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Variations in eruptive style and depositional processes of Neoproterozoic terrestrial volcano-sedimentary successions in the Hamid area, North Eastern Desert, Egypt

1464-343X/$ - see front matter � 2013 Elsevier Ltd. All rights reserved.http://dx.doi.org/10.1016/j.jafrearsci.2013.02.009

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Ezz El Din Abdel Hakim Khalaf Cairo University, Faculty of Science, Geology Department, Giza, Egypt

a r t i c l e i n f o

Article history:Received 6 June 2012 Received in revised form 29 November 2012 Accepted 20 February 2013 Available online 4 March 2013

Keywords:Hamid area Volcano-sedimentary successions Facies analysis Vent and eruption column dynamics Basin-fill architecture

a b s t r a c t

Two contrasting Neoproterozoic volc ano-sedimentary successions of ca. 600 m thickness were recognized in the Hamid area, Northeastern Desert, Egypt. A lower Hamid succession consists of alluvial sediments, coherent lava flows, pyroclastic fall and flow deposits. An upper Hamid succession includes deposits from pyrocla stic density currents, sills, and dykes. Sedimentological studies at different scales in the Hamid area show a very comp lex interaction of fluvial, eruptive, and gravitational processes intime and space and thus provided meaningful insights into the evolution of the rift sedimen tary environments and the identification of different stages of effusive activity, explosive activity, and relative quies- cence, determining syn-eruptive and inter-eruptive rock units.

The volcano-se dimentary deposits of the study area can be ascribed to 14 facies and 7 facies associations: (1) basin-border alluvial fan, (2) mixed sandy fluvial braid plain, (3) bed-load-dominate d ephemeral lake, (4) lava flows and volc aniclastics, (5) pyroclastic fall deposits, (6) phreatomagmatic volcanic deposits, and (7) pyroclastic density current deposits. These systems are in part coeval and in part suc- ceed each other, forming five phases of basin evolution: (i) an opening phase including alluvial fan and valley flooding together with a lacustrine perio d, (ii) a phase of effusive and explosive volcanism (pulsa-tory phase), (iii) a phase of predominant explosive and deposition from base surges (collapsing phase),and (iv) a phase of caldera eruption and ignimbrite-forming processes (climactic phase). The facies architectures record a chan ge in volcanic activity from mainly phreatomagmatic eruptions, producing large volumes of lava flows and pyroclastics (pulsatory and collapsing phase), to highly explosive, pumice-rich plinian-type pyroclastic density current deposits (climactic phase). Hamid area is a small-volume volcano, however, its magma compositions, eruption styles, and inter-eruptive breaks suggest, that it closely resembles a volc anic archite cture commonly associated with large, composite volcanoe s.

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1. Introduction

Explosive volcanic eruption s commonl y result in thick accumulation of pyroclastic debris over wide areas and instantaneous modification of the topography and drainage networks around avolcano (Manville et al., 2009; Németh et al., 2012; Smith,1991). Large volumes of pyroclastic debris are also delivered rapidly to nearby sedimentary basins and lowlands, forming ‘‘syneruption’’ (Smith, 1991 ) or ‘‘posteruption’’ (Manville, 2001 )deposits during and immediately after major eruptions. The influence of explosive volcanism on sedimentation , especially innonmarine settings, has been therefore investigated in a number of studies (Kataoka, 2005; Mcclaughry and Gaylord, 2005; Palmer and Shawkey, 1997; Smith, 1991; Smith et al., 2002 ). The volcaniclastic deposits in a sedimentary basin can in turn provide

valuable information on the location or direction of the source vent and the changing eruption styles of a volcano that has since been removed by erosion or buried in the subsurfa ce (Kataokaet al., 2001; Lipman, 1976; Németh, 2010; Riggs and Busby-Spera (1990)). They also provide excellent chronostratigr aphic marker horizons that can help correlate isolated sedimentary sections and unravel complex sedimentary facies architectures (Karaogluand Helvaci, 2012; Knott et al., 2005 ).

Upward volcano growth, combined with tectonic uplift pro- duces prograding coarsenin g-upwards volcanicl astic successions to P1 km thick in intra-arc or arc-marginal basins. Accumulatio nsare thickest close to each volcano and fine and thin distally, giving rise to proximal–distal facies patterns (Kano and Takarada, 2007;Smith, 1991; Zernack et al., 2009 ). Proximal facies corresponding to the volcanic edifice comprise lava flows, autoclastic, and pyroclastic breccias and hypabyssal intrusions and pass laterally into medial apron associations of pyroclastic-flow, debris-avalanche,

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debris- and hyperconcentrate d flow (lahar) deposits and then distal facies association of braided to meandering fluvial system deposits, overbank alluvium, and interbedded tephras (Kataoka,2005; Maeno and Taniguchi, 2009 ). In weakly extensional , low-relief intra-arc basins, andesitic–dacitic stratovolcanoe s may co-exist with intermediate-s ilicic composition calderas, resulting in complex intercalations of volcaniclast ic rocks, lavas, and intrusions ofvarying composition and depositional environment (Smith et al.,1993; White and Robinson, 1992 ).

The Hamid area (600 km2) is located between 26�580 and27�100N latitude and 32�500 and 33�020E longitude (Fig. 1). Wadi Hamid area, known for its Pb-mining activity, lies to the west ofGabal Dokhan and is occupied by the Neoproteroz oic Dokhan- Hammamat volcano-sed imentary successions (Fig. 1). These successions are unconformably overlain from the west by Phanerozoic

Fig. 1. Simplified geologic map showing the distribution of the Dokhan Volcanics and Haresources including Abdel Rahman, 1996; Hassan and Hashad, 1990; Grothaus et al., 19Desert (NED), Central Eastern Desert (CED), and South Eastern Desert (SED) according to

sandston e of Nubia facies. Publications on the North Eastern Desert (NED), including the study area, are few in number if compare dwith the published works on the central and southern parts ofthe Eastern Desert (e.g., Dardir and Abu Zied, 1972; Ghanem et al., 1973; Khalaf, 1999; Mohamed et al., 2000 ). Previous works focused on the larger scale implication s involving geochemistr y,geotectonic setting, and very few radiogenic dating of the volcano-sed imentary successions in Egypt (Breitkreuz et al., 2010;Willis et al., 1988; Wilde and Youssef, 2002 ). Comparativ ely little is known about the internal lithofacies subdivision s of the eruptive sequence , volcanological , and sedimentologi cal facies analysis and the implication s for the eruption style, transport, and depositional processes . New field observations and a stratigraphy for the internal subdivision of the Neoproteroz oic volcano-sedimen tary successions in the Hamid area, based on regional mapping, extensive

mmamat Group in the North Eastern Desert, Egypt (data were collected from several 79; Wilde and Youssef, 2002 ). The approximate boundary between North Eastern Greiling et al. (1994).

76 E.E.D.A.H. Khalaf / Journal of African Earth Sciences 83 (2013) 74–103


detailed stratigraphi c logging, and petrographic analysis are pre- sented. Hence, this paper aims to describe a detailed sedimentological and volcanological facies analysis for the volcano-s edimentary successions and interpret them in terms of depositional systems and external controls on deposition, proposing a model for the stratigraphi c evolution of continen tal basins dominated by subaerial deposits. In the following, this study focuses on the most representative part of the N–S oriented volcano-sed imentary succession in the study area (Fig. 2), where the continuity and massive character of the exposure and the preservation of many original igneous and sedimentary features permit a nearly complete study ofits anatomy.

Fig. 2. Geological map of the Hamid area, North Eastern Desert, Egypt showing major volc1995).

2. Geological setting

Based on the field relations and radiogenic ages, Stern and Hedge (1985) geologically subdivided the Egyptian Eastern Desert into a threefold division namely: northern (NED), central (CED),and southern (SED) parts (Fig. 1). They mentioned that the oldest rocks are concentrated in the southern part, while the youngest units comprising the Dokhan volcanics, the Hammamat sediments,and young granites as well as dyke swarms are common in the northern part. The NED–CED boundary is generally represented by a straight, curved or complexly curvilinear N60 �E trending thrust (dipping NW) or dextral strike-slip fault (Greiling et al.,

anic units, major structures, and basement rock distribution (modified from Khamis,



1994; Stern and Hedge, 1985; Stern and Gottfried , 1986 ) (Fig. 1). Anumber of small-scale Late Neoproterozoi c sedimentar y Hamma- mat basins, including Hamid Basin formed in the NED of the Egyp- tian Eastern Desert (Fig. 1). Most of these basins were underwent arifting after the collision between East and West Gondwana and were later filled by kilometer-th ick successions of non-marine siliciclastic fluvio-lacustrine sediments together with abundant andesitic and rhyolitic Dokhan volcanic deposits. Based on a detailed structural study of the Eastern Desert of Egypt, Greilinget al. (1994) stated that this collision ended at 615–600 Ma and extensional collapse occurred within the 600–575 Ma time span,followed by transpressiona l tectonism along major shear zones until 530 Ma.

2.1. The Dokhan Volcanics and the Hammamat Group

The Dokhan Volcanics and the Hammamat Group crop out insmall terrestrial post-amalgamati on basins in the Eastern Desert.They have been recognized in the Eastern Desert for many years (El Ramly, 1972 ) and are described by an extensive literature (Abdel Rahman, 1996; El Gaby et al., 2002; El Sayed et al.,2004; Khalaf, 1995, 1999; Mohamed et al., 2000; Stern and Gottfried, 1986 ). The term ‘‘Dokhan Volcanics was coined after the type locality of Gabal Dokhan (Fig. 1). The Dokhan Volcanics typically include basaltic andesite, andesite, dacite, and rhyolite that some consider to be a bimodal suite (Khalaf, 1995; Mohamed et al., 2000; Stern and Gottfried, 1986 ) although this conclusion has been challenged (Eliwa et al., 2006 ). The Dokhan Volcanics constitute an almost unmetamor phosed thick succession that vary in thickness from basin to basin, ranging from a few tens of meters to 1300 m (Akaad and Noweir, 1980; El Ramly, 1972 ).Field relations reveal that the Dokhan Volcanics postdate the calc-alkalin e I-type Older granites, but predate or are associated with the A-type Younger granites (Akaad and Noweir, 1980; ElRamly, 1972 ). Stern and Gottfried (1986) argued that the Dokhan rhyolites west of Safaga are extrusive equivalents of the Younger granites. The Dokhan Volcanics are either subduction-relate d(Abdel Rahman, 1996; Hassan and Hashad, 1990 ), or extension- related after crustal thickening (Fritz et al., 1996; Stern, 1994 )or transitional between subduction and extension (El-Bialy,2010; Eliwa et al., 2006; Mohamed et al., 2000; Ressetar and Monrad, 1983 ). However, their undeformed character, their temporal and spatial association with post tectonic A-type granite,and their high Zr/Y ratio suggest that their emplaceme nt followed the cessation of subducti on in the Eastern Desert in an extensional, within-plate setting (Johnson et al., 2011; Khalaf, 1995 ).Recent SHRIMP zircon dating gave weighted U–Pb ages of593 ± 13 and 602 ± 9 Ma for two andesites from the Gabal Dokhan volcanics (Wilde and Youssef, 2000 ).

