reaction zone above a subvolcanicintrusion and hot (>300°C) buoyant fluids rise towards the surface in focused ordiffuse discharge zones (modified after Alt, 1995a). Not to scale.

FIGURE 6.4 | Cross-section of the Solea graben, Troodos ophiolite, Cyprus,showing surfaces of equal whole-rock d180. Regionally these surfaces are sub-horizontal, but in the central epidosite zone they are nearly vertical indicatingup-flow of hot, modified seawater during convection. Modified after Schiffmanand Smith (1988).

142 | CHAPTER 6

by Mg-Fe enrichment and Na-Ca depletion and assemblagesthat include chlorite, sericite, quartz or rare talc (Lydon,1984; Eastoe et al., 1987; Skirrow and Franklin, 1994;Brauhart et al., 1998). The characteristics and compositionalchanges associated with discordant footwall alteration pipesare discussed in Section 7.3.

Although discordant altered zones typically cut acrossthe regional, deep, semi-conformable altered zones (Galley,1993; Brauhart et al., 1998), in some successions, they gradelaterally into deep, semi-conformable altered zones (Skirrowand Franklin, 1994; Hudak et al., 2000). Gibson et al.(2000) suggested that whether or not deep, semi-conformablealtered zones are cut by or transitional with pipe-like alteredzones, depends on whether the host succession (footwall)is dominated by coherent volcanic or volcaniclastic faciesrespectively, or timing of alteration.

Deep, semi-conformable altered zones

Since they were first discussed by Franklin et al. (1981), deep,semi-conformable altered zones have been documented inthe footwall beneath VHMS deposits in a variety of districtsincluding: Matagami, Snow Lake, Noranda and SturgeonLake districts in Canada; Bersglagen and Skellefte districts inSweden; Iberian pyrite belt in Spain and Portugal; TroodosOphiolite Complex in Cyprus; Panorama district in Australia;and the Sirohi district in India (MacGeehan, 1978; Gibson etal., 1983, 2000; Lagerbald and Gorbatschev, 1985; Galley,1993; Skirrow and Franklin, 1994; Tiwary and Deb, 1997;Brauhart et al., 1998; Bailes and Galley, 1999; Hannington etal., 2003a, 2003b). They have not been documented in easternAustralia, possibly because of structural complexities. Thusthe following discussions on deep, semi-conformable alteredzones are largely based on Canadian examples. Figure 6.5depicts the characteristics and typical zonation of deep, semi-conformable altered zones in the documented examples.

Deep, semi-conformable altered zones typically extendfor up to 20 km laterally and 1-4 km depth beneath paleo-seafloors and VHMS deposits (Gibson et al., 1983, 2000;Cathles, 1993; Galley, 1993; Skirrow and Franklin, 1994).They comprise vertically stacked, sub-horizontal altered zones(Galley, 1993; Skirrow and Franklin, 1994). Generally theseare (Fig. 6.5): an upper background K-Mg metasomatic zone;a transitional Na-Mg metasomatic zone; a central silicifiedzone; and a basal Ca-Fe metasomatic and base metal-leachingzone (Galley, 1993). In many systems only one or two ofthese altered zones are recognised. The alteration minerals inthe semi-conformable altered zones reflect the primary hostrock composition, bulk-rock composition established duringsynvolcanic hydrothermal alteration, and the subsequentmetamorphic grade (Paradis et al., in press). In greenschistfacies metamorphosed felsic rocks, these zones are typically,from base to top: albite or carbonate zone, silica or sericitezone and sericite or chlorite zone (Gibson et al., 1983; 2000).In mafic rocks, the zones are: albite zone, silica zone andclinozoisite/epidote + quartz zone (Galley, 1993; Skirrowand Franklin, 1994; Gibson et al., 2000; Hannington et al.,2003a). At higher metamorphic grades, such as in the SnowLake District, mineral assemblages can include kyanite,staurolite, sillimanite, chlorite, biotite, quartz, plagioclasecordierite, amphibole, epidote and garnet (Paradis et al.,1993, in press; Bailes and Galley, 1999).

The semi-conformable altered zones are interpreted tobe synvolcanic because they have undergone the same degreeof tectonic deformation as the surrounding rocks, haveprograde mineral assemblages, are spatially associated withVHMS deposits, and are commonly truncated by unalteredsynvolcanic intrusions (Gibson et al., 1983; Paradis et al.,1993; Skirrow and Franklin, 1994). At Snow Lake, Paradiset al. (1993) recognised that deep, semi-conformable alteredzones were superimposed on low-temperature (possiblydiagenetic) altered zones and also cross cut by discordantfeldspar-destructive zones associated with VHMS deposits.

PROCESS ANDCOMPOSITIONAL

CHANGES

K-Mg metasomatism+ Mg, K, Fe-Na, Ca, Cu, Pb, Zn, Si

Na-Mg metasomatism+ Na, Mg-Ca, Fe, Zn.Cu±Si

Silicification or sericitisation+ Si, Na-Fe, Mg, Mn,Zn±Ca

Ca-Fe metasomatism+ Ca-Mg, Mn, Na, K± Fe, Si

ASSEMBLAGE INMAFIC ROCKS

ASSEMBLAGEIN FELSIC ROCKS

Mg clays + chlorite+ zeolites +Fe-oxides ±K-feldspar

Albite + quartz +sericite +Mg-chlorite ± calcite

Quartz + albite

Clinozoisite/epidote+ quartz± actinolite± carbonate

Mg-clays + zeolites ±cristobalite ± adularia± analcime ±K-feldspar

Albite + quartz + sericite+ Mg-chlorite

Quartz ± albite ± sericite

Sericite + quartz± Mg-chlorite or chloritoid+ Fe-chlorite

FIGURE 6.5 | A schematic compilation of regional-scale, deep, semi-conformable altered zones and their characteristics. There is a progression, with increasingdepths in submarine volcanic successions, from the background Mg-K metasomatic zone to a transitional Na-Mg metasomatic zone characterised by feldsparalteration, a central silicified zone and a basal Ca-Fe metasomatic and base-metal leaching zone that typically includes epidote or chlorite, After Gibson et al. (1983),Galley (1993), Skirrow and Franklin (1994), and Brauhart etal. (1998).

They suggested that regional hydrothermal alteration post-dated the onset of diagenetic alteration, and pre-dated footwallalteration associated with hydrothermal alteration andmineralisation. In some Canadian examples and at Panoramain western Australia, deep, semi-conformable altered zones aregradational with discordant footwall alteration pipes suggestingthat regional hydrothermal alteration was synchronous withthe VHMS-related alteration (Gibson et al., 1999).

Deep, semi-conformable altered zones are commonlyspatially and temporally associated with subsurface syn-volcanic intrusions (Galley, 1993, 2003). These intrusionscan be individual granitic or porphyritic plutons or sheeted


dyke swarms (e.g. Gibson et al., 1983; de-Ronde et al., 1994;Brauhart et al., 1998). The tops of the subvolcanic intrusionsand associated dykes may be included in the basal semi-conformable altered zone (e.g. Galley, 1993; Brauhart et al.,1998).

One of the best-documented examples of the spatialassociation between a subvolcanic intrusion, regional-scalesemi-conformable altered zones and VHMS deposits comesfrom the Panorama district in Western Australia. Discordantchlorite + quartz zones directly beneath the VHMS deposits,are spatially associated with feldspar-destructive sericite+ quartz zones in the top of the Strelley Granite pluton

FIGURE 6.6 | Geology and altered zones within part of the Strelley succession, Panorama district, Western Australia (modified afterBrauhart et al., 1998).

144 | CHAPTER 6

FIGURE 6.7 | Schematic section of the geology and altered zones in the

Kangaroo Caves footwall succession, Panorama district, Western Australia

(modified after Brauhart et al., 1998). See Figure 6.6 for legend.

(Figs 6.6 and 6.7; Brauhart et al., 1998). Faults boundingthe discordant chlorite + quartz zone in the footwall of theKangaroo Caves deposit (Fig. 6.7) controlled the distributionof the feldspar-destructive sericite + quartz zone in the StrelleyGranite (Brauhart et al., 1998).

The variations in alteration mineral assemblage downthrough the semi-conformable altered zones correspond togeochemical gradients in which there are gradual decreasesin the Mg/Ca, Mg/Na and Na/Ca ratios of the altered rockswith increasing depths (Galley, 1993). Oxygen-isotopecompositions suggest that the altered rocks are 18O enrichedwith respect to unaltered volcanic rocks (Munha and Kerrich,1980; Barringa and Kerrich, 1984). The geochemicalgradients and O-isotope data are consistent with metasomaticalteration resulting from the interaction of volcanic rockswith seawater (Muehlenbachs and Clayton, 1972; Lagerbaldand Gorbatschev, 1985; Cathles, 1993).

Although modified seawater is interpreted to be the maincomponent, magmatic fluids may have also contributed tothe hydrothermal fluid (Lagerbald and Gorbatschev, 1985).The spatial association between altered zones and subsurfaceintrusions suggests a genetic link where intrusions may haveprovided heat and or fluid to the hydrothermal system.

The extent and intensity of deep, semi-conformablealtered zones implies that very large volumes of fluid must havereacted with the host volcanic rocks (Skirrow and Franklin,

1994). In the Snow Lake district Skirrow and Franklin(1994) estimated that approximately 1.1 x 107 metric tonsof SiO2 was added by a minimum 12 km3 of hydrothermalfluid at 1—2 km depth. Interactions between large volumesof modified seawater and volcanic successions at depth aresupported by geochemical and geophysical research at activeocean spreading ridges (e.g. Spooner and Fyfe, 1973; BischoffandDickson, 1975).

Deep, semi-conformable altered zones are characterisedby mineral assemblages that reflect the reactions of glassand both primary and secondary minerals with seawater attemperatures up to 400°C (Galley, 1993).

Background K-Mg metasomatic zones

These zones are often described as the least-altered ordiagenetically altered zones. At low temperatures (50-140°C)in the shallow subseafloor, the interaction of abundantseawater with the volcanic succession produces Mg-K-richalteration assemblages (Seyfried and Bischoff, 1977; Galley,1993). In felsic rocks these mineral assemblages includeadularia and Mg-smectite, whereas in mafic rocks they aredominated by zeolites and Mg-smectite. Seawater becomesenriched in Si, Fe3+, Mn and lesser amounts of Ca, Mg andsulfur (Seyfried and Bischoff, 1977).

Current mineral assemblages reflect the regionalmetamorphic grade. For example, at Snow Lake the dia-genetically altered zone is characterised by quartz + biotite +garnet, Fe2O3, MgO and K2O gains and CaO, Na2O, Cu, Pb,Zn losses (Paradis et al., in press). These compositional changesare consistent with low-temperature seawater-dominateddiagenesis of felsic volcanic facies to clays and zeolites (Section5.3). The current mineral assemblage reflects the overprint ofamphibolite facies metamorphism. In the Panorama districtthe background alteration mineral assemblage includesfeldspar + calcite ± ankerite + quartz + pyrite ± sericiteconsistent with greenschist facies metamorphism of clays andzeolites in felsic volcanic rocks (Brauhart et al., 1998).

Transitional zone or Na-Mg metasomatic zones

With increasing stratigraphic depth there is a transition fromK-rich zones to Na-rich zones (Munha et al., 1980; Munhaand Kerrich, 1980; Lagerbald and Gorbatschev, 1985;Schiffman and Smith, 1988; Brauhart et al., 2001). The Na-Mg metasomatic zones are characterised by the occurrenceof feldspar, usually albite. Greenschist facies assemblagestypically include chlorite, sericite, albite, epidote and quartz inmafic rocks, and albite, quartz, sericite ± chlorite ± carbonate(calcite or dolomite) in felsic rocks (Gibson et al., 2000). Inthe Panorama district, felsic rocks in the feldspar zone have theassemblage K-feldspar or albite + sericite + quartz + ankerite +leucoxene ± pyrite (Brauhart et al., 1998).

The transition to Na-rich zones reflects the behaviour of Naand K in seawater at elevated temperatures. Between 140° and200°C there is a transition between K- and Na-metasomatism(Seyfried and Bischoff, 1977). Munha et al. (1980) suggestedthat at lower temperatures (<150°C), Na in glass is exchangedfor K in seawater, resulting in precipitation of K-rich zeolites

and possibly K-feldspar. At higher temperatures, K in the rockis exchanged for Na in seawater, resulting in the formationof albite. Thus at moderate temperatures (140—300°C),metasomatic reactions between modified seawater and thevolcanic succession result in Na-Mg alteration assemblages(Seyfried et al., 1988). Regardless of the rock type, alterationmineral assemblages include Mg-smectite + chlorite + quartz+ albite, and compositional changes are Na2O and MgOgains, and CaO, Zn and Cu losses (Gibson et al., 2000). Theremoval of Mg from seawater lowers the pH of the fluid andseawater evolves from a moderately alkali, Mg-K-Na-SO4-rich fluid to a hot acidic Si-Na-Ca-rich hydrothermal fluid(Bischoff and Seyfried, 1978; Seyfvied et al., 1988).

Silicified zones

In greenschist facies felsic and some mafic rocks, the centralaltered zone is typically silicified, with assemblages of quartz


+ plagioclase or albite (Skirrow and Franklin, 1994; Gibsonet al., 2000). In some mafic rocks, the central zone is sericitic,dominated by sericite + quartz ± chlorite (Gibson et al.,2000). Central silicified or sericite zones overprint regionalalbite zones (Galley, 1993). At Snow Lake the silicified zoneis spatially and temporally associated with VHMS depositsand is zoned laterally from a silica zone to epidote and Fe-Mg-metasomatic zones (amphibolite grade; garnet + chlorite± biotite ± staurolite) (Skirrow and Franklin, 1994). Silicifiedzones are typically spatially associated with synvolcanicintrusions and the intensity and pervasiveness of alterationincreases with proximity to the intrusions (Skirrow andFranklin, 1994; Paradis et al., in press).

Silicified zones commonly contain patches of quartz +feldspar-altered rock, quartz-altered clasts in volcaniclasticfacies, and quartz veins (e.g. Fig. 6.8A: Gibson et al., 1983;Skirrow and Franklin, 1994). The patches of quartz + feldspar-altered rock are restricted to flow-top breccias, and flow-

A. Central silicified zoneModerate, patchy quartz alteration in this andesite fromthe central silicified zone resulted in a fine-grained, palerock, which resembles a rhyolite.Amulet Formation, Noranda district, Buttercup Hill,Canada.

B. Epidote + quartz zoneThis approximately one metre-wide patch of epidote+ quartz alteration facies in the upper Amulet andesitehas an irregular shape typical of patchy alteration in thebasal episite + quartz zone. The groundmass has beenpervasively epidote + quartz altered.Amulet Formation, Noranda district, Canada.

C. Epidote + quartz zoneAmygdales in this patch of epidote + quartz-alteredandesite from the basal epidote + quartz zone haveamoeboid shapes and were lined with Fe-oxides andfilled with epidote + quartz.Amulet Formation, Noranda district, Canada.

FIGURE 6.8 | Photographs from deep semi-conformable alteration zones in

the Noranda district, Canada.

146 | CHAPTER 6

banded and vesicular lavas (Gibson et al., 1983). In maficvolcanic rock the originally glassy groundmass, elsewheretypically altered to chlorite, is altered to quartz in this zone(Gibson et al., 2000). This led to intensely silicified andesitesin the Noranda sequence being misinterpreted as rhyolite: theAmulet rhyolite (Gibson et al., 1983).

With increasing depth, seawater carries larger amounts ofSi as Si solubility increases with temperature and pressure,and is enhanced in NaCl solutions or where the fluid is incontact with free Si or glass (Kennedy, 1950; Fournier,1985). The Si-rich hydrothermal fluid is rapidly heatedbeyond the temperature of the quartz solubility maximum:340-400°C at pressures below 900 bars (Fig. 6.9: Kennedy,1950; MacGeehan, 1978; Skirrow and Franklin, 1994). Theresult is gains in SiO2 and Na2O, due to the precipitation ofsilica within pore spaces and albitisation, and losses in FeO,MgO, CaO, K2O, MnO and other metals from the volcanicrocks (Gibson et al., 1983, 2000; Lagerbald and Gorbatschev,1985; Galley, 1993; Skirrow and Franklin, 1994). The Fe3+,Mg and possibly Zn leached from the silicified zone mayhave been transported laterally away from this environment,thereby producing semi-conformable Fe-Mg-metasomatisedzones (Skirrow and Franklin, 1994).

A second silicified zone is common directly beneath theseafloor in the Snow Lake, Noranda and Matagami Lakedistricts, where it is directly overlain by exhalites. This near-seafloor, silicified zone is related to low-temperature silici-fication during the hydrothermal alteration and devitrificationof glass in cooling pillow basalts and andesites (Skirrow andFranklin, 1994; Galley et al., 2002).

FIGURE 6.9 | Calculated solubilities for quartz in water up to 900°C atpressures between 200 and 1000 bars (after Fournier, 1985). The shaded areaoutlines the conditions for retrograde solubility.

Basal Ca-Fe metasomatic zones

Mineral assemblages in basal semi-conformable altered zonesare dependent on the host-rock composition and porosity. Infelsic rocks, basal zones are typically sericite or chlorite zones,whereas in mafic rocks they are clinozoisite or epidote +quartz zones (Gibson et al., 2000; Hannington et al., 2003a).Sericite zones are characterised by sericite + quartz ± Mg-chlorite assemblages (Gibson et al., 2000). Chlorite zonesare characterised by chloritoid + Fe-chlorite ± Fe-carbonateassemblages (Gibson et al., 2000). Epidote + quartz zonesare characterised by epidote + quartz + calcite + actinolite +chlorite assemblages (Galley, 1993).

Two alteration textures are persistent in epidote + quartzzones: pervasive and patchy. Pervasive epidote + quartz occursas selective pervasive replacement of plagioclase phenocrystsor the groundmass by epidote, crystallisation of fine quartzpatches in the groundmass (e.g. Fig. 6.8B), and replacementof Fe-Ti-oxide grain rims by sphene (Skirrow and Franklin,1994). Patchy epidote + quartz occurs as less than 1 cm to2 m irregular ovoids or amoeboid patches that infill vesiclesand gas cavities within mafic lavas (e.g. Fig. 6.8C: Gibson etal., 1983; Skirrow and Franklin, 1994). These patchy texturesare similar to the epidote + quartz metadomains described inspilites by Smith (1968, 1974, 1977; Smith et al., 1982).

These basal zones may grade into the discordant(discharge) altered zones that cut through the overlying semi-conformable and background altered zones (Brauhart et al.,1998).

Epidote + quartz zones are enriched in CaO and Sr anddepleted of MgO, Na2O, K2O, FeO ± MnO, Ba and basemetals (MacGeehan, 1978; Gibson et al., 1983; Richardson etal., 1987; Schiffman and Smith, 1988; Skirrow and Franklin,1994). Unlike the Canadian examples, the epidote + quartzzone in the Panorama district does not appear to have beenthe source of leached base metals (Brauhart et al., 2001).

At high temperatures (300-500°C), Ca-Fe-S-base metal-rich hydrothermal fluid reacts with the volcanic successionand possibly also with parts of the subsurface intrusionforming mineral assemblages typical of the basal Ca-Fe-metasomatic zones. Experimental work suggests that epidote+ quartz alteration involved modified seawater (Mg-depleted,Ca-Na-K-Cl fluid) at temperatures of 35O-5OO°C and lowwater-rock ratios of less than three (Gibson et al., 2000).

This zone is interpreted to represent the high-temperatureinteraction between modified seawater and the host volcanicfacies to form metal-rich hydrothermal fluid at the deepest partof the hydrothermal convection system. Hence, it representsthe roots of up-welling fluid discharge zones (Galley, 1993).

Alternatively, Smith (1968, 1977) interpreted thesedistrict-scale zones of albitised basalt with Ca-rich epidote +quartz and pumpellyite + quartz metadomains to have formedduring heterogenous burial metamorphism where local fluidflow promoted redistribution of elements. In some cases, henoted that the alteration was focussed suggesting that it wasrelated to local hydrothermal systems and subseafloor fluidcirculation.

Altered zones as part of a regional hydrothermalsystemDeep, semi-conformable altered zones superficially resembleregional diagenetic or metamorphic facies. This is becausethey are regionally extensive, vertically stacked alteredzones with mineral assemblages similar to those formedduring high-temperature diagenesis, regional greenschistfacies metamorphism and hydrothermal seafloor alteration(Galley, 1993; Paradis et al., 1993). Discriminating betweenthese processes and their products is difficult in submarinevolcanic successions. The differences are essentially related totiming, temperatures, and fluid-rock ratios. In reality, there isa progression from diagenesis to isochemical metamorphismwith increasing temperature and depth during burial (Fig.6.10A). Porosity, permeability and fluid-rock ratios decreasewith depth in diagenetic-metamorphic systems, therebyinhibiting the degree and pervasiveness of metasomatism attemperatures above 150°C. Typically once the temperatureand pressure realm of metamorphism has been reached, theporosity and permeability of the host succession has beendramatically reduced, fluid flow inhibited and metasomaticreactions ceased. Gibson et al. (2000) suggested seawater-dominated diagenesis might also progress to deep regionalhydrothermal alteration with increasing temperature anddepth in shallow subseafloor hydrothermal systems (Fig.6.1 OB). Deep regional hydrothermal alteration is interpretedto involve metasomatic reactions between seawater andthe volcanic succession at temperatures transitional withdiagenesis and greenschist facies metamorphism (i.e. 150—400°C) (Galley, 1993). Although the processes of diagenesisand deep regional hydrothermal alteration are very similar,and both involve reactions between seawater (or modifiedseawater) and volcanic successions at increasing temperaturesand depths, deep, semi-conformable altered zones areinconsistent with the diagenetic-metamorphic system. They


have anomalous mineral assemblages and alkali contents forigneous rocks (e.g. Fig. 6.11; Hughes, 1973), which suggestmetasomatic rather than metamorphic origins (Gibson et al.,1983; Galley, 1993).

Gibson et al. (1983) documented a vertically stackedsequence of altered zones in the Noranda sequence, from topto bottom: albite zone (spilites), silicified zone, and epidote +quartz zone (Fig. 3.16). This is consistent with a progressionfrom low-temperature diagenesis to moderate- and high-temperature metasomatism with depth in the stratigraphy.Munha and Kerrich (1980) referred to this process oftemperature and hence depth dependent metasomatismas 'hydrothermal metamorphism', a term that reflects thecombined processes that operated in the subseafloor.

In areas of volcanic and hydrothermal activity, it is probablethat there is a spectrum of alteration between diagenesis andhydrothermal alteration where these processes operate incombination. Hydrothermal activity in the depositional basinwould accelerate and intensify the process of diagenesis bycontributing additional fluid and heat, and by promotingconvection (Iijima, 1974; Marsaglia and Tazaki, 1992).

Deep, semi-conformable altered zones are assumed tobe the products of hydrothermal alteration within regionalsubseafloor hydrothermal systems (Gibson et al., 1983;Galley, 1993). These hydrothermal systems involve the large-scale convection of modified seawater through the permeablevolcanic successions (Spooner and Fyfe, 1973; Galley, 1993).The distribution of altered zones and spatial association withsubsurface intrusions suggests that subsurface intrusions,augmented by heat from the cooling volcanic succession, maybe the driving force for hydrothermal convection (Campbellet al., 1981; Lesher et al., 1986; Cathles, 1993; Galley, 1993;Skirrow and Franklin, 1994). Where the upper contacts ofsubsurface intrusions are sub-parallel to the volcanic-strata,the overlying isotherms are also semi-conformable with thestrata and progressive temperature-dependent seawater-rock

FIGURE 6.10 j The relationships between fluid convection, diagenesis, metamorphism and regional hydrothermal alteration in submarine volcanic successions thathost VHMS deposits. (A) Diagenetic-metamorphic system, where there is a progression from diagenesis to isochemical metamorphism with increasing temperatureand depth in the subseafloor. This transition reflects the maximum depth to which seawater circulates and reacts with the host rocks. (B) Diagenetic-hydrothermalsystem, where a subsurface intrusion promotes deep circulation of fluid via the recharge, reservoir and discharge zones. The depth progression from diagenesis toregional hydrothermal alteration (deep, semi-conformable alteration) is dependent on temperature and fluid circulation.

1 4 8 | CHAPTER 6

FIGURE 6.11 | Alkali ratios for altered andesite samples from the Amulet

rhyolite, Noranda district, Canada (after Gibson et al., 1983). Fields for the

primary and metasomatised (albite-altered) andesites and basalts are from

Hughes (1973).

reactions form a series of semi-conformable altered zones(Galley, 1993).

In axial mid ocean ridge hydrothermal systems, down-welling seawaters traverse extremely steep temperaturegradients in the upper crust, from less than 50°C near theseafloor to more than 250°C at 1-2 km depth (Mottl, 1983).Thus, vertically stacked deep, semi-conformable alteredzones result from metasomatic reactions that take place atprogressively higher temperatures with depth in the succession(Galley, 1993). The decrease in pervasiveness of alteration,from widespread nearly uniform diagenesis to more restrictedand patchy silicification and Ca-Fe metasomatism, may reflectdecreasing permeability with depth.

The distribution, relative timing, and spatial associationwith VHMS deposits suggest a genetic link between regionalhydrothermal alteration and mineralisation. Some authorshave proposed that the deep, semi-conformable altered zonesacted as reservoirs from which metals and sulfur were leached(e.g. Gibson et al., 1983; Lagerbald and Gorbatschev, 1985;Galley, 1993; Skirrow and Franklin, 1994). As such theyrepresent much larger exploration targets than the discordantaltered footwall zones (Galley, 1993).

6.3 | ALTERED ZONES WITHININTRUSIONS

Intrusions are commonly modified by deuteric or hydrothermalalteration associated with emplacement and may subsequentlyundergo diagenesis, regional metamorphism or hydrothermalalteration.

Deuteric alteration

Deuteric alteration, also referred to as autohydration orautometamorphism, is the alteration of recently crystallisedmagma by trapped magmatic fluid exsolved from the samecooling magma (Honnorez et al., 1979; Destrigneville et al.,1991). It has been recorded in intrusions and lavas in both

subaerial and submarine successions (Honnorez et al., 1979;Bohlke et al., 1980; McConnell et al., 1995). Sederholm(1929) originally defined deuteric alteration as the alterationthat takes place 'in direct continuation of the consolidationof the magma' and thus it is considered a magmatic alterationprocess. It is the earliest alteration style and is a short-lived process, typically occurring as intrusions cool fromtemperatures of several hundred degrees centigrade (Ade-Halletal., 1968; Honnorez etal., 1979). Small volume synvolcanicintrusions typically cool too rapidly to experience deutericalteration (cf. Gromme et al., 1969). In contrast, granitoidsand large-volume sills may have altered zones in their upperparts resulting from reactions between rising magmatic fluidsand the cooling intrusions (e.g. Fig. 6.1A).

Deuteric textural changes are minimal (Wilshire, 1959).Phenocrysts, particularly feldspars and mafic minerals, such aspyroxene or olivine, may be pseudomorphed by amphibole,chlorite or smectite (Fuller, 1938; Bohlke et al., 1980;Destrigneville et al., 1991). Open spaces, such as vesicles,mariolitic voids, and quench fractures, are lined or filledwith smectite, zeolites, carbonate, biotite, chlorite and oxides(Wilshire, 1959; Furbish and Schrader, 1980; Destrignevilleet al., 1991). High-Ti minerals, such as titanomagnetite, areoxidised and altered to low-Ti minerals, such as ilmenite ±hematite (Butler and Burbank, 1929; Ade-Hall et al., 1968;Surdam, 1968; Sherwood, 1988).

Deuteric alteration does not involve major chemicalcomposition changes; some components may be locallyredistributed at sub-millimetre scales or undergo oxidationstate changes (e.g. Fe2+/Fe3+ ratio, Scott and Hajash, 1976)that may alter rock thermomagnetic properties (Ade-Hall etal., 1968; Sherwood, 1988). The changes are quantitativelyunimportant when compared to the products of long-liveddiagenesis and hydrothermal alteration and may be difficultto distinguish from those of other alteration styles.

Hydrothermal alteration

Altered zones within synvolcanic intrusions may also resultfrom reactions with seawater or modified seawater circulatingthrough the intrusion, either during the prograde high-temperature stage of hydrothermal activity or during cooling.If the hydrostatic pressure is high enough (at sufficient depths)seawater will be forced into thermal contraction fractures inthe cooling intrusion (Burnham, 1979). The time interval forseawater-intrusion interaction may be limited by the rapiddevelopment of a local intrusion-related hydrothermal systemin the host succession, which would result in the lithificationand filling of primary pore space inhibiting fluid flow.Thereafter, episodic seawater-intrusion interaction wouldoccur only if the fluid pressure exceeds the tensile strengthcausing the altered rock adjacent to the contact to fracture(Secor, 1965, in Fournier 1985; Phillips, 1973; Henley andMcNabb, 1978). Alteration follows the advancing frontof brittle fracturing to deeper and deeper levels within theintrusion (Burnham, 1979; Giggenbach, 1997).

The resulting altered zones may be pervasive, occur alongcooling fronts or more commonly as selective-pervasivealteration adjacent to fractures or veins. Alteration mineralsfill vesicles, miarolitic voids and fractures, cement hydraulic

breccias, and pseudomorph primary magmatic minerals(Mevel and Cannat, 1991; Gillis et al., 1993; Kelley andGillis, 1993; Nehlig, 1993; Davidson, 1998; Galley, 2003).

Polya et al. (1986) and Davidson (1998) describedproximal zones of hydrothermal alteration within theCambrian Murchison Granite, western Tasmania. Theydescribed narrow, texturally destructive intense K-feldsparzones, associated with calcite veins, irregularly distributedin the margins of the granite and patchy selective-pervasivechlorite zones with chlorite pseudomorphs after biotite andhornblende or patches of chlorite ± pyrite ± sphene.

The significant mineralogical, compositional andisotopic changes associated with proximal hydrothermal alter-ation within intrusions are consistent with seawater-rockinteraction (Gregory and Taylor, 1981; Stakes and Taylor,1992; Cathles, 1993; Galley, 2003). Galley (2003) identifiedthree types of early hydrothermal and magmatic alterationwithin subvolcanic intrusions in the Snow Lake, Noranda andSturgeon Lake districts. The earliest, a greenschist alterationfacies in quartz diorite intrusions, is manifest as pervasiveepidote + quartz, and epidote + actinolite + quartz + albite+ magnetite ± sulfides, which replaced primary minerals andinfilled miarolitic cavities, vesicles and fractures. Mass changecalculations suggest CaO, Sr, Pb and CO2 were gained andFe2O3, MgO, Cu, Zn, Mo, Na2O, K2O and Ba lost. Galley(2003) interpreted this facies to be the product of high-temperature hydrothermal-magmatic alteration resulting fromemplacement of quartz-diorite intrusions into a seawater-saturated succession. The second chlorite-rich alteration faciesis characterised by quartz + chlorite + sericite and chlorite+ sulfide-filled fractures and vein selvages. It is most intensenear the margins of intrusions and directly beneath VHMSdeposits. The chlorite-rich zones gained Fe2O3, MgO, Cu ±Pb, K2O, Ba, and Zn, and lost CaO, Na2O, Sr ± SiO2, Baand CO2 and are interpreted to result from hydrothermalalteration (Galley, 2003). Overprinting both of these zonesis a biotite-rich alteration facies associated with silicification,and Cu-Mo-rich veins and breccia, which is interpreted asa magmatic alteration facies associated with late stage dykes(Galley, 2003).

Hydrothermal alteration may also result from theabsorption of fluid from and assimilation with sedimentinclusions incorporated into the magma as it was emplacedinto wet unconsolidated sediment. Wilshire and Hobbs(1962) described hydrothermal alteration in the margin ofa peperitic latite intrusion in a submarine volcaniclasticsuccession, near Port Kembla in New South Wales. Alkalifeldspar + chlorite + carbonate-rich zones coincide withabundant sediment inclusions and quench fractures in themargin of the intrusion, and the sedimentary inclusions havebeen chlorite ± zeolites ± carbonate altered. The altered latitegained Na2O and volatiles, and lost SiO2, A12O3, Fe2O3, K2O,MgO and CaO, whereas the sediment inclusions lost Na2Oand volatiles, and gained CaO and MgO ± K2O.


6.4 | CONTACT ALTERED HALOSAROUND INTRUSIONS

All magmatic intrusions transfer heat; they have thermalimpacts on the enclosing rocks or sediments and may inducecompositional changes. Contact alteration is used here as anon-genetic term referring collectively to the processes ofcontact metamorphism and contact hydrothermal alteration.

Contact or thermal metamorphism involves changes inrock texture and mineralogy of the immediate host rock as aresult of temperature increase (Yardley, 1989). The increasedtemperature drives dehydration and decarbonation reactions,and fluid migration away from the intrusion (Blatt et al.,1972; Manning and Bird, 1991). Only small volumes ofH2O- and CO2-rich fluids are generated from these reactionsand thus the metasomatic effect of contact metamorphismis negligible (Rose and Burt, 1979). Contact metamorphismtypically results in only local remobilisation but extensivestatic recrystallisation of existing minerals or components.

Contact hydrothermal alteration involves a substantialvolume of heated fluid, typically comprising trapped seawaterand pore fluid with or without magmatic fluid derived fromthe intrusion, which circulates through and reacts with thehost facies (MacGeehan, 1978; Taylor and Forester, 1979;Polya et al., 1986; Galley, 2003). This promotes textural,mineralogical and compositional changes in the host facies.

In submarine volcanic successions, abundant trappedseawater means that isochemical thermal metamorphism israre. In addition, interaction between hot magma and wetunconsolidated sediment can result in: peperitic contacts,fluidisation of sediment (e.g. Schmincke, 1967; Kokelaar,1982), fluid expulsion, induration (e.g. Einsele et al., 1980),secondary welding (e.g. Ito et al., 1984, in Kano, 1989),brecciation of host rock, local or regional hydrothermalalteration, quenching of the intrusion and magma-host rockassimilation (e.g. Wilshire and Hobbs, 1962; Puffer andBenimoff, 1997; WoldeGabriel et al., 1999).

Contact altered zones

Contact altered zones are spatially associated with intrusionmargins and may surround the intrusion as halos or aureoles.Successive contact altered zones reflect progressive changesin temperature or temperature and chemical conditions inthe host succession away from the intrusion (Rose and Burt,1979; Einsele et al., 1980; Yardley, 1989).

Contact altered halos may vary in thickness from a fewmillimetres at the margins of thin, shallow synvolcanic sills(e.g. Einsele, 1985; Boulter, 1993; Skirrow and Franklin,1994) to several kilometres around large subvolcanic plutonsor intrusive complexes (e.g. Boulter, 1993; Schweitzer andHatton, 1995; Galley, 2003). They may comprise one low-grade altered zone, a sequence of roughly concentric alteredzones, a series of asymmetric altered zones or overprintingaltered zones (Fig. 6.12).

Grades and mineral assemblages of contact altered zonesvary considerably, reflecting: temperature and compositionaldifferences between the host succession, the intrusion, and anyfluid; duration of the alteration system; emplacement depth;

150 | CHAPTER 6

FIGURE 6.12 | Cartoons of the variety of contact altered zones aroundintrusions. (A) Cross-section showing two sills that were emplaced into wetunconsolidated sediment at shallow levels beneath the seafloor. The sills bothhave single low-grade indurated zones, which have a lower porosity than thesurrounding host turbidites (after Einsele et al., 1980). (B) Schematic cross-section through roughly concentric zeolite and clay zones around a graniteemplaced into felsic volcanic fades in the Green Tuff Belt, Japan. These zonesare, from the intrusion outwards: a zeolite zone, devitrified zone, and least-altered zone (after Utada, 1991). (C) Schematic section of the asymmetricaltered halo around the Rustenberg Layered Suite intrusions in the Rooibergfelsic volcanic rocks of the Bushveld Complex, Africa (after Schweitzer andHatton, 1995). Above the intrusion is a thick (>1.4 km) halo comprising biotitehornfeis and quartz + sericite + albite zones, which are enriched in K2O, MgOand base metals. Beneath the intrusion is a thinner (<400 m) granoblastic zone,in which primary volcanic textures are overprinted by metamorphic textureswithout significant compositional changes. Schweitzer and Hatton (1995)postulated that the reason for the asymmetrical zonation was that heated fluidconvected freely above the intrusion and hydrothermal alteration dominated,whereas buoyant convection was inhibited beneath the intrusion acting as a seal,and thermal metamorphism dominated. (D) Map view of the prograde olivine,pyroxene and actinolite + chlorite zones associated with emplacement of theSkaergaard intrusion into mafic volcanic rocks, east Greenland (after Manningand Bird, 1991,1995). In the outer pyroxene zone and adjacent to fracturesin the pyroxene and olivine zones, high-temperature mineral assemblagesare overprinted by actinolite + chlorite, suggesting retrograde metamorphismoccurred as temperatures dropped and cooler hydrothermal fluids migratedinwards through fractures.

rate of cooling; and the subsequent regional metamorphicgrade. Although it is difficult to generalise about themineralogy of contact altered zones, there are a few indicatorminerals that are almost exclusively generated by intrusion-related hydrothermal alteration (i.e. minerals associated withmagmatic systems such as biotite, diaspore, fluorite, kaolinite,magnetite, pyrophyllite, rutile, topaz and tourmaline).

Typically altered halos associated with small-volumeintrusions emplaced at shallow depths below the seafloorcomprise low-grade altered zones adjacent to the intrusion,which grade into partially altered zones at the peripheries.Low-grade altered zones may be manifest as induratedsediment (e.g. Einsele et al., 1980; Kano, 1989), fused glass(e.g. Ross and Smith, 1960; Smith, I960; Christiansen andLipman, 1966; Schmincke, 1967; McPhie and Hunns,

1995), devitrified glass (e.g. Schweitzer and Hatton, 1995;WoldeGabriel et al., 1999), palagonite or clay minerals inmafic volcanic rocks (e.g. Upton and Wadsworth, 1970;Jakobsson, 1972; 1978) or zeolite and clay minerals in felsicvolcanic rocks (e.g. Iijima, 1978; Utada, 1991).

In contrast, high-grade altered zones tend to be associatedwith large volume subvolcanic intrusions and include high-temperature (up to 1000°C) mineral assemblages. Forexample, Seki et al. (1969) reported five high-grade alteredzones around a large intrusion in the Neogene Green Tuff Belt,Japan. From the contact to the margin they were: amphibolezone, actinolite zone, pumpellyite + prehnite + chlorite zone,laumonite + mixed-layer chlorite zone, and clinoptilolite +vermiculite zone.

SYNVOLCANIC INTRUSION-RELATED ALTERATION I 1 5 1

Indurated or fused zones

Thin contact altered zones of fused or secondary weldedvolcanic glass are common adjacent to intrusions and lavasin subaerial volcanic successions (e.g. Ross and Smith, I960;Smith, 1960; Christiansen and Lipman, 1966; Schmincke,1967; WoldeGabriel et al., 1999), and also occur aroundintrusions in ancient submarine volcanic successions (e.g. Itoetal., 1984, in Kano, 1989; McPhie and Hunns, 1995).

Christiansen and Lipman (1966) used the term fused forthe induration and deformation of glassy clasts resulting fromheating by adjacent lava, but emphasised that the term shouldnot be taken to imply that melting (fusion) had occurred. Theydescribed altered subaerial tuffs adjacent to the Combs Peakrhyolite lavas and domes near Fortymile Canyon, southernNevada. Three altered zones were developed parallel to thelava contact: an outer red zone characterised by the oxidationof glass, a middle indurated or partially fused zone and aninner densely fused zone characterised by fiamme and eutaxitictexture (Figs 6.13 and 6.14). In this case, eutaxitic texture wasinterpreted to result from the re-heating and accompanyingload compaction of originally glassy pumice clasts in tuffsbeneath the lava as a result of its emplacement (Christiansenand Lipman, 1966). The minimum temperature required forthis partial welding is 535°C (Smith, 1960).

Typically, indurated or fused zones closely parallel lavaor intrusion contacts and may be several millimetres totens of metres thick (e.g. Christiansen and Lipman, 1966;Einsele, 1985; Keating and Geissman, 1998). They arecommonly associated with thin (< 1—100 m) intrusions thathave peperitic or irregular contacts indicating emplacementinto wet unconsolidated sediments (Kokelaar, 1982; Branneyand Suthren, 1988; McPhie and Hunns, 1995; Keating andGeissman, 1998). Induration of sediment adjacent to contactsand around juvenile clasts in peperite is typically accompaniedby changes in colour associated with thin (cm scale) carbonate,quartz or Fe-oxide altered halos (Fig. 6.15A: Schmincke,1967; Kokelaar, 1982; Kano, 1989; Hunns and McPhie,1999). The most significant textural changes in this zone arecontact-parallel fiamme and eutaxitic textures in pumice-richfacies (Fig. 6.15B: McPhie and Hunns, 1995).

Devitrified zones

Devitrified zones are characterised by high-temperaturedevitrification textures such as spherulites, lithophysae andmicropoikilitic texture (Christiansen and Lipman, 1966;McPhie and Hunns, 1995). It is important to note thatdevitrification textures generated from re-heating of glassyvolcanic facies by intrusions are indistinguishable from thoseformed during first cooling (Lofgren, 1971a, 1971b).

Narrow zones oriented parallel to intrusion contactsmay be completely or partially devitrified (e.g. Keating andGeissman, 1998; WoldeGabriel et al., 1999). They mayoverprint fused zones. For example, Christiansen and Lipman(1966) described superposition of three devitrified zones on tothree fused zones in bedded rhyolitic tuffs (Fig. 6.14): an outerporous glassy zone, a middle dense glassy zone (vitrophyrethat is commonly perlitic), and an inner crystalline zone withmicrolites, spherulites and lithophysae.

FIGURE 6.13 | Distribution of the fused zone adjacent to the Combs Peak

rhyolite, near Fortymile Canyon, southern Nevada (modified after Christiansen

and Lipman, 1966).

FIGURE 6.14 | Idealised relationships between the Combs Peak rhyolite

(Nevada), the three fused zones and overprinting devitrified zones (modified

after Christiansen and Lipman, 1966).

Compositional changes associated with devitrificationare usually negligible. WoldeGabriel et al. (1999) found thatdevitrification in felsic volcaniclastic rocks in a 10m thickcontact zone around a basaltic intrusion at Grants Ridge,New Mexico, was associated with minor gains in K2O andlosses in H2O, Na2O, F, Fe2O3. The margins of the intrusionwere slightly enriched in SiO2, K2O and P2O5 and depletedin Fe2O3. They concluded that the thermal effects of theintrusion induced devitrification, dehydration and vapour-phase expulsion in the contact zone. Vapour-phase expulsionof fluoride, chloride, hydroxide, sulfide, and CO2 from silicicglass may have been responsible for the subtle chemicalvariations during devitrification (cf. Weaver et al., 1990).

Zeolite, clay or palagonite zones

Low-temperature altered zones characterised by palagonite,zeolite and clay minerals in mafic volcanic rocks (e.g. Uptonand Wadsworth, 1970), and zeolite and clay minerals in felsicvolcanic rocks (Utada, 1991) are common around shallowsynvolcanic intrusions in submarine volcanic successions.Mineral assemblages in these zones typically reflect thehost rock compositions. For example, altered zones aroundgranitoids in the felsic volcanic rocks of the Green TuffBelt contain calcic zeolites (Iijima, 1978; Utada, 1991). Incontrast, palagonite dominates altered zones around dykes in

152 | CHAPTER 6

A. Indurated siltstone in peperiteThe irregular clasts of indurated and silicified siltstone(grey) are mixed with feldspar-phyric rhyolite (green)clasts in this peperitic contact between rhyolite andsiltstone. Away from the rhyolite contact, the hostsiltstone is green-grey, but fades to cream or palegreen silicified siltstone in a zone about 1-2 cm wideadjacent to the rhyolite clasts in the peperite. This localcolour change and silicification result from the thermalmetamorphism of the unconsolidated silt in contactwith hot rhyolite.Early Permian Berserker beds, Mount Chalmers district,

Queensland.

B. Fused pumice brecciaWell-developed fiamme (F) and eutaxitic texturecharacterise the fused zone in this pumice brecciaimmediately adjacent to a rhyolitic sill. Away from therhyolite, fiamme in the pumice breccia are indistinctand parallel to bedding, whereas in the fused zonethey parallel the pumice breccia-rhyolite contact. Thefiamme and eutaxitic texture result from the partialwelding and compaction of glassy pumice clasts duringheating associated with emplacement of the rhyolite(McPhie and Hunns, 1995).Early Permian Berserker beds, Mount Chalmers district,

Queensland.

Figure 6.15 | Photographs of hand specimens from the indurated and fused zones adjacent to rhyolite sills near the IVIount Chalmers VHMS deposit, Queensland.

submarine basaltic hyaloclastite at Surtsey (Jakobsson, 1972,1978; Jakobsson and Moore, 1986), and chabazite, analcime,thomsonite, mesolite, phillipsite and natrolite characterise thezeolite zone associated with a swarm of sills in basaltic lavasand breccias at Piton des Neiges volcano on Reunion Island(Lacroix, 1936; Upton and Wadsworth, 1970). Zeolitesin these zones may be accompanied by chlorite, epidote,carbonate and clay minerals.

Compositional changes in the zeolite, clay or palagonitezones include K2O and MgO gains, and SiO2 and CaO losses(Hart, 1969; Thompson, 1973; Honnorezetal., 1979). Theseare consistent with low-temperature (<150°C) reactionswith seawater promoting oxidation, fixation of alkalis, andexchange of seawater-Mg for rock-Ca to form smectite (Alt,1999).

Greenschist facies zones

Synvolcanic granitoids, large composite intrusions andclusters of sills or dykes commonly have altered zones withepidote-, chlorite-, sericite-, biotite- or K-feldspar-bearingmineral assemblages characteristic of greenschist facies meta-morphism (Polya et al., 1986; Boulter, 1993; Neuhoff et al.,1997; Galley, 2003).

For example, Schweitzer and Hatton (1995) describeda 1.4 km thick greenschist facies aureole above the mafic

Rustenburg Layered Suite in the felsic Rooiberg volcanicrocks of the Bushveld Complex (Fig. 6.12C). The asymmetricaureole contains a biotite hornfels zone (immediately abovethe Rustenburg Layered Suite), and an overlying quartz+ sericite + albite zone, which grades up into least-altered,devitrified volcanic rocks. In the quartz + sericite + albitezone, hornblende or chlorite replaced mafic phenocrysts, andquartz + chlorite + epidote replaced the glassy groundmass.

Primary compositions may be significantly modifiedin greenschist facies zones. They are commonly enrichedin K2O and MgO and depleted in CaO, Fe2O3, Na2O andMnO (Schweitzer and Hatton, 1995; Large et al., 1996).The behaviour of SiO2 is variable. The mineralogical andcompositional changes reflect high-temperature (300-450°C)seawater-rock interactions similar to some proximal alteredzones associated with VHMS ore deposits (Galley, 2003).

Silicified zones

Silicified zones are typically pale grey in colour and can bemassive pervasive or patchy in texture, filling vesicles andfractures, or cementing breccias (Humphris and Thompson,1978; Skirrow and Franklin, 1994; Gifkins and Allen, 2001).They comprise chalcedony, cristobalite, quartz, quartz +feldspar, or quartz + sericite dominated alteration mineralassemblages.


Skirrow and Franklin (1994) described 10 cm to 2 mthick silicified contact aureoles associated with unalteredplagioclase- and quartz + plagioclase-phyric porphyry dykesin the submarine volcanic rocks beneath the Chisel LakeVHMS deposit in the Snow Lake district. The weakly silicifiedmottled zones consist of irregular light grey patches of quartz+ plagioclase + hornblende + magnetite ± biotite, whichcoalesce into massive quartz + plagioclase rock in intenselysilicified zones.

Compositional changes include gains in total mass andSiO2, which may be accompanied by gains in K2O or Na2O,and losses in Fe2O3, MgO, CaO and Zn (MacGeehan, 1978;Skirrow and Franklin, 1994; Gifkins and Allen, 2001).

Silicification is a common feature of hydrothermalalteration and incorporates both the addition of Si (largelyas vein infill) and the redistribution of Si that was originallyin glass or cristobalite (Henley and Ellis, 1983). Circulatingheated seawater can leach Si from the intrusion or felsicvolcanic glass in the host succession, resulting in a solutionsupersaturated with Si. Silica precipitation from this solutioncan occur by several mechanisms: cooling by conduction ormixing, decompression associated with boiling, heating intothe temperature range for retrograde Si solubility, or a pHchange (Dickson and Potter, 1982; Fournier, 1985). Thesolubility of Si generally increases with increasing temperature(Fig. 6.9); however, if a supersaturated solution is heated atconstant pressure (<900 bars) it will either boil or reach asolubility maximum and may precipitate quartz upon furtherheating (Fournier, 1985). A supersaturated saline fluid mayprecipitate quartz at temperatures between 300° and 55O°C(Fournier, 1985). Thus Si-saturated seawater would depositquartz on encountering temperatures greater than 300°Cin the intrusion or the immediate host rocks adjacent tothe intrusion. MacGeehan (1978) proposed this process,of Si leaching from volcanic glass and heating of the fluidinto the retrograde solubility temperature range, to explainsilicification in pillow basalts adjacent to synvolcanic gabbrosills in the Matagami district.

Genesis of contact altered zones

Contact altered zones may develop adjacent to an intrusionas heat is transferred from the cooling intrusion and heatedmodified seawater reacts with the host succession (Fig. 6.16).Vapour or fluid exsolved from the crystallising magma maycontribute both heat and elements to the hydrothermal fluid(Norton, 1984). Hydrothermal fluid temperatures are partlydetermined by the depth of emplacement, volume of theintrusion and the temperature and volume of contributedmagmatic fluid (Polyaetal., 1986; Eastoeetal., 1987; Cathles,1993; Galley, 2003). For example, two active hydrothermalsystems are recognised in the Guaymas Basin (Geiskes et al.,1982; Kastner, 1982). One is a low temperature (<300°C)hydrothermal system associated with the emplacementof shallow sills into unconsolidated sediments below theseafloor. The other is a deep high-temperature hydrothermalsystem associated with dykes or magma chambers that fed theoverlying sill complexes.

FIGURE 6,16 | Development of a contact metamorphic-hydrothermal systemin a submarine volcanic succession after the emplacement of a synvolcanicintrusion. (A) Initial fluid expulsion and migration away from the intrusion asheat from the intrusion drives dehydration and decarbonation reactions in thehost succession. A combination of thermal metamorphism and hydrothermalalteration, by seawater and magmatic volatiles and fluid, may produce a contactaltered zone. Seawater heated by the intrusion is buoyant and rises towards theseafloor either by diffuse flow or along fractures and faults. (B) In response, coldseawater is drawn down and heated in the vicinity of the intrusion, promotinghydrothermal convection and alteration between the intrusion and the seafloor.(C) Hydrothermal convection collapses as the intrusion cools. Cold seawatermay be drawn down along fractures to produce proximal zones of hydrothermalalteration within the intrusion, and retrograde zones that overprint highertemperature contact altered zones adjacent to the intrusion.

The volume of an intrusion influences the temperature andlongevity of the alteration system. Large volume intrusions(e.g. plutons and thick sills) influence the temperature of thehost rocks and the alteration system for longer than smallervolume intrusions (e.g. synvolcanic sills, cryptodomes anddykes). A small volume sill (-30 m thick) may cause thetemperature at the sill-sediment contact to rise as high as400°C (Einsele et al., 1980). However, calculations suggest

1 5 4 | CHAPTER 6

that the temperature at the contact will drop below boilingwithin five years. This will significantly reduce convectionand remaining heat will be lost mainly by conduction throughthe contact. Generally relatively small volume intrusionshave thermal effects restricted to several metres or hundredsof metres from the contacts (Utada, 1973). Although thecontact temperatures may be high, high-temperature alteredzones are rare because the isotherms dip sharply away fromsmall volume intrusions (Reyes, 1990).

In contrast, large volume intrusions, such as the Skaergaardintrusion in east Greenland, which had an estimated volumeof 180 km3, may take 500,000 years to cool to ambienttemperatures (Norton and Taylor, 1979; Norton, 1984).They may result in thick, high-grade contact metamorphic-hydrothermal altered zones (Seki et al., 1969) and may driveregional convection of modified seawater.

6.5 | CONTACT ALTERED ZONESASSOCIATED WITH THE DARWINGRANITE

Cambrian granites along the eastern margin of the MountRead Volcanics (Fig. 1.5) are extensively altered andsurrounded by concentric altered zones (Polya, 1981; Polya etal., 1986;Eastoeetal., 1987; Abbott, 1992; Large etal., 1996;Davidson, 1998; Wyman, 2001). Well-developed K-feldspar,chlorite and sericite zones have been mapped around themargin of the Darwin Granite and its northward extension inthe southern Mount Read Volcanics (Fig. 6.17: Jones, 1993;Large et al., 1996; Wyman, 2001).

The Darwin Granite alteration halo is of particularinterest because of its close spatial and possibly temporalrelationships with several small tonnage but high-grade Cu-Au prospects (Fig. 6.17: Jones, 1993; Large et al., 1996).These prospects occur in the Central Volcanic Complex alongthe western margin of the granite and above its northernsubsurface projection from Mount Darwin towards MountLyell. The deposit styles vary systematically with increasingdistance from the granite: from Fe-oxide veins, stockworksof pyrite + chalcopyrite ± hematite ± magnetite and quartz +pyrite + chalcopyrite veins, disseminated pyrite + chalcopyrite± covellite, to veins containing quartz, bornite, chalcopyriteand hematite (Wyman, 2001). Large et al. (1996) suggestedthat the Darwin Granite provided heat, metals and magmaticfluid to form VHMS deposits in the southern Mount ReadVolcanics, such as those in the Mount Lyell field.

This section summarises the setting, altered zones andgenesis of the Darwin Granite system and presents data sheetsof typical alteration facies (DG1 to DG6).

FIGURE 6.17 | Geological map of the Jukes-Darwin area in the southern MountRead Volcanics (western Tasmania), showing the limited surface extent of theDarwin granite and the thick hydrothermally altered halo around the granite(modified after Wyman, 2001). The locations of the six data sheets are shown.

Geological setting

The Darwin Granite is an I-type magnetite series equigranulargranitoid pluton dominated by pink granite intruded bysubordinate white granite, microgranite and quartz porphyryphases (Wyman, 2001). The surface extent of the pluton isapproximately 5 x 1 km (Fig. 6.17). However, modellingof gravity and aeromagnetic data along the eastern marginof the Mount Read Volcanics has suggested that a semi-continuous body of granite extends subsurface approximately100 km northwards to the Murchison Gorge (Leaman andRichardson, 1989; Payne, 1991; Large et al., 1996).

In the Darwin-Jukes area, the Central Volcanic Complexincludes feldspar-phyric dacite, quartz + feldspar-phyricrhyolite (e.g. data sheet DG1), pumice breccia, tuffaceoussandstone, blocky rhyolite breccia, and minor sedimentaryfacies (Jones, 1993; Wyman, 2001). A thick, columnarjointed, micropoikilitic or spherulitic rhyolite hosts theJukes Cu-Au Prospect and altered zones. The emplacementage of Darwin Granite is constrained to the Cambrian as itintruded the Middle Cambrian Central Volcanic Complex,and both the granite and Central Volcanic Complex areunconformably overlain by late Middle Cambrian TyndallGroup volcaniclastic rocks, which contain pebbles of granitenear Mount Darwin (Corbett, 1979, 1981, 1992; Jones,1993; Wyman, 2001).



The altered zones associated with the Darwin Granite cover a15x3 km area that extends north to Jukes Prospect (Wyman,2001). The altered zones at Mount Darwin and Jukes Prospectrepresent two parts of the same hydrothermal system: thoseadjacent to the granite at Mount Darwin record the lateralextent of the altered halo, whereas those at Jukes Prospectoccur at least 1 km above the north-plunging pluton (Fig.6.18).

Immediately adjacent to the granite at Mount Darwin isa thin (10—20 m) K-feldspar ± biotite (now chlorite) hornfelszone (Table 6.1). This grades outward into a 400 m wideK-feldspar + chlorite zone, which locally hosts coarsebreccias with magnetite ± tourmaline matrices (Jones, 1993;Wyman, 2001). The K-feldspar + chlorite zone grades intoa 300 m thick chlorite + magnetite zone (Wyman, 2001).A discontinuous wedge-shaped silicified zone separatesthe chlorite + magnetite zone from the outer K-feldspar +quartz zone. At Mount Darwin the K-feldspar + quartz zoneoccurs between 800 and 1000 m from the granite contact.At Jukes Prospect the K-feldspar + quartz zone is the centralaltered zone and is enclosed in chlorite, sericite and regionaldiagenetic-metamorphic zones. It is associated with Cu-Aumineralised rocks (Doyle, 1990; Large et al., 1996). Theperipheral sericite zone merges with the regional diagenetic-metamorphic albite + sericite zone (Wyman, 2001).

K-feldspar ± <20 K-feldspar ± chlorite Pervasive, hornfels

biotite hornfels (after biotite) ± quartz ±

sulfides

K-feldspar + 400 K-feldspar + quartz Intense Magnetite ± tourmaline fill in SiO2, AI2O3, K2O and net DG2

chlorite + chlorite ± sericite hydraulic breccias and veins mass gains

± hematite (after

magnetite) Na2O losses

Silicified ' Quartz ± sericite ± pyrite Strong to Texturally destructive, Large gains in SiO2 and DG3

± hematite intense cryptocrystalline and net mass

microcrystalline

K-feldspar + 200 K-feldspar + quartz Strong to Moderately texturally K2O, SiO2 and Fe2O3 DG4

quartz + sericite + chlorite + intense destructive, pervasive, gains

pyrite ± magnetite ± cryptocrystalline ,

chalcopyrite pseudomorphs plagioclase, Na2O losses

veins and vein envelope

Chlorite + 300 Chlorite + sericite + Moderate Moderate preservation, chlorite Fe2O3, MgO and K2O DG5

magnetite magnetite ± dolomite to intense pseudomorphs plagioclase, gains

± apatite ± pyrite, domainal replacement of

chalcopyrite veins groundmass and matrix SiO2, AI2O3, Na2O and

textures, infill in hydraulic net mass losses

breccias and veins, vein

envelope

Sericite Sericite ± chlorite ± Weak to Sericite partially to completely K2O gain and Na2O DG6

pyrite strong replaces plagioclase and

is disseminated in the CaO losses

groundmass and matrix

156 | CHAPTER 6

FIGURE 6.18 | Schematic cross-section of the Darwin Granite and altered

zones, illustrating relationships between the Mount Darwin and Jukes Prospect

alteration systems, and surface maps presented in the data sheets (Large et al.,

1996; Wyman, 2001).

Genesis of the alteration system

The proximity of intense hydrothermal alteration mineralassemblages to the granite contact and the northwardextending cupola region above the buried pluton supportthe interpretation that the hydrothermal system was drivenby heat from the intrusion (Eastoe et al., 1987; Large et al.,1996; Wyman, 2001). Overprinting relationships between K-feldspar, sericite and chlorite facies indicate multiple alterationstages in which low-temperature mineral assemblagesoverprinted initial high-temperature assemblages (Wyman,2001).

Initial sericite and chlorite alteration assemblages wereassociated with fracturing and vein formation around andabove the granite. As the fracture system evolved, the Jukeshydrothermal system became part of the discharge zone. Thewell-defined zones of hydrothermal alteration are interpretedto have formed from diffuse circulation of hydrothermal fluidthrough the volcanic rocks. Fluid access was enhanced byhydrothermal brecciation, and intense K-feldspar alterationand silicification was confined to the fracture zone above thegranite (Wyman, 2001).

Magnetite and tourmaline veins and breccias in the K-feldspar-rich zones immediately adjacent to the granitecontact, and in the centre of the Jukes alteration system,demonstrate that magmatic-hydrothermal fluids were exsolvedduring crystallisation (Large et al., 1996). The mass changes(i.e. gains in K2O, Fe2O3, Ba, Sr, Cu, Mo, W, Th and U) inthe altered zones adjacent to the granite are consistent withhydrothermal alteration of the volcanic rocks by magmaticfluid (Wyman, 2001). These fluids mixed with modified

seawater in reaction zones around the hotter portions of thedischarge zones to form K- and Fe-rich alteration assemblagesabove the granite (Wyman, 2001). The K-feldspar, chloriteand sericite zones in the Jukes area all show depletions inNa2O and CaO, which reflect the breakdown of plagioclaseand mafic minerals during alteration by modified seawater.

Large et al. (1996) suggested that the distribution,composition and zonation of alteration facies around theDarwin Granite, regional zonation of metals with respect tothe granite, distribution of Cu-Au-rich VHMS deposits in theMount Lyell field and pre-Tyndall Group timing of both thegranite and mineralisation support a genetic link between thegranite and these deposits. Furthermore, magnetite + apatite± pyrite veins in the Prince Lyell deposit are similar to thoseadjacent to and within the Darwin Granite and are consistentwith magmatic fluid contributing to their formation. Theseauthors proposed a model for the genesis of the Mount LyellCu-Au and related Pb-Zn-Cu massive sulfide deposits (Fig.6.19), which involves seawater convection deep into theCentral Volcanic Complex where it mixed with Fe, Cu, Au andP-rich magmatic fluids exsolved from the granite. The mixedmagmatic-seawater hydrothermal system produced alteredzones, magnetite veins (e.g. Jukes Prospect) and subseafloorCu-Au deposits (e.g. Prince Lyell) close to the granite, andcontemporaneous Pb-Zn-Cu massive sulfide deposits on theseafloor (e.g. Lyell-Comstock).

FIGURE 6.19 | Model of the zones of hydrothermal alteration and ore depositsin the Mount Lyell field and their relationship to the Darwin Granite, southernMount Read Volcanics, western Tasmania (after Large et al., 1996).

Weak, regional, pervasive albite + sericite alteration facies


DG1


Sample no. 143401

Alteration fades weak, regional, pervasive albite + sericite

Location Jukes Road

Formation Central Volcanic Complex

Succession Mount Read Volcanics

Volcanic facies massive, feldspar + quartz-phyric rhyolite

Relict minerals feldspar (3%, 2 mm), quartz (1%,<1 mm)

Relict textures micropoikilitic, porphyritic

Primary composition rhyolite

Lithofacies columnar jointed, massive

Interpretation sill

Alteration minerals albite + sericite + chlorite + hematite

Alteration textures recrystallised groundmass, sericitepseudomorphs after feldspar, pervasive-selective sericite > chlorite, disseminatedhematite

Distribution regional

Preservation good

Alteration intensity weak

Timing early

Alteration style regional diagenetic and metamorphic

Geochemistry

SiO2 76.17 Na2O 2.29 Rb 129 Zr 280

TiO2 0.31 K2O 3.66 Sr 22 Nb 13

AI2O3 13.30 P2O5 0.04 Ba 822 Y 39

Fe2O3 3.35 S 0 Cu 4

MnO 0.10 Total 100.00 Pb 20 Al 64

MgO o.67 (voLfree> Zn 39 CCPI 38

CaO 0.11 Th 21 Ti/Zr 6 .6

1 5 8 | CHAPTER 6

Intense, pervasive K-feldspar + chlorite alteration facies DG2

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

143278

intense, pervasive K-feldspar +chlorite

Mount Darwin


Mount Read Volcanics

massive, feldspar + quartz-phyricrhyolite

quartz + plagioclase (5%, 1-2 mm)

porphyritic

rhyolite

massive

unknown

quartz + K-feldspar > chlorite + sericite> hematite + pyriteplagioclase replaced by sericite,recrystallised groundmass of quartz+ K-feldspar > chlorite + sericite,disseminated hematite + pyrite,selective chlorite

local

none

intense

syn- to post-intrusion

proximal intrusion-relatedhydrothermal alteration


Geochemistry ^ 223

SiO2 71.35 Na2O 0.30 Rb 259 Nb 1g

TiO2 0.27 K2O 9.27 Sr 80 y 46

AI2O3 14.59 P2O5 0.04 Ba 2463

Fe2O3 2.96 S 0 Cu 8 A] g7

MnO 0.01 Total 99.72 Pb 31 c c p | 2J

M g O 0.92 (vol. free) Zn 89 T i / Z r 7 3

CaO 0.01 Th 25

Strong, foliated quartz + sericite + pyrite alteration facies


DG3

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution/zonation

Preservation


Timing

Alteration style

143237

strong, foliated quartz + sericite + pyrite

Slate Spur



feldspar + quartz-phyric rhyolite schist

quartz + plagioclase (15%, 1-5 mm)

porphyritic, microcrystalline, partlyspherulitic and possibly perlitic

rhyolite

foliated

unknown

quartz + sericite + pyrite + hematite

plagioclase replaced by sericite,

pervasive microcrystalline groundmassof quartz + sericite + pyrite, hematitestylolites, quartz overgrowths,schistosity

local

poor

strong


proximal intrusion-related hydrothermalalteration


Geochemistry

SiO2 81.77 Na2O 1.54 Rb 145 Zr 142

TiO2 0.17 K2O 4.51 Sr 35 Nb 11

AI2O3 10.23 P2O5 0.03 Ba 889 Y 33

Fe2O3 1.37 S 0 Cu 3

MnO 0.01 Total 99.92 Pb 16 Al 76

MgO 0.29 (vol. free) Zn 22 CCPI 20

CaO 0.01 Th 19 Ti/Zr 7.2

160 | CHAPTER 6

Intense, pervasive K-feldspar + sericite alteration facies DG4


Sample no. 143360

Alteration fades intense, pervasive K-feldspar + sericite

Location Jukes Road



Volcanic facies massive, feldspar + quartz-phyric rhyolite

Relict minerals plagioclase (2%, 2 mm), quartz (1 %, 1mm)

Relict textures micropoikolitic, porphyritic


Lithofacies columnar jointed, massive

Interpretation sill

Alteration minerals K-feldspar + quartz + sericite + chlorite +pyrite + magnetite

Alteration textures pervasive recrystallised groundmass ofK-feldspar + quartz + sericite > chlorite,cleavage defined by sericite, plagioclasealtered to K-feldspar > chlorite +sericite, K-feldspar overgrowths,chlorite pseudomorphs of feldsparmicrophenocrysts

Distribution/zonation local

Preservation moderate

Alteration intensity intense

Timing syn- to post-intrusion

Style intrusion-related hydrothermal alteration

Geochemistry

SiO2 73.74 K2O 7.60 Rb 161 Zr 257

TiO2 0.27 P2O5 0.04 Sr 38 Nb 11AI2O3 12.10 S 0 Ba 2449 Y 40

Fe2O3 5.51 Total 99.97 Cu 215MnO 0.03 (vol. free) Pb 17 Al 98

MgO 0.53 Zn 57 CCPI 42

CaO 0.01 Th 17 Ti/Zr 6.3Na2O 0.13

Strong, pervasive sericite alteration facies


DG5


Sample no. 143366

Alteration fades strong, pervasive sericite

Location Jukes Road



Volcanic fades massive feldspar + quartz-phyric

rhyolite

Relict minerals quartz (2%, 1 mm), plagioclase (2%,3 mm), hornblende (<1%)

Relict textures spherulitic, glomeroporphyritic


Lithofacies massive

Interpretation sill

Alteration minerals chlorite > K-feldspar + sericite > pyrite+ magnetite

Alteration textures pervasive K-feldspar + quartz +chlorite, chlorite + sericite + magnetitepseudomorphs after plagiociase,chlorite pseudomorphs afterhornblende, recrystallised spherulites

Distribution local



Timing syn- to post-intrusion

Alteration style intrusion-related hydrothermalalteration

Geochemistry

SiO2 72.77 Na2O 0.15 Rb 167 Zr 284

TiO2 0.37 K2O 5.06 Sr 18 Nb 13

AI2O3 13.23 P2O5 0.05 Ba 1323 Y 27

Fe2O3 7.11 S 0 Cu 234

MnO 0.04 Total 99.97 Pb 6 Al 97

MgO 1.15 (vol. free) Zn 114 CCPI 59

CaO 0.05 Th 19 Ti/Zr 7.8

162 | CHAPTER 6

Strong, pervasive sericite alteration facies DG6

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

143400

strong, pervasive sericite

Jukes Road



massive feldspar + quartz-phyric rhyolite

plagioclase (2%, 3 mm), quartz (1%,1 mm)

porphyritic, spherulitic?

rhyolite

massive

unknown

sericite + K-feldspar + quartz + hematite+ pyriteplagioclase replaced by sericite,recrystallised groundmass of sericite +K-feldspar + quartz, hematite stylolites,disseminated pyrite, weakly developedcleavage

local

poor

strong


intrusion-related hydrothermal alteration

Geochemistry

SiO2

TiO2

AIAFe2O3

MnO

MgO

CaO

75.23 Na2O 0.99 Rb 173 Zr 278

0.30 K2O 5.18 Sr 22 Nb 13

13.52 P2O5 0.05 Ba 1232 Y 41

3.67 S 0 Cu 4

0.02 Total 100.00 Pb 3 Al 86

0.99 (vol. free) Zn 35 CCPI 41

0.04 Th 21 Ti/Zr 6.5


I 163

7 | LOCAL HYDROTHERMAL ALTERATIONRELATED TO VHMS DEPOSITS

An understanding of the mineralogical and chemical zonationof hydrothermally altered rocks around submarine massivesulfide deposits is vitally important to both ore genesis studiesand to assist and focus mineral exploration. It has led to avast amount of literature on hydrothermal alteration aroundVHMS deposits, which various workers have summarised(e.g. Franklin et al., 1981; Barriga et al., 1983; Urabe et al.,1983;Lydon, 1984; 1988; Large, 1992; Madeisky and Stanley,1994; Barrett and MacLean, 1994b; Galley, 1995; Carvalhoet al., 1999; Sanchez-Espana et al., 2000, 2002; Gemmell andHerrmann, 2001; Large et al., 2001c).

Since the pioneering work in the 1950s and 1960s, whengeologists first recognised the critical link between volcanic-magmatic processes and massive sulfide genesis (e.g. Stanton,1955, 1959; Oftedahl, 1958; Gilmour, 1965; Horikoshi,1969), it became widely accepted that these deposits formon the seafloor from hydrothermal activity generated duringperiods of local quiescence between volcanic eruptive cycles(Sangster, 1972; Solomon, 1976; Franklin et al., 1983;Ohmoto and Skinner, 1983). The discovery of seafloorblack smokers and related sulfide chimneys on the present-day seafloor has further stimulated research and contributedto an improved understanding of ore forming processes inthe ancient deposits (Rona and Scott, 1993). Over the last15 years, many researchers have questioned whether allmassive sulfide deposits form by exhalation on the seafloor.Although it has been recognised for some time that stringerzones and the lower parts of some massive sulfides formed byreplacement (e.g. Large, 1997), several authors now considerthat replacement of particular volcanic units below theseafloor maybe be a key process for massive sulfide formation(e.g. Barriga and Fyfe, 1988; Khin Zaw and Large, 1992;Allen, 1994b; Bodon and Valenta, 1995; Hannington et al.,1999; Doyle and Allen, 2003).

This chapter is not a summary of VHMS ore genesis, butit highlights the features of alteration halos associated withVHMS deposits, particularly in the Mount Read province,western Tasmania, and fits the deposits and their alterationhalos in to the broad range of ore deposits that are foundin volcanic and volcano-sedimentary successions. In the laterpart of this chapter we provide descriptions, including datasheets depicting typical alteration facies, of examples from theMount Read province and Mount Windsor Subprovince in

eastern Australia that illustrate the range in alteration stylesand zonation associated with the spectrum of submarinevolcanic-hosted base metal ores.

7.1 COMMON FEATURES OF VHMSDEPOSITS

VHMS deposits display the following features:• They are hosted by submarine volcanic or volcano-

sedimentary successions.• They are the same age as the host volcanic succession

(i.e. the deposits are approximately synvolcanic and/orsynsedimentary).

• The host rocks vary from coherent to clastic volcanic orsedimentary facies and range in composition from basaltthrough andesite and dacite to rhyolite.

• Most deposits are hosted in thin volcaniclastic units(<100 m thick) between major volcanic formations.

• The economic parts of the deposits typically comprisemassive sulfide, principally pyrite, subordinate sphalerite,chalcopyrite and galena. The term massive implies greaterthan 80 wt% sulfides (Sangster, 1972).

• Massive sulfide lenses are commonly, but not always,aligned parallel to volcanic strata.

• Stringer (or stockwork) sulfide zones commonly underliethe massive sulfides and may contain economic Cu grades.

• Metal contents and metal ratios vary considerably. Depositsinclude Cu-rich, Au-rich, Cu-Zn, and polymetallic (Cu-Zn-Pb-Ag-Au) types, but all contain more Zn than Pb.

• Ore metals are typically vertically zoned within sulfidedeposits from Cu at the stratigraphic base to Zn, Pb, Ag,Au and Ba in general order towards the top. Nevertheless,there are many exceptions to this zonation pattern andsome deposits have no Ba.

• Intense hydrothermal alteration of the footwall volcanicrocks stratigraphically below the massive sulfide, tochlorite, sericite and quartz is common. By comparison,the hanging wall rocks are weakly altered or unaltered.Over 700 VHMS deposits have been recorded around the

world. They range in size from less than 100,000 tonnes to

1 6 4 | CHAPTER 7

over 510,000,000 tonnes (RioTinto, Iberian pyrite belt). Thetop 50 deposits, in terms of tonnes of contained Cu + Zn + Pbmetal, are listed in Figure 7.1. Most of these deposits are fromseven major VHMS provinces or districts (Fig. 7.2), fromoldest to youngest: Abitibi belt in Canada, Skellefte district inSweden, Mount Read province in Australia, Bathurst miningcamp in Canada, Southern Urals in Russia, Iberian pyrite beltin Spain and Portugal, and Hokuroku district in Japan.

Cu+Zn+Pb metal content (million tonnes)10 40

FIGURE 7.1Pb tonnes.

The 50 largest VHMS deposits in terms of contained Cu + Zn +

FIGURE 7.2 | Locations of the major VHMS provinces around the world.

7.2 | HYDROTHERMAL ALTERATIONHALOS ASSOCIATED WITH VHMSDEPOSITS

Hydrothermally altered zones proximal to VHMS depositsmay include footwall alteration pipes, stratabound alteredfootwall zones and altered hanging wall zones.

Previous studies (e.g. Franklin et al., 1983; Lydon,1988) emphasised the pipe-like hydrothermal altered zonesin the footwalls of many massive sulfide deposits. They arecommon in Archaean deposits in the Abitibi belt in Canada(e.g. Sangster, 1972) and in the Miocene Kuroko depositsof the Hokuroku district in Japan (e.g. Urabe et al., 1983).However, in other districts such as the Mount Read province,Mount Windsor Subprovince and Lachlan Fold Belt ineastern Australia, the Iberian pyrite belt, and the Bathurstmining camp, well-defined alteration pipes are less common,and stratabound altered footwall zones dominate (Large,1992). These two styles of altered footwall zones are describedbelow, with emphasis on the footwall alteration pipes, becausethey have received considerable attention from researchersand their alteration mineral zonation and genesis are betterunderstood.

Footwali alteration pipes

Figure 7.3 is a schematic cross-section of the geology, sulfideand alteration zonation related to a typical VHMS deposit.It is based on our understanding of the Hellyer deposit inthe Mount Read province (Gemmell and Large, 1992), butalso incorporates information on other Australian (Large,1992), Canadian (Franklin et al., 1981; Lydon, 1988; Lentzand Goodfellow, 1996), Japanese (Date et al., 1983; Urabeet al., 1983) and Spanish-Portuguese deposits (Barriga etal., 1983; Leistel et al., 1998; Sanchez-Espana et al., 2000).Immediately below the thickest part of the massive sulfideore the footwall alteration pipe, which may be oval in planbut is more commonly elongate along a synvolcanic fault,contains a concentric series of altered zones. These are, fromthe centre of the pipe outwards: siliceous core zone, chloritezone, sericite zone, and albite zone, which grades into least-altered volcanic rocks.

FIGURE 7.3 | Cross-section of idealised mineralisation and alteration zonationpatterns in a footwall alteration pipe beneath a typical VHMS deposit (modifiedafter Gemmell and Large, 1992; Lydon, 1997). (A) Sulfide mineral zones andgeology. (B) Hydrothermally altered zones.