The Hammamat Group typically comprises greenish-gray silt- stone, lithic sandstone, and polymict conglomerate containing pebble-sized clasts of quartz, foliated granite, purple Dokhan type andesite, felsic volcanic rock, basalt, quartz porphyry, and undeformed pink granites. The Hammamat Group ranges from about 4000 m thick (Abd El-Wahed, 2009 ) to about 7500 m (Fritz and Messner, 1999 ). The proportions of the Dokhan Volcanics and the Hammamat Group vary from basin to basin. Some contain only volcanic rocks; others are entirely sedimentary; yet other basins contain both volcanic and sedimentary rocks. Because of the varied distribution s of these volcanic and sedimentary rocks and different relationship s in basins where both rock types occur, the stratigraphy of the Dokhan Volcanics and the Hammam at Group is debated.Some workers consider that the Hammamat Group underlies the Dokhan Volcanics (e.g., Stern and Hedge, 1985 ; Willis et al.,1988); others believe that the Hammam at overlies the Dokhan (e.g., Akaad and Noweir, 1980; Hassan and Hashad, 1990; El Ramly,

1972; Ries et al., 1983 ); yet others infer that the two interfingerand are essentially contemporane ous (El-Gaby et al., 1989; Eliwa et al., 2010; Khalaf, 2004; Ressetar and Monrad, 1983; Stern et al., 1984 ). The later conclusion is consistent with the overlap of whole rock Rb–Sr ages; 610–560 Ma for the Dokhan Volcanics,600–585 Ma for the Hammamat Group, and 600–550 Ma for the Younger granites (Beyth et al., 1994; Jarrar et al., 2003; Willis et al., 1988 ). Several authors concluded that the Hammamat Group is texturally immature, and reflects rapid uplift, erosion, transport,and deposition in alluvial fans and braided streams within a series of more or less isolated inter-montane basins (Khalaf, 2004; Grot- haus et al., 1979; Ries et al., 1983 ). Wilde and Youssef (2002) sug-gest that the Hammamat Group was deposited in a major fluvialsystem of continental proportio ns that linked the various basins,and possibly linked to similar successions in Sinai and Jordan;other workers infer that the group was deposited in isolated,fault-bounded basins (Abdeen and Greiling, 2005; Grothaus et al.,1979). These are variously classified as a foreland basin, (Fritzet al., 1996 ), intermontane basins, a strike-slip pull-apart basin,(Shalaby et al., 2006 ), and fault-bounded basins (Abdeen and Gre- iling, 2005 ). The range of inferred structural controls included thrusting , normal faulting, strike-slip faulting, N–S to NW–SEextension , and magmatic doming. Both the Hammam at and the Dokhan units were affected by rapid hinterland uplift at about 595–588 Ma (Fritz et al., 1996; Loizenbauer et al., 2001 ) and subse- quently intruded by the 585 Ma younger granite (Andresen et al.,2009). Stern et al. (1984) and Hassan and Hashad (1990) proposedthat the Hammamat and the Dokhan rocks were deposited and erupted in a down-faulted graben trending NE–SW. A problem isthat the Hammamat Group and the Dokhan Volcanics are definedon the basis of facies, whereas deposition took place in a dynamic setting around isolated volcanic centers and basin systems with different structura l controls and different ages (Breitkreuz et al.,2010), so that the two facies should not be expected to occur inthe same relative stratigraphic position in every basin.

2.2. Granitic rocks

Granitic rocks predominate in the Eastern Desert and Sinai and belong to two main stages in the geotectonic developmen t of the Egyptian Shield. The older stage, known as Old granitoids (900–650 Ma), comprises calc-alkalin e syn-tecto nic diorite, tonalite,trondhje mite, and granodiorite intrusions . These rocks occupy the southeastern and the eastern part of the study area that are tra- versed by a considerable amount of dyke swarms (Fig. 2). They underlie the Hammamat conglomerates with an erosion unconformity surface. The younger stage, known as Young granites (590–520 Ma), comprise s late - to post-tectonic granodiorite, granite and alkali granite (Jackson et al., 1984; Stern and Hedge, 1985;Stern and Gottfried, 1986 ). These granites intrude Older granites,Dokhan Volcanics, and Hammamat Group in the form of dykes,off shoots, and veins.

2.3. The Hamid Volcano-Sed imentary Succession s (HVSS)

The Hamid Basin is situated in the NED (Fig. 1), where a complete succession of Neoproteroz oic NE–SW-trending extension alvolcano-s edimentary successions are preserved directly overlying the basement. Basement units are made up of the igneous and metamor phic rocks, represented by gneissose granites and low grade metasedim ents and metavolcan ics (Abdel Rahman, 1996;Mohame d and El-Sayed, 2008; Osman et al., 2001 ). The volcano- sedimentary deposits in this area are known as the Dokhan Volcan- ics-Hamm amat Group and are mainly composed of lava flows,pyroclast ics, and fluvial siliciclastic sedimentary rocks, which rest unconfor mably upon older granites and are intruded by younger



granites and subvolcanic intrusions of sill- and dyke-type. These volcano-sed imentary successions are considered a part of the Egyptian Pan-African belts in the NED. U–Pb in zircon dating ofthe Dokhan Volcanics in the same area has yielded an age of616 ± 4 and 615 ± 5.4 Ma (Ediacaran age) (Breitkreuz et al., 2010 )and a 585 Ma age has been reported for the depositio n of the Ham- mamat sediments (Wilde and Youssef, 2000 ).

The Hamid Basin in NED is a pull-apart basin bounded by both strike-slip and normal faults on the basin margins (Fig. 2). This basin records the recurrence of regional extensional and strike-slip tectonics after the orogenic events that formed the Pan-African Belts (Stern, 1994 ). It is bounded by NE- or ENE-trending dextral strike slip faults along the northeastern and southwe stern margins and by NW-trending normal faults along the northweste rn and southeastern margins (Fig. 2). Wadi Umm Guruf represents one of the main wadies following this trend. The NE–SW trend is one of the predominant trends crossing the Precambrian terrains ofthe NED. One of the major shear zones belonging to such a trend is the Qena-Safaga shear zone that separates the NED from the CED (Fig. 1). This shear zone is characterized by a lack of the ophio- litic ultramafics, mélange, and the BIF-bearing metavolcan ics, inaddition to the absence of the northwest trending megashear s ofthe Najd system (El Gaby et al., 1990; Greiling et al., 1994 ). El-Gabyet al. (1984) suggested that this trend is responsible for the uplift ofthe large granitic plutons in this region and the uplift was followed by upheaval of the older rock units and expose the granites.

The Hamid Basin is divided into the southweste rn and northwestern subbasin s, separated by an intrabasinal strike-slip faults that trend NE (Fig. 2). Some of these faults are responsible for intense deformation, marked by the upturning of beds and a high density of minor faults in broad fault zones, but these structures rarely juxtapose different units of the HVSS. Near the major faults,the beds have higher dips, locally reaching 90�. The Hammamat Group that interfingers with mafic volcanics and pyroclastics of

Fig. 3. Synthetic S–N geological cross section (A–B) showing the distribution of the varioulocation and Table 1 and Fig. 6 for detailed description.

the Dokhan Volcanics crop out mostly along the southweste rnmargins of the southwestern blocks in the northeastern subbasin (Fig. 2). The Hammamat Group, constituting the lowermost stratigraphic unit of the basin, consists of polymictic conglomerate and fluvio-lacustrine deposits that show marked variations in stratal patterns along the basin-boundin g faults. The polymictic conglomerate therefore provides an opportun ity to investigate the initial basin-form ing processes and tectonics that cause the variation sin stratal patterns and architectur e. The strata of these sediments have mainly N50 � to 55�E strike and mostly dip toward the northwest with an average dip angle of 40� (Fig. 2). In the southweste rnsubbasin , transverse folds are recognized in different stratigrap hic levels near the NE-trending border faults (Fig. 2). The axes of the transverse folds, forming synclines and anticlines concurrentl y,are roughly perpendicul ar to the northeastern border faults, plung- ing 40–60� toward the NW. The axes of these folds generally trend N50�E–S50�W. In the northweste rn subbasin, the felsic Dokhan Volcanics and interbedded pyroclast ics with distinctiv e layering constitute the uppermost startigraphi c unit of the Hamid Basin.At Wadi Hamid Pb-mine location (Fig. 2), these volcanics have 20–25� tillting towards the east, remarking angular unconformity between the strata in southwe stern and northwestern subbasin (Fig. 3). NW–SE trending fault system controls the main course ofWadi Hamid and has both dextral and sinistral strike slip sense of movements but the later is predominant. In the northweste rnpart of the mapped area, the NW–SE faults form normal faults constituting a graben system enclosing the Nubia Sandstone in its downthrow n side (Khamis, 1995 ). The presence of normal faults bounding the north sub-basin to the north can be recognized inthe rock record (Fig. 4). The bedding measurements and lithological similarities suggest that the rock sequence in the northweste rnsubbasin was deformed and folded into asymmetrical syncline fold with an axis trending about 55�E. The northern limb of the syncline is dipping southeast (33–35�) and forms a thick stratified sequence

s facies in the Neoproterozoic Hamid volcano-sedimentary successions, see Fig. 2 for

Fig. 4. A. Panorama view of pyroclastic density current deposit (FA7) near the Pb-mine showing stretched ignimbritic rocks (Inw/Iw) with ramp structure, which overlie the bedded conglomerates (Cb), and in turn underlie by laminated vitric tuffs (Tv). B. Field sketch of the facies architecture in figure A. Note the younger granite (YG) and subvolcanic rhyolites as well as normal faults crosscut these deposits. Older granites (OG) lie at the base of these deposits in the left of the panorama.



of alternating felsic Dokhan Volcanics and volcanog enic sediments (Fig. 3). In spite of the intrabasinal faults, the strata of the Hamid basin could be correlated by the conglomerate and fluvio-lacus-trine deposits at the base, the lava flow and pyroclastics in the middle, and the felsic volcanics at the top.

3. Methods and terminology

The lithofacies types are defined on the basis of field characteristics including rock/clast type, texture, and depositional structures, following (McPhie et al., 1993 ) and (Cas et al., 2008 ).Primary volcanic (pyroclastic) deposits and volcaniclastic sediments, accordin g to volcanic stacking system, palaeo-deposit ional processes, and relative-timing with respect to a volcanic eruption were grouped into facies associations (FA), each representing different depositio nal environments via using process names based on Fisher and Schmincke (1984) and Cas and Wright (1987). Five profiles were measured in the field (Fig. 2). Boundaries between facies associations, are often gradation al. A lithofacie s code system,originally introduced by (Miall, 1978 ) and Rust (1978) to standard- ize descriptive classification of lithologies and sedimentary structures that are applicable to most fluvial deposits, was adapted inthis study. Table 1 lists the lithofacies codes of Miall, 1978 appro-priate for this study, and those used by Mathisen and Vondra (1983), Orton (1991), Smith (1987) and for some non-marine and shallow marine volcaniclast ic sediments. The facies associations include alluvial fan, fluvial, lacustrine, pyroclastic density current (PDC) and volcaniclast ic mass flow deposits. The sedimentary lithofacies associations are intercalated with different volcanic lithofacies varieties. The applicati on of the facies analysis concept on the both sedimentary and volcanic rocks of the Hamid area helped to define the overall evolution of the volcano-s edimentary basin (Baltazar and Zucchetti , 2007 ).