Siliceous core zones

The siliceous core zones are composed of quartz + pyrite andquartz + pyrite + sericite ± chlorite assemblages (e.g. datasheets HE6, RB4, TH4, WT7, HR8, HN6 and HN7). Theymay not always be present, and have only been describedfrom a few deposits (e.g. Hellyer, Gemmell and Large, 1992;Brunswick No. 12 and other deposits in the Bathurst miningcamp, Zhang et al., 2003). Siliceous core zones are the mostintensely altered rocks in the centre of the pipes, and arecommonly intersected by networks of pyrite + chalcopyritestringer veins (Fig. 7.3). All primary rock textures withinthese zones have been completely destroyed due to theintensity of alteration. In some cases, quartz-rich alterationassemblages have overprinted earlier chlorite-rich alterationassemblages creating pseudobreccia textures. Mass-changecalculations indicate that gains of 50-100 g/100 g, mainlydue to Si addition, are common within siliceous core zones(Gemmell and Large, 1992, Fig. 11). Lentz and Goodfellow(1996) reported SiO2 gains of up to 300% in the siliceouscore zone in the centre of the alteration system below theBrunswick No. 12 massive sulfide deposit. The siliceous corerepresents the zone of maximum hydro thermal fluid flow andhighest temperatures.

Chlorite zones

Chlorite zones are dominated by chlorite (>50 wt% andcommonly >80 wt%), with subordinate quartz + pyrite +sericite ± carbonate (e.g. data sheets HE4, RB5 and HR7).

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS | 1 6 5

These zones are typically fine grained, dark and massive,preserving no volcanic or sedimentary textures. In manydeposits these zones host pyrite + chalcopyrite stringerveins (e.g. Woodlawn deposit in eastern Australia and mostNoranda district and Iberian pyrite belt deposits). Chloritezones are commonly deformed during tectonic events,resulting in chlorite schist zones, which may be strung-outor dislocated from the massive sulfide ore (Sangster, 1972).Studies of chlorite composition from Canadian, Australianand Japanese deposits indicate that the inner chlorite zonesare dominated by Mg-rich chlorite, with a general increasein Fe/Mg ratio passing from the inner chlorite zone to theouter edge of the sericite zone (e.g. Riverin and Hodgson,1980; Urabe et al., 1983; Paulick et al., 2001). Nevertheless,reverse trends have been recorded, where chlorite becomesmore Fe-rich towards the core (e.g. Eastoe et al., 1987; Lentzet al., 1997). Variations in chlorite composition are discussedin more detail in Section 4.2.

Sericite zones

Sericite zones surround the inner chlorite zones and arecharacterised by assemblages of sericite + chlorite + quartz +carbonate + pyrite (e.g. data sheets HE3, RB3, WT3, WT6,HR6 and HN5). At Hellyer, rocks in the sericite zone arestrongly to intensely altered, with up to 70 wt% sericiteand sparse relict primary textures. In other deposits, thealteration intensity in the sericite zone decreases towards theouter margin, where altered rock grades into the least-alteredfootwall rocks. In many cases, the sericite zones are laterallyextensive and merge with stratabound altered zones awayfrom the central pipe (e.g. Mount Chalmers, Large and Both,1980). Minor disseminated sphalerite or stockwork Zn mayoccur in the sericite zone, whereas Cu-enrichment is morecommon in the chlorite zone.

Albite zones

Some authors have described weakly altered zones of albite+ chlorite ± sericite that surround the main sericite zones(e.g. Iijima, 1974; Green et al., 1981; Urabe et al., 1983;Relvas et al., 1997; Goodfellow and McCutcheon, 2003).Although albite zones have not been widely described oraccepted in all districts they are discussed here because oftheir significance to mineral exploration. It may be difficult todistinguish hydrothermal albite facies from the backgrounddiagenetic alteration facies, as within albite zones primaryvolcanic textures are commonly preserved. In the Hokurokudistrict this zone has an albite + sericite + chlorite assemblage(Iijima, 1974; Urabe et al., 1983). In some deposits of theIberian pyrite belt an outermost halo of Na-sericite has beenrecognised (Relvas et al., 1997). In the Bathurst mining camp,Goodfellow and McCutcheon (2003) described an outermostaltered zone of albite + Mg-rich chlorite surrounding thefootwall alteration pipe. They described an increase in patchyalbite, which replaced K-feldspar phenocrysts, in proximityto the pipe. Similar albite-altered rocks have been recognisedadjacent to the Hercules footwall alteration pipe in the MountRead province (Large et al., 1996). Further research is required

166 | CHAPTER 7

to characterise the features of these outermost albite + chlorite+ sericite altered zones and to distinguish them from regionaldiagenetic albite alteration facies.

Variations in alteration zonation

The three main hydro thermally altered zones (siliceous core,chlorite and sericite) are not present in all footwall alterationpipes beneath VHMS deposits, and in some cases additionalzones, such as carbonate or talc zones, exist. Some variationsfrom the idealised alteration pipe model outlined in Figures7.3 and 7.4A are:

FIGURE 7.4 | Different patterns of mineral zonation in footwall alteration pipes.(A) Generalised model. (B) Hokuroku district model. (C) Noranda district andIberian pyrite belt model. (D) Mattagami district model.

The Kuroko deposits do not have an inner chlorite zone.Shirozu (1974) described intensely altered volcanic rocksbelow the massive sulflde ore and surrounding the siliceousCu-stockwork ore, as strongly silicified with abundantsericite and very little chlorite (Fig. 7.4B).Deposits in the Noranda district and the Iberian pyrite beltcommonly have intense chlorite core zones surrounded bysericite zones (Fig. 7.4C), but without siliceous core zones(Franklin et al., 1983; Carvalho et al, 1999; Sanchez-Espana et al., 2000).In the Iberian pyrite belt, Relvas et al. (1997) recogniseda Na-bearing sericite (paragonite) zone extending beyondand above the main sericite zone at both the Neves Corvoand Aljustrel deposits.In the Mattagami district in Canada, the intensely alteredcentral core of the footwall alteration pipe is talc-rich (Fig.7.4D) and surrounded by chlorite and sericite zones (e.g.Large, 1977; Roberts and Reardon, 1978).In several eastern Australian deposits, Mg- and/or Fe-bearing carbonates are common in the alteration mineralassemblages (Large et al., 2001c). At Hellyer a chlorite+ dolomite zone occurs below the massive sulflde nearthe top of the chlorite zone (Fig. 7.4A, data sheet HE5,Gemmell and Large, 1992). A more massive dolomitezone is developed at the western margin of the alteredstringer pipe at Mount Chalmers (Large and Both, 1980).Carbonate-rich zones are common in the footwall of manyIberian pyrite belt deposits, either marginal to the massivesulflde or distributed throughout the footwall alterationsystems (e.g. Rio Tinto, Williams et al., 1975, Solomon etal., 1980; La Zarza, Strauss et al., 1981; Tharsis, Tornos etal., 1998).Lydon (1988, Fig. 8) included an Fe-oxide zone at the top ofthe pipe below the massive sulflde in his footwall alterationpipe model, compiled from a number of deposits, but thereare few examples of this facies.

Stratabound altered footwall zones

Many massive sulfide deposits, possibly half, do not havefootwall alteration pipes, but are underlain by stratabound orsemi-conformable altered zones, which extend laterally for upto several kilometres away from the deposits (Figs 3.16B, Cand 7.5). These stratabound zones may extend for between 30and several hundred metres below the massive sulfide. Theyare typically developed around sheet-like deposits, and aremainly associated with Zn-Pb-rich deposits (e.g. Rosebery,Scuddles and Bathurst mining camp deposits). Sangster(1972) and Goodfellow and McCutcheon (2003) considerstratabound altered footwall zones to originally have beenpipes that were sheared and transposed parallel to stratigraphyduring tectonic deformation. However, this does not explainstratabound footwall zones in weakly to moderately deformedvolcanic successions, such as the Mount Read province andIberian pyrite belt. In the Iberian pyrite belt, Tornos (in press)notes that although many previous studies have claimed pipe-like morphologies to the footwall stockworks and alteredzones (e.g. Costa et al., 1995; Carvalho et al., 1999; Saez etal., 1999), recent studies have defined irregular to strataboundmorphologies for the footwall alteration associated with most

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS | 167

deposits (e.g. Tharsis, Tornos et al., 1998; Neves Corvo,Relvas et al., 2000).

Stratabound altered footwall zones have similar alterationmineral assemblages to footwall alteration pipes, but the zonesare distributed parallel to stratigraphy, rather than at rightangles. In some cases, the siliceous core and chlorite zones areconfined to the immediate footwall of the thickest Cu-richpart of the massive sulfide lens (e.g. Rosebery and Thalanga).In contrast to the pipes, the sericite zones of strataboundaltered footwall zones are the volumetrically dominantzones being both laterally and vertically extensive. Massivecarbonate zones are more common in stratabound thanpipe-like alteration systems (Large et al., 2001c), typicallyoccurring immediately along strike from the massive sulfideore lenses (e.g. Rosebery, Thalanga and Mount Chalmers).From an exploration perspective stratabound altered footwallzones typically present broader targets than footwall alterationpipes; however, they tend to be more diffuse and thuscreate challenges when searching for the associated VHMSdeposits.

Altered hanging wall zones

Compared to footwall alteration, hanging wall alteration istypically less intense and therefore has not received muchattention in the ore-deposit literature. Visible hanging wallalteration mineral assemblages are commonly sericite-richand restricted to a few metres above the massive sulfide ore.However, detailed petrographic and geochemical studies haveextended some altered zones to several tens of metres intothe hanging wall. There are a number of exceptions to thegenerally limited altered hanging wall zones.

Copper-Au-rich VHMS deposits that form by replacementbelow the seafloor may exhibit extensive sericite-rich alteredhanging wall zones (Fig. 7.6A). Mount Lyell in the MountRead province and Highway-Reward in the Mount WindsorSubprovince are good examples of subseafloor sericite zonesand are described in detail in Sections 7.7 and 7.10.

In stacked ore systems, such as Millenbach and Amuletdeposits in the Noranda district, and Que River in the MountRead province, the lower ore body in the stack has an intensealtered hanging wall zone, similar to that in the footwall (Fig.7.6B). This is due to hydrothermal fluids moving through thelower ore lens and hanging wall on their way to depositingthe upper ore lenses.

Significant altered hanging wall zones may have developedin situations where the hanging wall volcanic or sedimentaryunits were deposited while the massive sulfide was still formingon the seafloor. This is the case for the Hellyer deposit, whichhas a hanging wall alteration plume that comprises a coreof fuchsite + carbonate (e.g. data sheet HE10), surroundedby successive halos of chlorite + carbonate, quartz + albite,and finally patchy sericite (Fig. 7.6C, Gemmell and Fulton,2001). The interpretation that the hanging wall alterationzones formed after the seafloor massive sulfide is not in doubtbecause of sulfide and barite clasts in the directly overlyingvolcaniclastic debris-flow unit (McArthur and Dronseika,1990;Sharpe, 1991).

Although primary plagioclase destruction is a key processin the hydrothermal alteration associated with VHMS

FIGURE 7.5 | Examples of stratabound altered footwall zones (modified afterAshley et al., 1988; Large, 1992; Large et al., 2001c). (A) Scuddles, WesternAustralia, in plan view. (B) Teutonic Bore, Western Australia, in cross-section.(C) K lens at Rosebery, western Tasmania, in cross-section. (D) Thalanga,Queensland, in cross-section. Abbreviations are: FW = footwall and HW =hanging wall.

deposits, there are a few examples where albite zones exist inthe hanging wall of the deposit. At the Henty gold depositalbite + quartz is a common hanging wall alteration mineralassemblage (see Section 7.8, data sheet HN3), and at Hellyeralbite forms one of the altered zones in the hanging wallalteration plume (Fig. 7.6C, data sheet HE8).

In addition to these examples of obvious alteredhanging wall zones, some recent detailed lithogeochemical

1 6 8 | CHAPTER 7

FIGURE 7.6 | Examples of cross-sections through altered hanging wall zonesassociated with VHMS deposits. (A) Cu-Au deposit (modified after Large et al.,2001c). (B) A stacked ore system (after the Millenbach deposit, Knuckey etal.,1982). (C) Hellyer hanging wall alteration plume (modified after Gemmell andFulton, 2001).

studies have defined subtle geochemical halos in otherwiseleast-altered hanging wall volcanic rocks, which may extendseveral hundreds of metres above the massive sulfide ore(Large et al., 2001b). For example, at Rosebery a hanging wallalteration halo can be defined using three geochemicalparameters that may be applied during exploration: (1) whole-rock Ba/Sr ratio, which outlines a halo extending about 100 minto the hanging wall; (2) Mn content of carbonate, whichis anomalous over the same interval; and (3) whole-rock Tlcontent, which forms a halo that extends over 200 m intothe hanging wall volcanic rocks. More detail of the Roseberyalteration system is provided in Section 7.6.

Chemical reactions and mass changes

Footwall alteration results from the reaction of hydrothermalfluid (principally composed of heated seawater) with volcanicrocks. The temperature of the fluids that form VHMS depositsare estimated to vary from about 200° to 350°C based onfluid inclusion evidence (Pisutha-Arnond and Ohmoto,1983; Khin Zaw et al., 1996), the study of present day blacksmoker systems (Goldfarb et al., 1983) and thermodynamiccalculations of mineral stabilities (e.g. Sato, 1973; Large,1977; Ohmoto et al., 1983). Fluid salinities approximate thatof seawater, although values of up to four times seawater havebeen recorded (de Ronde, 1995; Solomon et al., 2002). ThepH varies from about 3 to 7 based on mineral assemblage,thermodynamic considerations and measurements at blacksmoker vents (Huston and Large, 1989; Scott, 1997).

The principal result of interaction of this hot, mildly acidicto neutral fluid with volcanic rocks as it ascends towards theseafloor is the breakdown of feldspars and volcanic glass, andtheir replacement by sericite, quartz, chlorite and carbonate.Petrographic evidence of these reactions is shown in Figure2.5 which depicts a series of progressively hydrothermallyaltered pumice-rich rocks from the footwall to the Herculesdeposit in the Mount Read province.

Reactions that describe these footwall alteration processesmay include reaction R7.1 (from Sanchez-Espana et al., 2000)in the sericite zone:

3NaAlSi3O8 + K+ + 2H+

albite-» KAl3Si3O10(0H)2 + 6SiO2 + 3Na+

sericite quartz(R7.1)

This reaction is typical of sericite replacing albite in theouter part of the alteration system. The reaction involves again in K from hydrothermal fluid and loss of Na in the rockas albite is replaced. Silica is conserved by the deposition ofquartz. Overall, the reaction leads to an increase in fluid pHdue to the consumption of H+. Reactions in the sericite zonealso involved sericitisation of K-feldspar and plagioclase inaddition to albite.

In the chlorite zone, there are two potential reactions(Pisutha-Arnond and Ohmoto, 1983):

Note, these reactions can be written with H4SiO4(aq) orSiO2 (quartz) + 2H2O.

Both of these reactions involve the addition of Mg andFe2+ from the fluid and the loss of alkalies (Na or K) and Hfrom the rock. However, the replacement of albite by chlorite

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS I 1 6 9

involves Si loss as H4SiO4, compared to the replacement ofsericite by chlorite, which involves Si gain. Mass transfercalculations in chlorite zones invariably indicate significantloss of Si (e.g. Gemmell and Large, 1992; Barrett andMacLean, 1994b) suggesting that reaction R7.2 is the keyreaction. Also both of these chlorite replacement reactions(R7.2 and R7.3) involve a release of H+ to the fluid and willcause an increase in fluid acidity. This means that continuedchloritisation of volcanic rocks over an extensive area mightproduce moderately acidic fluids, which may subsequentlycause silicic (quartz + sericite), or in the extreme case, argillic(pyrophyllite ± kaolinite ± sericite) alteration assemblages up-flow from the chlorite zone. Examples of this occur at theNeves Corvo and Lagoa Salgada deposits in the Iberian pyritebelt where Relvas et al. (1994, 1997) have reported dombassiteand pyrophyllite in the chlorite-rich central stockwork zonesbelow the massive sulflde ores. Other reactions in the chloritezone, not listed here, include chloritisation of mafic mineralssuch as biotite and amphibole.

In the siliceous core zone at the centre of the hydro thermalsystem, three reactions are proposed:

Quartz is the main alteration mineral in this zone andis associated with minor sericite and chlorite. Mass balancecalculations indicate that considerable SiO, gains. Aluminiumis commonly immobile (Gemmell and Large, 1992; Barrettand MacLean, 1994b), but is significantly diluted by largeSi gains. Quartz is probably deposited in the siliceous corezone according to reaction R7.4, as this involves mass gain ofSi without loss of Al. Enrichment of Si in the hydrothermalfluid may be due to leaching of Si from the footwall volcanicrocks during chloritisation (reaction R7.2). Leached Si is thendeposited in the hydrothermal vent immediately below theseafloor. Silica deposition may be caused by rapid conductivecooling (Fournier and Potter, 1982), mixing with seawater,or intense fluid-rock interaction at high-fluid-rock ratios.Replacement of both chlorite and albite by quartz (reactionsR7.5 and R7.6) requires a strongly acidic fluid and resultsin loss of Al as Al(OH)3(aq). Aluminium mobility of thistype is rare in VHMS systems, but may occur in intenselysilicified zones associated with acid alteration. Mass balancecalculations suggest this was the case in the siliceous core zoneof the Henty volcanogenic gold deposit (Callaghan, 2001, seealso Section 7.8).

In summary, by writing simple chemical reactions todescribe replacements in the major altered zones, it is possibleto gain some idea of the chemical processes, elemental gainsand losses, and variations in pH of the hydrothermal fluid.

Most reactions in the altered footwall zone involve thebreakdown of feldspar and loss of Na (and usually Ca) to thefluid. Major gains include K in the sericite zone, Mgand Fe inthe chlorite zone and Si in the siliceous core zone. In additionto Na and Ca, other losses include Si in the chlorite zone,and very rarely Al in the siliceous core. The fluid pH does notshow any systematic unidirectional change. Initially mildlyacidic fluids will become less acidic during sericitisation, butmore acidic during chloritisation. Intense siliceous core zonesmay be related to rapid cooling, fluid mixing or intense fluid-rock interaction of Si-saturated fluids during the peak of thehydrothermal activity.

Alteration box plot trends in altered footwallzones

The AI-CCPI Alteration box plot (Section 2.5) is a simpleway of tracking whole-rock compositional changes andrelating them to alteration mineralogy and position in thealtered footwall zones. In Figure 7.7, the fields of the majoraltered zones are shown on the Alteration box plot in relationto two alteration indices, Al and CCPI. Line AD representsan array of altered felsic volcanic rock samples passing fromthe outer edge of the altered footwall zone into the core zoneproximal to massive sulfide ore. Line ED represents a similarsample array for mafic volcanic rocks.

Let us first consider line AD. Point A represents a rhyoliticrock outside the altered zone. Sericite and weakly chloritealtered rocks in the margins of the altered zone increase theAl due to Na and Ca depletion, whereas the CCPI remainsrelatively constant. Altered samples plot progressively alongthe AB segment of the trend toward the plotted position ofsericite (phengite) on the perimeter of the Alteration box plot.In the inner part of the sericite zone, the Al is commonlygreater than 90 and the trend becomes vertical due to a strongincrease in the CCPI caused by gains in Fe and Mg related to

FIGURE 7.7 | The AI-CCPI Alteration box plot showing trends for altered

footwall zones. These data are based on case studies presented in Large et al.

(2001a).

170 | CHAPTER 7

increasing pyrite and chlorite in the rock (segment BC). Thefinal segment CD represents the chlorite zone, where CCPIreaches its maximum (80-100), due to the abundance ofchlorite and pyrite proximal to massive sulfide ore. Althoughquartz is not plotted on the Alteration box plot, due to theabsence of SiO2 from the two alteration indices, samples fromthe siliceous core zone commonly plot in the CD segmentdue to the presence of minor chlorite and pyrite. If carbonateis present in the chlorite zone, as at Hellyer and Thalanga, thesamples typically plot between F and D.

In cases where the footwall comprises mafic volcanicrocks, ED is the common trend from the edge to centre ofthe footwall alteration pipe. This difference is due to the factthat mafic rocks generally have higher Fe and Mg contents,and thus greater initial values of CCPI compared with felsicrocks (Fig. 2.10). Consequently, chlorite is generally moreabundant in the outer sericite zone and the combination ofincreasing chlorite and pyrite gives a trend along ED towardthe core of the footwall alteration pipe (e.g. Hellyer, Section7.5).

The genesis of footwall alteration pipes

The presence of tightly constrained altered zones, in acircular or more commonly elongate pipe, suggests thathydrothermal fluids were focussed along synvolcanic faults orfault intersections, and massive sulfide deposition occurredwhere the faults intersected the seafloor. The zonation in thefootwall alteration pipe is commonly interpreted to reflect adecreasing thermal gradient away from the fluid conduit (e.g.Large, 1977; Riverin and Hodgson, 1980).

up, and becomes more acidic due to fluid-rock interaction,metals are leached from the volcanic succession (e.g. Kajiwara,1973; Spooner and Fyfe, 1973; Solomon, 1976; Large, 1977;Ohmoto, 1996). Alternatively, metal-rich magmatic fluid maybe derived from the crystallisation of a magma, which is alsothe source of volcanism (e.g. Urabe and Sato, 1978; Henleyand Thornley, 1979; Sawkins and Kowalik, 1981; Stanton,1985, 1990). For supporting evidence and relative merits ofthese two models the reader is referred to recent discussionsby Lydon (1996) and Ohmoto (1996).

Recent research suggests that both fluids and metals areprobably derived from magmatic and seawater sources (e.g.Fig. 7.8C). Distinguishing criteria for the source includesthe deposit style, proximity to volcanic centres, alterationmineralogy, metal ratios of the deposits and the salinityand composition of primary fluid inclusions. For example,Large (1992) suggested that relatively soluble chloride-metalcomplexes, such as Zn, Pb and Ag, are probably derivedprincipally from seawater leaching of the volcanic succession,whereas the less soluble metals, such as Cu, Bi and Sn, maybe sourced directly from the magma chamber. Gold couldbe derived either by seawater leaching of volcanic rocks as abisulfide-Au complex, or directly from magma as a chloridecomplex or in a volatile phase (Fig. 7.8C). Goodfellow andMcCutcheon (2003) proposed a similar dual metal sourcefor the massive sulfide deposits of the Bathurst mining camp,with the largest deposits having a major magmatic-metalcomponent. Recently Solomon et al. (2004) have comparedthe salinity and composition of fluid inclusions in the stringerzones of the Hellyer VHMS deposit with those of porphyryCu deposits, and concluded that the metals at Hellyer had amagmatic source.

Source considerations Fluid-rock interaction in the alteration pipe

There are two competing models for the source of fluids The concept that the alteration zonation in footwall alterationand metals that form VHMS deposits (Fig. 7.8A and B): (1) pipes is a function of decreasing temperatures and fluid-rockevolved seawater, and (2) magmatic fluid. The first model ratios has recently been tested by Schardt et al. (2001) usinginvolves seawater convecting through a volcanic succession a thermodynamic model of fluid-rock interaction betweenabove an intrusion or magma chamber. As the seawater heats heated evolved seawater and an andesitic precursor (Fig.

FIGURE 7.8 | Models for fluid flow and metal source in VHMS hydrothermal systems (after Large, 1992). (A) Metals are derived from deep seawater leaching ofthe volcanic succession and basement. (B) Metals are sourced directly from the magmatic vapour plume, with no significant leaching of volcanic rocks. (C) A mixtureof volcanic and magmatic sources, with low-solubility metals (i.e. Cu and Au) provided from magma and high-solubility metals (i.e. Zn, Pb and Ag) from seawaterleaching of volcanic rocks.

7.9). This modelling was based on geochemical data fromthe Hellyer deposit. The classical sequence of altered footwallzones observed in many VHMS deposits (from the core tothe margin of the pipe: quartz —* chlorite —> sericite) wasreproduced by simulating the reaction between a 250—350cCfluid, with a pH of 4.5—5.0, and andesite under conditionsof decreasing fluid-rock ratio and temperature. Simulatedcooling from 350° to 100°C reproduced the full range offootwall alteration mineral assemblages. The pH of the fluidshowed little variation, from 4.5 to 4.0 (Schardt et al., 2001).Mg-rich chlorite formed in the inner chlorite zone, and Fe-rich chlorite developed in the outermost part of the sericitezone, similar to the pattern observed at many massive sulfidedeposits. This modelling was carried out using a Mg-bearing


fluid, with the assumption that a component of seawater-derived Mg was incorporated into the fluid at depth.

The modelling has shown that sericite zones form attemperatures below 250°C from the reaction of andesiticrocks with mildly acidic solutions (pH = 4.0—4.5). ExtensiveMg-chlorite zones are favoured by higher temperatures (250—300°C) and less acidic fluids (pH = 4.5-5.5). At lower pH,kaolinite and pyrophyllite are likely to develop in the sericitezone. At higher pH and lower temperatures (<200°C), K-feldspar is developed at the outer margin of the sericite zoneand in least-altered andesitic rocks (Schardt et al., 2001).Although carbonate alteration was not taken into account bythis modelling, it is likely that chlorite + carbonate assemblages,such as those developed adjacent to massive sulfides or at the

FIGURE 7.9 | Thermodynamic model of fluid-rock interaction between heated evolved seawater and Hellyer andesite(modified after Schardt et al., 2001). (A) Modelled mineralogical variations resulting from fluid-rock interaction withdecreasing temperatures. (B) Schematic representation of simulated water-rock interaction as a function of temperature.

172 | CHAPTER 7

periphery of many footwall alteration systems, are indicativeof more alkaline conditions. These may develop where hot,near-neutral hydrothermal fluids have mixed with and heatedseawater, leading to saturation of carbonate at the margins ofthe hydrothermal up-flow zones (Large et al., 2001c).

A model for the development of footwall alteration pipes

Using the results of the thermodynamic modelling of Schardtet al. (2001), numerical fluid-flow modelling by Yang andLarge (2001), and previous thermodynamic modellingof metal-sulfide growth by Huston and Large (1989), it ispossible to speculate on the progressive development ofsubseafioor zoned alteration pipes (Fig. 7.10).

Stage 1 (Fig. 7.10): Initial hydrothermal fluid flow isupwards along a sub-vertical permeable fault zone towardsthe seafloor. As the rising hydrothermal plume approacheswithin 1 km of the seafloor, secondary near-surface seawaterconvection above the plume head may enhance normaldiagenetic reactions in volcanic rocks adjacent to the fault,causing increased formation of zeolites, smectites and Mg-richchlorite, and albite replacement of primary feldspars. Duringongoing diagenesis and subsequent metamorphism this willproduce an outer albite zone (albite + sericite + chlorite),which is commonly difficult to distinguish from regionaldiagenetic and metamorphic mineral assemblages.

Stage 2 (Fig. 7.10): As low-temperature and mildly acidichydrothermal fluids, (T<250°C, pH = 4.0-4.5) continueto move upwards to the seafloor, sericite-rich alterationoverprints the early albite zone and expands out from thefluid conduit to form a pipe-like sericite zone. During thisstage sphalerite + galena + pyrite massive sulfides deposit onthe seafloor above the sericite zone. Minor pyrite + sphalerite+ galena stringer mineralisation may also occur in the core ofthe sericite zone at these temperatures (Eldridge et al., 1983).Convective near surface reflux of seawater leads to an albite +chlorite zone extending laterally into the least-altered volcanicrocks surrounding the sericite zone.

Stage 3 (Fig. 7.10): As the hydrothermal system intensifiesand the temperature of the discharging fluid rises above250°C, Mg-chlorite is stabilised adjacent to the main conduitand a chlorite zone develops in the core of the footwallalteration pipe. Under these high-temperature conditions,the fluid is capable of carrying significantly more Cu (e.g.Huston and Large, 1989), which is deposited in stringer veinsin the central chlorite zone and in the base of the massivesulfide. Within the massive sulfide mound, Cu progressivelydisplaces Zn upwards. Eldridge et al. (1983), Huston andLarge (1989) and Hannington and Scott (1989) describedthis zone refining process. As the alteration system evolves,the pH of the fluid initially increases during formation ofthe sericite zone (from 4 to 5.5) due to consumption of H+

(reaction R7.1), and subsequently falls during the formationof the chlorite zone (reactions R7.2 and R7.3).

FIGURE 7.10 | Model for the evolution of the footwall alteration pipe in a mound-style massive sulfide deposit. Stage 1: an initial low-temperature hydrothermal system produces an albite zone. Stage 2: increasing temperature results in the development of the sericite andZn + Pb-rich sulfide zones. Stage 3: higher temperatures produce the chlorite and Cu + Zn +Pb-rich sulfide zones. Stage 4: maximumtemperatures and low pH result in the siliceous core and Cu + Pb + Zn-rich sulfide zones. This model does not apply to all VHMSdeposits, some of which may form in brine pools (Solomon and Groves, 1994; Solomon and Quesada, 2003).

Stage 4 (Fig. 7.10): With increasing temperature (300° to35OCC) during chloritisation, Si is continually leached fromthe volcanic rocks (reaction R7.2) and the fluid becomessupersaturated in Si. Consequently, quartz is precipitated(reaction R7.4) in the upper-central part of the alterationpipe, forming a siliceous core zone. Maximum metalprecipitation in both the stringer zone and massive sulfideis commonly associated with this stage. Subsequently, thehydrothermal system wanes and collapses with an influx ofheated near-surface seawater that leads to overprinting bylower temperature mineral assemblages, which are commonlydominated by carbonates or barite, depending on theoxidation level of the overlying water column.

Some previous workers have suggested that the zonationin alteration pipes, from Mg-Fe chlorite in the core to sericiteat the margins, relates to the entrainment and mixing ofMg-bearing seawater with a Mg-poor hydrothermal fluidbelow the massive sulfide (e.g. Roberts and Reardon, 1978;Lydon and Galley, 1986). However, Riverin and Hodgson(1980) suggested that the presence of Mg-rich chlorite in thecentral and most intensely altered zone of the alteration pipe,and the abundance of Mg-chlorite in pyrite + chalcopyriteveins in the stringer zone, is consistent with Mg derivedfrom hydrothermal fluid rather than seawater. In the modeloutlined in Figure 7.10, we have assumed the hydrothermalfluid is Mg-bearing, possibly either due to entrainment ofseawater at considerable depth (>1 km) below the seafloor ordue to the leaching of Mg from mafic volcanic rocks deep inthe volcanic succession.

Although near-surface entrainment of seawater isconsidered to be important in stages 1 and 2 of our model,it is likely to result in an increase in the rate and consequentgrade of diagenetic alteration. This would lead to Na-Mgmetasomatism and albite + chlorite formation at the marginsof the alteration pipe, rather than Mg metasomatism andchlorite development within the core of the pipe, as previouslyproposed (cf. Franklin et al., 1981).

In the thermodynamic modelling of Schardt et al. (2001),temperature and pH were shown to be the principal factorscontrolling the balance between chlorite and sericite zones inthe footwall alteration pipe. However, two other factors alsoneed consideration: (1) the composition of the immediatefootwall volcanic rocks, and (2) the initial chemistry of themodified seawater as it rises up the conduit (e.g. Large, 1977).In the first case, particularly at low fluid-rock ratios, chloritealteration is favoured in mafic host rocks and sericite alterationin felsic host rocks. However, at high fluid-rock ratios typicalof the central parts of the hydrothermal pipe, the alterationmineral assemblage is controlled by fluid chemistry rather thanrock chemistry. In the second case, seawater-rock interactionsin deep rhyolite-dominated footwall volcanic successions,similar to those in the Hokuroku district and southern MountRead province, will generate modified seawater hydrothermalsolutions enriched in K and Si, but generally depleted inMg, Fe and Ca. These fluids will result in sericite zones asthey approach the seafloor. In contrast, footwall volcanicsuccessions dominated by andesite and basalt, similar to thosein the northern Mount Read province and the Abitibi belt,will generate evolved seawater fluids enriched in Mg, Fe andCa, with lesser K and Si, and are more likely to develop zonedchlorite-sericite alteration pipes.


In summary, the footwall alteration mineral assemblagesin VHMS systems are probably controlled by three factors:(1) the initial composition of the convective seawater-dominated ore fluid, which is constrained by the relativeabundance of mafic versus felsic volcanic rocks deep in thesuccession; (2) the temperature and pH regime during fluid-rock interactions in the footwall discharge zone (i.e. lowertemperature, acidic conditions favour sericite development,whereas higher temperature and/or more neutral pHconditions favour chlorite formation); and (3) the compositionof the immediate footwall host rocks.

Genesis of stratabound altered footwall zones

Stratabound altered footwall zones (e.g. Rosebery, Scuddlesand Teutonic Bore; Fig. 7.5) are interpreted to result fromhydrothermal-fluid flow parallel to volcanic strata (Fig. 7.11),rather than at right angles to the stratigraphy as in the casefor footwall alteration pipes. Alteration pipes are commonlydeveloped in relatively impermeable footwall volcanic rocks(e.g. the coherent or clastic facies of lavas and synvolcanicintrusions) where fluids are focussed along sub-vertical syn-volcanic faults (Fig. 7.10). In contrast, stratabound alteredfootwall zones are more commonly developed in volcanicrocks with moderate- to high-stratal permeability (e.g.volcaniclastic facies such as pumice breccia and volcanic

FIGURE 7.11 | Genetic models for the formation of stratabound alteredfootwall zones related to VHMS mineralisation. Fluid flow below and parallelto the seafloor and stratigraphy is controlled by the distribution of permeablevolcanic facies (e.g. volcaniclastic units), or impermeable cap-rocks (e.g. sillsor lavas). (A) Stratabound subseafloor replacement mineralised and alteredzones (e.g. Mount Lyell deposit, Mount Read province and TAG deep Cu zone,Middle Valley, Juan de Fuca Ridge). (B) Stratabound ore lens and alteredzones confined below an impermeable volcanic unit such as a sill (e.g. K lens atRosebery, Mount Read province).

1 7 4 | CHAPTER 7

sandstone). In high-permeability rocks, hydrothermal fluidsmove laterally along the strata, sub-parallel to the seafloor,and metals are deposited due to the mixing of hydrothermalfluids with seawater and/or cooling. In these cases, fluids arepoorly focussed and alteration tends to be of lower intensityand greater in lateral extent (Fig. 7.11). Sericite-altered rocksdominate stratabound altered footwall zones, and chloriticand siliceous zones are restricted to the immediate proximityof massive sulfides, where temperatures and fluid-rock ratioswere at a maximum. At K lens in the Rosebery deposit,Allen (1994b) proposed that rising hydrothermal fluids wereconstrained to flow laterally below an impermeable quartzporphyry sill. As a result, stratabound massive sulfide andassociated stratabound altered zones developed beneath thesill by replacement of more permeable and chemically reactivefelsic pumice breccias (Fig. 7.1 IB).

Significance of pyrophyllite and kaolinite in VHMSsystems

Pyrophyllite and kaolinite are generally rare in VHMSaltered zones; however, because they are only stable underrelatively acidic conditions their presence warrants somediscussion. Kaolinite has been reported from the sericiteand montmorillonite zones of some of the Japanese Kurokodeposits (e.g. Iijima, 1974; Ohmoto, 1996). It has alsobeen reported in the altered footwall zone of the MountChalmers Cu-Au VHMS deposit in Queensland (McLeod,1987). Pyrophyllite exists in the sericite zone of the WesternTharsis VHMS deposit in the Mount Lyell field, MountRead province (e.g. data sheets WT4 and WT5: Huston andKamprad, 2001), and is also reported in the stockwork zonesof several VHMS deposits in the Iberian pyrite belt (Relvasetal., 1997).

Figure 7.12 shows that the stability relationship betweenkaolinite and pyrophyllite is temperature dependent:pyrophyllite being stable above 280°C, whereas the stabilitybetween muscovite and pyrophyllite is controlled by the3.K+/3.H+ ratio. For a fluid with 3.K+ varying from 0.1 to 0.01,which is considered the range for VHMS-related fluid, thenpyrophyllite is stable at pH values of less than 3—4 units.

FIGURE 7.12 | Stability relations among selected silicate minerals at 500 barspressure (modified after Beane and Titley, 1981).

There are two ways that acidic fluids may be generatedto stabilise kaolinite or pyrophyllite in VHMS systems. Thefirst is by acid-producing fluid-rock reactions, such as thereplacement of albite (and volcanic glass) by Mg-Fe chlorite(reaction R7.2). This may result in a reduction in fluid pHby about 1 unit (from 5.5 to 4.5); however, buffering bysericite + chlorite assemblages in the rock will generallyprevent the pH dropping below 3.5, which is needed forkaolinite and pyrophyllite formation. These acidic alterationminerals can only form by this method in volcanic rocks thatcontain negligible K (e.g. tholeiitic basalts) and thus containno sericite to buffer the pH. The second method is by theintroduction of magmatic volatiles at some stage during thelife of the hydrothermal system (Ohmoto, 1996). Coolingof a magmatic gas containing SO2 at temperatures below400°C will increase fluid acidity to levels below pH = 3.5 by areaction similar to R7.7 (Burnham and Ohmoto, 1980):

4SO2(g) + 4H2O(1) -» H2S(aq) + 3H+ + 3HSO4~ (R7.7)

Researchers have recently argued that the presence ofpyrophyllite or kaolinite in VHMS altered zones supportsthe theory that these deposits are not simply the products ofseawater convection, but that their genesis involves input ofa magmatic-derived, low-pH fluid (e.g. Sillitoe et al., 1996;Huston and Kamprad, 2001).

Metamorphism of altered zones

Few detailed studies have been published on the effects ofcontact and regional metamorphism of VHMS-relatedaltered zones. Medium- to high-grade metamorphism ofchlorite and sericite zones leads to assemblages containingcordierite, anthophyllite, garnet, biotite, andalusite, staurolite,gahnite, hornblende and plagioclase, depending on the bulkcomposition of the altered zones (Franklin et al., 1981).Cordierite + anthophyllite assemblages commonly result frommetamorphism of chlorite-rich altered zones, with a spottedtexture due to cordierite porphyroblasts, leading to the termdalmatianite (Fig. 3.7A: Riverin, 1977). Some examples ofmineral assemblages from metamorphosed hydrothermallyaltered zones are provided in Table 7.1.