Following the Branney and Kokelaar (2002), the term ‘‘pyroclastic density current’’ has been applied for inflated mixtures of hot volcanic particles and gas that flow in variable concentrations and varying velocity along the ground. There are three end member types of pyroclastic density currents and their deposits, namely high density PDCs such as pumice-and- ash flows and block and- ash flows and low density PDCs, i.e., the pyroclastic surges, such

as base surges, ground surges and ash cloud surges. ‘‘Fiamme’’ isused here as a descriptive term for elongate lenses or domains ofthe same mineralogy, texture, and composition, which define apre-tecto nic foliation, and are separated by domains of different mineralogy, texture and composition (Bull and McPhie, 2007 ). Apre-tecto nic foliation is defined by the alignment and flatteningof fiamme parallel to bedding, known as ‘‘eutaxiti c texture’’ (cf.Ross and Smith, 1961 ). The author uses the term cooling unit has been applied as defined by Smith (1960) for the simultaneou sly cooled ignimbrite rock body. The ignimbritic cooling unit is subdivided into smaller units, called flow units, based on several characteristics such as grain size, grading, and abrupt boundaries asdetailed in the text. These units can be of two major types:source-co ntrolled and emplacement controlled flow units (Schmincke, 2004; Wilson and Hildreth, 2003 ).

Due to a lack of data, however , this different iation could not bedone in the study area. Due to silicification, alteration, and meta- morphism, the distinctio n between pyroclastic units and lava flowswas often difficult in the field. The focus of this study is on the least-alter ed outcrops where the pyroclastic units and the lava flows could be clearly distinguishe d on the basis of their accepted diagnost ic characterist ics (cf. Best and Christiansen, 1997; McPhie et al., 1993; Manley, 1996 ).

4. Stratigraphy and facies architectur e

The whole volcano-sedimen tary successions exhibit a large lateral and vertical variations in lithology and thickness in the S–Ntrend (Figs. 2 and 3). They begin in the south with the older granites and lava flows intercalated with volcanogeni c sediments and end in the north with the felsic pyroclastic rocks. Both successions show differences in lithology, degree of deformat ion, and ‘‘green- schist’’ metamorph ism. The sedimentary attributes especiall y inthe S–N trend are however, well preserved for high-resoluti on sed- imentolo gical analysis and primary contacts are locally preserved giving clues for stratigraphi c relationship s. Based on grain size,sedimentary structure s, and bedding features, combined with thin section studies, fourteen sedimentary facies were defined (Table 1).These facies were grouped into seven facies associations (FA) in the HVSS (Figs. 3 and 5 and Table 1) representing alluvial-fan deposits

Table 1Description and and interpretation of the facies in the Neoproterozo ic HVS succession, Hamid area, North Eastern Desert, Egypt.

Facies and facies code Characteristics Interpretation Facies association

Disorganized polymictic conglomerate (Cm)

Disorganized, massive, ill-sorted, clast-supported pebble-boulder conglomerate with poorly sorted,subrounded to subangular, clast size 5–50 cm.polymictic (granites, volcanics); granule-rich ormuddy sand matrix; laterally discontinuous and rare amalgamated with shallow scours; rare parallel alignment of elongate clasts; rare normal or inverse grading at base and top of beds

Debris flow,hyperconcentrated flow, orhigh-magnitude flood flows

Alluvial fan

Massive Sandstone (Sm) Faintly laminated, fining upward, cross bedded structure has been observed in some places.Moderately to poorly sorted, quartz-feldspar-rich arenites, occasionally volcanic clasts-rich with size up to 6 cm

Sandy braided rivers Fluvial braid plain

Gravely sandstone (Sg) Poorly sorted, faintly laminated or no structure,arkose to lithic arenite, occasionally volcanic clasts-rich in some varieties

Fluvial channel fill Fluvial braid plain

Laminated mudstone (Ml) Parallel lamination, with no desiccation cracks,silt-mud intercalation

Deep lacustrine deposits Lacustrine

Lava flow (Lf) Grayish-purple, massive, aphyric, with rubbly,vesicular bases and tops. Columnar joints iscommon. These individual sheets contain plagioclase-hornblende set in pilotaxitic to felted micolites, abundant opaque and epidote –richmatrix

Subaerial lava flows Coherent volcanic bodies

Bedded coarse tuffs (Tcb) Moderately to poorly sorted, mantle bedding,normal and inverse grading, vitric tuffs, crystals and lithics-rich, occasionally interbedded with volcaniclastic conglomerate (Cv) and hyaloclastic rocks (Hy)

Pyroclastic fall for Tcb and phreatomagmatic deposits for Cv and Hy

Pyroclastic fall and phreatomagmatic deposits

Lithophysae-rich ignimbrite (Il) Moderately welded, indurated and silicifiedlithophysa-rich. These nodules are quartz-filled,circular to star-shaped internal cavities exhibiting axiolitic, spherulitic, and pectinate texture

Welded pyroclastic flowdeposits (unconfined plain facies)

Pyroclastic density current deposits;spherulites with large central cavities,in-filled with quartz during diagenesis (cf. Holzhey, 1999, 2001 )

Massive breccias (Bm) Poorly sorted, clast-supported, no bedding and graded bedding, lithics-rich (granites, volcanics)

Co-ignimbritic breccias lag deposits

Pyroclastic density current deposits

Laminated vitric tuffs (Tv) Pervasive planar to low-angle truncating lamination with sigmoidal ripples/dunes. quartz- feldspar-rich with vitric clasts and occasionally accretionary lapilli. Soft-sediment deformation has been observed in some places

Phreatomagmatic fall of surge deposits

Base-surge of pyroclastic density current deposits; subaerial or very shallow water eruption and deposition

Bedded polymictic conglomerate (Cb) Indistinctly bedded or low-angle cross-stratifiedwith coarse to very coarse sand matrix, which iseither muddy or not; shallow to deep scours atbase; weak imbrication; indistinct normal and inverse grading in the upper and lower parts,clast-supported, or matrix-supported pebble- cobble conglomerate, angular to sub-rounded,clast size 10–30 cm, polymictic (granites,volcanics, pelites)

High-concentration flood flowon alluvial fans; or accretion of coalescing bars within braided channels

Alluvial fan

Nonwelded to welded ignimbrite (Inw/Iw)

Fiamme-rich, incipiently to densely welded,parataxitic or eutaxitic texture, porphyritic, lithic- rich at base to pumice and crystals-rich at top. Nolithophysae or nodules has been observed

Non welded to welded pyroclastic flow deposits (unconfined plain facies)

Hot pyroclastic density current deposits

Subvolcanic dykes/sills (Sd/SS) Concordant to discordant tabular rhyolitic bodies,porphyritic, quartz-K feldspar-rich

Intrusive coherent subvolcanic bodies

Coherent hypabyssal volcanic bodies



(FA1), fluvial braid plain deposits (FA2), lacustrine deposits (FA3),coherent volcanic bodies (FA4), pyroclastic fall deposits (FA5),phreatomagm atic volcanic complexes (FA6), and pyroclastic density current deposits (PDC, FA7). The first six facies associations form the lower part of the HVSS succession where there is inter- mixing of volcanic rocks with sediments, whereas the next seven facies form the upper part of the HVSS succession (Fig. 5). These successions, through which five geological profiles have been stratigraphi cally distinguished and measured, have a total thickness of 600 m (Fig. 6).

4.1. Facies analysis of the lower part of the HVSS

This succession occupies a large part of the mapped area in the southeastern subbasin, forming high mountainou s ridges with

irregular topography. The rocks of this succession crop out in amap-scal e anticline/syncl ine folds (Fig. 2). They comprise fluvialsiliciclast ic sediments together with abundant lava flows and volcaniclast ic rocks. These rocks are capped by lithophysae-rich ignimbri te (Il) of pyroclastic density current facies association.

4.1.1. FA1: Alluvial-fan deposits 4.1.1.1. Description. The basal part of FA1 is developed unconformably on the older granitoids (Fig. 7A). The FA1 crops out mostly along the southwe stern margins of the southwestern subbasin. Itis mostly composed of disorganized (Cm) to crudely stratified(Cb) conglomerates . The conglomerate units are less than 20 mthick and laterally persistent (>10 m in lateral extent). They are commonl y amalgamated into tens of meters thick units with sharp and erosional lower surfaces. They are composed of

Fig. 5. Generalized stratigraphic column of the Hamid area depicting two contrasting successions.



clast-suppo rted, dominantly pebble to boulder rounded gravel sized clasts in a poorly sorted (muddy) coarse sand matrix. The disorganized conglomer ates (Cm) are either nongraded or subtly in- versely graded in the basal part. The upper surfaces are generally irregular due to protruding clasts and are always gradational with the overlying sandstone–mudstone beds (Fig. 6, profile I). Large boulder clasts are floating in the middle part of beds and are mainly of granodioritic composition, but clasts of the adjacent and underlying volcanics and sediments are also present. Sand- stone beds (Sm), intercala ted within the conglomerates , are less than 0.5 m thick and have wedge or lenticular geometri es. They drape and fill the irregular space between the protruding clasts of the underlying conglomerates (Fig. 7A). The crudely stratifiedconglomerates (Cb) (Fig. 7B) are lenticula r and are similar to the disorganized conglomerates in texture but have a slightly scoured lower surface. They overlie the vitric tuffs (Tv) and occasional ly ap- pear as an intercala tion in the ignimbri te succession in profile V

(Fig. 6). Their upper contact with the ignimbrite beds is always dif- fused or gradational (Fig. 7C). Some domains of Cb display rounded outsized clasts of up to 1 m (Fig. 7D).

4.1.1.2. Interpretatio n. The disorganized conglomerates (Cm) with common floating or protruding outsized clasts and basal inverse grading suggest deposition from debris flows, whereas the crudely stratified conglomerates with slightly erosional lower surfaces (Cb)suggest turbulent flows or hypercon centrated flood flows (Nemecand Steel, 1984; Smith, 1986; Sohn et al., 1999 ). The crudely stratified conglomerates (Cb) with the undulating coarse-grained gravelly and laterally discontinuo us beds, indicate a channel-filling nature and an origin in a braided river system (Brierley et al.,1993; Todd, 1989 ). Poor developmen t of internal erosion surfaces suggests that the channels were rapidly filled with sediment during floods (Karcz, 1972 ). Rounded clastic components in debris flow deposits strongly suggest an elevated source (e.g., Schneider

Fig. 6. Five profiles showing the various stratigraphic successions in Hamid area from the south to the north. Note that the numbers refer to the different evolutionary phases characterizing the Hamid Basin.



and Fisher, 1998 ). The similarity in clast composition and size between these conglomer ates suggests either that the source region of the coarse debris remained unchanged, or that there was active intrabasinal erosion and resedimentation. The sandstone interbeds were probably deposited by sand-rich flood flows that accompanied the debris flows or by hypercon centrated flood flows during the waning stage of a mass-flow event (Nemec and Steel, 1984;Pierson and Scott, 1985; Smith, 1986 ). The overall characterist ics of the deposits suggest a mass-flow-dominated alluvial fan along the fault-controlle d basin margins (Benvenuti, 2003; Blair, 1999;Kim et al., 2009; Sohn and Son, 2004 ). Similar Hammam at-type conglomerates cropping out in the CED and NED have been interpreted as alluvial fan deposits (Grothaus et al., 1979 ). The signifi-cant presence of basement-d erived clasts (i.e., plutonic and volcanic rocks) coincides with the compositions expected for deposits which originated during periods of reduced explosive volcanic activity (Haughton, 1993; Riggs et al., 1997 ).

4.1.2. FA2: fluvial braid plain deposits 4.1.2.1. Descriptio n. Fluvial deposits occur in the form of laterally continuous beds and lenses in the southwe stern block throughout profiles I, II, and V, representi ng periods of stream and river reworking and re-establish ment (Fig. 6). These deposits are composed of reddish brown to red massive sandstones (Sm), and gravelly sandstones (Sg). They are commonly underlain and overlain byhorizontally stratified mudstone beds (MI) with gradational contact, showing an upward-fining trend (Fig. 6, profile I). The massive sandstone layers (Sm) are 0.1–0.5 m thick and consist of moder-

ately to poorly sorted, fine to very coarse sand. They are crudely stratified, low-angle cross-stratified, or massive (Fig. 7E). They are compose d almost exclusively of volcanic clasts, quartz, and feldspar crystals (Fig. 7F). Some rock fragments are also observed,varying from angular to sub-roun ded shape and maximum particle size up to 6 cm. Their matrix is fine sand and cemente d by iron oxides and clay minerals. The Sm facies resembles arenites, and lithic arenites. The gravelly sandstones (Sg) are characterized by alterna- tions of gravelly and sandy layers. The gravelly layers consist ofgranule- to pebble-si zed clasts, forming discontinuo us (3–4 mlong) stringers or streaks. Elongate clasts are generally aligned parallel to bedding plane or occasionally imbricated. The sandstone layers are 0.1–0.5 m thick and consist of moderately sorted, medium to very coarse sand. The gravelly sandstones (Sg) are mineral- ogically immature, varying from arkoses to lithic arkoses depending on the local provenance . The compositional ly immature Sg comprise mainly sub-rounded to subangular quartz and feldspars with occasionally volcanic and chert clasts (Fig. 7G).

4.1.2.2. Interpretatio n. This facies association is thought to represent the deposits of an aggradin g sandy braided-stream where poorly channelized, shallow stream flow conditions promote d the developmen t of low-relief bars (Kwon et al., 2011 ). Long-term aggradati on of the braidplain deposits without incised deep channels implies that high-sedi ment-load deposition occurred during aeustatic highstand (Smith, 1987 ). The indistinct alternation ofgravelly and sandy layers may have resulted from flow fluctuationor intermitten t supply of gravel during deposition (Rhee and

Fig. 7. Deposit features of facies associations (FA 1 to 3). (A) Disorganized pebble-boulder conglomerates (Cm) and intervening massive sandstone lens (Sm) of FA1 overlying the older granites with an erosional unconformity. Note sharp and concave-up base and diffuse and convex-up top of the conglomerates. (B) Crudely stratified pebble- toboulder conglomerates (Cb) and intervening massive sandston layers (Sm). Note the roundness of the clasts (arrows). (C) General view of horizontally stratifiedconglomerates (Cb). Note the gradational contact between the bedded conglomerates and overlying massive ignimbrite (Iiw/dw and the eroded, red-oxidized ignimbritic top near the Pb-mine. (D) Outsized jointed and rounded rhyolitic clast in bedded conglomerates (Cb). (E) Medium-to thick-bedded massive sandstones (Sm) are generally laterally persistent over a few tens of meters. The lower contact is mostly sharp with the underlying massive conglomerates (Cm). (F) Plane polarized light photomicrograph showing poorly sorted massive sandstone (Sm) composed of quartz (Qz), feldspar (fsp), and volcanic fragments (VF). (G) Cross polarized light photomicrograph displaying coarse-grained gravely sandstone containing subangular feldspar (fsp), quartz (Qz), rock fragments (RF), and chert clasts (Ch) set in recrystallized groundmass of quartz and feldspar. (H) Rock slab of mudstone showing parallel lamination. Note segregation of mud (black-colored) and silt (light-colored) in discrete lamina.



Chough, 1993 ). As there is little difference in velocities between flows transporting sand grains as suspended load and those transporting gravels as bedload, insignificant change in flow velocities can induce a considerable shear stress fluctuation, resulting innearly contemporane ous depositio n of sands and gravels (Walker,1975). Thus stratified sand layers and discontinuo us gravel layers might have formed at the same time on a stream bed without any significant flow fluctuations under upper flow regime condition (Wasson, 1977 ). The overall characteristics of this facies association are shallow unchanneliz ed braided-stream s, sand-bedload

regime deposits, similar to ancient arc-adjacen t alluvial plains induced by explosive volcanism (Palmer, 1997; Smith, 1987 ).

4.1.3. FA3: lacustrine deposits 4.1.3.1. Description. This facies association occurs dominantly atthe top of profile I (�25 m thick). The thickness of the FA3 increases up section with a decrease in the thickness of the sandstone beds in the amalgamated facies sequence. It is compose d ofmassive to laminated mudstones and silt layers intercalations.Parallel lamination ranges from mm to cm thick (Fig. 7H). The



mudstones are several meters in thickness and reddish gray color.The silt beds are laterally persistent and have distinct layer boundaries. Locally interlayered sandston es, less than 0.5 m thick, consist of well to moderate ly sorted, medium-to- fine sand. They have distinct lower and upper boundaries and lenticula r or wedge-shaped geometries.

4.1.3.2. Interpretatio n. This facies association is indicative of slow suspension settling of fine grains in a lacustrine environm ent.The laminate d mudston e facies that forms the coarsenin g- and thickening- upward siliciclastic succession possibly represents out- er fan deposits (Walker, 1984 ). The laminated character of the mudstone is the manifestation of a waning phase of the turbidity current deposition or storm surge. This facies may be interpreted as low-energy deep lacustrine basinal siliciclast ic turbidites (cf.Postma, 1986 ) related to the subaerial input system, below the wave base, where turbidity currents can be generated. The absence of desiccation cracks in the mudston e beds suggests that the flood-plains were always water-saturated (Elliott, 1974 ). The sandstone layers indicate intermitten t input of sand during floods, followed by water stagnation and subaerial exposure. This kind of lacustrine system is common in different volcanic environments (e.g., Renautet al., 1998, 2002 ).

4.1.4. FA4: coherent volcanic deposits 4.1.4.1. Description. The coherent volcanic deposits comprise all effusive or intrusive sub-volcani c rocks piercing the effusive phase of the volcano. Among these, lava flows and shallow intrusions are the most common ones (Table 1). The lava flows constitute a massive coherent core with relatively thin carapace of blocks formed breccias, which have rubbly, vesicular bases and tops. These volcanics are irregular rock bodies of tabular geometry with variable thickness, spanning from 3.0 to 20 m. The geometri es observed inthe lava flows have conspicu ously convex tops, while the bases are relatively planar (Fig. 8A). Hydrothermal alteration is pervasive throughout the lava flows, occurring along flow boundaries, cooling cracks, and tectonic fractures. These lava are greenish black to pale purple, strongly jointed, and porphyritic, including plagioclase phenocrysts and less abundant amphibo le and clinopyroxe nethat are semi-aligned in the matrix. This matrix shows a hyalophy- litic, sometimes trachytic texture, comprised of plagioclase micro- lites and an ore phase.

Sub-volcanic intrusions were observed in all the profiles with thicknesses of individua l bodies ranging from 1.0 to 10 m. They occur as NE-SW-tren ding necks and plugs with irregular sheets (Fig. 2) that cut the entire all successions within the study area and the neighbori ng granitoids. The most conspicuous feature isthe varicolored nature ranging from gray to red and deep purple.The sub-volca nic bodies have sharp contact with the vitric tuffs (Fig. 8B) and are commonly porphyriti c, blocky jointed, massive or flow-banded and contain quartz and K-feldspar phenocrysts (1–4 mm in size), as well as few and small-siz ed plagioclas e crystals with an abundant glassy matrix.

4.1.4.2. Interpretation. The effusive rock bodies were defined as lava flows, the intrusive rocks as plugs and necks (McPhie et al., 1993 ).The volcanic facies, represented by intermediate rocks (Table 2) are interpreted as viscous, slow moving blocky lava flows (Mueller,1991) as they are associate d with lava domes and coulées (Orton,1996). The massive to brecciated units display the attributes of acoherent flow in which autobrecciatio n processes were prevalen tand produced top and basal coarse breccias probably representing the proximal source of debris flow or volcanic debris avalanche deposits (e.g., Oregon Cascades; Bonnichsen and Kauffman, 1987;Fierstein et al., 2011 ).

The quartz–feldspar subvolcanic facies is an originally assemblage of coherent viscous lavas, lobes, sills/dykes, phenocryst-r ich,highly viscous silicic magma and its emplacement into the whole HVSS indicates that these rocks represent the youngest facies and the last intrusive event of the silicic volcanism. The facies can be characterized as subvolcani c, near-vent and non-explosi ve.Their presence is of special significance as they possibly represent feeder dykes implying that the volcanic edifice had been partly orcomplete ly constructed on the Hamid basement.

4.1.5. FA5: Pyroclastic fall deposits 4.1.5.1. Descriptio n. Pyroclastic fall deposits are represented by few outcrops in the study area where they were only found in profilesII, III, and V as single beds, up to 2.5 m thick, and each time in asimilar stratigraphi c position (Fig. 6). The fall units show characteristics indicative of atmospheric suspension settling including bedding (Fig. 8C), normal and inverse grading, and the absence oftractional features such as trough- and ripple-cr oss laminations.They exhibit grain sizes of fine to medium ash and are compose dof poorly-sor ted bedded coarse tuffs (Tcb) formed of volcanic and chert fragments, pumice clasts, and crystals set in altered fiam-me-rich matrix (Fig. 8D). In addition, lithic-rich horizons with dense clasts of andesitic compositi ons and older volcaniclast icfragments have been observed.

4.1.5.2. Interpretation. The FA5 is interpreted to be a distal plinian ash fall deposit, based on its pyroclast ic origin and characteristics such as bedding, normal grading, and the presence of juvenile fragments. This facies association was formed in the sustained eruption column that experienced variability in eruption intensity over time (Kataoka, 2005 ). The occurrence of the lithic clasts, pre- dominan tly lava flows sourced from the original edifice suggests ashallow fragmentation level. The presence of lithic clast-rich horizons at the base of graded intervals reflect periods of vent widening, increased magma discharge rate and/or conduit wall rock instabilit y (Bear et al., 2009 ). The fall deposits typically drape surfaces interpreted as low-gradient floodplains of a meander ing river system.

4.1.6. FA6: Phreatomagm atic volcanic complex deposits 4.1.6.1. Description. This facies association is observed in the upper and lower part of profiles II and V (c. 10 m) respectively . It is intercalated with the lava flows and pyroclastic deposits at various stratigrap hic levels. Their beds extend laterally for a few meters and show pinch-and-swel ling or lenticula r geometries. They are dominate d by volcanicl astic conglomerates (Cv) and hyaloclastic rocks. Individual Cv units are composed of medium sand- to granule-size volcanic lithics, crystals with hairline cracks and pumice fragments immersed in a sandy-sil ty matrix (Fig. 8E). These deposits are predominantl y integrated by fine-grained, moderate- towell-sort ed greenish tuffaceous sandstones. The hyaloclastics include pumice, vitrophyric (obsidian), and crystal fragments . Be- tween these fragments, the matrix is dominantly hyaline and perlite, showing devitrification structures and includes very fine-grained fragments of the same phenocrysts. Hydrothe rmal alteration is marked with a strong devitrification and obsidian flow ismarked by relict, perlite concentr ic angular glasses (Fig. 8F).