7.3 | THE SPECTRUM OF VOLCANIC-HOSTED DEPOSITS ANDASSOCIATED ALTERATIONPATTERNS

Our research into the Palaeozoic VHMS deposits of easternAustralia (Large, 1992; Gemmell et al., 1998; Gemmelland Herrmann, 2001; Large et al., 2001c) has revealedconsiderable variation in terms of volcanic environment, orebody shape, metal ratios, metal zonation, alteration mineralassembalges and zonation, and ratio of massive to stringerand disseminated styles of ore. Large (1992) identified 10major styles of deposits in Australia (Fig. 7.13), only one ofwhich was the classic mound style and associated footwallalteration pipe depicted in Figure 7.3. A number of factors

including temperature and salinity of the hydrothermal fluid;oxidation and H2S/SO4 characteristics of the hydrothermalfluid and seafloor environment; composition of volcanicand sedimentary rocks deep in the succession; permeabilityof the footwall volcanic rocks; and depth of seawater controlthe metal carrying capacity of the fluid, and the chemicalreactions that occur beneath and at the seafloor, leading toa broad range of deposit styles and associated local alterationhalos.

The spectrum of VHMS deposits found in submarinevolcanic successions indicates that there may be continuumof deposit styles between the end members that form thebasis for the current deposit classification in volcanic arcand rift environments: VHMS Cu-Zn-Pb, porphyry Cu-Au,epithermal Au-Ag and SEDEX Zn-Pb-Ag deposits. Recentworkers have emphasised the possible continuum betweenVHMS and epithermal deposits (e.g. Lydon, 1996; Sillitoeet al., 1996). Large (2000, 2004) and Large et al. (2001c)extended this approach to include porphyry and SEDEXdeposits. Sillitoe et al. (1996) introduced high-sulfidationand low-sulfidation VHMS deposits based on mineralisationstyle, hypogene alteration and sulfide mineralogy, seawaterdepth, and volcano-magmatic setting. High-sulfidationVHMS deposits develop in shallow-water environments,proximal to volcanic centres, and tend to be associated withzones containing argillic alteration minerals (e.g. pyrophyllite,allunite, kaolinite, diaspore) and high-sulfidation sulfideminerals (e.g. enargite, luzonite, bornite, tennantite). For mostgeologists the high-sulfidation—low-sulfidation classificationscheme has proven difficult to embrace because these termswere originally based on the chemistry of the ore fluid andenvironment of mineral deposition, rather than a series ofgeological criteria that could be measured and applied inthe field or in drill core. For this reason we do not endorseadoption of the high-sulfidation—low-sulfidation terminologyfor VHMS deposits.

An alternative approach, suggested by Large et al. (2001c)and expanded here, is to place individual deposits within arange of features defined for ores in volcanic arcs and rifts. Adiamond-shaped diagram (Fig. 7.14) shows deposits plotted

in terms of their attributes relative to end-member depositmodels for VHMS, epithermal Au, porphyry Cu and SEDEXZn-Pb deposits. The main eastern Australian depositsdescribed in this chapter, and the Bathurst 12 deposit fromthe Bathurst mining camp, are plotted on the diagram.

Hellyer plots very close to the ideal VHMS deposit endmember. This is because the deposit exhibits most of thefeatures of the idealised VHMS alteration-mineralisationsystem outlined in Figure 7.3.

Mount Lyell (Western Tharsis, Section 7.7) and Highway-Reward (Section 7. 10) plot toward the porphyry end of thespectrum with a significant magmatic component. This isbecause these deposits are Cu-Au-rich subsurface replacementores that formed in proximal volcanic environments dominatedby synvolcanic rhyolitic intrusions. Mount Lyell also containsalteration minerals typical of an acid fluid or high-sulfidationepithermal environment (pyrophyllite and zunyite), whichmay suggest that magmatic fluid was involved (Huston andKamprad, 2001). These Cu-Au-rich massive sulfide depositsand others like them (e.g. Mount Morgan, Boliden, Bousquet)are considered to be hybrid VHMS-epithermal-porphyrydeposits that are not easily classified as end members in theVHMS-epithermal-porphyry spectrum (Large, 2004).

Henty is a gold-rich, base-metal-poor volcanic-hosteddeposit within the Mount Read province (Section 7.8). Thegold ore occurs in a stratabound subseafloor replacement zonesurrounded by concentric altered zones dominated by quartz,sericite, carbonate and albite. The deposit is neither a typicalVHMS nor an epithermal deposit, but has some features ofboth, and is best described as a hybrid VHMS-epithermaldeposit.

Rosebery is a Zn-Pb-Ag-Au massive sulfide deposit(Section 7.6). It differs from the classic VHMS deposit in itslow Cu content, sheet-like stratiform nature with no stringerzone, and lack of a well-defined footwall alteration pipe. It ishosted in proximal and distal volcanic facies dominated bypumice breccia, volcanic sandstone and siltstone, and blackshales. Rosebery and other sheet-like Zn-Pb-rich deposits,such as those in the Bathurst mining camp, have many featuressimilar to SEDEX deposits even though they are in volcanic

Mattabi Chloritoid + siderite + andalusite Chlorite Franklin et al. (1977)

Mattabi Quartz + chloritoid + andalusite + kyanite Silica core Franklin et al. (1977)

Coronation Cordierite + anthophyllite Chlorite Whitmore (1969)

Anderson Lake Mg-chlorite + biotite + kyanite Chlorite Franklin etal. (1981)

Amulet A Anthophyllite + cordierite Chlorite Beaty and Taylor (1979)

Geco B i °J t e ! I f ? +. chl0|r!te + muscovite + almandine + cordierite + Chlorite Stanton (1984)

anthophylnte + staurolite x '

Balcooma Chlorite + quartz + staurolite + biotite + cordierite + garnet Chlorite Huston etal. (1992)

Balcooma Quartz + muscovite + biotite Sericite Huston etal. (1992)

Dry River South Quartz + muscovite + biotite + staurolite + andalusite Sericite Huston et al. (1992)

Skellefte district Chlorite + cordierite + andalusite Chlorite Weihed et al. (2000)

Boliden Sericite + quartz + andalusite + corundum Central zone Nilsson (1968)


TABLE 7.1 | Mineral assemblages recorded in medium- to high-grade metamorphosed hydrothermally altered zones local to VHMS deposits.

176 | CHAPTER 7

FIGURE 7.13 | Schematic representation of

the various shapes and alteration zonation

pattens associated with VHMS deposits

(modified after Large, 1992).

successions. Recognising their hybrid natures and range infeatures we have plotted the Rosebery and Brunswick No. 12deposits on the boundary between VHMS and SEDEXdeposits (Fig. 7.14).

Hydrothermal alteration related to the spectrum ofdeposits

Deposits in the VHMS spectrum exhibit a continuum ofalteration zonation patterns that are depicted in Figure 7.15.The shapes of the alteration halos, their mineral assemblagesand zonation, change progressively along the spectrum.

Porphyry Cu (-Au) deposits (Fig. 7.15A) exhibit a seriesof very extensive roughly concentric altered zones. Theseinclude potassic zones in the cores (K-feldspar and/or biotite),

enveloped by phyllic zones (quartz + sericite + pyrite), andfinally propylitic zones (carbonate + chlorite + epidote) at themargins, which merge with the regional diagenetic or meta-morphic facies.

Hybrid massive sulfide Cu-Au deposits (Fig. 7.15B) exhibitsimilarly shaped, but less extensive, altered hanging wallzones compared to porphyry Cu deposits. Although theylack potassic zones, they comprise siliceous and/or chloritecore zones, which contain Cu-Au ore, and are surroundedby sericite zones and propylitic halos. Alteration mineralscharacteristic of highly acidic alteration (e.g. pyrophyllite,kaolinite, zunyite, topaz) may occur in the sericite zone in asimilar pattern to porphyry systems.

Classic mound or lens-shaped Cu + Zn ± Pb-rich VHMSdeposits have both massive sulfide and footwall stringer zoneswith well-zoned footwall alteration pipes (Fig. 7.15Q and


subordinate hanging wall altered zones, unlike the previousmembers of the spectrum where altered rocks dominate thehanging wall. Mg-bearing altered zones (Mg-chlorite or talc)are common features of these deposits dependent on thechemistry and temperature of the fluids.

Sheet style VHMS deposits are Zn-rich, strata paralleland have extensive stratabound altered zones (Fig. 7.15D).Stringer zones are less common, but where present theyare stratabound rather than pipe-shaped. Carbonates arecommon alteration minerals, particularly around the marginsof the massive sulfide ore.

SEDEX Zn-Pb-Ag deposits are at the distal end of themassive sulflde spectrum in terms of proximity to volcaniccentres, ratio of sedimentary to volcanic host rocks, andtemperature of formation (Fig. 7.15E). Alteration halos are

KEY ALTERATION MINERALSchloritesericitecarbonate

FIGURE 7.15 | Variations in alteration halos for the spectrum of deposits fromporphyry Cu-Au to SEDEX Zn-Pb-Ag. (A) Classical porphyry Cu-Au deposit (e.g.El Salvador, after Gustafson and Hunt, 1975; McMillan and Panteleyev, 1998).(B) Hybrid Cu-Au massive sulfide deposit (e.g. Mount Morgan, Mount Lyell, orHighway-Reward, after Large et al., 2001c). (C) Classic mound-style VHMSCu-Zn or Cu-Zn-Pb deposit (e.g. Hellyer, after Gemmell and Fulton, 2001). (D)Sheet-style VHMS deposit (e.g. Rosebery, after Large et al., 2001b). (E) ClassicSEDEX deposit (e.g. HYC, after Large et al., 2000).

1 7 8 | CHAPTER 7

commonly stratigraphically controlled and dominated bycarbonate minerals (ferroan-dolomite, ankerite and siderite,Large et al., 2001c). Chlorite and siliceous zones are rare,and a sericite zone is usually restricted or entirely absent.Manganese and Tl are common alteration halo indicators inSEDEX and sheet style VHMS deposits, but are less commonin the mound-style VHMS, hybrid Cu-Au deposits andporphyry Cu deposits (Large et al., 2001c).

7.4 | COMPARISONS BETWEENARCHAEAN, PALAEOZOIC ANDCAINOZOIC VHMS ALTERATIONSYSTEMS

Australian Palaeozoic VHMS alteration halos

The altered zones around eastern Australian PalaeozoicVHMS deposits have diverse morphologies and mineralassemblages related to variations in their volcanic settings andmodes of formation (Large, 1992; Large et al., 2001c). Well-defined footwall alteration pipes are relatively uncommon, orperhaps unrecognised because of subsequent deformation.In Australian deposits, footwall alteration pipes are mainlyassociated with synvolcanic faults (e.g. Mount Morgan, Taube,1986) and with relatively impermeable footwall rocks (e.g.Hellyer and Highway-Reward, Gemmell and Large, 1992;Large, 1992; Doyle, 2001). Laterally extensive strataboundaltered footwall zones are more typical, especially beneathsheet-like Zn-rich polymetallic deposits, such as Rosebery,Hercules and Thalanga. Stratabound altered footwall zonesare attributed to non-focussed discharge and lateral migrationof hydrothermal fluids in permeable volcaniclastic units inthe footwall (e.g. Rosebery, Green et al., 1981).

Quartz + sericite + pyrite assemblages are volumetricallydominant in all types of alteration halos. The proximalaltered footwall zones are typically quartz rich, containingless than 20% phyllosilicates but greater than 5% pyrite indisseminations and veins. Much broader footwall feldspar-destructive altered zones, with mineral assemblages dominatedby sericite or sericite + chlorite, a lower proportion of quartz,and a few percent of disseminated pyrite, envelop them.Chlorite-rich assemblages tend to be restricted to small zonesimmediately beneath ore lenses (e.g. Rosebery and Thalanga,Green et al., 1981; Paulick et al., 2001), and in several casesare associated with carbonate (e.g. Hellyer, Thalanga andWoodlawn, Davis, 1990; Herrmann and Hill, 2001) . Wherethere are footwall alteration pipes, chlorite exists in the medialzones, usually between a quartz-rich core and a surroundingsericite zone (e.g. Hellyer and Highway-Reward, Gemmelland Large, 1992; Doyle, 2001). However, the chlorite-richfootwall alteration pipes that are characteristic of manyCanadian Archaean deposits do not seem to be present inthe Australian Palaeozoic deposits. Several of the Tasmanianexamples are virtually devoid of chlorite, such as Henty,Boco and Chester (Boda, 1991; Green and Taheri, 1992;Callaghan, 2001). These are base-metal-poor systems; some

of them contain aluminous minerals such as pyrophyllite andkaolinite. Recent interpretations suggest they are analogous toseafloor acid-sulfate epithermal systems, and possibly involvesignificant magmatic fluid in their formation (Large et al.,2001c).

The sheet-like Zn-rich deposits do not have extensivevisually recognisable hydrothermally altered hanging wallzones, although there may be subtle hanging wall geochemicalhalos (e.g. Rosebery, Large et al., 2001b). However, somedeposits with vertical pipe-like footwall alteration and/ormineralised zones exhibit altered hanging wall zones thatextend for several tens to hundreds of metres above thedeposit. For example, the 200 m thick basalt unit overlyingthe Hellyer deposit contains an upward flaring plume ofdistinctive green, Cr-bearing muscovite (fuchsite), which ismore or less concentrically enclosed by discontinuous chlorite+ carbonate and quartz + albite and sericite zones (Gemmelland Fulton, 2001). A stratabound quartz + albite (± chlorite)altered zone up to 100 m thick exists above the pyritic zones atHenty (Halley and Roberts, 1997). The altered hanging wallzones at Hellyer and Henty contain only traces of pyrite,which suggests that they formed from fluids of very differentcomposition to those in the footwall. In contrast, the quartz+ sericite + pyrite zones extending above the Highway andReward massive sulfide bodies are essentially similar to thefootwall stringer zones. This supports the interpretation thatthese deposits formed entirely below the seafloor (Doyle andHuston, 1999).

Carbonate alteration facies are common features of theAustralian Palaeozoic polymetallic deposits. They are typicallythin stratabound zones that enclose, lie immediately above,or are laterally equivalent to, the sulfide lenses (e.g. Rosebery,Henty and Thalanga, respectively). Except at Henty, thehydrothermal carbonates are generally not calcic; they havevarious Ca-Mg-Fe-Mn compositions, which in some casesvary systematically towards ore (e.g. Rosebery, Large et al.,2001b).

FIGURE 7.16 | Idealised cross-section of the four altered zones around Kuroko-type massive sulfide deposits, Japan (modified from Franklin et al., 1981).

Japanese Cainozoic VHMS alteration halos

The Miocene deposits of the Hokuroku district in northernJapan generally have similar hydrothermal alteration facies tothe Australian deposits, although the transitions to diageneticfacies in the host succession is clearer because they are notdeformed or metamorphosed. The idealised Kuroko depositmodel has four altered zones surrounding the sulfide deposit(e.g. Fig. 7.16: Franklin et al., 1981). These include an innerquartz + sericite zone immediately beneath the sulfide deposit.The inner quartz + sericite altered zone typically encloses aquartz + pyrite stockwork zone (Keiko ore) that underliesthe massive sulfide ore (Oko and Kuroko ores) (Ohmotoand Skinner, 1983) and is analogous to pyritic stringer zonesbeneath some Australian deposits. The inner quartz + sericitezone is laterally surrounded by a sericite + Mg-chlorite +montmorillonite zone that extends in a thin layer over thetop of the deposit. It is succeeded outward by mixed-layerclay alteration facies (mineral assemblages of sericite + inter-layered illite/smectite + chlorite + albite + K-feldspar), whichmay extend for up to several kilometres laterally and 200 minto the hanging wall. The outermost zeolite zone typicallycontains relict plagioclase in mineral assemblages progressingfrom analcime + montmorillonite + quartz ± calcite, throughmordenite + montmorillonite + quartz ± inter-layeredillite/smectite, to the background diagenetic clinoptilolite +mordenite alteration assemblage.

There is some deposit-specific variation within thatidealised Kuroko alteration zonation pattern. For example, atthe Fukuzawa deposits the sericite + Mg-chlorite zone in thehanging wall contains relict plagioclase, and analcime seemsto exist in more distal parts of the zeolite zone than mordenite(Date et al., 1983). Alteration mineral assemblages in thealtered footwall zones around the Uwamuki deposits broadlyconform to the idealised zonation but include peripheralzones of kaolinite in (presumably disequilibrium) mineralassemblages of sericite + chlorite + quartz + albite + pyrite(Urabe et al., 1983). These authors noted that kaolinite isnot otherwise common around Kuroko deposits and, wherepresent, usually occurs in the core zones with pyrophyllite+ diaspore. Subsequent work around the Uwamuki depositby Shikazono et al. (1998) indicated that the kaolinitezones are greater than 200 m from ore and extend into thehanging wall and therefore may not have been directly relatedore deposition. Marumo (1989) also found kaolin mineralsin the hanging wall of the small Inarizawa sulfide depositsand concluded that they formed during a low-temperaturewaning phase of the ore-related hydrothermal system. Theexistence of kaolinite ± pyrophyllite ± diaspore assemblages,characteristic of low pH, acid-sulfate systems, suggests thatthe Hokuroku district also contains a spectrum of volcanic-hosted deposits similar to those recently recognised in theearly Palaeozoic belts of Tasmania and north Queensland(Large et al., 2001c).

Carbonate-bearing assemblages have not been widelydescribed in the Hokuroku district. Nevertheless, Shikazonoet al. (1998) reported that magnesite, siderite, dolomite andcalcite were common and characteristic in the ore horizonand hanging wall rocks. Based on isotopic and fluid inclusiondata from Uwamuki they concluded that carbonatesprecipitated in a post-ore hydrothermal stage by interaction

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS |

of hydrothermal fluids with biogenic marine carbonates. Theerratic distribution and the post-ore formation of carbonates,and indications of complex overprinting of different systems(VHMS and acid-sulfate) leaves some doubt about thegenetic relationships between the massive sulfide deposits andcarbonate assemblages.

Canadian and Australian Archaean VHMSalteration halos

There are two major classes of altered zones associated withCanadian Precambrian massive sulfide deposits: (1) well-defined narrow footwall alteration pipes, and (2) broadirregular altered footwall zones that are transitional to deepsemi-conformable alteration facies, with or without localisedpipes (Morton and Franklin, 1987; Kerr and Gibson, 1993;Gibson et al., 1999). The former are commonly associatedwith small (<5 Mt) Cu-Zn deposits and are interpreted to haveformed in deep water in dominantly coherent mafic volcanicrocks. The latter generally exist beneath larger deposits ofvariable metal associations formed in relatively shallow water(^500 m) and in dominantly felsic volcaniclastic rocks.

Pipe-like altered footwall zones are epitomised by the Cu-Zn deposits of the Noranda district in the Abitibi belt. Thesecharacteristically have upward flaring footwall alteration pipesthat are roughly circular in plan view, generally with slightlysmaller diameter but greater vertical extent than the overlyingmassive sulfide lenses. They are commonly recognisable forup to 1 km below the deposits (Franklin et al., 1981). Theupper part of the footwall alteration pipe (Fig. 7.17) enclosesa stringer zone or stockwork of pyrite ± chalcopyrite ±pyrrhotite veins in a core dominated by Fe-chlorite passinglaterally and upward through Mg-chlorite to an outer zonedominated by sericite ± chlorite ± quartz (Lydon, 1984; Kerrand Gibson, 1993). Lydon (1996) noted the existence of talc-bearing or aluminous assemblages in the upper parts of somefootwall alteration pipes.

Depletions of Si, Na, Ca and K and additions of Mg andFe generally characterise the alteration of the chloritic core,

179

FIGURE 7.17 | Idealised cross-section of a typical zoned footwall alterationpipe beneath Noranda-type massive sulfide deposits, Abitibi belt, Canada(modified from Lydon, 1984; Kerr and Gibson, 1993).

1 8 0 | CHAPTER 7

whereas small additions of K and possibly Si occur in thesericitic shell (Barrett and MacLean, 1994b). This generallyresults in significant net loss of mass, due to Si loss, fromthe overall alteration pipe (Barrett and MacLean, 1991). Ina few unusual cases there may be net mass gains (e.g. Norbecdeposit: Barrett and MacLean, 1999). Footwall alterationpipes of this type may represent zones of hydrothermaldischarge that were focussed by synvolcanic faults. Theirvertical extent suggests that the hydrothermal fluid sourceswere very deep. Overprinting relationships indicate that thefootwall alteration pipes were initially zones of sericite ±quartz altered rock. As the hydrothermal system intensified,sericite was replaced by chlorite concurrent with metal zonerefining in the sulfide lenses (Kerr and Gibson, 1993).

Non pipe-like, broad altered footwall zones have morevariable morphologies and mineral assemblages as exemplifiedby the differences in the Home and Mattabi deposits. TheHome deposit has a poorly defined altered footwall zone ofquartz + sericite ± chlorite that is many times wider than themassive sulfide bodies (MacLean and Hoy, 1991; Kerr andGibson, 1993). Calculations by MacLean and Hoy (1991)indicate that the Home footwall alteration was accompaniedby significant net mass gains; mainly gains of Si, Fe andK, slightly offset by losses of Na, Ca and Mg. Beneath theMattabi deposit are siderite + chloritoid ± andalusite, kyaniteand pyrophyllite zones, which narrow with depth and aretransitional downward and laterally into an extensive semi-conformable ankerite + chlorite + sericite + quartz zone(Franklin et al., 1975; Morton and Franklin, 1987). Gibsonet al. (1999) suggested that the aluminous assemblages atMattabi (and several other deposits that are notably Au rich)were analogous to the advanced argillic assemblages formedby low pH fluids in acid-sulfate epithermal systems.

The Archaean massive sulfide deposits of the Panoramadistrict in the Pilbara of Western Australia occur near the topof a 2 km thick basaltic to rhyolitic volcanic succession abovea large synvolcanic granite pluton (Brauhart et al., 2001).Large, semi-conformable altered zones of feldspar-destructivesericite + quartz and chlorite + quartz alteration assemblagesoccupy the lower and middle parts of the volcanic successionand extend almost the entire exposed strike length (20 km).Locally transgressive chlorite + quartz altered zones, boundedby synvolcanic faults, extend upwards from the semi-conformable altered zones to beneath the massive sulfideprospects. Mass changes in the feldspar-destructive alteredzones were modest: small gains of Si and losses Ca, Na, Feand K in the lower sericite + quartz zones, and small gains ofMg, Fe, Si and losses of K, Na, Ca in the transgressive chlorite+ quartz zones.

In the Golden Grove district of the Archaean Yilgarncraton, Western Australia, the volcanic succession that hoststhe Scuddles and Gossan Hill massive sulfide deposits alsoexhibits the effects of regional-scale, intense feldspar-destructivealteration. The entire footwall succession of altered andesiticto rhyolitic volcaniclastic units, although preserving primaryvolcanic textures, is composed essentially of quartz + chlorite(± minor sericite). The alteration process, interpreted as asyn-depositional or early hydrothermal regional metasomaticevent, virtually removed all Ca, Na and K from the rocksand added substantial Si, Fe and Mg (Sharpe and Gemmell,2001). At Gossan Hill the regional quartz + chlorite alteration

facies is overprinted by two alteration facies related to sulfidemineralisation. A narrow stratabound chlorite (± siderite,ankerite talc and andalusite) zone envelops the lower Cu-richmassive magnetite + sulfide ore lens. An intense quartz zoneunderlies the upper Zn-rich massive sulfide lens and enclosesa discordant zone of sulfide stringer veins that connects theupper and lower lenses (Sharpe and Gemmell, 2001).

There is considerable diversity among alteration faciesaroundArchaean massive sulfide deposits. Features that appearto be common to most Archaean districts are large semi-conformable altered zones and localised discordant alteredfootwall zones. Brauhart et al. (2001) highlighted some of thedifferences in mineral assemblages and mass changes betweenPanorama and the Canadian semi-conformable altered zones.However, the well-defined discordant footwall alterationpipes are typically chlorite rich (if not metamorphosed tohigher grades) and characterised by significant net mass loss,which is attributable to major Si loss and only partly offset byMg and Fe gains.

Comparisons

Despite the many variations in mineral assemblage, morphologyand extent of alteration facies associated with VHMS deposits,both within districts and across geologic time, there is onefeature that is common to all: proximal altered footwall zonesdo not contain feldspar. Feldspar destruction is usually manifestin the presence of sericite, chlorite or smectite clays, or theirhigher grade metamorphic equivalents. One or more of theseAl-bearing phyllosilicates is almost invariably present because,although VHMS-type hydrothermal fluids readily transportsilica, alkalis and other cations, Al is generally immobile inthese moderately acidic systems. One suspects that pyrite is anequally ubiquitous component of proximal altered footwallzones but it is frequently not mentioned in alteration mineralassemblages, due to the unnatural distinctions that manyauthors make between alteration and mineralisation. It is alsobecoming increasingly apparent that alumino-silicates suchas kaolinite, pyrophyllite, andalusite and others exist locallyin altered zones across the entire age spectrum of VHMSdeposits, from Archaean to Cainozoic, and that there mayalways have been continua between moderate and low pHsubmarine hydrothermal systems. Because of the consistentfeldspar destruction, alteration indices such as Al (Ishikawaet al., 1976), CCPI and S/Na2O (Large et al., 2001a) shouldbe effective indicators of alteration intensity in all kinds ofVHMS districts, at least at prospect scales.

Footwall alteration facies of the Australian Palaeozoic andJapanese Cainozoic deposits are typically more sericitic thanchloritic, and their proximal zones tend to be quartz rich.The limited mass change data available indicate they havetypically undergone moderate to large net mass gains, whichare dominantly attributable to gains in Si that significantlyoutweighed losses and gains in other components. Thispredominance of sericite is probably not due to low availabilityof Fe and Mg. Iron may be significantly added as pyrite andMg commonly appears in carbonates or chlorite in upper orupper-peripheral altered footwall zones.

In contrast, chlorite dominates the well-defined footwallalteration pipes that underlie many Canadian Archaean

deposits. These zones are characterised by significant net massloss, in which the large loss of Si outweighed addition of Mgand Fe. Although there are some exceptions (e.g. Panorama),major Si and net mass losses are indicated wherever chloriteis dominant in an alteration mineral assemblage. Thisgeneralisation also applies to Australian Palaeozoic systems;for instance the small chlorite-rich zones in the Hellyeralteration pipe (Gemmell and Large, 1992) and Thalangafootwall (Herrmann and Hill, 2001).

In terms of mass change, the major difference betweenArchaean deposits with chloride footwall alteration pipes,and Palaeozoic to Cainozoic deposits with quartz + sericite-dominated altered footwall zones, is that the former lost massand the latter gained mass. In addition, in all cases, the majorcontributor to net mass change was Si.

This difference in the behaviour of Si is probably relatedto the evolution of the hydrothermal systems and particularlythe compositions of hydrothermal fluids, which originatedas seawater in both cases. Evidence of chlorite overprintingsericite ± quartz assemblages in the Canadian footwallalteration pipes suggests that fluid compositions changed asthe hydrothermal system intensified. The initial fluid wasprobably over-saturated in Si and deposited quartz alongthe discharge zone to the seafloor as it cooled, whereas thelater fluid, possibly of higher temperature and associated withCu enrichment of the sulfide deposit, was undersaturatedand leached Si from the core of the discharge zone. Thischange may be explained in terms of the regional deep semi-conformable altered zones associated with Archaean deposits.The lower semi-conformable altered zone is typically a zoneof silicification at temperatures greater than about 400°C(Kennedy, 1950; MacGeehan, 1978; Fournier, 1985; Galley,1993; Skirrow and Franklin, 1994), attributed to down-goingmodified seawater being heated up to the range of retrogradeSi solubility at 400-600°C (Fournier, 1985). If fluid depositedSi in the deep semi-conformable altered zone, it would thenbe undersaturated in Si as it ascended the discharge zone,even if it cooled as it ascended.

The link between Archaean deposits and regional semi-conformable altered zones, which are generally not recognisedin the younger VHMS districts, suggests that Archaeancrustal conditions (thin crust and large high-level plutonicintrusions) favoured large, intense and presumably long-livedsystems. Palaeozoic and younger VHMS districts are nottypically associated with large high-level plutons analogousto the Flavrian pluton of the Noranda district (Kerr andGibson, 1993) or the Strelley granite at Panorama (Brauhartet al., 2001). Their absence may account for the less extensive,perhaps less evolved, altered footwall zones associated withSi and net mass gains that are most common beneath thePalaeozoic and younger massive sulfide deposits.


7.5 | HELLYER: A MASSIVE ELONGATEPOLYMETALLIC LENS

The Hellyer deposit is located in the northern part of theCambrian Mount Read province, western Tasmania (Fig.1.5). The pre-mining resource was 16.2 Mt of 13.9% Zn,7.1% Pb, 0.4% Cu, 168 g/t Ag and 2.5 g/t Au (Gemmell andLarge, 1992; McArthur, 1996). The deposit is a single elongatelens of massive sulfide about 800 m in length (north—south)by 200 m in width (east—west) and with an average verticalthickness of 45 m (Fig. 7.18). It occurs in the mainly calc-alkaline, intermediate to mafic Que-Hellyer Volcanics atthe base of the Mount Charter Group, which is equivalentto the western volcano-sedimentary sequences (Corbett and

Sericite + quartz zone

Sericite zone

Chlorite zone

Quartz zone

FIGURE 7.18 | Hellyer plan showing the altered zones immediately below themassive sulfide ore (approximately 400 RL). The black line is the outline of thebase of the massive sulfide. Modified after Gemmell and Large (1992).

1 8 2 | CHAPTER 7

Komyshan, 1989). The massive sulflde lens is bisected andoffset by a major north-south trending fault, the Jack fault(Figs 7.18 and 7.19). Beneath the massive sulflde lens isan elongate, carrot-shaped, zoned, footwall alteration pipe(described in detail by Gemmell and Large, 1992). Above thedeposit is a moderately well developed altered hanging wallzone (Gemmell and Fulton, 2001).

The massive sulflde lens exhibits classical metal zoning,with minor Cu concentrated in a pyritic core, followed bylow grade Zn + Pb, and then high grade Zn + Pb + Ag inthe upper parts of the massive sulflde lens (McArthur andDronseika, 1990; Large, 1992; McArthur, 1996). The centreof the massive sulflde deposit is capped by a quartz + pyritezone, which is flanked by thin irregular lenses of barite that

FIGURE 7.19 | Cross-section of the Hellyer deposit showing the distributionof rock types, mineralised zones, and altered footwall and hanging wall zones.(A) The ore lens and altered zones are offset along the Jack fault (modified afterGemmell and Large, 1993). (B) Reconstructed 10740 N/10870 N cross-sectionshowing the massive sulfide and footwall alteration pipe prior to folding andfaulting (modified after Downs, 1993).

are directly over the high-grade massive sulflde (Sharpe,1991). Barite and massive sulfide clasts occur in volcaniclasticmass-flow units flanking the deposit.

Geological setting

The Hellyer ore body occurs above a footwall comprisingfeldspar-phyric andesitic and basaltic lavas and sills thatconsist of coherent and autoclastic facies, which are mainlyhyaloclastite and peperite (Fig. 7.20: Waters and Wallace,1992). Basalt (Hellyer basalt) and black mudstone (QueRiver Shale) dominate the hanging wall (Komyshan, 1986).The abundance of basalt-mudstone peperite indicates thatmost of the basalt units are sills that intruded the blackmudstone (McPhie and Allen, 1992; Waters and Wallace,1992). Very thick, graded quartz-bearing rhyolitic pumiceousand volcanic lithic breccias interbedded with turbidites andmudstones of the Southwell Subgroup occur in the upperparts of the hanging wall (Corbett, 1992; McPhie and Allen,1992). The ore lens position is marked by a 0—40 m thickvolcaniclastic unit, which mainly consists of coarse polymicticvolcanic breccia, sandstone and laminated volcanic siltstone(Waters and Wallace, 1992).

Trilobites in the Que River Shale, very thick sections ofblack mudstone and the abundance of graded mass-flow unitscollectively indicate that the Hellyer massive sulfide formed ina moderate to deep (>1000 m) submarine setting (Large et al.,2001a). The volcanic facies association indicates proximity tointrabasinal vents for effusive, basaltic and andesitic eruptionsand synvolcanic intrusions.


Gemmell and Large (1992), and Gemmell and Fulton(2001) provided detailed description of both the footwall andhanging wall alteration facies, and zonation at Hellyer. Thefollowing section summarises their work.

Footwall alteration facies and zonation

A zoned carrot-shaped footwall alteration pipe extends forat least 500 m beneath the Hellyer deposit (Figs 7.6C and7.20). At the centre of the alteration pipe, immediately belowthe massive sulfide lens, is a siliceous core zone dominated byintense, pervasive quartz + sericite + pyrite alteration facies.This zone is progressively enclosed in chlorite, sericite andstringer envelope (or sericite + quartz) altered zones (Fig.7.6C).

The moderate, selective sericite + quartz alteration facies(e.g. data sheet HE2 in the stringer envelope zone) is the10-50 m wide outermost part of the alteration pipe andgrades outward into weak, selective-pervasive albite + chloritealteration facies (least-altered footwall, data sheet HE1).Primary volcanic textures are preserved (although modified),lithic fragments exhibit sericite-altered margins, and feldsparphenocrysts are partly altered to sericite. The AI shows anincrease from background values of around 30—55 to valuesof 60-70 (Fig. 2.8B).

intense footwallv " alteration plume

V V / V

V ,i I V n V


Southwell Subgroup: crystal-richvolcaniclastic shale, greywacke and minorfelsic lava

Que River Shale: black shale and minorsandstone

Hellyer basalt: massive to pillowed basalt,pillow breccia, hyaloclastite and andesiticlava

Mixed sequence: polymict volcanic breccia,massive and auto-brecciated dacite andmassive sulfide ore

Lower andesites and basalts: andesitic,dacitic and basaltic lavas, hyaloclastitesand minor volcani-clastic facies

FIGURE 7.20 | Schematicstratigraphic section through theQue-Hellyer Volcanics showing the

Lower basalt: massive to pillowed basalts, ex ten t of a| tered z o n e s at the He||yerhyaloclastites and pillow breccias , ~ n. , . . . ..,. .' and Que River deposits. Modified

after Waters and Wallace (1992).

Sericite ± chlorite dominates the strong, selective-pervasive alteration facies in a 10—15 m wide zone markingthe outer extent of an intense hydro thermal alteration system,recognised by obliteration of volcanic textures, presence ofminor sulfides (mainly pyrite) and complete replacement offeldspar phenocrysts and feldspathic groundmass by sericite ±chlorite (e.g. data sheet HE3). The AI is typically between 70and 85. On the Alteration box plot samples from this faciesplot along a line from the least-altered footwall field towardthe chlorite corner.

In the intense, pervasive chlorite alteration facies, allprimary minerals and glass in the footwall andesitic rocksare completely replaced by fine-grained chlorite with minorpyrite, sericite, quartz and carbonate (e.g. data sheet HE4).The AI is between 90 and 100 and the CCPI between 80 and90. In the upper parts, immediately below the massive sulfide,this zone includes an intense, spheroidal chlorite + carbonatealteration facies (Fig. 7.6C), which has up to 50% dolomitein a fine-grained matrix of chlorite (e.g. data sheet HE5).This alteration facies has a lower AI (50—80) than the intensechlorite alteration facies due to the elevated whole-rock CaOrelated to the dolomite component in the rock.

In the siliceous core zone, all volcanic textures arecompletely destroyed and the rock is composed of a fineintergrowth of quartz + sericite + pyrite + chlorite (intense,pervasive quartz + sericite + pyrite alteration facies, data sheetHE6). This zone also contains a series of sub-vertical pyrite +quartz + sphalerite + galena ± chalcopyrite ± carbonate ± bariteveins, interpreted as hydrothermal feeders below the ore body.Alteration indices in the siliceous core zone are extremely highwith values of both AI and CCPI exceeding 90.

Hanging wall alteration facies and zonation \

Hanging wall alteration facies at Hellyer are less welldeveloped than the footwall alteration facies; however,recent detailed studies by Gemmell and Fulton (2001) haverecognised an upward flaring zoned alteration system that iscentred above the thickest part of the massive sulfide lens. The

altered hanging wall zone extends through the hanging wallpillow basalts up to the contact with the overlying Que RiverShale (Fig. 7.19). Data sheet HE7 is an example of the least-altered hanging wall andesite. The distribution and intensityof alteration facies in the altered hanging wall zone is patchy,with pillow margins more intensely and pervasively altered(e.g. data sheet HE9) than the pillow interiors. The outermargin of the altered hanging wall zone is defined by weaksericite alteration facies grading inwards to a pink-white,strong, pervasive albite alteration facies (e.g. data sheet HE8),moderate, pervasive chlorite + carbonate alteration facies (e.g.data sheet HE9), and in the centre of the system, a distinctivegreen, strong, pervasive fuchsite alteration facies (e.g. datasheet HE 10). There is no systematic trend in the alterationindices in the altered hanging wall zones. AI and CCPI valuesare commonly low in the albite alteration facies due to highNa2O, and low MgO and FeO values (e.g. data sheet HE9).

Ore genesis

Based on geological, textural, and metal zonation studies,McArthur (1989, 1996), Large (1992) and Gemmell andLarge (1992) concluded that the Hellyer massive sulfidedeposit grew as a mound in a seafloor depression. The metalzonation from Fe -» Cu —» Pb-Zn —> Ba was interpreted tobe an expression of hydrothermal zone refining (Large, 1992),which developed similarly to that described by Eldridge etal. (1983) for the Kuroko deposits. Solomon and Khin Zaw(1999), however, presented fluid inclusion data (indicatinghigh ore fluid salinities: averaging 11 wt%) to propose thatsulfide deposition occurred in a seafloor depression brine-pool, directly above the footwall alteration pipe. Solomonand Gaspar (2001) provide textural evidence in support ofsulfide accumulation in a brine pool. Solomon and Groves(1994) and Solomon et al. (2004) suggest that the abnormallyhigh salinity and other chemical characteristics of the Hellyerfluid inclusions, are strongly suggestive of involvement ofmagmatic fluids in the hydrothermal system.