4.1.6.2. Interpretati on. These units are composed of volcaniclast icmaterial reworked by epiclastic processes (Cas and Wright,1987). The absence of tractional structures or diffusive stratifica-tion with gradational contacts suggests that Cv bodies result from the rapid, progressive aggradati on of hypercon centrated sheet flows (Smith, 1986 ), which represents the main transport and depositio n process involved within these units. The predominance of hyperconcentr ated sheet flow deposits suggests deposition in an

Fig. 8. Deposit features of facies associations (FA 4 to 7). (A) Coherent lava flows with convex-upper surface and planar-base. Note autobrecciated structure with vesicles-rich top. (B) A sharp intrusive contact between subvolcanic dykes and planar vitric tuffs (Tv). (C) Primary pyroclastic fall units display typical bedding structure. (D) Cross polarized light photomicrograph of bedded coarse tuffs (Tcb) consisting of feldspar (fsp), quartz (Qz), and volcanic fragments (VF) set in a devitrified fiamme-rich matrix. (E)Crossed polarized light photomicrograph showing volcaniclastic conglomerate (Cv) comprising quartz and feldspar crystals with hairline cracks as well as volcanic and pumice fragments immersed in a sandy-silty matrix. (F) Crossed polarized light photomicrograph showing remnants of relict concentric perlitic fractures with complete devitrification of microcrystalline quartz in the cracks (Qz). Note the decomposition of most probably mafic phenocrysts into epidote–opaque–quartz products. (G) Close-up of the fabric in the massive breccias (Bm) with polyhedral blocky clasts in fine-grained matrix. Note the vesicles-rich top in the juvenile clasts. (H) Laminated vitric tuffs (Tv)overlie younger granite. Note Pb–Zn mineralizations filling fractures.



alluvial context (sensu Blair and McPherson, 1994 ). An alternativ einterpretation of Cv as lahar or peperite would imply the presence of an elevated volcanic edifice from which peperites could originate (Németh and Martin, 2007 ). The provenance of this FA6,which mainly comprises pyroclastic and effusive volcanic clasts,indicates high affinity with the volcanic landscape, as it is almost unrelated to the country rocks (i.e., older granites). The presence of juvenile magma clasts within a sedimentary host and jig-saw fit textures that were mainly defined by developmen t of hairline cracks in crystals (Fig. 8E) support peperites (Doyle, 2000; Erkület al., 2006 ). Hyalocla stic rocks have abundan t volcanic glassy groundmass, and lower crystal contents that caused more fluidbehavior (e.g., McBirney and Murase, 1984 ). Relict perlitic glassy

fractures are commonly present in ancient, altered, formerly glassy volcanic rocks (Allen, 1988 ). These fractures are accentuated byrecrystal lization and devitrification of secondary microcrysta lline quartz in the cracks during hydrothe rmal stage.

4.2. Facies analysis of the upper part of the HVSS

The upper part of the HVSS is compose d of a series of pyroclastic density current deposits (FA7) and volcanogenic sediments that typically differ in fabric and sedimentary structures. These rocks are widespread in the northweste rn parts of the outcrops of the HVSS studied here. They cover a vast area, including the famous Hamid Pb-mine in its boundaries (Fig. 2). FA7 is compose d of four

Table 2Representative X-ray fluorescence (XRF) major and trace element analyses of lava flow, pyroclastic, and subvolcanic dyke facies of the HVSS. XRF analyses were performed on aPhilips PW 1510/20 spectromet er at the Nuclear Material Authorit y (NMA), Egypt using stand ard procedures. Major elements were analyzed on fused glass disks and trace elements on pressed power pellets.

Exposure area Wadi El Mesdar (32�500E/27�100N) Wadi Hamid near Pb-zinc mine (32�000E/27�100N)

Facies Lava flows Pyroclastics Pyroclastics Subvolcanic dyke

Sample No. H1 H2 H3 H4 H5 H6 H7 H8 H9 H10 H11 H12 H13 14Type A A A D RD RD R R D D R R R R

Major oxides element (Wt.%)SiO2 60.7 61.61 62.31 64.42 68.27 68.51 69.51 70.05 65.41 66.24 69.58 70.51 72.05 74.05 TiO 2 0.86 0.96 1.24 0.95 0.4 0.5 0.35 0.46 0.65 0.57 0.4 0.31 0.17 0.16 Al2O3 16.99 17.83 16.41 14.88 14.17 13.19 12.97 12.4 14.19 14.3 13.98 12.23 12.55 11.01 Fe2O3 7.38 6.19 6.79 5.98 3.99 5.59 4.99 3.99 4.6 3.79 3.79 2.64 2.39 1.99 MgO 2.3 2.3 1.41 2 0.9 1.4 0.76 0.6 2.3 1.4 0.6 1.2 0.8 0.4 CaO 4.4 3.4 3.4 4.2 2.96 1.68 1.68 1.12 4.11 3.39 1.4 2.24 2.24 1.12 Na2O 3.73 3.74 4.34 4.29 4.83 4.28 4.2 4.75 4.95 4.54 4.76 4.75 4.53 4.37 K2O 1.8 1.8 2.09 2.21 3.34 3.22 4.17 4.71 3.39 3.82 4.19 4.97 4.02 5.35 P2O5 0.62 0.71 0.4 0.35 0.14 0.12 0.12 0.08 0.11 0.11 0.05 0.09 0.26 0.07 LOI 0.84 1.33 0.74 0.72 0.77 0.93 0.93 0.86 0.32 1.14 0.72 1 0.92 0.49

Total 99.62 99.87 99.13 100 99.77 99.42 99.68 99.02 100.03 99.3 99.47 99.94 99.93 99.01

Trace elements (ppm)V 129 120 16 143 68 91 43 46 14 60 16 21 25 7Cr 13 30 23 20 13 16 12 11 27 15 33 34 26 17Ni 18 19 10 27 7 14 3 4 4 7 6 5 4 3Cu 31 25 10 38 38 48 28 29 11 25 12 10 12 33Zn 99 52 46 80 53 71 97 100 83 50 81 51 91 43Pb 28 18 24 34 52 34 32 32 23 45 21 18 112 137 Ga 21 21 26 19 21 16 22 21 23 21 24 18 20 19Rb 47 65 71 37 66 40 45 50 147 60 145 108 165 145 Sr 674 613 427 899 503 406 217 267 176 553 75 172 88 149 Ba 692 419 319 936 534 882 871 978 229 500 138 410 248 205 Zr 181 209 278 131 333 194 275 307 355 233 557 202 192 289 Y 14 27 37 11 20 16 23 27 59 20 59 23 32 24Nb 5 8 10 3 9 5 7 9 28 9 30 10 12 10

Code Fe2O3: Total iron as ferric oxides. LOI: Loss on ignition. Rock type: A = Andesite, D = Dacite, RD = Rhyodacite, R = Rhyolite.



lithofacies (Table 1) that grade laterally and vertically into each other. These facies involve: (1) massive volcanic breccias (Bm),(2) laminated vitric tuffs (Tv), (3) lithophysae-rich ignimbrite (Il),and (4) nonwelded to welded ignimbrite (Inw/Iw).

4.2.1. Massive volcanic breccia facies (Bm)4.2.1.1. Description. This lithofacies is only observed in the northern region, near the Pb-mine (profile IV, Fig. 6). It is composed of angular to subangular clasts in a pinkish red matrix of fine ash without displaying jigsaw-fit texture. They occur in laterally sheets with planar bases and eroded tops. Average thicknesses of single units are ca. 2 m. However , vertical amalgamati on surfaces between stacked units are rarely visible, resulting in deposits up to 5 mthick without any visible bounding surfaces. The deposits show aweak-bedde d, no signs of grading, and poorly-sorted , loose compaction. The clast sizes usually range from pebbles to cobbles,not exceedin g diameters of 20 cm. However , single outsized clasts of 2 m in diameter have been observed. These clasts encompass immature lensoid rhyolite boulders (probably juvenile), andesitic,and rarely granitic fragments (Fig. 8G). The matrix of the deposits is commonl y composed of lithic and pumice fragments, crystals and glass shards, showing significant alteration to clay minerals.The fragments do not show any alignment within the matrix.

4.2.1.2. Interpretati on. Massive breccias are interpreted to be a co- ignimbrite lag breccias, similar to the lithic breccias that occur within the pyroclast ic flow deposits (Druitt et al., 1999; Sparks and Wilson, 1990 ). These breccias are thought to represent a proximal ignimbrite facies, formed by near vent deposition and segregation of dense lithic clasts during pyroclastic flow emplaceme nt(Druitt and Sparks, 1982 ). Co-ignimbr ite lag breccias are defined

by Druitt and Sparks (1982) and Walker (1985) as thick (1–20 m), coarse-grained (up to meter size blocks), stratified, poorly sorted, variably fines-depleted, lithic and dense juvenile clast-rich deposits formed in the deflation zone of a collapsing eruption column (1–8 km from vent) at the onset of a caldera collapse. The initial eruption column collapsed when its bulk density became greater than the ambient atmosph ere. In the deflation zone, the expanded mixture moved outward, deflating as gases escaped, forming a pyroclast ic flow (Walker and McBroome, 1983). The compositi on and abundance of the dominant lithic clast types ofthis lithofacie s, i.e., andesitic and rhyolitic lava clasts suggest that the original volcanic edifice was the source for majority of the lithic clasts ejected into the eruption column. The coarse, crudely stratified, poorly sorted nature, and depletion in fine ash exhibited bymassive breccias, all suggest depositio n from an expanded system in a layer-upon-layer fashion. The low fine-ash fraction was likely caused by elutriation of fines from the deflation zone at the site oferuption column collapse. As the deflating mixture (gas + particles )expanded outwards from the site of collapse, gases escaped carrying off fine vitric material and depleting the remaining mixture.Subseque ntly fines may have also been lost during flow by elutriation caused by upflow of gases during transport (Cas and Wright,1987; Druitt and Bacon, 1985).

4.2.2. laminated vitric tuffs facies (Tv)4.2.2.1. Descriptio n. This lithofacies appears in sequences of different thickness, up to 40 m thick and crops out in profiles III, IV and Vnear the base of the Pb-mine in the north (Fig. 6). It exhibits a sharp contact with the underlying granites (Fig. 8H). Its contact with the surroundi ng rocks is poorly exposed because of weathered outcrops, but appears to be an undulatory surface that is discordant



with the bedding in the overlying ignimbrites (Fig. 9A). In addition to the basal erosional discordances associated with this facies,there is also a discordant surface within this facies. This surface comprises a relatively steep (35–45�), sharp, planar surface that truncates the underlying units and is draped by succeedi ng deposits. In outcrop, the vitric tuffs are light brown to light purple in col- or and are characteri zed by planar to low-angle lamination (Fig. 9B). The deposits of this facies infill topograp hic lows in U-shaped gullies (Fig. 9C). These beds display well-developed antidune structures indicating the direction of the transport. At least two different dune bedding types are observed within the vitric tuffs; (1) type d (Cole, 1991 ; Fig. 9D) or type II (chute and pool;Schmincke et al., 1973 ); (2) type b (Cole, 1991 ; Fig. 9E) or type III(Schmincke et al., 1973 ) antidunes. These dune structures commonly have wavelengths ranging between 1.0–3.0 m with ampli- tudes of 9.0–13 cm. In some areas, syn-depositi onal structures (such as cross-strati fication, Fig. 10A) are common in these deposits. Soft-sedim ent plastic deformat ion structures, like convolutes and contorted beds with slumping have been observed in these deposits (Fig. 10B).

Under the microscop e, the Tv facies is lithic-poor, poorly sorted and matrix-rich. It has a framework of feldspars, quartz, pumice and occasional vitric fragments. The quartz and feldspar grains are subangular to subround ed, and the pumice fragments have ragged boundari es. The feldspars are variably altered to sericite with associated haematite and chlorite, and the pumice fragments are completely replaced by the same minerals. Occasionally, the accretionary lapilli are scattered within a matrix of structureless fineash. These lapilli are spheroidal, less commonl y discoidal, 1 mmto 4 cm in size. These lapilli have been subjected to diagenetic and metamorph ic changes which make a determination of the original morphology of the fine ash difficult that composed them.However, well preserved devitrified shards containing a quartz- filled vesicles occur in the core of some accretionary lapilli. The matrix surrounding them consists of fine-grained quartz and sericite that are probably products of devitrified and recrystal lized finevitric ash.