1 8 4 [ CHAPTER 7

Weak, selective-pervasive albite + chlorite alteration faciesLeast-altered footwall

HE1

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict mineralogy

Relict texture

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

Hfw-LAA(FPS-I)

weak, selective-pervasive albite + chlorite

footwall

Que-HellyerVolcanics


monomictic mafic breccia

feldspar

feldspar phenocrysts, perlitic fractures andcurvi-angular cm-scale clasts, areas ofjigsaw-fit clastsandesite

massive, matrix supported and poorlysorted

andesitic hyaloclastite

albite + chlorite + sericite + calcite

selective-pervasive chlorite + calcite +

sericite in matrix and chlorite + sericite inclasts, chlorite infill in perlitic fractures,sericite + calcite-altered plagioclase

regional

good

weak

syn volcanic?

diagenetic


GeochemistrySiO2 54.69 K2O 1.55 Rb 68 Zr 125TiO2 0.64 P2O5 0.12 Sr 299 Nb 7.0AI2O3 17.92 S 1.92 Ba 500 Y 19Fe2O3 7.65 CO2 Cu 0MnO 0.09 Total 97.96 Pb 0 Al 36MgO 3.79 LOI 3.96 Zn 0 CCPI 57CaO 3.14 Sb Ti/Zr 30.9Na2O 6.45 Tl

Moderate, selective sericite + quartz alteration faciesFootwall


HE 2

ISample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

Hfw-SEZ

moderate, selective sericite + quartz

footwall

Que-Hellyer Volcanics


polymictic mafic breccia

feldspar

porphyritic and perlitic clasts, areas ofjigsaw-fit clasts, clasts with curviplanarmargins

andesite-basalt

massive, clast supported, poorly sorted

resedimented polymictic hyaloclastite

sericite + chlorite + quartz + albite + calcite+ pyrite

selective domainal, calcite vein infill,disseminated pyrite, albite + sericite +calcite-altered feldspars

alteration zone around pipe

moderate

moderate

synmineralisation

peripheral hydrothermal

GeochemistrySiO2 57.38 K2O 3.34 Rb 127 Zr 123TiO2 0.58 P2O5 0.13 Sr 97 Nb 7.0AI2O3 15.33 S 4.15 Ba 6700 Y 26Fe2O3 7.76 CO2 Cu 300MnO 0.10 Total 96.27 Pb 2700 Al 63MgO 3.51 LOI 7.89 Zn 4500 CCPI 69CaO 2.71 Sb Ti/Zr 28.3Na2O 1.29 Tl


18 6 | CHAPTER 7

Strong, selective-pervasive sericite + chlorite alteration faciesFootwall

HE 3

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

Hfw-SZ

strong, selective-pervasive sericite +chlorite

footwall



massive polymictic breccia

rare feldspar

deformed feldspar phenocrysts andandesite clasts

dacite-basalt

massive, matrix supported, poorly sorted

resedimented polymictic hyaloclastite

sericite + chlorite + quartz + pyrite +ankerite + (albite)selective pervasive, vein-halo (pyrite etc.),disseminated pyrite, and infill (carbonate)

pipe

poor

strong

synmineralisation

hydrothermal

GeochemistrySiO2 54.50 P2O5 0.10 Cu 500 Al 79TiO2 0.52 S 6.02 Pb 5100 CCPI 77AI2O3 14.51 CO2 Zn 7800 Ti/Zr 33.4Fe2O3 10.93 Total 97.92 SbMnO 0.21 LOI 8.11 TlMgO 5.16 Zr 94CaO 1.47 Nb 6.0Na2O 0.90 ™ ]]' Y 18K?° 3'58 Ba 4100


Intense, pervasive chlorite alteration faciesFootwall


HE 4

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

Hfw-CLZ

intense, pervasive chlorite

footwall



coherent feldspar-phyric andesite

nil

porphyritic, perlitic fractures

andesite

massive

indeterminate

chlorite + pyrite + (quartz + sericite +calcite + galena)

pervasive, chlorite infill in perlitic fractures,chlorite + quartz-altered feldspars

pipe

moderate

intense

synmineralisation

hydrothermal

GeochemistrySiO2 37.69 K2O 1.82 Rb 79 Zr 140TiO2 0.59 P2O5 0.12 Sr 39 Nb 8.0AI2O3 16.08 S 8.65 Ba 2000 Y 25Fe2O3 18.86 CO2 Cu 400

MnO 0.41 Total 96.35 Pb 5200 Al 95MgO 11.38 LOI 12.29 Zn 9300 CCPI 94CaO 0.64 Sb Ti/Zr 25.2Na2O 0.12 Tl


CHAPTER 7

Intense, spheroidal chlorite + carbonate alteration faciesFootwall

HE 5

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

135756

intense, spheroidal chlorite +carbonate

footwall



coherent, feldspar-phyric andesite

nil

porphyritic

andesite

massive

indeterminate

chlorite + dolomite + (quartz + sericite)

nodules-spheroids

local

poor

intense

synmineralisation

hydrothermal

Geochemistry

SiO2 34.88 P2O5 0.13 Cu 300 Al 75TiO2 0.60 S 4.13 Pb 7200 CCPI 93AI2O3 15.63 CO2 Zn 11500 Ti/Zr 29.0Fe2O3 12.89 Total 88.72 SbMnO 0.80 LOI 15.40 TlMgO 13.07 Zr 124CaO 4.86 Rb 66 Nb 9 0

Na2O 0.01 Sr 58 1 28K2O 1.73 Ba 1200


Intense, pervasive quartz + sericite + pyrite alteration faciesFootwall

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS I 189

HE 6

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

Hfw-SCZ

intense, pervasive quartz + sericite + pyrite

footwall



feldspar-phyric andesite breccia

nil

porphyritic, periitic fractures, jigsaw-fitclasts

andesite

massive


quartz + sericite + pyrite + (chlorite)

pervasive, pseudomorphs of feldspar± pyroxene?, quartz and sericite veinsdisseminated pyrite, feldspar overgrowthson phenocrystscore of pipe

moderate

intense

synmineralisation

hydrothermal

GeochemistryGeochemistrySiO2 67.42 K2O 2.17 Sr 15 Nb 4.0TiO2 0.30 P2O5 0.05 Ba 14800 Y 28AI2O3 8.13 S 8.87 Cu 2200Fe2O3 12.34 CO2 Pb 8000 Al 91MnO 0.07 Total 101.12 Zn 9800 CCPI 85MgO 1.42 LOI 7.74 Sb Ti/Zr 24.7CaO 0.32 TlNa2O 0.03 Rb 79 Zr 74


1 9 0 | CHAPTER 7

Subtle, pervasive aibite + chlorite + calcite alteration faciesHanging wall

HE 7

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

142562

subtle, pervasive aibite + chlorite + calcite

hanging wall

Que-HellyerVolcanics


massive, feldspar-phyric amygdaloidalandesite

plagioclase

massive, porphyritic, weakly amygdaloidal

andesite

massive

lava

chlorite + aibite + calcite + quartz +(chalcopyrite)selective-pervasive, chlorite or quartz infillin amygdales, quartz + calcite veins

regional

good

weak

synmineralisation

diagenetic to metamorphic

GeochemistrySiO2 51.14 K2O 0.42 Rb 21 Zr 151TiO2 0.55 P2O5 0.60 Sr 125 Nb 7.1AI2O3 14.77 S 0.06 Ba 226 Y 21Fe2O3 8.82 CO2 4.76 Cu 745MnO 0.21 Total 99.91 Pb 6 Al 38MgO 5.56 LOI 8.17 Zn 147 CCPI 75CaO 5.44 Sb 1.1 Ti/Zr 21.8Na,0 4.19 TI <0.5


Strong, pervasive albite alteration faciesHanging wall


HE 8

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

142622

strong, pervasive albite

hanging wall



feldspar-phyric basalt breccia

altered plagioclase

porphyritic

basalt

massive and jigsaw-fit breccia

lava or sill

albite + chlorite + calcite + (pyrite)

pervasive, albite + chlorite, massive

chlorite + pyrite veins, patchycalcite domains possibly irregularpseudomorphs

local, plume?

poor

strong

post mineralisation

diagenetic-hydrothermal?


192 I CHAPTER 7

Moderate, pervasive chlorite + carbonate alteration faciesHanging wall

HE 9

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

142593

moderate, pervasive chlorite + carbonate

hanging wall



monomictic basaltic andesite breccia

altered feldspars

porphyritic, jigsaw-fit clasts

basaltic andesite

massive

pillow lava

chlorite + caicite + sericite + (albite +quartz)

pervasive, spheroidal and rhombic caicite,caicite veins with chlorite vein-haloalteration

local, plume?

good

moderate

post mineralisation

hydrothermal

GeochemistryGeochemistrySiO2 42.34 K2O 1.92 Rb 55 Zr 140TiO2 0.61 P2O5 0.35 Sr 281 Nb 8.8AI2O3 11.78 S 0.07 Ba 737 Y 21Fe2O3 5.41 CO2 13.00 Cu 109MnO 0.12 Total 99.46 Pb 4 Al 28MgO 4.34 LOI 15.99 Zn 51 CCPI 76CaO 15.44 Sb 0.6 Ti/Zr 26.2Na2O 1.06 Tl 0.7


Strong, pervasive fuchsite alteration faciesHanging wall


HE 10

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

142643

strong, pervasive fuchsite

hanging wall



massive feldspar-phyric basalt

feldspar

porphyritic, minor amygdales

basaltic andesite

massive

lava or sill

sericite (fuchsite) + calcite + ankerite +chlorite + pyrite

pervasive, disseminated pyrite, sericitecleavage, calcite ± sericite pseudomorphsafter feldspar

plume

poor

strong

post mineralisation

hydrothermal

GeochemistrySiO2

TiO2

AIAFe2O3

MnO

MgO

CaO

Na2O

26.32

0.54

17.35

3.38

0.25

1.88

22.67

0.00

K2O

S

CO2

Total

LOI

5.00 Rb

0.21 Sr

0.08 Ba

19.48 Cu

99.15 Pb

21.13 Zn

Sb

Tl

143 Zr

176 Nb

4473 Y

162

2 Al

43 CCPI

8.1 Ti/Zr

18.3

74

4.9

18

23

50

43.9


194 | CHAPTER 7

7.6 | ROSEBERY: A POLYMETALLICSHEET-STYLE DEPOSIT

The Rosebery massive sulfide deposit is a sheet-stylepolymetallic Zn-Pb-Cu-Ag-Au VHMS deposit in the northernCentral Volcanic Complex of the Mount Read Volcanics,western Tasmania (Fig. 1.5: Green et al., 1981; Large, 1992).The mining resource is 32 Mt at 14.7% Zn, 4.5% Pb, 0.6%Cu, 146 g/t Ag and 2.3 g/t Au (data from Pasminco Miningand Exploration). Compared with the Hellyer deposit, whichcomprises a single ore lens (described in the previous section),Rosebery is composed of at least 16 separate ore lenses (Fig.7.21). These vary in size from 0.1 to 5 Mt. Unlike the carrot-shaped footwall alteration pipe at Hellyer, the Roseberyore lenses are enclosed in strata-parallel altered zones. Theore lenses are principally composed of massive and bandedsulfides of sphalerite, galena, barite, pyrite and chalcopyrite,with minor tetrahedrite-tennantite, arsenopyrite, pyrrhotite,hematite and magnetite. In some sections of the mine (e.g.A and B lenses, Huston and Large, 1987) barite-rich lensesoverlie the Zn-Pb-Cu ore lenses.

Geological setting

The Rosebery, Hercules and South Hercules polymetallicore bodies are hosted by the same stratigraphic sequence inthe upper part of the Central Volcanic Complex, west ofthe Henty fault (Solomon, 1964; Green et al., 1981). Thefootwall comprises a thick (up to 500 m), poorly stratifiedrhyolitic-dacitic succession of weakly graded, feldspar-phyricpumice breccia, which is interpreted to be the product oflarge volume submarine caldera-forming eruptions, and

rhyolitic and dacitic sills (Hercules Pumice Formation: Lees,1987; Allen, 1994b; Large et al., 2001b). The ore lenses occurin the 5 to 10 m thick, finely stratified pumiceous siltstone,sandstone, crystal-rich sandstone and claystone top (hostrocks) of the footwall pumice breccias (Lees, 1987; Corbettand Solomon, 1989; Allen, 1994b; Large et al., 2001b). Thehost rocks are overlain by black mudstone, which representsa hiatus in volcanism marked by non-volcanic sedimentation.The hanging wall comprises a 5 to 400 m thick successionof massive to stratified, feldspar + quartz-phyric rhyodaciticvolcaniclastic units of the White Spur Formation, interbeddedwith black mudstone (Lees, 1987; Allen, 1994b). The footwallrhyolitic pumice breccias haveTi/Zr of 7-9 (Fig. 4.5), whereasthe host interval porphyry sill has Ti/Zr of 12-14, and thevolcaniclastic facies 10-30 (Large et al., 2001b).

Bedforms and textures within the footwall pumicebreccias and host rocks are consistent with deposition fromvolcaniclastic turbidity currents, debris flows and suspensionin a below-wave-base environment (McPhie and Allen, 1992).The footwall pumice breccias are interpreted to represent thesubmarine deposits from a large, felsic explosive eruption(Allen, 1994a). The host rocks may have been derivedfrom water-settled suspension sedimentation or the influxof volcaniclastic turbidites from distal rhyolitic volcaniccentres (Large et al., 2001b). The hanging wall volcaniclasticunits probably comprise the medial to distal facies from anextrabasinal felsic volcanic centre (Allen, 1994a).

A below-wave-base submarine setting for the Rosebery-Hercules succession is indicated by the presence of sedimentarystructures in the footwall pumice breccias consistent withdeposition from cold water-supported gravity flows andwater-settled fall, rare intercalated black pyritic mudstone andthe associated VHMS deposits (Gifkins and Allen, 2002).

FIGURE 7.21 | Long-section of the Rosebery mine, western Tasmania, showing the drives and main ore lenses, labelled alphabetically (provided by

Zinifex Rosebery mine, 2004). K lens is at depth at the north end of the mine.



Four strata-parallel altered zones enclose the Rosebery orelenses. From the periphery to the core of the alteration systemthese are: sericite zone, chlorite zone, Mn-carbonate zone,and quartz + sericite zone (Fig. 7.5C: Large et al., 2001b).

Outside the altered footwall zone, least-altered rhyoliticvolcanic rocks contain plagioclase crystals in a pumice- andshard-rich matrix (e.g. data sheet RBI). This matrix is weaklysericite + chlorite + quartz altered, commonly enhancing theshard and pumice textures. Plagioclase crystals are weaklyaltered with disseminated fine-grained sericite and albiteovergrowths. These rocks have AI = 30—60 and CCPI = 15—40 and plot in the least-altered box of the Alteration box plot.Data sheet RB2 is an example of the least-altered hanging wallrocks.

The outer part of the hydrothermal alteration system is abroad sericite zone with scattered Mn-carbonate blebs, whichextends up to 300 m into the footwall, but less than 25 minto the hanging wall rocks (Fig. 7.5C). It extends along theupper contact of footwall pumice breccia for at least 1000 mbeyond the ore lenses (Large et al., 2001b). One textural-compositional variant of this enveloping zone is representedin alteration facies data sheet RB3. The white mica contents ofthe facies varies from about 20 to 60% as plagioclase crystalsare increasingly replaced by carbonate and sericite, and theglass shard-rich matrix by fine sericite, with proximity tomassive sulfide. The AI increases from 60 to 95 as the sericiteproportions increase.

Intense, schistose chlorite alteration facies (e.g. data sheetRB5) is concentrated in the immediate footwall of the orelenses forming a thin (typically less than 5 m thick) chloriticzone, which is commonly thickest (5-10 m) beneath the Cu-rich sulfide lenses at the south end of the mine. This alterationfacies has variable chlorite (15—50 wt%) and sulfide (10—30 wt%) contents. The AI is between 95 and 100, and theCCPI between 70 and 90.

Commonly overlying the ore lenses is a zone of intense,proximal Mn-carbonate alteration facies up to 10 m thick(e.g. data sheet RB6), which is closely associated with massivesulfide, but locally extends several tens of metres beyond thelimits of known sulfide lenses. The intense, proximal Mn-carbonate alteration facies typically has a spotty texture, with25—60% Mn-carbonate spots in a sericitic, or locally chloritic,matrix with low sulfide content. Carbonate composition in thisfacies varies from rhodochrosite (MnCO3), to manganosiderite((Mn,Fe)CO3) and kutnahorite (CaMn(CO3)2) (Braithwaite,1974; Large etak, 2001b).

Typically the massive and semi-massive sulfides occurin strata-parallel zones of intense quartz + sericite alterationfacies. This alteration facies has a bleached appearance withtextures that vary from massive-pervasive to spotty and augen-schist textured (e.g. data sheet RB4). The latter comprises ananastomosing fabric of strongly foliated sericite dominateddomains wrapping around siliceous knots of quartz withminor sericite. The intense quartz + sericite alteration faciescontinues laterally beyond the margins of the ore lenses, whereit contains 1-10% disseminated sulfides (pyrite, sphalerite,galena).

Genesis of the ore lenses and alteration system

Most previous workers have interpreted the Rosebery depositto be synvolcanic exhalative in origin (e.g. Braithwaite, 1974;Green et al., 1981; Huston and Large, 1987; Green and Iliff,1989; Large, 1992; Khin Zaw et al., 1999). Solomon andGroves (1994) consider that the sheet-like form, stratiformsulfide banding, large size and high Zn-Pb metal contentindicate that Rosebery formed within a brine pool fromrelatively high-salinity fluids, similar to the genesis of manySEDEX deposits. However, the ore lenses are not associatedwith well-developed stringer sulfide zones or alteration pipestypical of seafloor systems. Instead, there are footwall zonesof disseminated sulfides, with altered zones that are alignedparallel to the ore lenses and the volcanic strata. Thesefeatures, combined with textures in the massive sulfidesindicative of replacement, suggest that the ore lenses didnot form immediately above hydrothermal vents, but mayhave formed from lateral fluid flow, either on the seafloor,or below the seafloor by replacement of the fine-grained topsof permeable pumice breccia units (Fig. 7.11: Allen, 1994a;Doyle and Allen, 2003; Martin, 2004).

196 | CHAPTER 7

Weak, selective-pervasive albite + quartz * sericite alteration faciesLeast-altered footwal!

RB1

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

139602

weak, selective-pervasive albite + quartz+ sericite

footwall

Kershaw Pumice Formation (CVC)

Mount Read Volcamics

feldspar-phyric pumice breccia

plagioclase

plagioclase crystals and cm-sizedplagioclase porphyritic tube pumice clasts

rhyolite

massive, clast supported, poorly sorted

subaqueous mass flow deposit

albite + sericite + quartz > chlorite + pyrite+ hematite

selective-pervasive, disseminated, foliated,sericite ± chlorite fiamme, sericite +hematite stylolites, albite + sericite alteredplagioclase, feldspar overgrowths onplagioclase

regional

good

weak

synvolcanic to burial

diagenetic

GeochemistrySiO2

TiO2

AI2O3

Fe2O3

MnOMgO

CaONa2O

73.26

0.32

14.13

2.11

0.03

0.97

0.99

3.87

K2O

P 2 O 5

S

CO2

Total

LOI

2.38

0.04

0.01

98.20

RbSrBa

CuPbZn

SbTl

110257866

13

26

0.00.0

ZrNbY

AlCCPI

Ti/Zr

25714.2

33

4131

7.5



Subtle, selective albite * quartz + sericite alteration fadesLeast-altered hanging wall

RB2

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

139586

subtle, selective albite + quartz + sericite

hanging wall

White Spur Formation


feldspar > quartz crystal-rich pumiceoussandstone

plagioclase and quartz

clastic (feldspar and quartz crystals, andpumice shards)

Primary composition rhyolite-dacite

Lithofacies massive

Interpretation subaqueous mass-flow deposit

Alteration minerals albite + sericite + quartz + chlorite

selective clast alteration, disseminatedsericite, albite + sericite-altered plagioclase

regional

good

subtle

synvolcanic

diagenetic

Alteration textures

Distribution

Preservation


Timing

Alteration style

GeochemistrySiO2 74.40 K2O 1.99 Sr 261 Nb 10.8TiO2 0.35 P2O5 0.05 Ba 1472 Y 33AI2O3 14.34 S 0.03 Cu 3 :

Fe2O3 2.06 CO2 0.27 Pb 62 Al 34MnO 0.04 Total 99.26 Zn 142 CCPI 27MgO 0.57 LOI Sb 0.0 Ti/Zr 10.6CaO 0.58 Tl 0.0Na2O 4.41 Rb 75 Zr 197


198 I CHAPTER 7

Moderate, foliated sericite alteration faciesFootwall

KB 3


Sample no. 139747

Alteration facies moderate, foliated sericite

Location footwall

Formation Kershaw Pumice Formation


Volcanic facies feldspar-phyric pumice breccia

Relict minerals plagioclase

Relict textures porphyritic, fibrous tube pumice clasts


Lithofacies massive, normally graded

Interpretation syneruptive, mass-flow-emplaced deposit

Alteration minerals sericite + albite + quartz + carbonate

Alteration textures foliated, schistose, stylolites, fiamme?,fractured and albite-altered plagioclase,

quartz veinlets

Distribution local

Preservation poor

Alteration intensity moderate

Timing synmineralisation

Alteration style hydrothermal and metamorphic GeochemistrySiO2 71.46 K2O 4.16 Rb 182 Zr 229.9TiO2 0.32 P2O5 0.06 Sr 56 Nb 12.6AI2O3 12.93 S 0.06 Ba 1042.5 Y 32Fe2O3 2.44 CO2 1.87 Cu 3MnO 0.11 Total 98.02 Pb 5 Al 67MgO 1.65 LOI 3.89 Zn 29 CCPI 41CaO 1.49 Sb 2.7 Ti/Zr 8.3Na2O 1.36 Tl 0.9

Intense, augenFootwal!

schistose quartz + seriate alteration facies


RB4

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

139778

intense, augen schistose quartz + sericite

footwall



feldspar-phyric pumice breccia

nil

foliated clasts (pumice)

rhyolite

massive

indeterminate

sericite + quartz + sulfides

augen schistose, sericite + sulfidecleavagelocal

poor

intense

synmineralisation

hydrothermal and metamorphic

GeochemistrySiO2

TiO2

Al2O3

Fe2O3

MnOMgO

CaONa2O

72.950.27

12.86

2.580.17

1.14

0.270.01

K2O

P 2 O 5

S

co2TotalLOI

4.240.041.94

0.4098.19

3.57

RbSrBa

CuPbZn

SbTl

20218

1767

23943006900

7.54.5

ZrNbY

AlCCPI

Ti/Zr

22413.2

35

9545

7.2


2 0 0 I CHAPTER 7

Intense, schistose chlorite alteration faciesFootwall

•RB5

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

139743

intense, schistose chlorite

footwall



feldspar-phyric breccia

nil

porphyritic?

rhyolite

massive

indeterminate

chlorite + pyrite + sphalerite + quartz +sericite

foliated, schistose

local

poor

intense

synmineralisation

hydrothermal and metamorphic

GeochemistrySiO2

TiO2

MAFe2O3

MnOMgO

CaONa2O

44.22

0.33

14.48

16.58

1.49

2.22

0.10

0.05

K2O

P 2 O 5

S

co2Total

LOI

3.37

0.06

7.18

0.95

91.03

7.55

RbSrBa

CuPbZn

SbTl

17414

948

1678

604

68200

3.97.1

ZrNbY

AlCCPI

Ti/Zr

24712.0

35

9783

8.0



Intense, proximal In-carbonate alteration faciesHanging waff

RB6

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

139740

intense, proximal Mn-carbonate

hanging wall

White Spur Formation


feldspar-phyric pumice breccia?

nil

rare feldspar crystals

dacite

massive

indeterminate

rhodochrosite + sericite + pyrite

nodular-spheroidal rhodochrosite,disseminated pyrite, sericite-alteredfeldspar

local, immediate hanging wall ofsulfidelens

poor

intense

synmineralisation

hydrothermal

GeochemistrySiO2

TiO2

AI2O3

Fe2O3

MnOMgO

CaONa2O

16.79

0.39

10.73

7.50

29.36

1.97

2.04

0.05

K2O

P 2 O 5

Sco2Total

LOI

Rb

3.64

0.13

3.24

22.20

98.64

23.35

183

SrBa

CuPbZnSb

TIZr

274297

206749958.8

24.6

142

NbY

AlCCPI

Ti/Zr

7.220

7370

16.5


I CHAPTER 7

7.7 | WESTERN THARSIS: A HYBRID Cu-Au VHMS DEPOSIT

Western Tharsis is a stratabound disseminated pyrite + chalco-pyrite deposit in the northwestern part of the Mount Lyellmining field, western Tasmania (Fig. 1.5). It contains around12.4 Mt at 1.3% Cu and 0.3 g/t Au (Huston and Kamprad,2001). Although discovered in 1897, it is the only non-exploiteddeposit of at least 22 deposits in the Mount Lyell field, whichtogether produced a total of 113 Mt of ore at average grades of1.36% Cu, 6.8 g/t Ag and 0.4 g/t Au (Corbett, 2001).

The mineralised zone is a sub-vertical stratabound lens,up to 150 m thick and narrowing towards the surface, witha down dip extent of greater than 1000 m and strike extentof about 300 m. Most of the deposit consists of disseminatedpyrite + chalcopyrite in a gangue of quartz + sericite ± chloriteand, locally, magnetite. Smaller bornite-rich mineralisedzones similar to the North Lyell ores exist in the upper parts,particularly associated with quartz and quartz + pyrophyllitealteration assemblages. The bornite zones also contain minorchalcocite, chalcopyrite, mawsonite, digenite, enargite, molyb-denite, woodhouseite and barite.

Geological setting

The Lyell ore bodies and their altered halos are focussed alongthe Great Lyell fault and occur at a variety of stratigraphicintervals in the Central Volcanic Complex and in theoverlying Mount Julia Member of the Comstock Formation,Tyndall Group (Corbett, 2001). Western Tharsis is situatedin a steeply west-dipping, overturned, east-facing successionof altered intermediate to felsic volcanic rocks assigned to theCentral Volcanic Complex of the Middle Cambrian MountRead Volcanics. In the Mount Lyell area, these volcanic rocksare reverse-faulted to the east against the late Cambrian toEarly Ordovician siliciclastic conglomerate and sandstone ofthe Owen Group. All the Mount Lyell field deposits lie withina 6 x 1 km pyritic altered zone adjacent to the complex faultcontact (Corbett, 2001).

Two units of rhyolitic volcaniclastic rocks with subordinateinterbedded volcanogenic sandstone and siltstone comprisethe immediate stratigraphic footwall and host rocks atWestern Tharsis (Huston and Kamprad, 2001). These units,each several hundred metres thick, contain some ash- tolapilli-sized clasts but primary volcanic textures are typicallyobscured by alteration and their volcaniclastic origin is largelyinterpretative. They are separated by a 10-50 m thick groupof andesitic volcaniclastic rocks and locally amygdaloidalcoherent lavas or sills. The stratigraphic hanging wall consistsof a 200—300 m thick complex of intercalated felsic andintermediate volcaniclastic rocks. A thin unit of felsic quartzporphyry, possibly a correlate of the lower Tyndall Group,occurs between the altered Central Volcanic Complex and thefaulted contact with the Owen Group. In this part of the field,the (North Lyell) fault contact dips at 70° to the southwest.Deep drilling indicates that the Western Tharsis mineralisedzone may intersect the fault at around 1500 m below surface(Corbett, 2001).

The presence of thick, graded beds in the Central VolcanicComplex at Lyell indicates a subaqueous environment of

deposition for the host succession. The occurrence of exhala-tive massive sulfide bodies, limestone with shallow marinefauna, and welded ignimbrite in the Tyndall Group atComstock suggest a shallow marine setting for mineralisation(Jago et al., 1972; Corbett et al., 1974; White and McPhie,1997; Corbett, 2001).


Corbett (2001) showed that the deposit is enclosed by a400-500 m wide zone of quartz + sericite + pyrite schistadjacent to the North Lyell fault. This is part of a pyritic corezone extending 4 km from the Lyell Highway to the LyellComstock mine. This proximal, strong to intense, feldspar-destructive altered zone grades outwards to less intense sericite+ chlorite alteration facies in felsic and intermediate volcanicrocks. At surface above Western Tharsis and around the uppermineralised zone, the proximal quartz + sericite + pyritealteration facies includes numerous bodies up to 20 m across,of microcrystalline quartz ± pyrite, termed silica heads.

Huston and Kamprad (2001) subdivided the WesternTharsis system into five main alteration facies. An intense,pervasive, proximal quartz + chlorite + pyrite ± sericite alter-ation facies (e.g. data sheet WT8) exists in the chalcopyrite +pyrite mineralised zone at depths greater than 350 m belowsurface.

An intense, proximal quartz + pyrophyllite + pyritealteration facies (e.g. data sheets WT4 and 5) occurs in a150 m wide zone associated with the bornite + chalcopyritemineralised zone between 100 and 400 m below surfaceand in a 50 m thick zone extending along the stratigraphicfootwall to 750 m below surface. This facies includes narrowzones of quartz + topaz assemblages (particularly in the upperparts, which may correspond to Corbett's silica heads) andlocally minor phases including fluorite, barite, zunyite andwoodhouseite.

A strong, pervasive, medial quartz + sericite + pyrite alter-ation facies (e.g. data sheets WT3, 6 and 7), which enclosesthe two proximal facies above, and extends up to 150 moutwards into the stratigraphic footwall and through most ofthe hanging wall. The outer margins, adjacent to the weak,medial chlorite + sericite ± carbonate alteration facies andin upper part of hanging wall succession, contain minordisseminated carbonate.

A weak, pervasive, medial chlorite + sericite ± carbonatealteration facies (e.g. data sheet WT2) occurs in the peripheralzones. It exists in both footwall and hanging wall, about 150-200 m outwards from the mineralised zone, particularly inandesitic volcanic rocks with minor pyrite or hematite.

A weak, selective quartz + chlorite + sericite + carbonatealteration facies (e.g. data sheet WTl) grades westward intoleast-altered rocks composed of quartz + albite + chlorite (±sericite and carbonate).

Sericite compositions in the outer zones are slightlyphengitic (up to 0.5 Fe + Mg atoms per formula unit) gradingto essentially non-phengitic and slightly sodic (molecular Na/Na+K <0.15) in the proximal to medial zones. This variationhas potential as a deposit-scale exploration vector, which canbe effectively measured by short wavelength infrared (SWIR)spectral analysis (Herrmann et al., 2001). SWIRspectrometry

is similarly effective in delineating pyrophyllite, topaz andzunyite-bearing zones (see Section 8.2 for more detail).

Chlorite compositions are moderately Fe-rich (molecularMg/Mg + Fe ratios = 27-48) with a subtle trend to Fe enrichmenttowards the mineralised zone. Carbonates in the distal tomedial alteration facies are ankeritic to sideritic in composition.They show a distinct trend of Fe enrichment from the footwalltowards the mineralised zone (Huston and Kamprad, 2001).Carbonates in the hanging wall are moderately manganiferous(up to 0.2 Mn atoms per formula unit).

Ore genesis

Metallogenic interpretations of the Mount Lyell depositshave fuelled geological debate for over a century and remaincontroversial today (Corbett, 2001; Huston and Kamprad,2001). Early models that related mineralisation to Devonianor Cambrian intrusions were succeeded, during the 1960s,by acceptance of Cambrian synvolcanic origins. In the 1980sand early 1990s, the North Lyell type bornite ores werepopularly attributed to re-mobilisation during Devoniandeformation. Large et al. (1996) revived the magmaticconnection, interpreting Cambrian granites to be the sourceof hydrothermal fluids and metals. In recent years, a magmaticconnection has been further supported by wider recognitionof advanced argillic type assemblages, which are consistentwith the involvement of magmatic volatiles and acidic fluids.

Nevertheless, there is still disagreement over the timingof mineralisation. Huston and Kamprad (2001) pointedto an apparent (Pb-isotopic) 40 Ma age difference betweenstratiform synvolcanic Pb + Zn + Cu sulfide lenses at LyellComstock and the disseminated Cu + Au deposit at PrinceLyell. They suggested a two event history: Middle Cambrian

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS | 20 3

syngenetic stratiform Pb + Zn + Cu mineralisation followedby a 40 Ma period of tectonism that culminated in high-sulfidation type Cu + Au mineralised zones derived fromdeep Ordovician granites. However, the more extensivefield evidence gathered by Corbett (2001) indicates that allthe alteration and mineralisation was restricted to MiddleCambrian, and ceased during deposition of the lower part ofthe Tyndall Group.

Rather than a temporal overprinting of different styles,Corbett (2001) envisaged a single, vertically extensive,submarine volcanic, hybrid magmatic-seawater hydrothermalsystem. It produced disseminated chalcopyrite + pyrite (andlocally magnetite + apatite) mineralised zones in the deeperparts, high-sulfidation type bornite mineralised zones andintense siliceous altered zones in the upper subseafloor zones,and deposited exhalative Pb + Zn + Cu massive sulfide lensesat the seafloor. The Western Tharsis zone encompasses thetransition between deep chalcopyrite + pyrite and upper high-sulfidation types of mineralisation (Fig. 7.22).

Corbett's diagrammatic representation shows the systemas sub-vertical, cutting through sub-horizontal volcanicstrata and focussed along or adjacent to the Great Lyell fault.However, the Western Tharsis deposit appears to be sub-vertical and stratabound. This is possibly a misinterpretation;primary volcanic textures and facies associations are largelyobscured in the intensely altered zones. Furthermore, Corbett's(2001) model suggests diapir-like upward movement of thephyllosilicate-rich altered volcanic rocks on the hanging wallside of the fault zone, which may have disrupted the volcanicsequence.

The arguments about Mount Lyell are not yet settled.Nevertheless, the emerging recognition of high-sulfidationore deposits may renew interest in exploration in westernTasmania.

pyntic core zone(senate + chlorite + pyrite

silica schists)

disseminatedchalcopyrite-pyrite bodie

FIGURE 7.22 | Cross-section model of the Mount Lyell,

vertically extensive, submarine, hybrid magmatic-seawater

hydrothermal, alteration and mineralisation system, western

Tasmania (modified after Corbett, 2001).

2 0 4 | CHAPTER 7

Weak, selective chlorite + sericite + quartz + carbonate alteration faciesLeast-altered footwall

WT 1

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

113084

weak, selective chlorite + sericite + quartz+ carbonate>250 m stratigraphicaliy below mineralisedzone



massive plagioclase + quartz-phyricrhyolitephenocrysts of albitised plagioclase >quartz

porphyritic

rhyolite

massive

rhyolite lava

sericite + chlotite + carbonate > quartz

selective-pervasive, matrix altered to 20-40 |jm chlorite, aligned sericite cleavage,albite + sericite-altered plagioclase

regional?

moderate

weak

synvolcanic and syndeformation

diagenetic and tectonic-metamorphic

GeochemistrySiO2 75.69 K20 3.33 Sr 37 Nb 13TiO2 0.26 P2O5 0.04 Ba 243 Y 40AI2O3 13.00 S 0.37 Cu 3Fe2O3 2.45 CO2 1.17 Pb 14 Al 65MnO 0.04 Total 99.32 Zn 28 CCPI 40MgO 0.78 LOI 3.12 Sb 0.7 Ti/Zr 5.4CaO 1.04 Tl 0.9Na2O 1.15 Rb 112 Zr 291



Weak, pervasive, medial chlorite + sericite ± carbonate alteration facies y\/T 2

Sample no. 113086

Alteration facies weak, pervasive, medial chlorite + sericite± carbonate

Location -200 m stratigraphically belowmineralised zone

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style



altered andesite

nil

nil

andesite

massive

indeterminate

chlorite + carbonate > sericite

pervasive, foliated, carbonate-alteredplagioclase, carbonate veins

poor

weak

synvolcanic and subsequentsyndeformationdiagenetic and tectonic-metamorphic

Geochemistry

SiO2 47.78 K20 1.47 Rb 46 Zr 72TiO2 0.51 P2O5 0.09 Sr 83 Nb 3AI2O3 14.56 S 0.04 Ba 367 Y 17Fe2O3 12.59 CO2 7.84 Cu 10

MnO 0.19 Total 97.30 Pb 6 Al 45MgO 4.63 LOI 10.71 Zn 199 CCPI 82

CaO 5.53 Sb 0.8 Ti/Zr 42.5Na2O 2.07 Tl 0.5


2 0 6 | CHAPTER 7

Strong, pervasive, medial quartz + sericite + pyrite alteration facies WT 3

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

113092

strong, pervasive, medial quartz + sericite+ pyrite-100 m stratigraphically below mineralisedzone



volcaniclastic rhyolite

nil

relict granular texture, possibly clastic

rhyolite

indeterminate

quartz + sericite + siderite

pervasive, mosaic of sutured 20-600 |jfractured quartz grains with interstitialsericite and patches coarse siderite

poor

strong

synmineralisation

hydrothermal

GeochemistrySiO2 70.58 K2O 3.63 RbTiO2 0.23 P2O5 0.04 SrAI2O3 12.12 S 0.02 BaFe2O3 8.03 CO2 3.11 CuMnO 0.11 Total 98.92 PbMgO 0.68 LOI 4.44 ZnCaO 0.16 SbNa,0 0.21 TI

126 Zr 246

19 Nb 13

758 Y 42

18

5 Al 92

63 CCPI 67

2.5 Ti/Zr 5.60.9



Intense, proximal quartz + pyrophyllite + pyrite alteration facies yyi 4


Sample no. 113102

Alteration fades intense, proximal quartz + pyrophyllite +pyrite

Location proximal to upper mineralised zone



Volcanic fades altered rhyolite

Relict minerals nil

Relict textures nil


Lithofacies

Interpretation indeterminate

Alteration minerals quartz + pyrophyllite > sericite + pyrite

Alteration textures pervasive, mosaic of sutured 40-400 prn, quartz grains with ragged patches of semi-

aligned pyrophyllite, minor sericite anddisseminated, fractured pyrite

Distribution

Preservation poor


Timing

Alteration style hydrothermal

Geochemistry

SiO2 80.78 P2O5 0.05 Cu 72 Al 80TiO2 0.24 S 1.61 Pb 7 CCPI 54AI2O3 12.25 CO2 0.40 Zn 6 Ti/Zr 5.8Fe2O3 2.21 Total 99.49 Sb 0.6MnO 0.01 LOI 3.08 Tl <0.5MgO 0.12 Zr 249CaO 0.03 Rb 31 Nb 13Na2O 0.36 Sr 117 Y 5K2O 1.43 Ba 1170

2 0 8 | CHAPTER 7

Intense, proximal quartz + pyrophyllite + pyrite alteration facies WT 5

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

113105

intense, proximal quartz + pyrophyllite +pyrite

-200 m straigraphicaily above mineralisedzone



altered rhyolite

rhyolite

indeterminate

quartz + topaz + pyrite > carbonate

pervasive, domainal 50-100 pmmicrocrystalline quartz and granular topaz,minor disseminated pyrite, irregular mm-scale patches > carbonate veinlets

poor

intense

hydrothermal


Geochemistry

SiO2 68.25 P2O5 0.07 Cu 10 Al 36TiO2 0.28 S 0.37 Pb 21 CCPI 96AI2O3 18.13 CO2 3.11 Zn 14 Ti/Zr 4.6Fe2O3 1.35 Total 95.14 Sb 0.1MnO 0.30 LOl 8.19 Tl <0.5MgO 1.12 Zr 364CaO 2.05 Rb 1 Nb 14Na2O 0.05 Sr 33 Y 10K2O 0.06 Ba 171


Strong, pervasive, medial quartz + sericite + pyrite alteration facies yyy g

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

113110

strong, pervasive, medial quartz + sericite+ pyrite-100 m stratigraphically above mineralisedzone



altered dacite

dacite

indeterminate

quartz + sericite + pyrite > chlorite andcarbonatepervasive, 50-100 |jm microcrystallinequartz with interstitial shreds and seamsof aligned sericite, disseminated euhedralpyrite, some highly deformed and re-crystallised quartz + carbonate > chloriteveins/patches

poor

strong

hydrothermal

Geochemistry

SiO2

TiO2

AI,0'2^3

Fe2O3

MnOMgOCaONa2O

60.290.35

12.8410.770.521.081.260.13

K2O

Sco2

TotalLOI

3.970.086.282.64

100.218.09

RbSrBaCuPbZnSb

120 Tl25 Zr

1737 Nb88 Y

214138 Al1.0 CCPI

Ti/Zr

1.0182

923

7872

11.5


2 1 0 I CHAPTER 7

Strong, pervasive, medial quartz + sericite + pyrite alteration facies WT 7

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

113284

strong, pervasive, medial quartz + sericite+ pyrite

deep mineralised zone



altered rhyolite

nil

nil

rhyolite

indeterminate

quartz + sericite + pyrite + chalcopyrite

pervasive: 50-100 pm microcrystallinequartz, interstitial shreds and seams ofsericite, disseminated euhedral pyrite,> chalcopyrite

poor

strong

hydrothermal

GeochemistrySiO2

TiO2

AI2O3

Fe2O3

MnOMgOCaONa2O

72.750.22

10.116.110.010.360.330.05

K2O

SCO2

TotalLOI

2.800.354.170.02

97.284.22

RbSrBaCuPbZnSbTl

64 Zr111 Nb

1330 Y13800

21 Al32 CCPI

99.2 Ti/Zr<0.5

200

11

12

676.6



Intense, pervasive, proximal quartz + chlorite + pyrite ± sericite alteration facies WT 8

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

113264

intense, pervasive, proximal quartz +

chlorite + pyrite ± sericite

deep mineralised zone



altered dacite

dacite

indeterminate

quartz + sericite + chlorite + pyrite chal-copyritepervasive: mosaic of 40-200 pm dustyquartz, interstitial and anastomosingsericite cleavage, irregular domains ofchlorite + sulfides

Distribution

Preservation


Timing

Alteration style

poor

intense

hydrothermal

Geochemistry

SiO2

TiO2

AI2O3

Fe2O3

MnO

MgO

CaO

58.74

0.49

13.6512.660.231.750.26

K2O

P2O5

S

co2TotalLOI

3.450.275.380.1797.08

6.73

Na2O 0.03

RbSrBaCuPbZnSbTl

92362097120001599

0.9

<0.5

Zr 167Nb 13Y 25

Al 95CCPI 79Ti/Zr 17.6


212 | CHAPTER 7

7.8 | HENTY: A VOLCANOGENIC GOLDDEPOSIT

The Henty volcanogenic gold mine and nearby Mount Juliagold prospect are hosted by Cambrian Mount Read Volcanics,near the junctions of the North and South Henty faultsand the Great Lyell fault (Fig. 1.5). The deposit comprisesat least six steeply dipping, thin, stratabound, disconnectedsiliceous lenses of up to a few hundred metres vertical extent,distributed over about 2.5 km of strike length (Callaghan,2001). The estimated total geological resource in December2001 was 2,154,000 tonnes at 12.1 g/t Au (838,800 oz.).This included production of 820,000 tonnes at 17.5 g/t Au(462,000 oz.).