4.2.2.2. Interpretation. Tv deposits that display structure s indicative of lateral transport (i.e., presence of ripples, dunes, low-angle cross bedding, etc.) have been interpreted as being deposited by base surges (Bull and Cas, 2000 ). Additional features consistent with abase-surge origin recognized in this study include: the U-shaped gully (Fig. 9C); the discordant surfaces (Fig. 9A); poor sorting;and the abundance of ash matrix in the tractional deposited units.Type ‘‘b’’, and ‘‘d’’ dunes of Cole (1991) were identified within the Tv beds. Type ‘‘b’’ dunes are asymmetric al and built from planar beds with progressive ly steepening layers. Sand-wave crests always migrate in the downstream direction, therefore they are pro-gressive. Similar structures are termed as ‘‘type III’’ dune structures by Schmincke et al. (1973). Type ‘‘d’’ sand waves comprise steeply dipping stoss-side layers sigmoida l in shape. They migrate in the upstream direction and may be considered as regressive. They are similar to the ‘‘type II’’ dune or ‘‘chute and pool’’ structure ofSchmincke et al. (1973). Close relation of both progressive and regressive types within the same deposit can be explained by the pulsatory nature and change in flow regime of the base surges, just like the dry and wet surges in Roccamonfina and Sugarloaf Moun- tain (Cole, 1991 ). Ash-dom inated Tv base surge facies displays lateral/vertical change in bedding style, i.e., from planar to wavy orfrom planar to dune beds. Units with similar facies change, and climbing dunes are interpreted to occur due to a decrease in suspended-load transport rate and/or an increase in bedload transport rate and the decrease of the surge flow power (Sohn and Chough,1989). Their occurrence within more massive parts of Tv suggests a relatively high particle concentratio n within the depositional

boundary layer of the proximal-m edial pyroclastic density current (cf. Druitt, 1992 ).

The relatively fine grain sizes (mainly coarse-ash), poor sorting,ash-cover ed pumice clasts, and vitric shards within facies Tv, col- lectively indicate that it was derived from a explosive phreatomagmatic fragmentation , similar to other documented phreatom agmatic and phreatoplini an pyroclastic density current deposits (Cas and Wright, 1987; Morrissey et al., 2000; Gertisser et al., 2010; Vespermann and Schmincke, 2000 ). Observed characteristics of the TV base surge tephra are typical for maar tephra such as: dune and antidune type bedding; thinning of individua lbeds; occurrence of accretionary lapilli; and plastic deformation of beds (Lorenz, 1985 ). Gevrek and Kazancı (2000) classified the maars as ‘‘small’’, ‘‘medium’’ and ‘‘large’’ based on their diameters being less than 500 m, 500–1000 m, and more than 1000 m,respectively . According to their classification scheme, the vitric tuffs can be considered as a small maar based on its restricted occurrences and pinching-ou t geometry.

4.2.3. Lithophysae-rich ignimbrite (Il)4.2.3.1. Description. The lithophysae- rich ignimbri te facies (Il) lays on top of lava flows intercala ted with volcanicl astic rocks, with no intermediate paleosol or deposit. This facies occurs in profilesII and III (Fig. 6). These deposits are massive, poorly sorted, with average grain sizes ranging from coarse ash to bombs and blocks with diameters up to 10 cm. The rocks of this facies form massive or flow-banded sheets, up to 30 m thick, compose d of pumice lapilli and lithic fragments supported in a matrix of devitrified vitric ash and shattered crystals. Welding compaction in this facies can be recognized by the presence of large nodules (1–10 cm in diameter) hosted in the rest more deeply weathered rock mass. Most ofthe nodules are silicified, indurated, flattened and elongated bydeformat ion. Some nodules show spheroid al concentr ic fractures resembli ng large-scale perlitic texture (macro-perlite, Fig. 10C).Each nodule contains a quartz-filled, circular and star-shap edinternal cavity. This cavity displays fine axiolites (acicular crystals)growing inward from their walls (Fig. 10D) and defining a pectinate texture (McArthur et al., 1998 ).

4.2.3.2. Interpretati on. The characteristics of the nodules suggest that these are lithophysae preserved in weathered ignimbri te. Lith- ophysae are large spherulites that happened to develop central gas cavities wherever the external pressure was low enough to allow dissolved gas to expand a cavity. During diagenesis, aqueous solu- tions of silica are carried to these cavities and deposited as agate,quartz, and sometimes common opal or jasper. This process ofsilicification was interpreted as being caused by vapor-phase alteration (sensu Cas and Wright, 1987; Streck and Grunder, 1995 ) and/ or deuteric alteration in the syn-volcanic stages. Lithophy sae, filledwith chalcedony and quartz, are quite common in the periphery ofboth subaqueous rhyolite lobes (cf. Kano et al., 1991 ) and subaerial lava domes (cf. Holzhey, 1999, 2001 ). High temperature crystallization domains include lithophys ae, spherulite and pectinate textures (McArthur et al., 1998 ). These units are deposited from pyroclast ic density currents with fluid escape-dominate d flow-boundary zones (Branney and Kokelaar, 2002 ). The units of pyroclastic accumulation constitute the record of explosive volcanic activity during the evolution of the lower succession.

4.2.4. Nonwelded to welded ignimbrite facies (Inw/Iw)4.2.4.1. Description. This facies is topograp hically and stratigraphi- cally higher than facies Tv deposited in the upper succession ofHVSS. Thus, this facies is the youngest primary volcanic deposit in the studied succession. Its deposits (see profiles IV and V,Fig. 6) partly drape the pre-eruptio n topography, thickenin g in val- leys and depressions. Their lower bounding surfaces are flat or

Fig. 9. Deposit features of facies associations (FA7). (A) An angular unconformity formed by rapid draping and burial of volcanotectonically disturbed vitric tuffs (dip to the right) by rheomorphic ignimbrites (dip vertically �90�). Note the welding zones in ignimbritic rocks, which range from incipient welded to densely welded at the top. (B)Planar to low-angle lamination in vitric tuffs. (C) U-shaped erosive gully infilled by deposits of the surge facies (vitric tuffs). (D) Chute and pool/type II (Schmincke et al., 1973 )or type d (Cole, 1991 ) dune structure. Note the flattened (?) tectonically compacted sigmoidal ripple/dune (arrowed). (E) Type III (Schmincke et al., 1973 ) or type b (Cole,1991) dune structure.



reflect the palaeosurfa ce, their tops are mostly eroded. In fact, their contact with the underlying facies Tv is marked by an unconformity surface (Fig. 9A). These rocks were truncated by NW–SW-trending faults (Fig. 4). The deposits of this facies occur as single units or as a series of stacked beds. They are always internally massive, poorly sorted and lack internal stratification. The color of the ignimbritic samples, varying from creamy beige to reddish purple,reflects the changes in the degree of welding and alteration. Thick- nesses of single units can vary from 2.0 to 70 m with an average of50 m. The most prominent feature of these facies types is the

occurrence of well-developed ramp structures (Fig. 4) consisting of sets of flow laminations defined by the alignment of juvenile fragments and variations in color and degree in vesiculari ty. The main flow directions are E–W, and SW to WSW. Most of the hydrothermal alteration in these rock types occurs as fracture and cavity fillings. Four grades of welding were defined from observati on ofspecimens at outcrop and their petrographi c characteri stics (Fig. 11). Grades of welding are defined as: a = incipiently welded,b = moderately welded, c = highly dense welded, and d = partially welded. The degree of welding varies from incipiently welded of

Fig. 10. Deposit features of facies associations (FA7). (A) Small-scale cross stratification in vitric tuffs. (B) Vitric tuffs show penecontemporaneous slumping and folding reflecting soft sediment deformation. (C) Plane polarized light photomicrograph showing circular and half moon shapes for lithophysae quartz-filled nodules (Qz). Note the flow folding of the devitrified fiamme (arrow). (D) Plane polarized light photomicrograph demonstrating pectinate texture defined by fine axiolites growing inwards from the walls of a juvenile pyroclast containing a quartz-filled vesicle (Qz). Note the reddish oxidation of the feldspar (fsp). (E&F) Scanning electron micro-images of juvenile clasts are highly vesicular with sub-spherical to smoothly irregular-shaped bubbles. (G) Plane polarized light photomicrograph exhibiting vesicles-rich glassy volcanic fragment (VF)enclosed in welded ignimbritic rocks. (H) Plane polarized light photomicrograph showing flow banding and folding in welded ignimbritic rocks.



rank II (Peterson, 1979; Streck and Grunder, 1995 ) through partially and moderate ly welded of rank III and IV (Quane and Russell,2005; Wilson and Hildreth, 2003 ) to the densely welded of rank V(Peterson, 1979; Branney and Kokelaar , 1992; Wilson and Hildreth,2003). These rocks exhibit columnar jointing and vary in welding degree both vertically and laterally, with welding most intense closest to the inferred source, decreasing rapidly downwind. Weld- ing facies are completely gradational and the entire deposit shows a gradual decrease in compaction features upwards. The welding isdependant upon compaction of the tuff components upon or soon after deposition and the temperature of particles and gases upon deposition (cf. Ross and Smith, 1961 ).

The basal facies of the studied ignimbrite is represented by a15–30 cm incipiently welded vitrophyr e (Fig. 11). The most common facies comprises 20–30-cm long vitrophyric sections of massive, perlitic, largely devitrified glass. Strongly flattened fiammetextures are locally preserved in a more altered matrix. Alkali feldspars show truncation joints due to the vertical compacti on. The basal boundary is accompanied by typically thin, 30 cm long,low-angl e domains of glassy vitrophyre penetrating the devitrifiedwelded facies. A 5–10 m thick, dense, moderate ly to densely welded deposit overlies the basal vitrophyre (Fig. 11). The juvenile fragments in these zones occur in many varieties consisting of well preserved , mm to cm-long, dense or pumiceo us lapilli, embedde d

Fig. 11. Schematic stratigraphic column of the upper ignimbritic cooling unit and its major welding facies (facies Inw/Iw). Left: stratigraphic column at Pb-mine locality; the width of the log corresponds to the average grain size, in mm. Photomicrographs are in plane polarized light.



in a matrix of recrystallized ash and glass shards. These particles show blocky shapes with abundan t irregular, spherical to sub- spherical vesicles (Fig. 10E). Individual pumice fragments exhibit very intense vesiculat ion and can be up to 12 cm long in pumice concentratio n zones (Fig. 10F). Lithoclasts and alkali feldspars be- haved as solids in a ductile matrix of deformed and compacted fiamme and finer-grained components. Lithoclas ts are rimmed byrotational ductitle shear marks and pressure shadows with asymmetric r-type tails of glass shards trailing away from lithic clasts (Fig. 11) (Schminck e and Swanson , 1967 ), whereas fiamme are preferentiall y flattened. White secondary quartz commonl y re- places porosity in small fiamme of this facies. Most of the welded facies displays subvertic al fractures normal to bedding; the

factures are up to 1 m wide and 20 m long. Highly welded and compacted ignimbrite with a eutaxitic texture dominates this 5–10 m thick facies (Fig. 11). The main feature of this facies is the rapid up-section increase in the degree of flattening and welding. The parataxiti c and eutaxitic texture is well-defined. Fiamme and matrix can be easily distinguishe d from each other. Spheruliti c texture in some fiamme also occurs. The deformed pumice and fiammerange from a few centimeters to a few tens of centimeters in long axis length. The shape of fiamme is blocky or lenticular with cuspate or ragged edges. The majority of fiamme shows bi-cuspate and tricuspate, plate-like or crescent-shape d morphologies , with at least one curved face that is demonstrably part of a burst bubble wall. Major lithic components of these welded zones are



subangular volcanic clasts composed of andesite and basalt, inwhich its size is up to 30 cm in long-axis length (Figs. 10G and 11). Micro-scale rheomorphic flow folds (harmonic, disharmoni c,sigmoidal, and isoclinals (Fig. 10H) and brittle shears within the steep deformation zones are observed in this facies. The top ofthe cooling units consist of poorly welded and weakly compacted ignimbrite (�5 m thick) with partial erosion (Fig. 11). Clasts are randomly oriented and deformed in the porous deposit. Highly welded and flattened ignimbritic fiamme consist mainly of microcrystalline alkali feldspars and quartz, which were probably formed during devitrification and post-emplac ement vapor phase crystallization.