Geological setting

The ore lenses lie in a laterally extensive but narrow strataboundaltered zone (A-zone of Callaghan, 2001) at the stratigraphicboundary between the Central Volcanic Complex and thebase of the Tyndall Group. In the Henty area these unitstrend NNW to NNE and face east, with steep easterly toslightly overturned steep westerly dips. The Henty fault zone,trending about 015° and dipping at 70° to the west, obliquelytruncates the volcanic succession. The stratabound altered andmineralised zones occur in the immediate footwall of the faultzone, extending about 200 m down-dip from the fault (Halleyand Roberts, 1997; Callaghan, 2001). The intersection of thefault and the favourable stratigraphic horizon plunges at alow angle to the south. This is a consequence of the gradualchange in trend and slight overturning of the host sequence,from NNW with steep easterly dip in the south, to NNEand steep westerly dip in the northern part of the mine area(Halley and Roberts, 1997). As in many Au deposits, gradecut-offs rather than lithological differences define the orezones (Callaghan, 1998). Most of the high-grade ore exists inthin lenses or sheet-like bodies up to 7 m thick in the intense,massive quartz (MQ) alteration facies (Halley and Roberts,1997) but this facies is not uniformly auriferous (Callaghan,2001). The stratigraphic upper part of the mineralised A-zonetypically has a high disseminated base-metal-sulfide content oris spatially associated with lenses of massive pyrite or massiveto banded sphalerite + galena (Penney, 1998). Discontinuousmassive pyrite lenses up to 2 m thick exist at this stratigraphiclevel for 600 m of strike but extend less than 150 m down-dipfrom the Henty fault.

Feldspar-phyric to aphyric dacitic lavas and rare basalticlavas intercalated with dacitic to basaltic hyaloclastite andpolymictic volcanic breccias of the Central Volcanic Complexdominate the stratigraphic footwall between the Hentyfault and A-zone (e.g. data sheet HN1: Callaghan, 1998).The footwall succession includes discontinuous calcareousvolcaniclastic units and hematitic fossiliferous limestone.Immobile element ratios indicate that the protoliths of thealtered and mineralised zone were compositionally uniformdacitic volcanic units.

The stratigraphic hanging wall, immediately east of theA-zone, comprises massive, andesitic, feldspar crystal-richvolcanic sandstone (e.g. data sheet HN2), dacitic volcanicbreccia, lavas and polymictic volcano-sedimentary breccia,

intercalated with calcareous volcanic sandstone, hematiticfossiliferous limestone and minor mudstone. This lithologicallydiverse part of the hanging wall succession, up to 200 mthick, is recognised as the Lynchford Member: the lowermostunit of the Tyndall Group (Callaghan, 2001). It is succeededeastwards by the Mount Julia Member consisting of gradedrhyolitic breccia, quartz + feldspar-rich volcanic sandstone andminor siltstone intruded by southward thickening quartz +feldspar porphyritic rhyolite sills, cryptodomes and associatedhyaloclastites. Overlying this is a thick succession of quartz-rich epiclastic sandstone and volcanolithic conglomerate (ZigZag Hill Formation), which passes conformably eastwardsinto siliciclastic and micaceous sandstone and conglomerate.

White and McPhie (1996) interpreted the massivecrystal-rich volcanic sandstone of the Lynchford and MountJulia Members as originating from large subaerial or shallowmarine explosive eruptions that produced pyroclastic flows,which transgressed into a shallow marine environment. Thefossil assemblage in the limestone units (Jago et al., 1972)and local welded ignimbrite units in the Mount Julia Member(White and McPhie, 1996) also indicate that the Henty hostrocks, or at least those immediately overlying the mineralisedzone, were deposited in a near-shore, shallow-marine setting.


The distribution of alteration facies in this elongate andstratabound alteration system reflects decreasing alterationintensity down-dip away from its intersection with the Hentyfault zone, and differing thermo-chemical conditions fromfootwall to hanging wall. Table 7.2 summarises the features ofthe Henty—Mount Julia alteration facies.

A moderate, footwall sericite + quartz ± carbonate alter-ation facies (MA) occupies the wedge of stratigraphic footwallbetween the Au-bearing A-zone and the Henty fault.

The A-zone has a discontinuous sheet-like inner zoneof intense, massive quartz alteration facies (MQ, e.g. datasheet 7), which is composed of microcrystalline quartz withmultiple generations of fine veinlets of quartz + calcite ±sulfides (pyrite, chalcopyrite and galena, which contain mostof the Au). It grades outwards (stratigraphically up and downas well as down-dip away from the fault) through intense,proximal, domainal quartz + sericite alteration facies (MV,e.g. data sheet HN6) to intense, foliated sericite + quartz +chlorite + pyrite alteration facies (MZ, e.g. data sheet HN5).These enveloping alteration facies have progressively lessquartz and greater phyllosilicate contents and exist in variableproportions in different parts of the deposit. They typicallyhave sericitic groundmasses with foliated to schistose fabricsanastomosing around small siliceous domains that possiblyrepresent silicified lithic clasts (Callaghan, 2001). Massivepyrite lenses at the stratigraphic top of the A-zone gradelaterally and down-dip to massive carbonate alteration facies(CB, e.g. data sheet HN4) in peripheral areas. This faciesincludes massive rocks composed of carbonate + chlorite,and thin bands and lenses of carbonate that are difficultto distinguish from fossiliferous limestone. Some of thecarbonate bands contain fragments of red jasper.

A zone of strong, selective, hanging wall albite + quartzalteration facies (e.g. data sheet HN3) lies either immediately

above (Callaghan, 2001) or 20-40 m stratigraphically abovethe mineralised zone (Halley and Roberts, 1997). In proximalareas, the hanging wall albite + quartz alteration fades is up to100 m thick, possibly extending up to the base of the Zig ZagHill Formation (Callaghan, 1998). It is considered to be ofhydrothermal origin, distinct from regional-scale diageneticalbite alteration facies.

Ore genesis

Halley and Roberts (1997) interpreted Henty as a Au-richvolcanogenic massive sulfide deposit because of its associationwith conformable pyrite and carbonate lenses, colloformtextures in pyrite, the presence of red jasper clasts that resemblesiliceous exhalites, and C-, O- and Pb-isotopic data thatindicate a Cambrian synvolcanic origin for the strataboundalteration system. They suggested that its unusual high Au/Agratios, extent of footwall silicification and high proportion of


base-metal sulfides disseminated in the upper altered footwallzone were due to the shallow-marine, near-shore setting whereinput of meteoric water produced low-salinity hydrothermalfluids, which boiled and cooled at some depth below theseafloor. Although the Au in the intense, massive quartzalteration facies exists largely in late-stage veinlets related toDevonian brittle deformation and remobilisation, there is noevidence for addition of metals during this event.

Callaghan (2001) proposed synvolcanic inputs ofmagmatic volatiles, fluids and metals to account for the Au+ Cu + Bi + Ag + Te metal association and Al mobility in theintense, massive quartz alteration facies, which are atypical ofseawater dominated VHMS systems. He envisaged a low pH,high salinity, submarine, subseafloor type of pulsed magmaticplus seawater, high-sulfidation epithermal system. This modelinvokes the proto-Henty fault as a magmatic volatile and fluidconduit that reactivated during Devonian deformation todislocate the eastern and western halves of the hydrothermalsystem.

Table 7.2 | The Henty-Mount Julia alteration facies and their defining characteristics (Callaghan, 1998).

Alteration facies

Intense, massivequartz

Intense, proximalquartz + sericite

Intense, peripheralsericite + quartz +chlorite + pyrite

Massive carbonate

Moderate, footwallsericite + quartz ±carbonate

Strong, hanging wallalbite + quartz

Code

MQ

MV

MZ

CB

MA

AS

Mineral assemblage

Quartz ± (carbonate, sericite,pyrite, chalcopyrite, galena,gold)

Quartz + sericite ±(carbonate, pyrite,chalcopyrite, galena,sphalerite)

Sericite + quartz + pyrite+ chlorite ± (carbonate,chalcopyrite, galena)

Calcite ± chlorite

Sericite + quartz ±(carbonate, pyrite)

Albite + quartz ± (chlorite)

Sulfides(%)

~2

0.1 to 5

2-10

<10

<2

0

Gold

(g/t)

Variable;average 36

0.1 to 1

0.5 to 2

?

?

0

Distribution

Thin lenses in core of A-zone.

Enclosing and gradational to intense massivequartz alteration facies.

Peripheral, enveloping the intense proximalquartz + sericite alteration facies.

Discontinuous stratiform lenses at stratigraphictop of A-zone in peripheral parts of the system;laterally equivalent to massive pyrite lenses.

Stratigraphic footwall, in felsic volcanic rocksbetween Henty fault and A-zone.

Directly adjacent to A-zone and extending up to100 m into hanging wall succession.

2 1 4 | CHAPTER 7

Moderate albite + chiorite + calcite alteration faciesLeast-altered host rock

HN1

Sample no.

Alteration Facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

255005

moderate albite + chlorite + calcite

down dipofA-zone



massive plagioclase-phyric dacite

albitised plagioclase phenocrysts

porphyritic

dacite

massive to brecciated

dacite lava

albite + chlorite + calcite + quartz

selective-pervasive in irregular chlorite and

calcite veinlets and blebs, albite ± sericite-

altered plagioclase

regional

moderate

moderate

synvoicanic plus subsequent fault-related

deformation

diagenetic and tectonic deformation

GeochemistrySiO2

TiO2

AI2O3

Fe2O3

MnO

MgO

CaO

Na2O

58.40

0.46

12.50

4.42

0.13

1.51

8.71

5.05

K2O

S

co2Total

LOI

0.97

0.17

0.01

6.82

99.15

7.91

Au 0.005

Sr

Ba

Cu

Pb

Zn

Sb

Tl

Zr

114

327

7

8

139

1.2

0.5

188

Nb

Y

AlCCPI

Ti/Zr

40.9

28

1548

14.7



Weak, selective aibite + chlorite + caicite alteration faciesLeast-aitered hanging wall

HN2

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

255024

weak, selective aibite + chlorite + caicite

lower hanging wall unit + down dip of A-

zone

Lynchford Member (Tyndall Group)


massive, feldspar crystal-rich volcanicastic

sandstone

albitised plagioclase crystals

abundant sand-sized crystals and sparse

lithic clasts, subangular

andesite

massive to crudely banded, moderately

well sorted, matrix supported

volcaniclastic mass-flow deposit

aibite + chlorite + caicite > sericite

selective-pervasive aibite + chlorite-altered

matrix, irregular discontinuous caicite

veinlets, domainal microcrystalline quartz +

chlorite + pyrite

regional

moderate

weak

synvolcanic

diagenetic and tectonic deformation


GeochemistrySiO2 54.90 K2O 0.29 Au 0.005 Zr 130TiO2 0.64 P2O5 0.09 Sr 134 Nb 29.6AI2O3 14.40 S 0.08 Ba 127 Y 22Fe2O3 3.73 CO2 6.97 Cu 7MnO 0.19 Total 98.89 Pb 5 Al 12MgO 1.81 LOI 8.72 Zn 132 CCPI 42CaO 9.05 Sb 0.8 Ti/Zr 29.5Na2O 6.74 Tl 0.5

2 1 6 | CHAPTER 7

Strong, selective, hanging-wall albite + quartz alteration faciesHanging waff

HN3

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

255038

strong, selective, hanging wall albite +quartz (AS)

hanging wall, + 70 m stratigraphicallyabove A-zone

Mount Julia Member (Tyndall Group)


massive volcanicastic sandstone

quartz and albitised plagioclase crystals

sparse crystals and few lithic clasts in10-20 urn matrix

rhyolite

massive to crudely banded

volcaniclastic mass-flow deposit

albite + quartz > sericite > chlorite + calcite

selective-pervasive altered matrix,microcrystalline albite + quartz matrix withminor interstitial chlorite, quartz + calciteinfill irregular veinlets, sericite in laterparallel veinlets

local and stratabound in hanging wallsequence

moderate

strong

albite + quartz probably syn volcanic,sericite veinlets syn deformation

hydrothermal?

GeochemistrySiO2

TiO2

AI2O3

Fe2O3

MnO

MgO

CaO

Na2O

77.70

0.16

12.10

1.13

0.02

0.36

0.75

6.57

K2O

S

co2

Total

LOI

0.29

0.01

0.00

0.99

100.09

0.90

Au 0.005

Sr

Ba

Cu

Pb

Zn

Sb

Tl

Zr

47 Nb

100 Y

4

2 Al

100 CCPI

0.9 Ti/Zr

0.5

182

14.6

26

17

5.3


Massive carbonate alteration faciesHost-rock equivalent?

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

255050

massive carbonate

upper A-zone, laterally equivalent tomassive pyrite

Lynchford Member (Tyndall Group) orCentral Volcanic Complex


marine limestone with minor volcaniclasticcomponent

calcite + plagioclase + quartz

plagioclase and quartz crystal fragments

massive to thinly bedded

impure marine carbonate

calcite?

10-50 |jm microcrystalline calcite, calciteveins, stylolites

local and stratabound in peripheral upperpart of A-zone

moderate

weak

synvolcanic diagenesis, syndeformationaldynamic recrystallisation

diagenetic and tectonic-metamorphic,doubtful hydrothermal carbonatecomponent


HN4


GeochemistrySiO2 22.80 K2O 0.58 Sr 263 Nb 1.0TiO2 0.18 P2O5 0.08 Ba 476 Y 16AI2O3 4.16 S 0.72 Cu 21Fe2O3 2.33 CO2 24.70 Pb 76 A! 4MnO 0.26 Total 94.44 Zn 100 CCPI 60MgO 0.85 LOI 29.11 Sb 10.0 Ti/Zr 22.0CaO 36.40 Tl 0.5Na2O 1.38 Au 0.017 Zr 49

2 1 8 | CHAPTER 7

Intense, foliated sericite + quartz + chlorite + pyrite alteration faciesFootwali?

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

255030

intense, foliated sericite + quartz + chlorite+ pyrite (MZ)

peripheral altered zone



volcaniclastic breccia

minor quartz

blocky clasts

dacite

indeterminate

indeterminate

sericite + quartz + pyrite + chlorite

foliated, semi-mylonitic, disseminatedpyrite, non-foliated domains of quartz +calcite > sericite

local, enclosing mineralised lens

poor

intense

synmineralisation

hydrothermal and tectonic-metamorphic

HNS


GeochemistrySiO2 65.30 K2O 5.35 Sr 15 Nb 29.4TiO2 0.53 P2O5 0.12 Ba 475 Y 25AI2O3 15.00 S 3.31 Cu 22Fe2O3 6.18 CO2 0.88 Pb 30 Al 83MnO 0.03 Total 99.32 Zn 100 CCPI 54MgO 1.28 LOI 4.85 Sb 3.1 Ti/Zr 18.5CaO 0.97 Tl 2.0Na2O 0.37 Au 0.424 Zr 172

I LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS | 2 1 9

Intense, proximal, domainal quartz + sericite alteration faciesFootwali?

Hi 6

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

255053

intense, proximal, domainal quartz +sericite (MV)proximal altered zone enclosing andtransitional to mineralised MQ



indeterminate

dacite

indeterminate

indeterminate

quartz + sericite + pyrite

10-40 |jm microcrystalline sutured mosaicof quartz, selective domainal sericite withcleavage, disseminated pyrite

local, enclosing mineralised lens

poor

intense

synmineralisation

hydrothermal

GeochemistrySiO2

TiO2

MAFe2O3

MnOMgOCaONa2O

Sco2

88.10 K2O0.44 P A6.291.210.01 Total0.31 LOI0.310.08 Au

1.940.030.540.37

99.631.44

0.318

SrBaCuPbZnSbTlZr

9100

91324

100

1.60.9178

NbY

AlCCPITi/Zr

12.214

8541

14.8


2 2 0 | CHAPTER 7

intense, massive quartz alteration faciesFootwall?

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

255044

intense, massive quartz (MQ)

central siliceous core of A-zone



indeterminate

sparse polycrystalline quartz crystals?,pseudomorphs of feldspar?

porphyritic?

dacite

indeterminate

indeterminate

quartz + calcite + pyrite + chalcopyrite

2CM0 (jm microcrystalline in coarseand fine equigranuiar domains of quartz,disseminated calcite blebs, possible calcitepseudomorphs after feldspar crystals,pyrite + chlorite ± calcite veinlets

local, ore lens

poor

intense

synmineralisation

hydrothermal

GeochemistrySiO2

TiO2

AI2O3

Fe2O3

MnO

MgO

CaO

Na2O

87.700.390.801.170.070.125.290.05

K2O

P2O5

Sco2

TotalLOI

0.24

0.04

0.22

4.25

100.35

4.33

AuSrBaCuPbZnSbTl

0.048 Zr23 Nb

100 Y1690

47 Al100 CCPI1.6 Ti/Zr0.5

163

2.3

12

680

14.3


7.9 | THALANGA: A POLYMETALLICSHEET-STYLE DEPOSIT

The Thalanga deposit is located near the western end of theEarly Ordovician Mount Windsor Subprovince (Fig. 1.7). Itwas the most economically significant deposit in the MountWindsor Subprovince. The sulfide lenses were up to 25 mthick and distributed over about 3000 m strike and 400 mvertical extent. The pre-mining resource estimate was 6.6 Mtgrading 1.8% Cu, 2.6% Pb, 8.4% Zn, 69 g/t Ag and 0.4 g/tAu.

Geological setting

The Thalanga deposit consisted of several semi-connected,thin, stratabound and stratiform massive sulfide lenses hostedin a distinctive quartz crystal-rich volcanic unit, whichis sandwiched between the underlying rhyolitic MountWindsor Formation and the overlying mixed andesitic-daciticTrooper Creek Formation (Fig. 7.23). The host unit (knownas the Thalanga horizon or favourable unit) is composed ofquartz + feldspar crystal-rich volcanic breccia, sandstoneand siltstone, and co-magmatic, peperitic quartz + feldsparintrusions (Paulick and McPhie, 1999).


The ores are massive to semi-massive lenses dominatedby pyrite and sphalerite with variable proportions of galena,chalcopyrite, pyrrhotite, magnetite and barite (Gregory etal., 1990). Barite-rich zones exist in the up-dip and westernperipheries of the West and East Thalanga lenses. Chlorite +tremolite + carbonate rocks, interpreted as metamorphosedchlorite + carbonate alteration assemblages (Herrmann andHill, 2001), are closely associated with the West Thalangaore lenses. Magnetite-bearing quartzite bodies in theperipheral or upper parts of some of the sulfide lenses, andalso intercalated with volcanic siltstone of the host unit tothe west, are interpreted to be metamorphosed exhalativesiliceous ironstones (Duhig et al., 1992).

The stratigraphic footwall is a laterally extensive,1200 m thick, submarine rhyolitic succession. It is domin-ated by sparsely quartz + feldspar-phyric coherent rhyoliticlavas (e.g. data sheet TH1) and domes that may haveformed a low volcanic rise in the Thalanga area (Paulick andMcPhie, 1999). Rhyolitic hyaloclastite breccias and volcanicsandstones are locally significant, particularly in the upperpart of the footwall beneath the western sulfide lenses. Thehanging wall succession is composed mainly of unalteredto weakly altered coherent lavas and sills of feldspar-phyricto aphyric dacite (e.g. data sheet TH2), and minor basaltic-andesite. At Thalanga, it includes minor volcaniclastic rocks

FIGURE 7.23 | Schematic facies architecture of the submarine volcanic succession, from the Mount Windsor Volcanics,through the Trooper Creek Formation, to the Rollston Range Formation, northwest of the Thalanga mine, Queensland.Modified after Hill (1996).

2 2 2 | CHAPTER 7

of mixed dacitic-rhyolitic derivation, including lithic mass-flow breccia, sandstone and massive to laminated chertysiltstone, which increase in proportion westward. Paulick andMcPhie (1999) interpreted the volcanic facies assemblageto indicate that the deposit formed in a below-storm-wave-base environment on an elevated, lava-dominated, rhyoliticcentre. The compositions of the footwall and hanging wallsuccessions, respectively, indicate that they were rhyoliticmagmas derived from crustal melting, and mixed-mafic-felsicmagmas from subduction-modified mantle, in an extensionalback-arc-basin setting (Stolz, 1995).

Regional deformation and metamorphism, related toMid-Late Ordovician granitoid intrusions produced uppergreenschist facies metamorphic mineral assemblages anda near-vertical foliation, particularly in phyllosilicate-richhydrothermally altered volcanic rocks.


Underlying the Thalanga deposit is an extensive zone of strong,pervasive quartz + sericite + pyrite ± chlorite alteration facies(e.g. data sheet TH3) characterised by 1—4% disseminatedpyrite and an absence of primary feldspars. This alterationstyle was pervasive in both clastic and coherent rhyolites,and typically produced pseudoclastic breccia and mottled,domainal alteration textures (Paulick and McPhie, 1999). Thezone extends beneath the entire strike length of the depositand is at least 200 m thick in the Central area, pinching outto less than 50 m near the lateral and down-dip extremities(Herrmann and Hill, 2001). It has a broad, upward flaringshape and gradational boundaries with the surrounding least-altered rhyolites.

Within the broad zone of feldspar destruction there aresemi-stratiform stringer zones of intense, pervasive quartz +pyrite alteration facies (e.g. data sheet TH4) up to 50 m thick.They extend obliquely up through the footwall at about 15°to the host unit and intersect it beneath the East, Central and"eastern edge of the West Thalanga ore lenses, suggesting thatthey were paths of maximum hydrothermal fluid flow. Thesezones are composed essentially of quartz and up to 20% pyritein disseminated grains and anastomosing veins. They typicallycontain less than 20% phyllosilicates (sericite ± chlorite).Intense, macrocrystalline quartz + K-feldspar alteration facies(e.g. data sheet TH5) exist in the immediate footwall, lateralto the sulfide lens and stringer zone at East Thalanga, and alsostratigraphically lower in the footwall succession at Centraland West Thalanga.

Stratabound chlorite + dolomite altered zones, sub-sequently metamorphosed to chlorite + tremolite ± carbonateassemblages, formed in permeable volcaniclastic footwallrocks close to the palaeo-seafloor and lateral to the WestThalanga ore lenses (stratabound alteration facies; data sheetsTH6, 7 and 8). Local zones of non-pyritic, foliated alteredrhyolite with relict plagioclase (moderate, foliated sericite+ chlorite alteration facies; data sheet TH9), exist withinthe least-altered rhyolite, mainly around the peripheries ofthe feldspar-destructive, strong quartz + sericite + pyrite +chlorite alteration facies, and may represent low-temperaturehydrothermal recharge zones.

Ore genesis

There is consensus amongst researchers that Thalanga isa sheet-like, synvolcanic, deformed and metamorphosedVHMS deposit formed in a deep-marine back-arc rift.Isotopic data suggests that the hydrothermal fluid and sulfurwere dominantly of seawater origin (Hill, 1996; Herrmannand Hill, 2001). The hydrothermal system pervasively altereda very broad zone in the mainly coherent rhyolitic footwallsuccession to quartz + sericite + pyrite + chlorite assemblages.The ore-forming fluids were focussed in low-angle quartz +pyrite stringer zones. These pyritic stringer zones have a semi-stratiform distribution, which suggests control by volcanicfacies related permeability contrasts. However, some appearto cut through coherent rhyolite units and thus may representdeformed synvolcanic fault zones (Paulick and McPhie,1999).

A major proportion of the massive sulfide ore wasdeposited in thin, extensive, stratiform and strataboundlenses, either directly on the palaeo-seafloor or a few metresbelow it. The distribution of massive pyrite and Cu-rich zonessuggests that the down-dip eastern parts of West and CentralThalanga ore bodies, and central part of East Thalanga, weresites of high-temperature hydrothermal discharge (Hill, 1996;Paulick et al., 2001). Subordinate stratabound semi-massiveore lenses were formed by subsurface replacement and/orinfilling of coarse volcaniclastic units of the host unit, whichwere deposited by syneruptive, synhydrothermal mass flowson top of the accumulating seafloor massive sulfide lenses.Chlorite + carbonate (pre-metamorphic) alteration mineralassemblages intimately associated with the West Thalangasulfide lenses, probably formed by mixing of hydrothermalfluid and cold seawater in permeable volcaniclastic unitsimmediately below the palaeo-seafloor, in proximal to medialparts of the hydrothermal discharge system. Apart fromdisseminated and vein-type pyrite in the altered footwallzones, all the sulfides were deposited at the top of the rhyoliticMount Windsor Formation, or in the quartz crystal-rich unitthat immediately overlies it. The massive, coherent dacitelavas and sills of the hanging wall succession are essentiallyunaltered and unmineralised; their emplacement appears tohave ended local hydrothermal circulation.

Weak, patchy quartz + sericite alteration faciesLeast-altered footwal!


TH1

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

140802

weak, patchy quartz + sericite

Thalanga footwall


Mount Windsor Subprovince

massive quartz + plagioclase-phyricrhyolite

quartz and albitised plagioclase

phenocrysts

porphyritic

rhyolite

massive

rhyolite lava

quartz + sericite + biotite + chlorite

selective-pervasive (patchy)microcrystalline quartz matrix, weaklyaligned sericite ± biotite in cleavage

regional, broadly stratabound, footwall

succession

moderate

weak

synvolcanic

diagenetic


Geochemistry

SiO2 76.40 K2O 4.04 Rb 110 Tl 0.5TiO2 0.11 P2O5 0.02 Sr 82 Zr 146AI2O3 11.90 S 0.01 Ba 1056 Nb 13Fe2O3 1.64 CO2 <0.1 Cu 5 Y 40MnO 0.04 Total 98.52 Pb 20MgO 0.67 LOI 0.60 Zn 48 Al 56CaO 1.42 Sb 0.2 CCPI 25Na,0 2.27 Ti/Zr 4.5

2 2 4 | CHAPTER 7

Subtle, selective-pervasive quartz 4- albite alteration fa.ciesLeast-altered hanging wall

TH2

Sample no.

Alteration fades

Location

Formation

Succession

Volcanic facies

Relict minerals

Relic textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

140799

subtle, selective-pervasive quartz +

albite

Thalanga hanging wall



massive sparsely plagioclase-phyric

dacite

plagioclase phenocrysts

weakly porphyritic, faintly flow banded

dacite

massive to flow banded

dacite lava or sill

quartz + albite ± (chlorite + actinolite ±epidote)

selective-pervasive; mosaic of 20 pmquartz + albite, dissemiated chlorite andacicular prisms actinolite defining weakrelict flow banding, quartz ± calciteveins

regional, broadly stratabound, hanging

wall succession

good

subtle

synvolcanic

diagenetic


GeochemistrySiO2 74.00 K2O 3.82 Rb 64 Tl 0.5TiO2 0.33 P2O5 0.06 Sr 76 Zr 164AI2O3 12.70 S <0.01 Ba 1285 Nb 9Fe2O3 1.47 CO2 Cu 3 Y 26MnO 0.03 Total 98.50 Pb 11MgO 0.52 LOI 0.47 Zn 29 Al 44CaO 1.47 Sb 0.2 CCPI 19Na2O 4.10 Ti/Zr 12.1


Strong, pervasive quartz + sericite + pyrite ± chlorite alteration fadesFootwall

TO 3

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

140808

strong, pervasive quartz + sericite +

pyrite ± chlorite

Thalanga footwall



massive quartz + plagioclase-phyric

rhyolite

quartz

porphyritic

rhyolite

massive

rhyolite lava

quartz + sericite + chlorite + pyrite

pervasive, microcrystalline quartz matrixwith strongly aligned sericite (cleavage),scattered 1-2 cm elliptical chlorite-richdomainslocal, broadly stratabound in footwallbeneath entire Thalanga system,>200 m thick, thinning laterally

poor

strong

synmineralisation

footwall hydrothermal


GeochemistrySiO2 75.70 K2O 2.39 Rb 64 Zr 128TiO2 0.07 P2O5 0.01 Sr 8 Nb 14AI2O3 11.40 S 0.65 Ba 326 Y 36Fe2O3 5.14 CO2 <0.1 Cu 18MnO 0.08 Total 98.03 Pb 11 Al 96MgO 2.38 LOI 2.75 Zn 77 CCPI 73CaO <0.01 Sb 0.2 Ti/Zr 3.3Na2O 0.21 Tl <0.5

2 2 6 I CHAPTER 7

Intense, pervasive quartz + pyrite alteration faciesFootwall

TH4

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

140724

intense, pervasive quartz + pyrite

Thalanga footwall




quartz phenocrysts

porphyritic

rhyolite

massive

rhyolite lava

quartz + sericite + pyrite > biotite ±chlorite

pervasive, mosaic of 100-200 pm quartzand pyrite with interstitial shreds of semi-aligned white mica > biotitelocal, broadly stratabound in footwall,discrete sheet-like zones at -15° to hostunit, intersecting it beneath suifide lensespoor

intense

synmineralisation

hydrothermal


GeochemistrySiO2 67.00 K2O 1.99 Sr 16 Nb 7

TiO2 0.05 P2O5 0.01 Ba 2300 Y 44AI2O3 6.60 S 10.61 Cu 22Fe2O3 14.34 CO2 <0.1 Pb 16 Al 97MnO 0.02 Total 102.40 Zn 74 CCPI 88MgO 1.67 LOI 8.24 Sb 1.1 Ti/Zr 3.8

CaO 0.05 Tl 1.9Na2O 0.06 Rb 68 Zr 79


intense, microcrystalline quartz + K-feldspar alteration faciesFootwall

TH5

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

140902

intense, microcrystalline quartz + K-

feldspar

Thalanga footwall




quartz and albitised plagioclase

phenocrysts

porphyritic

rhyolite

massive

rhyolite lava

quartz + K-feldspar > trace pyrite

selective-pervasive; matrix of 10-50 pmmicrocrystalline quartz + K-feldspar,albitised plagioclase, quartz veins withfeldspar-alteration selvage

local, lozenge shaped zones, broadly

stratabound in footwall

poor

intense

synmineralisation

hydrothermal


GeochemistrySiO2 84.20 K2O 5.87 Sr 42 Nb 9TiO2 0.06 P2O5 0.01 Ba 1936 Y 20AI2O3 7.40 S 0.48 Cu 3Fe2O3 0.60 CO2 Pb 70 Al 93MnO <0.01 Total 99.15 Zn 118 CCPI 9MgO 0.08 LOI 0.49 Sb 0.4 Ti/Zr 4.5CaO 0.11 Tl 1.5Na2O 0.34 Rb 94 Zr 80

2 2 8 | CHAPTER 7

Intense, pervasive, stratabound chlorite + tremolite alteration faciesFootwall

TH8

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

145401

intense, pervasive, stratabound chlorite

+ tremolite

Thalanga footwall



quartz + plagioclase-phyric rhyolitic

volcaniclastic breccia?

nil

nil

rhyolite

indeterminate

phlogopite + chlorite + tremolilte > minorpyrite, sphalerite, chalcopyrite

pervasive, foliated phlogopite ± chlorite,30% coarse prisms and bands ofrandomly oriented tremolite

local, stratabound, typically immediately

below sulfide lenses

nil

intense

synmineralisation

footwall hydrothermal


GeochemistrySiO2 44.57 K2O 3.48 Rb Zr 171TiO2 0.135 P2O5 0.04 Sr Nb 10AI2O3 13.06 S 1.66 Ba 38300 Y 32Fe2O3 3.88 CO2 0.30 Cu 1500MnO 0.15 Total 91.89 Pb 200 Al 75MgO 17.67 LOI 3.23 Zn 5900 CCPI 85CaO 6.76 Sb Ti/Zr 4.7Na2O 0.18 Tl


Intense, stratabound. pervasive chlorite + tremolite + calcite alteration faciesHost rock

TH7

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

145417

intense, stratabound, pervasive chlorite +

tremolite + calcite

Thalanga host rock



quartz + plagioclase-phyric rhyolitic

volcaniclastic breccia?

nil

nil

rhyolite

indeterminate?

chlorite + tremolite + calcite > pyrite +chalcopyrite + sphalerite + galena

pervasive; coarse interlocking prisms oftremolite with 10% interstitial sulfides andragged calcite patches

local, stratabound, proximal to medial,closely associated with or lateral to WestThalanga sulfide lenses

nil

intense

synmineralisation

hydrothermal

GeochemistrySiO2 33.10 K2O 0.58 Rb Zr 89

TiO2 0.060 P2O5 0.15 Sr Nb 6

AI2O3 6.48 S 4.82 Ba 8700 Y 13

Fe2O3 4.82 CO2 7.10 Cu 7500

MnO 0.57 Total 92.41 Pb 4200 Al 55

MgO 19.00 LOI 5.86 Zn 23800 CCPI 97

CaO 15.60 Sb Ti/Zr 4.0

Na2O 0.13 Tl


2 3 0 I CHAPTER 7

Intense, strataboynd tremolite + dolomite + calcite alteration faciesHost rock

TH8

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relic textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

145418

intense, stratabound tremolite + dolomite+ calcite

Thalanga host rock



quartz + plagioclase-phyric rhyoliticvolcaniciastic breccia?

nilnilrhyolite

indeterminate?

dolomite + calcite > minor tremolite+ chlorite + pyrite + chalcopyrite +sphalerite + galena

pervasive; mosaic of 1 mm suturedspheroidal dolomite and interstitial calcitedisseminated sulfides and sparse raggedtremolite prisms

local, stratabound, proximal to medial,closely associated with or lateral to WestThalanga sulfide lenses

nilintense

synmineralisation

footwall hydrothermal, seawater mixing?


LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS I 231

Moderate, foliated sericite + chlorite alteration fadesFootwaSI

TH9


Sample no. 140727

Alteration facies moderate, foliated sericite + chlorite

Location Thalanga footwall

Formation Mount Windsor Formation

Succession Mount Windsor Subprovince

Volcanic facies massive quartz + plagioclase-phyricrhyolite

Relict minerals quartz and albitised plagioclase

Relict textures porphyritic


Lithofacies massive, foliated

Interpretation rhyolite lava

Alteration minerals sericite + quartz + biotite + chlorite

Alteration textures selective-pervasive 50-100 pmmicrocrystalline quartz matrix in<1 mm lenses wrapped by foliatedsericite ± biotite (augen texture),broken grains, cleavage, sericitealtered plagioclase

Distribution local, stratabound in upper part ofmedial to distal footwall, particularlydown dip of ore zones


Alteration intensity moderate


Alteration style hydrothermal, tectonic-metamorphic

GeochemistrySiO2 70.90 Na2O 1.05 Rb 145 Zr 162TiO2 0.10 K2O 3.92 Sr 40 Nb 19AI2O3 14.50 P2O5 0.01 Ba 952 Y 54Fe2O3 1.69 S <0.01 Cu 3MnO 0.04 CO2 0.10 Pb 5 Al 87MgO 3.94 Total 96.37 Zn 231 CCPI 52CaO 0.12 LOI 2.67 Sb 0.2 Ti/Zr 3.7

Tl 2.6

2 3 2 | CHAPTER 7

7.10 | HIGHWAY-REWARD: A PIPE STYLECu-Au VHMS DEPOSIT

The Highway-Reward Cu-Au deposit, in the central part of theMount Windsor Subprovince (Fig. 1.8), represents a contrastin style of deposit and stratigraphic setting. It consists of twodiscordant, vertical pipe-like bodies of massive pyrite about200 m apart, hosted in the proximal facies of a non-explosive,submarine felsic volcanic centre located near the top of theTrooper Creek Formation (Fig. 7.23).

Geological setting

The lithofacies association at Highway-Reward representsa deep marine intrusion-dominated felsic volcanic centre.Doyle and McPhie (2000) recognised at least 13 coherentfeldspar- and quartz + feldspar-phyric dacitic to rhyoliticsynvolcanic sills, small cryptodomes (e.g. data sheets HR1and 4) and lavas in the immediate area. The abundance andcomplex overlapping relationships of coherent intrusive unitsindicate a proximal volcanic setting. Thin volcanic sandstoneand siltstone units and thicker units of crystal- and pumice-rich sandstone and breccia separate the intrusions. The crystal-and pumice-rich facies were mainly derived from explosiveeruptions and deposited in the submarine basin from water-supported gravity flows. The succession is upright and dips at20-30° to the southeast.

Massive pyrite ± chalcopyrite exists in two vertical pipe-like bodies about 150 m apart. Both are discordant to bedding,parallel to a locally developed northeast trending sub-verticalcleavage (S4) and have irregular-amoeboid outlines with plandimensions of about 200 x 75-150 m. The western pipe(Highway) has a vertical extent of 250 m and the eastern pipe(Reward) of at least 350 m (Doyle and Huston, 1999). Theyare dominantly composed of fine-grained (<0.5 mm) pyritewith interstitial chalcopyrite, minor tennantite, sphalerite,quartz and sericite, and traces of chlorite, galena, barite,hematite and aikinite (PbCuBiS3). The massive sulfide pipesare intersected by chalcopyrite, barite, quartz + carbonate andanhydrite veins, and contain inclusions of quartz + sericite +pyrite altered volcanic rocks in their margins. The Highwayand Reward massive sulfide pipes contain approximately2 Mt and 5 Mt of pyrite, respectively. They include hypogenesulfide resources estimated at 1.2 Mt @ 5.5% Cu, 1.2 g/t Auand 6.5 g/t Ag in the Highway pipe and 0.2 Mt @ 3.5% Cu,1 g/t Au and 13 g/t Ag in the Reward pipe.

The pipes are enveloped by a broad 200 x 500 m halo ofvein and disseminated low-grade Zn + Pb + Ba sulfides. Withinthat are several small zones of massive to laminated sphalerite+ pyrite + galena + chalcopyrite + barite. A 20-30 m thickstratabound lens of sphalerite-rich massive sulfide exists involcaniclastic rocks 50 m above and south of the Rewardpipe. It has a pyrite-rich base that thickens northwards intoa discordant lens of massive pyrite lying above the southernedge of the Reward pipe. Sphalerite-rich sulfides also existlocally in narrow discordant zones at the margins of the mainHighway and Reward massive sulfide pipes.


A discordant zone of feldspar-destructive hydrothermal alter-ation envelopes the massive sulfide pipes. It has an ellipticalarea of 500 x 250 m in plan and extends from 60 m aboveto at least 150 m below the massive sulfide bodies (Doyleand Huston, 1999). Doyle and Huston's (1999) alterationzonation is here simplified down to six alteration facies.

Intense, stringer quartz + sericite + pyrite alteration facies(e.g. data sheet HR8), locally flanked by intense, pervasivechlorite + pyrite alteration facies (e.g. data sheet HR7) occupyfeeder zones which extend vertically beneath both pipes andpossibly meet at depth. Zones of similar quartz + sericite +pyrite altered rocks extend into the hanging wall above thesouthern parts of both massive sulfide pipes (Doyle andHuston, 1999). These intensely altered zones pass laterallyoutwards to enveloping zones of intense sericite + quartz +pyrite (e.g. data sheet HR6) and strong, pervasive chlorite+ sericite + quartz + pyrite alteration facies (e.g. data sheetHR5). These locally enclose, and in turn grade laterally andupwards in to, non-pyritic zones of moderate, pervasivechlorite alteration facies (data sheet HR3). The weak, regional,selective albite ± hematite alteration facies (e.g. data sheetHR2) exists at greater than 50—200 m from the massive sulfidepipes. It comprises two sub-facies with mineral assemblagesof feldspar + carbonate ± quartz ± chlorite + sericite andhematite + quartz ± sericite ± chlorite ± albite, which areregionally distributed in the Trooper Creek Formation andare respectively attributed to alteration during diagenesis andsynvolcanic low-temperature fluid convection.

Ore genesis

The deposits were initially thought to have had a two-stageorigin (Beams et al., 1998). The stratiform Zn-rich zone wasinterpreted as a syngenetic Cambro-Ordovician sulfide lensand the pyrite + chalcopyrite pipes as Siluro-Devonian syn-deformational deposits, because of their discordance to hostvolcanic rocks, parallelism to the youngest cleavage (S4) andthe observation that anhydrite overprinted the dominantS3 cleavage. However, Doyle and Huston (1999) refutedthis microtextural relationship and argued for a syngeneticvolcanic-associated, subseafloor replacement origin for themassive sulfide pipes. Lead isotopic ratios, the gradation fromstratiform Zn-rich sulfides into discordant Cu-rich massivepyrite, relict framboidal sulfide textures, hydrothermal alter-ation facies and their relationships to primary volcanic facies,and the S3 tectonic overprint are all consistent with EarlyOrdovician synvolcanic formation of all the sulfide zones.

The Highway-Reward massive sulfide pipes have somesimilarities with disseminated to massive Cu-Au deposits inthe Mount Lyell field (Large et al., 2001c). The similaritiesinclude metal ratios, dominance of pyritic subseafloorreplacement style mineralisation and low 534S values; mostlyin the range 5 to 7.5%o at Highway-Reward and 5 to 10%o atMount Lyell (Solomon etal., 1969; Doyle and Huston, 1999).Given the emerging evidence for involvement of magmaticfluids at Mount Lyell (Corbett, 2001; Huston and Kamprad,2001) it is reasonable to similarly classify Highway-Reward asa hybrid seawater-magmatic hydrothermal system.


Weak, selective-pervasive quartz + sericite +albite alteration faciesLeast-altered rhyolite

HR1

Sample No.

Alteration Facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

137068

weak, selective-pervasive quartz + sericite

+ albite

upper medial, 50 m east of Highway pipe

(10075N)




rhyolite

quartz, plagioclase

porphyritic

rhyolite

massive

partly extrusive cryptodome

quartz + sericite + albite? > (calcite, pyrite)

selective-pervasive, microcrystalline

groundmass, albite + sericite or calcite-

altered plagioclase

regional

moderate

weak

diagenetic

GeochemistrySiO2

TiO2

AI2O3

Fe2O3

MnO

MgO

CaO

Na2O

75.08

0.30

12.75

1.57

0.08

2.02

0.61

2.16

K2O

S

co2

Total

LOI

2.22

0.06

0.51

97.37

2.88

Rb

Sr

Ba

Cu

Pb

Zn

Sb

Tl

50 Zr

46 Nb

1382 Y

43

14 Al

104 CCPI

2.2 Ti/Zr

1.0

158

9

22

60

44

11.4


2 3 4 | CHAPTER 7

Weak, regional, selective albite + hematite alteration facies HR2

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

137105

weak, regional, selective albite + hematite

upper periphery, 250 m east of Rewardpipe




quartz, piagiociase

porphyritic, giomeroporphyritic piagiociase,microcrystalline groundmass

rhyolite

massive

synvolcanic sill

quartz + albite + chlorite > (sericite +calcite + hematite)

selective-pervasive in groundmass; 20-60urn microcrystalline quartz + albite ±calcite-altered piagiociase, disseminatedchlorite and hematite patches

regional

good

weak

synvolcanic

diagenetic

GeochemistrySiO2

TiO2

Fe2O3

MnO

MgO

CaO

Na2O

70.58

0.33

13.75

2.41

0.05

2.26

1.78

6.51

K2O

S

CO2

Total

LOI

0.34

0.06

<0.01

98.07

2.42

Rb

Sr

Ba

Cu

Pb

Zn

Sb

Tl

8 Zr

142 Nb

92 Y

<2

5 Al

38 CCPI

0.8 Ti/Zr

<0.5

161

9

21

24

39

12.3


Moderate, pervasive chlorite alteration facies


HR3

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

137079

moderate, pervasive chlorite

upper proximal zone, between Highway

and Reward sulfide pipes




rhyolite

quartz, mafic phenocrysts?

porphyritic, amygdaloidal?

rhyolite

massive

partly extrusive cryptodome

quartz + sericite + chlorite

pervasive, microcrystalline mosaic ofquartz + chlorite + sericite, sericitepseudomorphs after plagioclasephenocrysts, anastomosing wispy sericitefoliation, recrystallised overgrowths onquartzlocal; medial to proximal zones laterally

equivalent to upper parts of sulfide pipes

moderate to poor

moderate

synmineralisation

hydrothermal

GeochemistrySiO2

TiO2

AI2O3

Fe2O3

MnO

MgO

CaO

Na2O

71.21

0.30

14.09

3.40

0.15

3.59

0.39

0.16

K2O

S

CO2

Total

LOI

3.27

0.05

0.01

96.61

3.61

Rb

Sr

Ba

Cu

Pb

Zn

Sb

70 Tl

21 Zr

1791 Nb

6 Y

7

120 Al

0.8 CCPI

Ti/Zr

1.5165

9

23

93

66

10.9


2 3 6 I CHAPTER 7

Weak, pervasive albite + sericite alteration faciesLeast-altered dacife

HR4

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

136919

weak, pervasive albite + sericite

Highway, medial footwall



massive, sparsely plagioclase-phyricmassive dacite

plagioclase <1mm

porphyritic and micropoikilitic

dacite

massive to weakly flow banded

cryptodome

albite + chlorite + sericite > (zeolite?,quartz)

pervasive groundmass, microcrystallinepartly preserving micropoikilitic texture,albite ± sericite-altered plagioclase,chlorite veinlets

regional

good

weak

synvolcanic

diagenetic

GeochemistrySiO2

TiO2

AI2O3

Fe2O3

MnO

MgO

CaO

Na2O

66.74

0.55

16.63

4.30

0.18

2.48

0.19

4.40

K2O

P2O5

S

co2Total

LOI

1.81

0.10

0.00

97.38

2.42

Rb

Sr

Ba

Cu

Pb

Zn

Sb

Tl

44 Zr

68 Nb

515 Y

1

2 Al

162 CCPI

0.3 Ti/Zr

<0.5

161

9

25

48

51

20.5



Strong, perfasiwe chlorite + sericite + quartz + pyrite alteration facies HR 5

Sample no.

Alteration facies

Location

Formation

Succession

Volcanic facies

Relict minerals

Relict textures

Primary composition

Lithofacies

Interpretation

Alteration minerals

Alteration textures

Distribution

Preservation


Timing

Alteration style

137127

strong, pervasive chlorite + sericite +quartz + pyrite

Highway footwall, 100 m east of stringerzone



massive quartz + plagioclase-phyricrhyodacite

quartz + altered plagioclase

porphyritic

rhyodacite

massive

cryptodome

quartz + sericite + pyrite + chlorite ± rutile

pervasive in groundmass, millimetre

patches of microcrystalline quartz andwispy domains of aligned sericite, somebroken quartz phenocrysts

local; medial to proximal zones laterallyequivalent to upper parts of sulfide pipes

poor

strong

synmineralisation

hydrothermal

GeochemistrySiO2

TiO2

AI2O3

Fe2O3

MnO

MgO

CaO

Na2O

64.55

0.44

17.02

5.54

0.04

1.87

0.20

0.18

K2O

S

co2Total

LOI

4.53

0.08

2.97

97.43

4.93

Rb

Sr

Ba

Cu

Pb

Zn

Sb

Tl

92 Zr

28 Nb

3683 Y

16

32 Al

68 CCPI

2.0 Ti/Zr

2.0

149

8

25

94

59

17.8


2 3 8 | CHAPTER 7

Intense sericite + quartz + pyrite alteration facies HR§


Sample no. 137080

Alteration fades intense sericite + quartz + pyrite

Location upper proximal, 20 m east of Highway pipe

Formation Trooper Creek Formation


Volcanic facies massive, sparsely plagioclase-phyricdacite

Relict minerals altered plagioclase

Relict textures porphyritic

Primary composition dacite

Lithofacies massive

Interpretation cryptodome

Alteration minerals quartz + sericite + pyrite

Alteration textures pervasive, polycrystalline quartzpseudomorphs after plagioclase,microcrystalline matrix of quartz + sericite> pyrite

Distribution local, proximal zone enveloping Highwaysulfide pipe

Preservation poor




GeochemistrySiO2 73.82 K2O 3.65 Rb 78 Zr 107TiO2 0.41 P2O5 0.06 Sr 14 Nb 6AI2O3 12.23 S 2.89 Ba 2222 Y 13Fe2O3 4.55 CO2 Cu 1385MnO 0.02 Total 98.56 Pb 10 Al 96MgO 0.76 LOI 4.09 Zn 90 CCPI 57CaO 0.09 Sb 1.3 Ti/Zr 23.0Na2O 0.08 Tl 5.3

Intense, pervasive chlorite + pyrite alteration faciesFootwall


HR7


Sample no. 137083

Alteration fades intense, pervasive chlorite + pyrite

Location Highway footwall, adjacent to stringerzone



Volcanic fades altered dacitic pumice breccia?

Relict minerals nil

Relict textures nil

Primary composition dacite

Lithofacies indeterminate

Interpretation indeterminate

Alteration minerals chlorite + pyrite > quartz

Alteration textures pervasive cryptocrystalline groundmassor matrix of chlorite, cleavage, 5%disseminated euhedral pyrite with quartzpressure shadows

Distribution local; narrow zones enveloping footwallquartz + pyrite stringer zone

Preservation nil




GeochemistrySiO2 24.90 K2O 0.01 Rb <1 Zr 169TiO2 0.61 P2O5 0.12 Sr 9 Nb 9AI2O3 18.79 S 6.37 Ba 21 Y 27Fe2O3 16.73 CO2 Cu 102MnO 0.35 Total 90.83 Pb 12 Al 99MgO 22.70 LOI 13.80 Zn 522 CCPI 100CaO 0.21 Sb Ti/Zr 21.7Na2O 0.04 Tl

2 4 0 | CHAPTER 7

Intense, stringer quartz + sericite + pyrite alteration fadesFootwali

HR8


Stacked SWIR spectra (hull quotient)

Photomicrograph (xn)

TiO2-Zr immobile element plot

Sample no. 137129

Alteration facies intense, stringer quartz + sericite + pyrite

Location Reward footwall stringer zone



Volcanic facies massive

Relict minerals nil

Relict textures nil

Primary composition rhyodacite

Lithofacies massive

Interpretation synvolcanic sill?

Alteration minerals quartz + sericite + pyrite

Alteration textures pervasive, irrregular sericite

pseudomorphs after feldspar in <10 pm

microcrystalline quartz, interstitial sericite

groundmass, disseminated 5-10%

euhedral pyrite

Distribution local; footwall stringer zones beneath

sulfide pipes

Preservation nil




GeochemistrySiO2 75.78 K2O 1.80 Rb 30 Zr 52

TiO2 0.18 P2O5 0.02 Sr 13 Nb 3

AI2O3 5.87 S 7.16 Ba 1404 Y 8

Fe2O3 9.89 CO2 Cu 11

MnO 0.02 Total 101.21 Pb 3 Al 90

MgO 0.26 LOI 5.81 Zn 12 CCPI 83

CaO 0.20 Sb 1.2 Ti/Zr 20.8

Na2O 0.03 Tl 1.2

I 241

8 | FINDING ORE DEPOSITS IN ALTEREDVOLCANIC ROCKS

Recognising alteration facies that may be genetically related toore deposits is an important step in mineral exploration. Evenmore helpful is the ability to identify alteration facies thatare likely to be associated with particular ore deposit typesand thus prioritise exploration targets. The characteristics ofalteration facies have the potential to be used as explorationvectors, guiding explorers to the most prospective alteredzones in a system, and thereby enabling efficient and earlytesting of the best targets, avoiding expensive, protractedexploration programs, and improving the chance of success.

The processes that alter volcanic facies and the range oftextures and mineral assemblages they produce are complexand challenging. As described in previous chapters, thereare a variety of alteration processes, which can produce abroad range of alteration mineral assemblages and textures.Ancient volcanic rocks commonly contain a complexassemblage of overprinting alteration minerals and textures,which reflect multiple episodes of alteration by a variety ofprocesses: diagenesis, hydrothermal alteration, deformation,metamorphism or weathering.

In early Palaeozoic volcanic regions, like the MountRead Volcanics, western Tasmania and the Mount WindsorSubprovince, north Queensland, patience and experienceare required to unravel the complexities of altered rocksand recognise those altered zones that are 'red-herrings'to mineral explorers. In fact, numerous geologists haveinitially doubted that the foliated, weathered and moss-covered rocks encountered in western Tasmania really were ofvolcanic origin. Several intensive, protracted and ultimatelyunsuccessful exploration programs have been conducted inthe Mount Read Volcanics on unfavourable altered zones.On the other hand, there may be altered zones that remainunder-explored because favourable alteration facies werenot recognised. Recognising the occurrence of altered rocksand identifying favourable or prospective alteration faciesand zones are important steps toward minimising risk andexpenditure during exploration in these environments.

This chapter draws together the descriptive andgeochemical techniques described in Chapters 2, 3 and 4,and an understanding of the different alteration processes thatmodify submarine volcanic successions. It proposes methodsfor discriminating alteration facies associated with particularprocesses, identifying favourable altered zones for mineral

exploration, and guiding exploration within those zonestoward potentially mineralised areas.

8.1 | PRINCIPLES OF DISCRIMINATINGBETWEEN DIAGENETIC,HYDROTHERMAL ANDMETAMORPHIC ALTERATIONFACIES

Diagenetic facies

As discussed in Chapter 5, the characteristics of diageneticfacies are:• They are typically widespread with district or regional-scale

distribution.• At local scales, they display variable alteration intensity

and patchy distribution. This is mainly controlled bydistribution of coherent versus clastic volcanic facies,and variations in the primary composition, permeability,porosity and the proportion of glassy to crystalline facies.

• They occur in vertically-stacked, extensive, sub-horizontalaltered zones, which have mineral assemblages that reflectincreasing temperature with depth.

• They have undergone relatively minor (< 10 wt%) chemicalchanges that are predominantly in response to hydration oralkali-exchange reactions between the volcanic facies andmodified seawater. Mass transfers are generally small, anorder of magnitude less than those in intense hydrothermalalteration facies. The scale of migration of elements is alsosmall (millimetres to tens of centimetres) and thus on alarger scale (i.e. basin scale) the changes are essentiallyisochemical.

• Their mineralogical and textural changes vary from subtleto strong. Quartz phenocrysts, for example, are relativelystable and commonly well preserved, whereas maficphenocrysts and volcanic glass are relatively unstable andtypically completely altered.These changes are commonly overprinted or obscured

2 4 2 I CHAPTER!

by subsequent alteration as diagenesis is often the earliestpreserved post-emplacement process.

Metamorphic facies

Metamorphic facies share some characteristics with diageneticfacies, but also differ significantly in these ways:• Distribution varies in scale: contact metamorphic facies

associated with intrusions may be as narrow as a fewcentimetres and up to several kilometres wide. Regionalmetamorphic facies (either burial metamorphism ormetamorphism associated with deformation) can be tensor hundreds of kilometres wide and several kilometresthick.

• Metamorphic facies are uniform and pervasive: they arenot typically patchy at a scale of metres to tens of metres.Primary volcanic textures have virtually no influence onhigh-grade metamorphic facies, which are principallydetermined by whole-rock compositions and metamorphicconditions.

• Chemical changes are minor; metamorphism is generallya process of phase-change in response to changingtemperature and pressure at low water-rock ratios, whichlimits the redistribution of chemical components in andout of the system. The most common metamorphicreactions are dehydration and decarbonation reactions. Thecomposition and mineralogy of metamorphic facies aregenerally strongly influenced by the primary compositionof volcanic facies.

• Mineralogical and textural changes vary from subtleto intense depending on the degree of metamorphism.Typically, primary volcanic quartz phenocrysts are wellpreserved up to about amphibolite grade, but fine-grainedor glassy facies and some mafic phenocrysts are unlikelyto survive even low grades of zeolite and greenschist faciesor contact metamorphism. Metamorphic re-crystallisationproduces a wide variety of distinctive textures, such asgranoblastic, porphyroblastic, decussate, schistose, andgneissic, which are not easily confused with primaryvolcanic or diagenetic textures.

Hydrothermal alteration facies

Hydrothermal alteration facies are unlike diagenetic andmetamorphic facies in their potential for major compositionalchange. This is because hydrothermal alteration typicallyinvolves large volumes of fluid, which facilitate large-scalemass transfers into, out of, or around hydrothermal systems.Depending on the intensity of alteration, this characteristicdetermines or limits the other characteristics of hydrothermalalteration facies.• Hydrothermal alteration facies generally have local

distribution, limited to tens or hundreds of metres andrarely exceeding a few kilometres.

• Hydrothermally altered zones commonly have high aspectratios (i.e. narrow lateral and great vertical extents) becauseconvecting, typically ascending, fluids produce them.

• Locally, on small-scales, the distribution of hydrothermalalteration facies is generally uniform, or pervasive. However,

the distribution is mainly dependent on permeability andporosity; therefore hydrothermal alteration facies may berestricted to fractures and vein selvedges in coherent orotherwise impermeable rocks.

• The degrees of mineralogical and textural preservation, andchemical changes are extremely dependent on alterationintensity and pre-hydrothermal alteration composition andtexture of the facies. Pre-existing textures and minerals areless likely to be preserved in proximal zones of hydrothermalsystems, through which hot reactive fluids are flushed,than in peripheral zones with lower temperature, partlyneutralised fluids and lower fluid-rock ratios. As in the othertypes of alteration, quartz crystals in felsic volcanic faciestend to survive intense alteration, except where major lossof silica is involved (e.g. in chlorite zones). Other primarycrystal phases are commonly progressively replaced (e.g.feldspars altered to sericite) and may be useful as indicatorsof alteration intensity. Hydrothermal alteration facies rarelypreserve primary textures in originally glassy facies.

• Hydrothermal mineral assemblages are largely controlledby fluid composition and physicochemical conditions, andare not noticeably influenced by primary compositions; atleast in the intensely altered zones, which had high fluid-rock ratios. Thus, an intensely hydrothermally altered zonemay cut across volcanic lithofacies of different primarycompositions and textures (e.g. coherent andesite andrhyolitic breccia) and comprise only one alteration faciesin which the protoliths are mineralogically and texturallyindistinguishable.

• Hydrothermal alteration commonly involves significantmass transfer of chemically mobile elements. Elementsmay be gained through precipitation or lost throughdissolution. These mass transfers may produce largepositive or negative net mass changes within particularalteration facies (generally with implications for volumechange) or balance each other out to produce negligible netchange. Significant mass changes are commonly evidentin composition data and derivative alteration indices. Forexample, Na depletion typically accompanies hydrolysisand sericitisation of plagioclase. However, substantial masschanges in some major elements are commonly obscuredby the constant sum effect; this applies especially to Si.

• Major chemical modifications are frequently reflected inexotic mineral assemblages. For example, VHMS-relatedalteration facies commonly contain disseminated pyriteor base-metal sulfides, and several types of Zn deposits areassociated with Mn-rich mineral assemblages.It is important for economic geologists to recognise

hydrothermal alteration facies, which may indicate the large-scale transport and deposition of economically valuableelements, and to discriminate these from alteration facies thatresult from other alteration processes that are unrelated toore deposition. In some cases, examination of an individualaltered sample can reveal important facts that help to identifythe alteration process. For example, a rock with gneissic fabricis metamorphic; a rock composed essentially of quartz andpyrite is probably of hydrothermal origin. However, alterationtextures and mineral assemblages may not easily distinguishsome weak hydrothermal alteration facies, perhaps inperipheral zones, from diagenetic or metamorphic facies.

One of the main criteria distinguishing hydrothermal

from other alteration facies is the distribution or extent ofthe altered zone. This cannot be determined by observationof an individual specimen or outcrop. It requires systematicprospect-scale mapping and knowledge of the district-scalegeological context. Unfortunately, in the last decade of the20th century there has been a significant decline in in-fieldand on-ground geological data collection, particularly in themineral exploration and mining industries. There is a trendtowards using high technology remote sensing to rapidlyexplore large areas at continuously improving resolutions.However, to ensure meaningful interpretation of these data,it is imperative that this virtual geology is not disconnectedfrom real rocks. The combination of a variety of criteria, andhigh-quality mapping, will lead to the best interpretation ofalteration facies.

8.2 | EXPLORATION VECTORS ANDPROXIMITY INDICATORS

Mineral zonation

Mapping of sulfide distribution, particularly pyrite, is animportant exploration technique. Sulfide abundances areeasily estimated by eye, even in weathered samples, and shouldbe applied at an early stage of exploration wherever geologicalexposure permits. VHMS deposits commonly have extensivefootwall zones of disseminated pyrite. For example, the MountLyell Cu-Au deposits (western Tasmania) all lie in a zone ofgreater than 1 % disseminated pyrite, which is 6 km long and1 km wide at the surface (Corbett, 2001). Such pyritic zonesprovide very large exploration targets for initial area selection.They have the potential to be delineated into high-abundancezones, in order to reduce the size of the targets for intensiveexploration and drill testing. Interpretation of sulfide vectorsis straightforward: more is better, and sulfide proportionsgenerally increase with proximity to sulfide deposits.

Other components of alteration mineral assemblages thatare easily recognisable in all sample types and may be spatiallyzoned around mineral deposits include silicates, carbonatesand Fe-oxides. The ratios of quartz to phyllosilicates,sericite to chlorite, and carbonate to silicates are commonlysystematically zoned around VHMS deposits, and recognitionof the zonation patterns can provide useful explorationvectors, at least on a prospect scale.

Unfortunately, interpretation of the patterns is rathercomplex. Australian VHMS deposits are typically associatedwith siliceous proximal zones (Section 7.4) but there are manyvariations even within mineral fields and districts. Thereforeit is unwise to be too strictly empirical or model-drivenin applying this approach. It is better to map out mineraldistributions and relate them to alteration intensities, ratherthan rely on the recognition of specific zonation patterns,which may relate to an ore deposit model. Carbonate +chlorite assemblages, for example, are indicators of oreproximity in some VHMS deposits, such as Rosebery, Hellyerand Thalanga deposits, but only occur in the peripheral orleast-altered zones of the Western Tharsis deposit (MountLyell field). Therefore, rigidly applying a carbonate + chlorite

FINDING ORE DEPOSITS IN ALTERED VOLCANIC ROCKS | 2 4 3

vector could be misleading and potentially guide explorationaway from some Cu-Au deposits in the Mount Lyell field.

Mapping of mineral zonation is effective where largesystematic datasets are available (i.e. where there are plenty ofoutcrops or drill cores) and mineral assemblages are visuallydistinctive or can be determined by simple field tests (e.g.effervescence in acid for carbonate or sodium-cobaltinitritestaining for K-feldspar). However some mineral assemblagesthat are not readily identified visually, are discretely zonedand may be diagnostic of a deposit style. New field-basedmineralogical tools, such as portable SWTR spectrometers(Section 2.4), will facilitate major improvements, which willnot only aid exploration, but also contribute to understandingthese deposit systems (Thompson et al., 1999). SWIR spectralstudies have recently shown some spectacular examples ofmineral zonation, particularly in acid-sulfate type systems(e.g. case studies in Thompson et al., 1999, and Huston andKamprad, 2001).

Major element lithogeochemistry

Although intense hydrothermal alteration frequentlyproduces simple alteration mineral assemblages, the mineralsare commonly fine grained. These minerals may be difficult torecognise visually, and it can also be difficult to estimate theirabundances. In these cases, lithogeochemistry can frequentlyhelp to identify minerals and quantify compositional changeseven in less intensely altered rocks that contain incipient,overprinting or domainal alteration minerals. Analysis ofwhole-rock samples to determine major element abundanceis a way of supporting and augmenting estimates of mineralproportions and alteration intensity, which have beendetermined visually or by other methods (e.g. Section 2.4).

Quantitative lithogeochemical data can be used in twoways: (1) to indicate alteration intensity, and (2) to estimatemineral proportions in mineral assemblages where the mineralspecies and their individual compositions are known.

Interpreting exploration vectors based on compositionaldata seems straightforward. The data can be plotted as contourmaps or cross-sections (e.g. Figs 2.7 and 2.12), down-holeline graphs (e.g. Fig. 2.14) and so on, and the vectors inferredbased on expected variations in mineral abundance orcomposition. Decreases in Na2O contents of volcanic rocks,for instance, are usually related to increasing sericite or chloriteat the expense of plagioclase. Na2O depletion is a popular andreliable vector used in VHMS exploration (e.g. Na2O halomaps of the Fukuzawa area in Date et al., 1983). Variations incarbonate content, both increases and decreases, are typicallyevident in CO2 data and in CaO, MgO or Fe2O3, dependingon the carbonate species. Sulfide content can be quantifiedby sulfur analyses. Weight percent sulfur is generally nearlyequivalent to volume percent of pyrite in felsic rocks, if pyriteis the only sulfur-bearing phase. This is due to pyrite havinga density of just under twice the density of felsic rock, andsulfur comprising just over half the mass in pyrite.

However, major element data are subject to distortionby closure, otherwise known as the constant sum effect.This phenomenon is more fully explained in Section 4.1.It particularly affects the dominant chemical components(e.g. SiO2, A12O3 and Fe2O3) and can be significant in

2 4 4 | CHAPTER!

hydrothermally altered rocks with large net mass gains.Although additions of exotic hydrothermal components suchas sulfur and CO2, and depletions of Na2O, are relativelyimmune to the effects of closure, it seriously compromisesthe use of some other major components as explorationvectors. For example, SiO2 may not provide effective vectorsin hydrothermal systems where mineralisation was associatedwith silicification. If closure in major element data is likelyto obscure the effects of alteration on the compositional dataand exploration vectors, then it is preferable to estimate theindividual component mass changes (by immobile elementtechniques, Section 4.1) and use those as exploration vectors.

The alternative approach is to convert quantitative majorelement lithogeochemical data to modal mineral proportionsusing a method such as MINSQ (Herrmann and Berry,2002) or GENMIX (Le Maitre, 1981). This does not removethe effects of closure, but is a way of quantifying mineralproportions, which can then be used as vectors in mineralexploration. This approach was used by Large et al. (2001b,Fig. 6) to demonstrate systematic variations in proportionsof alteration minerals around the Rosebery K-lens sulfidedeposit (Fig. 2.14).

Alteration indices

Alteration indices formulated from two or more componentsof major element analyses (Section 2.4) enhance thecompositional contrast between variably altered samples andthus are frequently more effective as exploration vectors thansingle component lithogeochemical data.

For example, sulfur and Na2O proportions in the footwallof the Rosebery K-lens deposit range from 0.01% (limit ofdetection) up to about 7.2% and 5.6%, respectively (Largeand Allen, 1997). However, the ratio S/Na2O ranges from0.002 to 194, because those components increase and decreaserespectively in response to increasing alteration intensity(Large et al., 2001b). Both components vary over two to three

orders of magnitude, whereas S/Na2O varies across about fiveorders of magnitude. Carefully formulated indices can in thisway amplify compositional changes and reflect variations inmore than one mineral composition or abundance.

Where systematic lithogeochemical data are available,plotting and contouring of alteration indices on plans and cross-sections provides numerical indications of alteration intensity(e.g. Fig. 2.7). Datasets of alteration indices are of assistancein guiding exploration towards potentially mineralised alteredzones, especially when used in combination with alterationfacies or mineral zonation maps. Mineral explorers haveincreasingly applied these techniques to VHMS explorationover the last two decades; however, few results or case studieshave been published.

Exploration data are commonly limited to a fewsamples or drill holes and are not suitable for contouring.Nevertheless, useful vectors can be inferred from sparse butstrategically or fortuitously located data. This is exemplified inlithogeochemical data from a few drill holes near the northernend of the Rosebery deposit (65R, 109R, 113R and 128R;Table 8.1). If, in a VHMS exploration scenario, the first twoholes were drilled in sequence (65R followed by 109R), thenthe lithogeochemical vectors would suggest that explorationwas heading away from the most favourable zone. Theintermediate third hole, 113R, would then be superfluous,merely reinforcing interpretation of vectors in the first twoholes. The anomalous values in the near-miss hole (65R)would encourage further persistence. If, on the other hand,the first hole in a greenfields exploration program was 109R,the major element or alteration indices data would not justifycontinuing exploration in that vicinity, even if the favourablestratigraphic setting was recognised. In this case, successwould depend on the explorer recognising other vectors orindicators of proximity, such as the distal trace element Tl andSb halos identified by Large et al., (2001b).

In favourable geologic settings, limited lithogeochemicaldata, even from a single drill hole, may yield useful vectors.For example, samples from a single hole, such as HL6 or

TABLE 8.1 | Selected major element data and alteration indices for samples of pumice breccia from the footwall to the Rosebery K-lensmassive sulfide deposit, western Tasmania. The values tabulated are (A) averages of three samples from the top 30 m of the footwall unitand (B) the uppermost sample of the footwall intersected in eachdriii hole. The alteration indices, S/Na2O and Al, generally show greaterincreases with proximity to ore than the changes in Na2O, S and Zn. Averaging the uppermost three samples smoothes the gradientstowards ore, but diminishes the anomalies in the medial intersection, 113R. Data from Large and Allen (1997). Locations of the drill holesare shown in Figure 2.7 of this volume and Figure 2 of Large et al. (2001b).

(A) Averages of three samples from the top 30 m of the footwall unit

109R 450 2.99 0.18 0.00 0 44

113R 250 1.66 0.37 0.13 2 61

65R 75 0.28 0.79 0.18 44 89

128R 0 0.01 0.49 0.49 49 89

(B) Uppermost sample of footwall unit

109R 450 1.53 0.29 0.01 0 51

113R 250 0.21 1.09 0.38 5 89

65R 75 0.08 1.02 0.02 13 95

128R 0 0.01 0.37 0.19 37 89

HL14 drilled through the footwall zones beneath the Hellyermassive sulfide deposit, generally exhibit gradually increasingalteration indices (Figs 9 and 20 of Gemmell and Large,1992). Recognition of these variations, in combination withthe visible alteration facies, confirms that an altered zone existsand indicates the direction of increasing alteration intensity,guiding further exploration.

Drill hole NC4, which intersected the Tyndall Group-Central Volcanic Complex contact south of Henty is anotherinteresting example. In this hole, an abrupt down-hole increasein the alteration indices is associated with a change of lithotype(Fig. 4.4). The lithogeochemical data support the recognitionof an extensive stratabound altered zone of which the upperboundary is most favourable for VHMS exploration.

Bivariate (x-y) plots of two alteration indices, such as theAI-CCPI Alteration box plot (Large et al., 2001a), are usefulin identifying compositional trends and different alterationfacies. This graphic approach simplifies the recognition of rockcompositions that lie outside the normal range of primaryvolcanic compositions (i.e. those that have been modified bychemical or depositional processes; Fig. 2.9). It also assistsclassification of different alteration facies and identifying thezones of greatest prospectivity (Fig. 2.11).

In recent CODES research projects, box plots of custom-designed alteration indices have been effective in severalother types of hydrothermal systems, including low- andhigh-sulfidation epithermal Au-Ag deposits (Williams, 2000)and Broken Hill type Pb-Zn-Ag deposits (Large, 2004).The Ishikawa et al. (1976) alteration index (AI) has beensuccessfully applied to many plagioclase-destructive and/or K-feldspar-bearing alteration styles, but there is scopefor more experimentation with new indices. As outlinedin Section 4.1, the formulae for alteration indices typicallyhave chemical components that were increased by alterationin their numerators, and components that were decreasedin the denominators. The gained or lost components canoften be inferred from the differences in alteration mineralassemblages. However, immobile-element-based mass changecalculations provide a more rigorous method of selectingcomponents for formulating alteration indices. Section 4.1summarises several techniques of estimating mass changesby comparing compositions of alteration facies to their least-altered precursor compositions and their potential applicationto exploration vectors is discussed below.

Mass change vectors

Hydrothermal alteration commonly involves major changesin chemical composition; in fact these changes are one ofthe characteristic features of hydrothermally altered rocks.Significant masses of mobile chemical components may havebeen gained or removed from an altered zone. However,closure in composition data will obscure or distort theamounts of these changes, except in the special cases wherethe mass gains and loses balance exactly, so that there is no netmass change. It is unsound to assume zero net mass change inan alteration facies, and in these cases the unquantified effectof closure on raw major element data limits their usefulness,or that of alteration indices based on them, as indicators ofalteration intensity.


The solution to the closure problem is to estimate themass changes of all mobile major-element components, usingan immobile-element-based method of the type described inSection 4.1. Spatially located mass change data can then beused as direct indicators of alteration intensity or as multiplecomponent alteration indices, in the same way as majorelement lithogeochemical data.

This approach has the potential to target favourable areasduring exploration. It provides closure-free quantification ofcompositional changes, which help delineate hydrothermalfluid pathways, zones of greatest alteration intensity andprospective areas. Furthermore, the quantification of absolutemass changes is a means of estimating the 'quality' of analtered zone.