4.2.4.2. Interpretati on. This lithofacies is interpreted as ash-flow deposit and is described by many authors as the most common ignimbrite lithofacies (e.g., Branney and Kokelaar, 2002; Boyce and Gertisser, 2012 ). It is relatively thick and poorly sorted, implying that the current comprised a relatively high particle concentratio nand in this respect, did not resemble a typical pyroclast ic surge.The columnar jointing, high consolidation and welding is thought to be highly related to the emplaceme nt thickness and conseque ntly with the cooling rate of the unit in the valley (Rotolo et al., 2013 ). The characterist ic red color of the ignimbri tic facies is interpreted as due to a vapor phase alteration that led to the high temperature oxidation during devitrification of the matrix and pumice fragments ,transforming them into a secondary mineral assemblage s. The predominance of vesicles- enriched juvenile fragments (Fig. 10E and F)suggests that the magma was vesiculated and fragmented during the eruption as a result of exsolution of magmati c volatiles. Juvenile pyroclast characterist ics, especially vesicularity, are very similar inPlinian-style eruptions (Allen and McPhie, 2009). Flow banding and folding are primary textures, characterist ic of rhyolites, and reflect laminar and folding flowage (Cas and Wright, 1987 ; Scutteret al., 1998 ). These syn-depositi onal deformat ion structure s show that the facies (Inw/Iw) was produced by particle to particle agglutination. The deposits usually show a normal coarse-tail grading ofthe lithic clasts while the pumice clasts show a reverse grading (Fig. 11). Basal normal grading of this facies suggests the action tur- bulence. On the other hand, the top inverse grading suggests constrained particle motion and rapid deposition from highly concentrated and laminar flow with minimal turbulent or granular shear close to the lower flow boundary (Branney and Kokelaar ,2002; Todd, 1996 ). The coexisten ce of these apparently contradic- tory depositional features in the same deposits is interpreted asthe result of nonuniformity and spatial–temporal variation s of the properties of the pyroclast ic density current, such as the degree ofturbulence, particle concentratio n, and suspended -load fallout rate (Branney and Kokelaar, 2002 ; Vazquez and Ort, 2006 ). The change in abundance of clast types from lithics-dominated to pumice-d om- inated suggests changing clast supply conditions at vent, possibly reflecting tapping of progressively deeper levels of the magma feeder system. The change in abundance of clast types from lithics- dominated to pumice-domin ated suggests changing clast supply conditions at vent, possibly reflecting tapping of progressively deeper levels of the magma feeder system. The presence of the andesitic lithics (Figs. 10G and 11) imply that the tuffs containing them were deposited on a volcanic substrate, which have been picked up by the pyroclastic flows (Buesch, 1991 ). The welding process involves sin- tering, compaction, and flattening of pyroclasts in response to suffi-cient load stress under high temperature s (Cas and Wright, 1987;Quane and Russell, 2005 ). The rapid emplaceme nt of the facies (Inw/Iw) and subsequent burial of the facies (Tv) would have provided sufficient load to promote flattening and compacti on of the juvenile clasts (Fig. 11). The occurrence of partial welding also has implication s for the temperat ure condition s. The minimum temperature required for welding is between 500 �C and 650 �C (Quane and

Russell, 2005 ) implying that this range of temperature s represents aminimum estimate for the emplaceme nt temperature of these deposits.

5. Geochemistr y

5.1. Bulk rock composit ions

Bulk rock analyses for major and trace elements were carried out for lava flow, pyroclastics, and subvolcanic dykes from the two successions (Table 2). Geochemical classification using K2Ovs. silica diagram (Peccerillo and Taylor, 1976 ) shows that the Ha- mid volcanics form a unimodal volcanic suite and have a moderate range of SiO 2 (60–74 wt.%) ranging from mafic (andesite) to felsic (dacite and rhyolite) members (Fig. 12A). As a function of SiO 2 con-tent, the variation of major and trace elements (Table 2) show good negative correlation (i.e., TiO 2, Al2O3, Fe2O3, MgO, Sr), whereas the total alkalics (Na2O + K2O) and high field strength elements (Zr, Y,Nb) show positive correlation regarded with SiO 2 content. On the Zr/TiO2 vs. Nb/Y classification diagram (Winchester and Floyd,1977) there is a clear distinction between the lava flow and pyroclastics that plot in the andesite, dacite/rhyodac ite and rhyolite field, indicating a moderate degree of differentiation in the magma source (Fig. 12B), whether as a fractional crystallization product oras melted continental crust. These rocks are relatively depleted incompatib le transition elements , enriched in LILEs (Rb, Ba, K) relative to high field strength elements (Nb, Zr, P, Ti) and show strong affinity to medium to high K calc-alkalin e subduction- related rocks (Table 2, Fig. 12A). This is further supported by low Nb/Y ratio (<1.0, Fig. 12B). The low Nb content in these rocks (Table 2) may be attributed to magmati c systems linked to subduction zone processes or more generally to crustal processes (Duncan, 1987; Hof- mann, 1988 ). On a Ti vs. Zr plot, the data for all volcanic facies lie along a linear trend, with a Ti/Zr ratio of �8; only the subvolcanic samples show some spread (Fig. 12C). This indicates that Ti and Zrwere indeed immobile during hydrothermal alteration and meta- morphism and that the pyroclastics, the lava flow, and the subvolcanic dyke rocks probably originate d from the same magma source. The deviation of the subvolcanics from the linear trend may be interpreted to reflect the consecutive tapping of a texturally homogeneous but chemically zoned magma (Paulick and McPhie, 1999 ). On the Nb vs. Y tectonic discrimination diagram (Pearce et al., 1984 ) the majority of data points plot in the volcanic arc granites field (Fig. 12D). Only two samples of facies Inw/Iw plot in the within plate granite field.

The arc-related geochemica l characterist ics for the Hamid volcanics suggest that subduction continued until near the end ofthe collision stage beneath the Dokhan Volcanics province or that the mantle source retained a geochemica l memory of subduction-related modifications (Khalaf, 2012 ). The temporal and spatial association of the Dokhan Volcanics with the Hammam at sediments and post tectonic A-type granite as well as their high Zr/Y suggest that their emplaceme nt followed the cessation of subduction in the Eastern Desert in an extensional, within-plate setting (El-Bialy, 2010; Khalaf, 2012; Eliwa et al., 2006 ).

5.1. Variations in the stratigraphy

The major elements like Fe2O3, TiO 2, and P2O5 decrease in concentration upward in the stratigraphy, whereas K2O show an increase (Fig. 13). The enrichment factor (most evolved/less evolved concentratio n) of bulk rock major elements broadly follows the stratigrap hy and remains very low (0.17 for TiO 2, 0.26 for Fe2O3, 0.11 for P2O5), except for K2O (2.97 for K2O). However,the concentr ation of many trace elements shows a strongly

Fig. 12. The geochemical characteristics for lava flows and pyroclastics in Hamid area. (A) On K2O vs. SiO 2 (wt.%) classification diagram (after Peccerillo and Taylor, 1976 ), the volcanic samples plot in high K-calc-alkaline andesite-rhyolite field. (B) On the Zr/Ti vs. Nb/Y classification diagram (after Winchester and Floyd, 1977 ), the volcanic samples plot in andesite-rhyolite field. All samples are subalkaline, based on the Nb/Y ratio (<1.0). (C) On the Ti vs. Zr diagram most samples have a Ti/Zr ratio of – 8 except subvolcanic dyke samples deviate towards higher Ti/Zr values. (D) On the Nb vs. Y tectonic discrimination diagram (after Pearce et al., 1984 ), a volcanic arc character is indicated for the volcanic rocks in two successions except two samples of facies Inw/Iw plot in within plate granitic field.



different behavior, with a pronounced decrease (enrichment factor of 0.13–0.30 for Sr and Ba, respectively) upward in the stratigraphy, except for HFSE (2.0 for Zr, 1.59 for Y and �2 for Nb).Some mobile and immobile elements such Fe2O3, Ba and Zrshow nice parallelism in their decreasing concentr ations, which emphasizes the boundaries of compositional groups A, B, and C(Fig. 13), which correspond to lava flows, Facies II, and Facies Inw/Iw, respectively.

Fig. 13. Compositional trends of bulk rock for major and trace eleme

6. Discussion

6.1. Magma chamber zonation

Fractiona l crystallization is clearly the principal differentiation process generating compositi onal zoning during magma accumulation and crystallization in a shallow magma body (Hildreth and Mahood, 1985; Hildreth and Wilson, 2007 ). Bulk rock behavior

nts, along reconstituted stratigraphy of the Hamid successions.



show that Hamid volcanics follows the general model of large silicic differentiate d magma chambers, with a progressive tapping of its compositi onally zoned magma reservoir from less evolved to a highly evolved layer. A cryptic zoning is observed by the strong variations in trace element concentration especially Sr and Ba(Wolff and Storey, 1984 ) whereas major and trace elements (suchas TiO 2, P2O5, Rb, Y, and Nb) are less affected (Fig. 13). This con- trasts with silicic systems, in which both major and trace elements characterize the fractionation. Although major element concentrations in the Hamid volcanics follow a specific increasing ordecreasing trend upward in the stratigraphy, they range narrowly in composition. In contrast, trace element concentrations vary more significantly and follow typical fractionation trends (Whiteet al., 2003 ). Caution should be exercised in interpreting Sr and Ba trends for bulk rock Neverthel ess, Ba broadly increases with stratigraphy , confirming a correlated elevation of magma temperature, pressure and H2O% (Guo and Green, 1989 ). The compositional groups A, B, and C are well defined with the variation ofconcentratio n in Fe2O3, K2O, Sr, Ba, and Zr elements upward inthe stratigraphy . Such a variation in these elements suggest that the fractionati on of feldspar, ziron, and iron oxides may be largely responsible for the developmen t of such zonation (Fig. 13).

6.2. Vent and eruption column dynamics

The Hamid Basin preserves complex volcanic and volcanicl astic successions associated with two distinct volcanic sucessions (Figs. 3 and 6). Their beds represent a discontinuo us, unconform- able volcanic successions as evidenced by the presence of unconfo- rmities or erosional discordan ces (Figs. 3 and 9A). The Hamid eruption evolved spatially over time from an early phase in the south to a later or possibly contemporane ous phase in the north.Their pyroclastic rocks reflect a spectrum of eruptive conditions,fragmentation mechanisms, and provide information on the dynamic reconstruction of eruptive systems on a detailed scale (e.g., De Rosa, 1999). Based on the stratigraphi c data and the results from facies analysis, the HVSS has been have interpreted asthe products of four eruption stages.