For example, let us consider a hypothetical program oflithogeochemical sampling over two altered zones of similardimensions in a VHMS district. Mass change estimates mightshow that the first altered zone involved negligible masstransfers and the second had significant mass gains, of theorder of 20—30 g/lOOg and equating to tens of millions oftonnes of altered rock (cf. Thalanga footwall zone, Herrmannand Hill, 2001). In this case, we would conclude that thesecond altered zone has greater mineral potential. Substantialmass changes demonstrate that a hydrothermal system hadthe intensity, and perhaps duration, to move large amountsof SiO2, CO2, S and other components into the alterationfacies. Therefore, it probably also had the capacity, if fluidcompositions were suitable, to move large amounts of baseand precious metals and potentially, if a favourable site andprocess for deposition is available, form an ore deposit. Thefirst altered zone in our hypothetical example was produced bynear-isochemical alteration and resulted in negligible changesto the whole-rock composition, suggesting that alterationinvolved less reactive or smaller volumes of fluid, perhapsover a short duration. The differences may be semi-evident inalteration mineral assemblages and intensities, and possiblyin the composition data despite distortion by closure, but theonly way to quantify the difference for objective explorationdecisions is by mass transfer techniques.

The major difficulty in this method is in determiningprecursor compositions to compare with the altered compo-sitions. Poor exposures, limited lithogeochemical data, lateralvariation in the primary composition of volcanic facies orstructural complexity make the pairing of alteration faciesand unaltered (or least-altered) precursors problematic, andfrequently impossible, in practical application. There are nopublished examples where mass change calculations haveled to a mineral discovery, probably because of the least-altered precursor problem and the only recent developmentof easy mass change calculation techniques. Nevertheless,the mass change approach will contribute to a higherlevel of lithogeochemical interpretation and explorationtargeting where host volcanic successions are compositionallyuniform and sufficiently understood to enable its confidentapplication.

Mineral chemistry vectors

As noted in Section 4.2, the main limitations to the wide use ofmineral chemistry in exploration have been that the analytical

2 4 6 | CHAPTER 8

tools — electron microprobe and X-ray diffraction — arecomplex laboratory-based instruments requiring considerableexpertise in operation and data interpretation. That has mademineral analysis slow and expensive relative to geochemicalanalyses of rocks, soils and sediments, and consequentlyexplorers have largely ignored the mineral chemistry vectorpossibilities. Researchers at CODES are currently developinglaser ablation ICP-MS techniques for micro-analysis of traceelements in sulfides. These are likely to provide explorationvectors but, for similar reasons, they may not ultimately bewidely applied by mineral explorers.

However, the advent during the last decade of portableshort wavelength infrared (SWIR) spectrometers, whichcan indirectly measure compositional variations in micas,clays and carbonates, could establish mineral compositionmapping as a viable exploration technique (Sections 3.1and 4.2, and references therein). SWIR spectrometers suchas PIMA are relatively inexpensive at about US$21,000 topurchase or US$70 per day for hire. They can analyse upto a few hundred samples per day of all types of geologicalmaterials, which require no preparation apart from drying.SWIR spectrometers are simple to operate and the PC-basedspectral recognition software now available has simplifiedspectral interpretation and data manipulation, so that anoperator can quickly become an expert interpreter.

White micas, chlorites and clays in altered zones aroundmineral deposits frequently show spatial compositional varia-tions that could be exploration vectors (Section 4.2). The easeof SWIR spectral analysis now enables explorers to rapidlytest for the existence of mineral composition vectors in a largeenough set of orientation samples. If the results are promising,the technique can be inexpensively applied on a routine basisto assist exploration targeting. If, on the other hand, SWIRspectral features are invariant or spatially erratic, then littletime and money will have been expended.

There are not yet many published mineral explorationcase studies involving portable SWIR spectral analysis becauseit is a relatively new technique (e.g. Denniss et al., 1999;Huston et al., 1999; Merry and Pontual, 1999; Herrmannet al., 2001; Jones et al., in prep.). Nevertheless, recent andcurrent research at CODES shows great potential for SWIR-determined white mica composition vectors, on scales of tensto hundreds of metres, in a variety of volcanic-hosted goldand base-metal deposits. Further work is required on spatialSWIR spectral variations in chlorites and clay minerals. It islikely that mineral explorers will rapidly adopt this techniqueover the next few years.

Part of the stimulus comes from very recent developmentsin airborne high-resolution visible-to-SWIR spectral scanningsystems, such as HyMap", which offer great promise for district-scale mineral mapping in exploration of well-exposed bedrockareas (Taranik, 2001). For example, mineral maps from a trialHyMap* airborne spectral survey of the Panorama VHMSdistrict, Western Australia, apparently 'show the completehydro thermal convective system' (Cudahy et al., 2000). AtPanorama, these authors consider that spectrally interpreteddistributions of white mica, pyrophyllite and topaz definealtered zones that formed at the boundary between magmaticfluid and seawater convection, in addition to seawater rechargezones, and hydrothermal discharge zones. The discharge zonesare prospective for massive sulfides. A similar HyMap* survey

of the Mount Lyell area in western Tasmania has producedmineral distribution and pyrophyllite abundance maps (e.g.Fig. 8.1). These illustrate the high spatial resolution nowavailable from airborne spectral surveys, and their enormouspotential for alteration mapping and using vectors duringexploration in well-exposed, thinly vegetated areas.

Remote sensing spectral systems are also findingapplications in regolith mapping (Craig, 2001) and explorationof partly covered areas. Bierwirth et al. (2002) used HyMapdata to map distributions of a range of minerals — includingpyrophyllite, white mica, Mg- and Fe-chlorite, calcite,dolomite, kaolinite, tourmaline, hematite and goethite — inaltered zones associated with epithermal and lode Au depositsin the poorly exposed, largely alluvium- and calcrete-coveredIndee District of the Central Pilbara.

These demonstrations of district-scale mineral andmineral compositional mapping by remote sensing toolsshould certainly encourage explorers to use spectral data inprospect-scale investigations. In addition, high-output, multi-purpose visible-SWIR and thermal infrared spectral, andlaser instruments such as CSIRO's HyLogging and HyChipssystems (Syddell, 2004) and the OARS prototype (CSIRO,2002), are being developed for routine logging of drill core,cuttings, soil and other geological sample materials.

Isotopic vectors

Section 4.3 introduces the potential for isotope geochemistryto yield interpretations of hydrothermal fluid sources,temperatures, water-rock ratios, and broad halos forexploration targeting.

Oxygen isotopes are particularly useful in explorationbecause oxygen is a major component of hydrothermal fluids,and it readily exchanges isotopes with silicate minerals atfractionation factors that are mineral specific and temperaturedependent. Furthermore, the 518O-depletion halos observedaround several deposit types typically extend further fromore than most other geochemical anomalies and may providedirect vectors to ore zones. For example, the 518O-depletionzone around the Fukuzawa deposits in the Hokuroku district,Japan, extends up to 1 km beyond the Na2O-depletionanomaly (Green et al., 1983). Waring et al. (1998) found618O-depletion anomalies in dolomitic shale at Mount Isa(Queensland), which extend up to 2 km beyond Cu ore zones,with low and uniform isotopic gradients (<2%o per 100 m)that allow estimates of the distance to ore. Most importantly,the O-isotopic anomalies produced in hydrothermallyaltered zones appear to survive subsequent deformationand metamorphism. For instance, Cartwright (1999)argued convincingly that a hydrothermally related regional-scale 618O depletion zone in Proterozoic metapelites in theBroken Hill district, NSW, had survived high-grade regionalmetamorphism up to granulite facies. The final section of thischapter summarises several VHMS-related alteration studiesand exploration programs, which have applied whole-rock O-isotope geochemistry.

Sulfur-isotope geochemistry has been widely appliedto interpretations of sulfur (and hence fluid) sources, andhydrothermal temperatures, which have been used in thedevelopment of VHMS genetic models. For example,

FINDING ORE DEPOSITS IN ALTERED VOLCANIC ROCKS I 2 4 7

Map projection: UTM zone 55, AGD66

FIGURE 8.1 | Mineral maps of the Mount Lyell mine area, western Tasmania, interpreted from HyMap® airborne hyperspectrai data. Map A shows the zonaldistributions of eight important alteration minerals. Map B shows relative abundance of pyrophyllite (warm colours = high abundance), and discriminates thepyrophyllite-rich facies at North Lyell, Western Tharsis and Glen Lyell from weaker responses in the Owen Group exposed on Mount Lyell. The spatial resolution (pixelsize) is about 5 m. Mineral spectral responses are partly restricted by vegetated areas, which appear as dark grey tones on the HyMap band (greyscale) backgroundairphoto images. These maps were created by K. Yang, M.A. Quigley and J.F. Huntington as part of the 2003 HyMap® mineral mapping project for Copper Mines ofTasmania and Mineral Resources Tasmania, carried out through the C-Vista strategic alliance between CSIRO and HyVista Corporation.

2 4 8 | CHAPTER 8

S-isotope compositions constrained some of the geneticinterpretations for formation of the Hellyer deposit (e.g.Gemmell and Large, 1992; Solomon and Khin Zaw, 1997).It also has exploration potential for discriminating differenttypes of deposits and hydrothermally altered zones, whichmay have different economic potential. The regional studyof sulfide deposits in the Mount Read province by Solomonet al. (1988) found considerable variation in S-isotopecompositions consistent with different geologic settings andmetal associations, and which contributed to interpretationsof hydro thermal geochemistry. Green and Taheri (1992)took both a genetic and discriminatory approach to theinterpretation of low 634S values of pyrite (-1.2 to +4.7%o)at the Boco prospect, western Tasmania. They suggested thataltered zones at Boco formed in a seawater-hydrothermalsystem, which leached sulfur from volcanic host rocks attemperatures that were too low to inorganically reduceseawater sulfate, and transported base metals to form an oredeposit. Subsequent recognition of advanced argillic alterationmineral assemblages at Boco and several other Tasmanian andVictorian prospects indicate possible involvement of magmaticfluids, and hence a magmatic source of sulfur (Herrmann etal., 2004). Regardless of the genetic uncertainties, sulfide634S values of less than 5%o could distinguish barren pyriticaltered zones from more prospective base and precious metal-rich VHMS systems in the Mount Read province.

There are few published accounts of S-isotope compositionsas direct vectors in mineral exploration. However, existing datafor Rosebery (Davidson et al., 2000) and Hellyer (Gemmelland Large, 1992) suggest broad halos of 634S enrichment indisseminated pyrite in footwall zones lateral to the main up-flow zones, which could be used to increase exploration targetsizes and zero-in on Zn-rich VHMS deposits, particularly inpermeable volcaniclastic successions (Large et al., 2001c).

At regional scales, two recent studies of deeply coveredareas have promoted S-isotope compositions of sulfates ingroundwater as potential indicators of buried oxidising Pb-Zn-Ag sulfide deposits, in the Broken Hill region of NewSouth Wales (Waring et al., 1998) and Gawler Craton inSouth Australia (Kirste et al., 2003). In the latter case, sulfatesfrom oxidising sulfide deposits with low 534S signatures (-2.5to +5.6%o) appear to have contributed to anomalous low 534Svalues in groundwater sulfates, detectable several hundredmetres downstream from the Menninnie Dam prospect.Background 634S values of sulfates in ground waters are 16 to18%o in the Gawler region and -13.5%o in the Broken Hillregion. The concept is probably less applicable to explorationfor sulfide deposits with higher 634S signatures (e.g. TasmanianVHMS deposits, 8 to 17%o, Solomon et al., 1988), whichwould provide less contrast against background groundwatercompositions. Furthermore there are many, typically difficultto determine, hydrological and geochemical factors thatcomplicate interpretations of local groundwater isotopicanomalies. This new application of S-isotope geochemistry isone that will probably appeal only to the most persistent ofunder-cover mineral explorers.

Carbon isotopes, like sulfur, are used for interpretingfluid sources and hydrothermal conditions but have notbeen widely applied as exploration vectors. Huston's (1999)review of stable isotopes in VHMS systems found carbonate613C values in most deposits occupy a narrow range of-5 to

0%o, consistent with seawater dissolved bicarbonate sources.Low fractionation factors, and the limited occurrencesof carbonates in massive sulfide deposits (Ohmoto andGoldhaber, 1997) restrict the applications of C isotopes,except in conjunction with O isotopes. For example, KhinZaw and Large (1992) interpreted a coupled positive trendof 613C and 618O data in Mn-rich carbonates at SouthHercules, Tasmania, as temperature-related, and then, withadditional fluid inclusion temperature data, estimated theisotopic compositions of the hydrothermal fluid. Althoughtheir paper did not describe spatial zonation of isotopic data,the genetic discussion speculated that mineralised and alteredfacies were zoned according to variations in temperature andhydrothermal fluid-seawater mixing ratios, controlled bypermeabilities in the volcaniclastic succession. In these typesof deposits associated with lateral carbonate facies, isotopicdata could provide prospect-scale exploration vectors if thehydrothermal temperature gradients were consistent.

Callaghan's (2001) study of the Henty-Mount Julia golddeposit, Tasmania, used carbonate 613C and 618O data ina boomerang shaped trend for intensive modelling of fluidcompositions and genetic concepts. The data, crudely dividedinto proximal and distal carbonates, lie on two trends joinedat an abrupt inflection. Both of the fluid mixing or fluid-limestone interaction models proposed by Callaghan (2001)to account for the trends offer potential for prospect to district-scale isotopic vectors, or at least methods of discriminatinghydrothermal and sedimentary carbonates.

Whole-rock O-isotope vectors in VHMS exploration

In most cases, the proximal altered zones of VHMS systemsshow significant 618O depletion, partly attributable to high-fluid temperatures and low-fractionation factors of someminerals (e.g. chlorite) in seafloor hydrothermal dischargezones, and partly due to the contrast with 618O enrichmentcaused by low-temperature seawater-rock reactions in normalsubmarine volcanic successions.

A classic semi-regional study by Cathles (1993) in theNoranda district, Canada, discovered a low 618O anomaly(<6%o) in volcanic rocks around the Flavrian felsic pluton. Thepluton is surrounded by a discontinuous halo of high whole-rock 618O anomalies (>9%o) 10-15 km from the intrusion.Several narrow finger-like zones of low 618O values extendradially from the inner 18O-depleted zone through the high618O halo, in the directions of most of the known VHMSdeposits in the district (Fig. 8.2). These low 518O zones recordareas of high hydrothermal-fluid flow and concentrateddischarge, which are favourable for mineral deposits. Theconcentric zones of 18O depletion and enrichment aroundthe pluton closely match the isotopic zonation patterns ofCathles' (1983) numerical model. He concluded that whole-rock 818O sampling, at 0.5 km intervals along traversesadjacent to the margins of plutons, could identify plutonswith sufficient energy to drive long-lived hydrothermalsystems, and favourable settings for detailed massive sulfideexploration.

Another district-scale study, in the Panorama area ofWestern Australia, showed a similar pattern of low whole-rock618O around the perimeter of a large subvolcanic intrusion


FIGURE 8.2 | Map showing spatialrelationships between felsic plutons, whole-rock618O anomalies, and massive sulfide deposits inthe Noranda area, Abitibi belt, Canada (modifiedafter Cathles, 1993).

(Brauhart et al., 2000). The granitoid pluton underlies a1.5 km thick mixed mafic to felsic volcanic succession thathosts several small polymetallic massive sulfide depositsand prospects along a single favourable horizon at thestratigraphic top of the sequence (Fig. 6.6). Narrow radialzones of low 618O point to most of the known deposits andprospects (Fig. 8.3). These low 618O zones coincide withintense feldspar-destructive sericite + quartz and chlorite +quartz zones. Brauhart et al. (2000) calculated hydrothermaltemperatures from the 618O data. They used fractionationfactors calculated to suit the specific modal mineralogy ofeach sample, an initial fluid 618O value of+2%o and assumedhigh water-rock ratios. The resulting calculated temperaturedistribution closely matched the O-isotopic pattern, the low618O zones coinciding with temperatures greater than about300°C (Fig. 8.3). This indicates that temperature was themain control on low whole-rock 818O. It is consistent withincreased temperature with depth in the volcanic succession,and in the transgressive discharge or feeder zones beneaththe sulfide deposits. The authors concluded that whole-rockO-isotope mapping could be used as a regional explorationvector, and pointed to additional favourable targets in thePanorama district.

Green and Taheri (1992) followed up the Hokurokuwork of Green et al. (1983) with several empirical isotopicstudies of alteration systems in the Mount Read province. Thealtered footwall zones beneath the Hellyer deposit exhibit asubtle whole-rock 618O anomaly with values ranging from8.3 ± 1.3%o in the central stringer zone, through 9.8 ± 1.7%oin the enclosing sericitic zone, to background values around11.3±0.9%o in adjacent least-altered footwall andesites.There is also a subtle depletion anomaly of 10.6 ± 1.2%oin the basalts immediately above the deposit, compared tobackground values of 11.8±2.2%o. However, the 518O-depletion zone is narrow, reflecting the strong fault or fracturecontrol on hydrothermal-fluid flow. This limits its utilityin exploration. The Hercules alteration system also showsa range of whole-rock 6!8O values from 6.8%o in footwallzones to background values of 14.0 to 15.5%o. There aresome unexpectedly high values (around 15%o) in relativelyproximal parts of the footwall and lowvalues (down to 6.8%o)

FIGURE 8.3 | Distribution of whole-rock 618O values and estimatedhydrothermal temperatures in the Panorama district of the Pilbra region, WesternAustralia (modified after Brauhart et al., 2000).

2 5 0 | CHAPTER!

in the apparent hanging wall rocks to the east. These maybe partly due to fault displacements that dismembered thealteration system. They highlight the difficulty of applyingbroad-scale geochemical exploration techniques in deformedand structurally complex terrains.

In contrast to Hellyer and Hercules, the apparently barrenBoco altered zone has whole-rock 618O values that are notsignificantly different to background values (9.9 ±1.0 and10.5 ± l.l%o, respectively). This is consistent with Green andTaheri's (1992) interpretation that the Boco alteration faciesformed in a low-temperature (<200cC) seawater hydrothermalsystem, incapable of transporting base metals and reducingseawater sulfate. Alternatively, the 618O values could indicatea higher temperature, isotopically heavier fluid (>280°C,--5%o), representing either evolved seawater or mixed seawaterand magmatic water. The presence of advanced argillicassemblages in parts of the Boco system implies highly acidicfluid conditions, which supports a magmatic fluid input(Herrmann et al., 2004).

The least-altered volcanic rocks in VHMS-hostingsuccessions typically have anomalously high backgroundwhole-rock 618O values (>9 or 10%o), which are attributableto low-temperature diagenetic alteration. The curves inFigure 4.17 indicate that re-equilibration with quite smallproportions of cold seawater can produce large positive shiftsin volcanic rock 618O values. On the other hand, zones oflow 618O reflect high-temperature hydrothermal alterationat high water-rock ratios. The empirical data from Norandaand Panorama show that low 818O zones may be regionally

extensive at depths of greater than 1 km below favourablehorizons, and in narrow finger-like zones that point towardfavourable sites for hydrothermal discharge. They may formrelatively broad halos around massive sulfide deposits.

Despite these promising research results, VHMSexplorers have been less than enthusiastic about O-isotopevectors and there are few examples of successful applicationin Australia. This may be largely attributable to the expenseof isotopic analysis (currently around US$150 per sample)and the recognition that interpretation of isotopic data is notstraightforward.

A notable exception is the case of the Thalanga West 45deposit, documented by Miller et al. (2001). These authors tooka similar approach to Brauhart et al. (2000), using estimatesof modal mineralogy to determine tailor-made fractionationfactors for each sample, to calculate isotopic equilibrationtemperatures from whole-rock 618O data, at assumed highwater-rock ratios and fluid isotopic composition. They foundthat zones of apparent high temperatures (>230°C) coincidedwith the known Central, East and Orient massive sulfide lenses.The existence of an additional isotopic-temperature anomaly,about 1 km west of the known resources, stimulated furtherexploration that turned up a favourable REE geochemicalanomaly in the same sector. Subsequent exploratory drillingdiscovered a 0.23 Mt polymetallic massive sulfide lens. Itremains sub-economic, but may represent the first successfulVHMS exploration application of O-isotope geochemistry inAustralia.

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I 271

INDEX

A page number in bold indicates that the reference is to afigure. A bold t indicates that the reference is to a table.

AI-CCPI Alteration box plot see Alteration box plotalbite alteration 39, 42-44, 61, 65, 69, 133-5, 165-7, 191,

214,216Alteration box plot 31-4, 36, 169-70, 245alteration distribution 22, 63-4alteration facies

describing and defining 15—36diagenetic 241—2distribution 63-4hydrothermal 242-3metamorphic 242variables 22

alteration fluids 93, 170-3Alteration Index (AI) 30-2, 34, 169, 245alteration indices 26-30, 34, 73, 169-70, 244^5alteration intensity

describing 27testimation, integrated approach 33—6explained 25—36illustrated 28-29, 36lithogeochemical indications of 243-4

alteration mineral assemblages see mineral assemblagesalteration nomenclature 16, 19-22, 2It, 22t, 23talteration pipes 63, 164-74, 176, 182-3alteration plumes 63, 167, 168, 191-3alteration processes 4—6alteration rates 6alteration textures

deformation textures 52, 54-6, 55-7described 37-63, 37t, 38t, 62tdissolution textures 41, 50-1, 52dynamic recrystallisation textures 52illustrated 39-40, 62, 103-4, 110, 111-13infill textures 41,48-9overprinting and false/pseudo textures 37, 54-63,

58-61, 62trecrystallisation textures 52, 53replacement textures 37-8, 41, 42-7static recrystallisation textures 52, 53

alteration timing 69-71, 7It, 172alteration zonation

boundaries 64contact altered zones 5, 64, 66-7, 67, 139, 149-56,

242diagenetic 64, 105-8facies model 3greenschist facies zones 115, 116, H6t, 131, 142,

144-5, 152halos 66-9Hellyer deposit 178,182-3,184-93Henty deposit 178, 212-3, 214-20Highway Reward deposit 14, 167, 178, 232, 233-40Hokuroku Basin 119-27hydrothermal 5-6, 66-8, 164-9, 243mapping 243metamorphic 64—5, 66—7Mount Read Volcanics 128-38patterns 64-9,98, 165-8regional deep semi-conformable 66, 142-8regional metamorphic 115-17,140Rosebery deposit 178,195,196-201scales described 64tThalanga deposit 14, 178, 222, 223-31veins and fractures 67, 69Western Tharsis deposit 202-3,204-11

amphibolite facies 115-17,140Amulet deposit see Noranda districtanalytical techniques

electron microprobe 19,25,88field observations 18ICP-AES (inductively coupled plasma atomic emission

spectrometry) 76-7ICP-MS (inductively coupled plasma mass

spectrometry 76—7isotope geochemistry 92—5HyMap® 246, 247lithogeochemical sampling 73-87mineral chemistry analysis 87-91NAA (neutron activated analysis) 76petrography 24-5, 33-4PIMA 25, 33, 245

2 7 2 | INDEX

SWIR spectroscopy 19, 24, 25, 33, 88, 90, 202-3,243, 245

X-ray diffraction (XRD) 19, 24, 25, 33, 88X-ray fluorescence spectrometry (XRF) 76-7

anhydrous minerals 97

BBathurst mining camp 164-5,170Boco prospect 248, 250burial-related alteration 97-138 see also diagenesis and

submarine environments

ccarbonates

diagenetic 105in exploration 90-1, 243-4hydrothermal 47, 91, 166, 178, 188, 201, 217,

229-30cataclastic texture 52CCPI (chlorite-carbonate-pyrite index) see also Alteration

box plotexplained 31, 34exploration, uses for 169,245

cementation 97, 102, 105, 108-10, 132Central Volcanic Complex see Mount Read Volcanicschlorite 19-22, 89, 138, 156, 165-8, 187, 239closure

and alteration indices 26constant sum effect 243—4explained 78-9,243-5mass change anomalies 81

compaction 97, 109-10, 132compositional nomenclature 20contact alteration 5, 149-62corrosion vugs 41, 50—1, 52crystallisation

primary 4textures 52of zeolite assemblages 105, 110, 114, 121

DDarwin Granite see Mount Read Volcanicsdata sheets

contents of 36Darwin Granite 157-62Hellyer deposit 184-93Henty deposit 214—20Highway-Reward deposit 233-40Hokuroku Basin 122-7Mount Read Volcanics 133-8Rosebery deposit 196-201Thalanga deposit 223-31Western Tharsis deposit 204—11

deep, semi-conformable altered zones 142—6deformation textures 52, 54-6, 55-7detection limit explained 75deuteric alteration 148devitrification

explained 4texture 37, 39, 62zones 151

diagenesisexplained 5Hokuroku Basin 118-27isotope geochemistry analysis 93-4and metamorphism 16, 98, 102, 114, 115Mount Read Volcanics 128-38in submarine volcanic successions 97, 102—14

diagenetic mineralscarbonates 105genesis of 108-14layered silicates 102,105other diagenetic minerals 105zeolites 105, 110, 118, 120-7

diagenetic zonesHokuroku Basin 118—27Mount Read Volcanics 128-38zonation 64-5, 105-8

discharge zone 141—2dissolution 41, 50-1, 52, 97, 102, 108-10, 114, 132dynamic recrystallisation textures 52

electron microprobe see analytical techniqueselement concentrations 32-3eutaxitic texture 54, 57exploration

Alteration box plot 31-4,169-70alteration identification as tool in 241—50Alteration Index (AI) 30-2isotope geochemistry in 92-5, 246-50lithogeochemistry in 73-87mineral chemistry in 87-91sulfide mapping 243use of chlorite in 89use of white mica in 90—1vectors and proximity indicators 32-3, 94-5, 243-50

false textures see pseudotexturesfiamme 54, 57fluid-rock interaction 169, 170-2, 171, 173foliation 52, 54, 70, 711footwall alteration 163-74, 179, 182-9fused zones 151

geochemistry see isotope geochemistry, lithogeochemistrygeothermal gradient 98geothermometers 92-3glass

alteration in submarine volcanic successions 15, 97-8common alteration minerals 19tcrystallisation 4diagenesis 102—14disequilibrium assemblages 24hydration 4-5, 98-102reactive quality, 6

Green Tuff Belt see Hokuroku district

INDEX | 2 7 3

Hhalos 5-6, 38, 41, 66-9, 149-56, 157-62, 163-74, 178-

81hanging wall alteration 163, 164, 167-8, 190-3Hellyer deposit see also Mount Read province

alteration fades and zonation 178, 182-3, 184—93Alteration Index (AI) 31, 183explained 11, 181-93exploration 245geological setting 181—2ore genesis 170, 183white mica 90

Henty deposit see also Mount Read provincealteration facies and zonation 178, 212-13, 214-20explained 12geological setting 212hanging wall alteration 167isotopic data 248ore genesis 213

Hercules deposit see also Mount Read provincealteration halo 178explained 12, 128geological setting 194

Highway-Reward deposit see also Mount WindsorSubprovince

alteration facies and zonation 14, 167, 178, 232,233-40

explained 14geological setting 12-14, 232hanging wall alteration 167,178ore genesis 232submarine facies associations 13

Hokoroku Basin see Hokuroku districtHokuroku district

alteration 64-5, 67, 118-27, 122-7geological setting 118—20Green Tuff Belt 52, 64, 107, 118, 150, 151Kuroko deposits see Kuroko depositsoxygen isotopes 94-5, 246size of VHMS deposits 164

hydration of volcanic glass 4-5, 98-102hydrothermal alteration

boundaries between zones 64chemical reactions 168and diagenetic alteration 128discharge zone 141-2,156discriminating 16—19explained 4halos 66-7, 164-74intensity measures 32-3intrusion-related 140-61 see also intrusionsmetamorphic assemblages 174, 175tplagioclase destruction 31, 167recharge zone 141subseafloor systems 140—1syntectonic 6tectonic deformation 6VHMS deposits 5-6, 163-240zones 5-6, 66-8, 164-78, 243

hydrothermal convection 1, 94, 140-1, 140

hydrothermal fluid 67, 172-3HyMap® 246, 247

Iberian pyrite belt 90, 91, 142, 164, 165, 166, 166, 174ICP-AES (inductively coupled plasma atomic emission

spectrometry) 76—7ICP-MS (inductively coupled plasma mass spectrometry

76-7indices

alteration 26-30, 34, 73, 244-5Alteration box plot see Alteration box plotAlteration Index (AI) 30-2, 34, 169, 245CCPI (chlorite-carbonate-pyrite index) 31, 34, 169,

245molar proportion alteration 30multi-component and normalised 26, 30simple ratio 26

induration 150, 151infill textures 41,48-9intrusions

halos 66—7cryptodomes 2, 6, 66, 128, 139, 143, 153, 212, 232dykes 2, 6, 66, 141, 139, 149, 152, 153plutons 139, 143, 153sills 66, 70, 100, 128, 148-9, 152-3, 150, 152, 174,

182, 194,202,212,221,232in submarine volcanic successions 2 ,3 , 139synvolcanic 139—62

isotope geochemistryapplications 92-5carbon 76, 77, 248exploration 94-5, 246-50hydrogen 76, 77, 93oxygen 94-5, 246,248-50stable isotopes 92-5sulfur 76, 77-8, 92, 141, 180, 246, 248water-rock ratios 93

Kkaolinite in VHMS altered zones 88, 150, 174-75, 178-80keratophyre 98K-lens see Rosebery depositKuroko deposits see also Hokuroku district

alteration 164, 179Alteration Index (AI) 31alteration model 166, 178—9

least altered see also alteration intensityexplained 26alteration indices 32

lithification 97, 108-10lithogeochemistry

analytical methods 73-8carbonates 76-7chemostratigraphy 79-81C-H-N elemental analyser 76closure 78-9, 81compatible elements 79—80

2 7 4 I INDEX

europium 87explained 73-8, 74tand exploration 243-5hydrous minerals 76-7ICP-AES (inductively coupled plasma atomic emission

spectrometry) 76-7ICP-MS (inductively coupled plasma mass

spectrometry 76-7immobile elements 79-81, 85, 87inaccuracies in 77incompatible elements 79limit of detection 75LOI (loss on ignition) 76mass change 73, 81-7, 85-6, 87, 97, 165, 180-1,

245-6NAA (neutron activated analysis) 76precision and accuracy required 75recalculating to volatile free 77-8REE (rare earth elements) 73, 79, 81, 87reporting data 77sampling methods 73—8summing elements 77use of reference materials 75XRF (X-ray fluorescence spectrometry) 76-7

Mmass change see lithogeochemistrymassive sulfide 163—5, 167Mattabi deposit 91, 180metamorphism

burial metamorphism and diagenesis 16, 97, 98, 102,114, 115-17

contact metamorphism 5, 11, 12, 64, 66-7, 149-54,242

explained 4, 5, 20, 24regional metamorphism 5, 64-6, 115, 139, 140-8of VHMS-related altered zones 174-5

metasomatic alteration 4, 5, 144-6microanalysis 24microprobe see analytical techniquesmineral assemblages

alteration assemblages 23—5, 34, 165burial effects on 97, 109common assemblages 21t, 22t, 109, 165disequilibrium 23-4, 63equilibrium 23-4in exploration 243igneous 16, 19isotopic studies of 92—5nomenclature 19-20

mineral chemistry 87-91, 245-6minerals defined 87-8Mount Lyell field see also Mount Read province

deposits 12,202,243geological setting 202halos 202hanging wall alteration 167HyMap system 246, 247mineral zonation 243, 247ore genesis 203Western Tharsis deposit see Western Tharsis deposit

Mount Read provincealteration 7-12, 163-164Chester deposit 90history 9, 11-12Hellyer deposit see Hellyer depositHenty deposit see Henty depositHercules deposit see Hercules depositMount Lyell field see Mount Lyell fieldMount Read Volcanics see Mount Read Volcanicsoxygen isotopic exploration 249—50Que River deposit see Que River depositRosebery deposit see Rosebery depositsize of VHMS deposits 164Western Tharsis deposit see Western Tharsis deposit

Mount Read Volcanics see also Mount Read provinceAI and CCPI ranges 32,34alteration 128-32, 133-8Central Volcanic Complex 9-10, 69, 128, 130-32,

157-62chemostratigraphic discrimination and correlation 80compaction effects at 110Darwin Granite 154-6, 157-62geology of 7-12, 128, 129Kershaw Pumice Formation 128,134—6Mount Black Formation 128, 133, 137-8pyritic alteration systems 89metamorphic assemblages 11Sterling Valley Volcanics 128

Mount Windsor Subprovincealteration 14, 164, 222, 232geology 12-14, 221-40Highway-Reward deposit see Highway-Reward depositThalanga deposit see Thalanga deposit

NNAA (neutron activated analysis) 76 see also analytical

techniquesnaming altered rock see alteration nomenclatureNoranda district 66, 142-7, 167, 179-81numerical fluid-flow modelling 172

ooverprinting

textures 37, 70-1, 70t, 711and false/pseudo textures 37, 54-63, 58-61, 62trelationships 69-71, 7It

palagonite 99, 99-100Panorama district 180, 246, 248-50paragenetic sequence 69perlite 37,40,54, 100-1, 101PIMA (portable infrared mineral analyser) 25, 33, 245 see

also analytical techniquesplagioclase destruction 31,167-9pseudotextures 16, 37, 54-63, 58-61, 62t, 63pyrophyllite in VHMS systems 174, 202, 207-8

Que River deposit 11-12, 70, 90, 167 see also Mount Readprovince

INDEX I 2 7 5

Rregional metamorphism see metamorphismrelict textures 14, 16, 19, 24, 25, 37-38, 141replacement textures 37-38, 41, 42-7Rio Tinto deposit see Iberian pyrite beltRosebery deposit see also Mount Read province

alteration 70, 128, 131t, 168, 173, 178, 195, 196-201

Chlorite-carbonate-pyrite index (CCPI) 195geochemical alteration parameters 168geological setting 194explained 11-12, 194-201hydrothermal carbonates 91hydrothermal fluid flow 174K-lens 30, 34, 68, 91, 174, 244lithogeochemical data in exploration 82, 244sericite 70-1

Scuddles deposit 180sericite alteration 20-2, 70-1, 165-78, 185-6, 198-9, 206-

11, 218,225, 231, 238 see also white micaShort-wavelength infrared spectroscopy (SWIR) see

analytical techniquessiliceous alteration 165-9, 182, 189, 199, 219-20smectites 102, 109, 117Snow Lake District 142-6sodium content in volcanic rocks 34—6, 243—4solution seams 51, 52spilite 98stable isotopes see isotope geochemistrystockwork see stringer zonesstringer zones 163, 165, 172-3, 179stylolites 51, 52submarine environments 2, 97—138submarine facies associations

Mount Read Volcanics 10-11, 214-20Seventy Mile Range Group 12-13

submarine volcanic successions 1-14, 3, 97-138, 241-9SWIR spectroscopy see analytical techniquessynvolcanic intrusions see intrusions

Thalanga deposit see also Mount Windsor Subprovincealteration facies and zonation 14, 178, 222, 223-31explained 14geological setting 12—14, 221—2mass change estimations 85—6ore genesis 222oxygen isotopic exploration 250

thermodynamic alteration model 170-2

vectorsexploration 243—50isotopic 246—50lithogeochemical 243-4mass change 245mineral chemistry 79, 89-91, 245-6sulfide 243-4

vein-halo alteration 38,41,43,69VHMS deposits

alteration patterns 164—78classification 163—6common features 163—4comparisons 178-81exploration 241-50footwall alteration 163-74halos 5-6, 67-9, 68, 163-8, 175, 178-81hanging wall alteration 163—4, 167—8Hellyer deposit see Hellyer depositHenty deposit see Henty depositHighway-Reward deposit see Highway-Reward depositkaolinite, presence of 174major VHMS provinces 164Mount Lyell see Mount Lyell fieldpyrite in 243pyrophyllite, presence of 174regional alteration zones 5-6, 139-62Rosebery deposit see Rosebery depositThalanga deposit see Thalanga depositWestern Tharsis deposit see Western Tharsis deposit

VMS deposits see VHMS depositsvolcanic facies

alteration processes 2—6alteration in submarine environments 97—138associations 2changes in 108clastic facies 1,19coherent facies 1, 6, 16, 19common clay minerals in 102crystalline facies 6volcaniclastic facies 1, 6, 12, 16—19

volcanic-hosted massive sulphide deposits see VHMSdeposits

vugs 41,50-1,52

Wwater-rock ratios 93—4Western Tharsis deposit see also Mount Read province see

also Mt Lyell fieldalteration facies and zonation 202—3, 204—11case study 202-11geological setting 202ore genesis 203white mica 91

white mica 20, 89-91, 165, 246Woodlawn deposit 178

XX-ray diffraction 19, 24, 25, 33, 88X-ray fluorescence spectrometry (XRF) 76—7

zeolites 105-6, 106t, 110, 114-16, 118, 120-1zones see alteration zonation

About the authors

Dr Cathryn Gifkins is a Research Fellow at the Centre for Ore Deposit Research at the University

of Tasmania. Cathryn brings to the publication a strong background in mapping, describing and

interpreting altered and deformed volcanic rocks in submarine successions. Her current research

focusses on the textural, mineralogical and compositional effects of alteration in glassy volcanic

rocks, the link between volcanic centres and mineralising hydrothermal systems, and the facies

architecture and stratigraphy of the Mount Read Volcanics. '

Walter Herrmann is a Research Fellow in economic geology at the Centre for Ore Deposit

Research. Wally's background in mineral exploration in Australian volcanic successions,

principally the Mount Read Volcanics and the Mount Windsor Subprovince is a valuable asset

to the book. He has a special interest in understanding hydrothermal alteration as a method for

discriminating and discovering VHMS and porphyry deposits.

Professor Ross Large is Director of the Centre for Ore Deposit Research and has a long and

celebrated academic and exploration career. Ross has a comprehensive knowledge of VHMS

deposits, and has actively promoted and developed the application of geochemical techniques to

mineral exploration. This innovative approach has recently produced the Alteration box plot, an

alternative way to relate alteration intensity, mineralogy and geochemistry.

Altered Volcanic Rocks

Documents

alteration studies

common alteration textures

timing of alteration

hydro thermal alteration

volcanic facies geology

altered halos local

amphibolite facies transition

ore deposits vein