6.2.1. Opening stage Following an initial period dominated by mass flows, re-estab-

lishment of fluvial systems began with the headward erosion of acanyon through the granitic rocks and lava flows, a process often associated with the break-out of a temporary lake, such as Miocene Eoil Basin and SE Korea (Jeong et al., 2008 ). Aggradat ional streams developed in these channels rapidly evolved from deeper, perma- nent braided rivers (FA1 and FA2) to shallow, ephemeral, sediment-laden outbursts associate d with flash flood events (FA3,Fig. 6). These fluvial deposits reflect many episodic influxes of volcaniclastic material rather than a perennial supply in a fluvial-lacustrine system. Typically the modern profile of many streams and rivers follow closely their pre-eruptio n profiles, and incision and erosion is overwhelmi ngly confined to the deposits of the eruption itself (Manville et al., 2009 ).

6.2.2. Pulsatory stage An increase in eruption intensity led to the onset of the lower

Hamid sequence , consisting of pulsatory activity associate d wit aseries of explosions with different magnitud e and intensity. After the deposition of facies associations FA1, FA2, and FA3, renewed volcanic activity caused the deposition of FA4 and FA5 contempo- raneously with FA6. These deposits were deposited over the southern region of the study area suggesting a predomin antly sustained eruption column that experienced fluctuations in eruption intensity over time (Figs. 3 and 6). In FA6, the presence of juvenile clasts within a sedimentar y host, secondary alteration of the friable

fragments , and jig-saw fit textures that were mainly defined bydevelopmen t of hairline cracks in crystals (Fig. 8E) support a po- tential phreatomagm atic explosive eruption that opened the vent in the braided river system. This is also supported by the absence of a polygenetic composition of the juvenile clasts (Doyle, 2000;Erkül et al., 2006 ). The phreatom agmatic eruption was probably facilitate d by ‘‘wet’’ eruptive conditions in the vent site (Kong,2000; Jeong et al., 2008 ). After that, there was the emplaceme ntof a series of dilute (as suggested by internal features structure;Branney and Kokelaar, 2002 ), juvenile-r ich ignimbrites (facies Il)along the southern slope of the Hamid volcanoes (Fig. 6). They were emplaced contempor aneously with the sedimentation ofFA4 to FA6, suggesting an unstable column with moderate explosive activity associated with a large mass discharge rate. These re- peated explosions led to progressive vent opening. The textural characteri stics of these deposits, such as angular juvenile clasts and monomictic volcanic components, show that the clasts, inthese early eruptive processes, were fragmented explosively.

6.2.3. Collapsing stage The eruption dynamics changed again to a more sustained

phase with the depositio n of Bm and Tv, which consists of volcanic breccias and multiple fallout beds with discontinuo us distribution ,underlain by and locally intercalated with ignimbri tic deposits. Bmis interpreted to be a co-ignimbrite lag breccia produced during vent widening. Vent widening in the south presumably occurred in response to vent wall rock instabilities and increased magma discharge rate during the early Plinian phase that destroyed a large portion of the southern side of the pre-existing stratovolcano edifice. The vertical gradation into an overlying fine vatic tuffs (Tv) reflects changing source condition s at the vent, a progressive increase in magma discharge rate and cessation of vent-widening activity (Németh et al., 2012 ). The discontinuous distribution suggests syn-erup tive erosion or contemporane ous depositions of fallout beds with the ignimbritic deposits during the final stage of the eruption (Dellino et al., 2004 ). The facies Tv is related to a partial collapse of the eruption column (see Maeno and Taniguch i,2009), which generated base surge deposits over a northweste rnpart of the study area (Fig. 6). It is important to mention a few additional points in order to fully understand the mechanis m offragmentation and deposition of the Tv. The small sizes and extremely low percentage of free crystals is inconsistent with crystallization-in duced degassing and decompress ion at low eruption rates (Mattsson, 2010; Taddeucci et al., 2004 ). The paucity on free crystals (and high percentage of juvenile clasts) in the primary volcanic facies are due to rapid ascent of the magma from source to vent,and then, fast quenching of magma in contact with water (Winter,2001). The fine grain size of the facies Tv, mainly fine vitric tuffs (Table 1), suggests that the eruption was also phreatomagm atic (Heiken and Wohletz, 1991; Wohletz, 1986 ). The restricted occurrence and abruptly pinching-ou t geometry of this unit suggest that the magnitud e of the eruption was small, and the pyroclastic density current was dense and nonturbulent so it could be ponded intopograp hic lows (Schumacher and Schmincke, 1990 ). Accretion -ary lapilli are especially typical of phreatomagm atic pyroclastic deposits (McPhie, 1986; Schumacher and Schmincke, 1995 ) but have also been observed in eruptions considered to be magmatic rather than phreatomagmati c, in conditions, however, of high atmosph eric humidity (Cole et al., 2002; Watanabe et al., 1999 ).Phreatomag matic fragmentation mechanis ms are not excluded here but they cannot be inferred purely from the existence ofaccretionary lapilli as recrystallization has obliterated the original morphology of the fine ash particles . Since accretionary lapilli are formed in the air, their presence in the stratigraphy of the HVSS implies that these pyroclastic deposits were formed in subaerial condition s or very shallow water (Asvesta and Dimitriadis, 2010 ).



Tv beds are a typical example of ‘‘wet’’ base surge deposits due towidespread occurrence of accretionary lapilli and presence ofdepositional structures such as antidunes and cross bedding. Pres- ence of both regressive and progressive antidune structures within the same deposit reveals the pulsatory nature and change in flowregime of the base surge (Kuscu et al., 2007 ). Most likely, a large amount of water was added to the erupting magma, resulting ina wet and dense pyroclastic density current that spread from a collapsing fountain rather than from a well established Plinian eruption column (Koyaguchi and Woods, 1996).

6.2.4. Climactic stage After deposition of the bedded conglomerates (Cb), the initial

large-scale plinian phase was progressively followed by a column- collapsing phase. Partial collapse of the eruption column produced a small volume, high temperat ure, pyroclastic flow directed to the northwest-s outheast (Inw/Iw, Figs. 3 and 6). The high temperature of the pyroclast ic flow was reflected in the welded zone that developed in the ignimbri te (Fig. 11) during post-emp lacement compaction. The magma that fed the high evolved group of ignimbrites has adacitic to rhyolitic in composition (Fig. 12) with pre-eruptiv e tem- peratures P900 �C (White et al., 2009 ). Excavation of the conduit wall rock during vent widening was accompanied by decompression of the magma system resulting in an increase in the magma discharge rate. This caused a change in clast supply at vent from lithic clast-domin ated to pumice-dom inated. This is preserved as a gradational change between lower horizons during lithic-rich source conditions to upper horizons during pumice-rich source conditions (Fig. 11). Sustained eruptive activity in the north and increased magma discharge rate likely created great instability in the magma chamber roof and surrounding wall rocks of the progressive ly drain- ing magma chamber. Further progressive collapse of the roof block and erosion of the conduit wall rock during vent excavation and fragmentation, tapping progressive ly deeper levels of the subsurface,generated vast quantities of lithic clast debris and signaled the onset of caldera collapse. Moreove r, these lithics may be derived from the breakup of the country rock due to conduit-wal l abrasion (Macedo-nio et al., 1994 ), or conduit pressure changes during magma ascent (Papale and Dobran, 1993 ). These clasts do not match any basement rocks underlying the Hamid area and are interpreted to have been vent-derive d. In addition, deeper metamorphic basement lithic clasts are absent and the lithic clasts within ignimbrite units are inferred to have been derived from the shallowest (<1 km) levels in the stratigraphy , intersected at the vent and conduit (Allen, 1998).Dense, boulder-s ize lithic clasts and pumice clasts were all mixed and ejected up along the widening conduit in the north as well asthrough newly created, concentrica lly distributed fissures formed by steadily increasing magma discharge rate and progressive collapse of the magma chamber roof.

The facies Inw/Iw is one of the most widespread eruption phases and is extensive ly preserved in north-western situation inthe mapped area (Fig. 2). It directly overlies vitric tuffs with anangular unconformity , separated only by bedded polymictic conglomerates (Cb) (Fig. 4). Similar examples have previously been observed on many proximal ignimbrite emplacements (Branney and Kokelaar, 2002; Cole et al., 1993; Druitt and Sparks, 1982; Rosi et al., 1996 ). Although the location of an eruptive vent is not well constrained due to tectonic uplift and erosion, it may have been located near the Pb-mine locality because surge and flow deposits are only distributed across the northern part of Hamid area (Haray-ama, 1992; Nagahashi et al., 1996; Nagahashi, 1998 ). Besides, the pyroclastic flow sheets unconformably overlie basinwar d-tilted base surge deposits around the caldera volcanoe s due to insurgen tuplift caused by growth of huge magma chambers preceding the caldera-form ing eruptions (Lipman, 1976 ; Riggs and Busby-Spera,1990; Smith and Bailey, 1968 ). The thick pyroclastic flow sheets

represent sudden aggradation during caldera-form ing eruption s(Yamamo to, 2009 ). The change from phreatom agmatic (facies Tv)to dry explosive eruptions (facies Inw/Iw) coincides with an overall trend of increasing discharge rate from stage 3 to stage 4, and from a vertical eruption column to the production of high particle concentration pyroclast ic density currents (Allen, 1998). Thus, it ismost probable that the alternating wet/dry style recorded by the upper part of the ignimbrite results from fluctuations in magma discharge rather than water supply. The volume estimate d for the ignimbrite units is large compare d to that of the base surge units, and similar to documented base surge deposits formed during phreatop linian eruption s (e.g., Layer C-Askja, Sparks et al.,1981). This stratigraphy can be interpreted as the result of progressive instability of the eruption column shifting from partial collapse to final total collapse dynamics (Di Muro et al., 2004 ). The transitional collapsing/buo yant behavior of volcanic plumes has al- ready been observed in Plinian eruptions fed by silica-rich magmas and often explained by a significant variation in mass discharge rate and/or water contents (e.g., Di Muro et al., 2008; Melnik,2000; Rosi et al., 2001; Wilson and Hildreth, 1997 ). In contrast,the waning phase of basaltic explosive eruptions is more commonly characterized by effusive activity (e.g., Andronico et al.,2009; Pioli et al., 2008; Scollo et al., 2007 ), or by further explosivity sustained by magma-water interaction (Taddeucci et al., 2002;Walker et al., 1984 ) or by a series of weak explosions (Mannen,2006). In addition, PDC (FA 7) deposits of the northernwest block are not commonl y associate d with dry basaltic or andesitic explosive eruption s. I therefore infer that the final (recorded) stages ofthe HVSS attest uncommon eruption dynamics for low viscosity magma, which might have been marked by an increase of mass discharge rate and/or a decrease in water in the eruptive mixture.In general, dense pyroclasts can be interpreted as formed by a partially degassed magma body residing in the shallow portion of the magmati c system or rising very slowly (Houghton et al., 2004; Pioli et al., 2008 ). In fact, collapsed vesicles are commonly interpreted tobe the result of shrinkage due to gas loss after significant growth and coalescence (Cashman and Mangan, 1994 ). In addition, the small size of these vesicles suggests open-system degassing conditions associated with rapid cooling of a small magma batch. On the contrary, higher vesicularitie s and smaller groundmass crystallin- ity suggest fast rise from depth with no significant residence time at shallow depths before the eruption. In the Hamid area, stratigraphic data indicate a general increase of the proportion of the vesicles- enriched juvenile clasts with time (Fig. 10E/F; facies Inw/Iw), which is not necessar ily accompani ed by an increase ofthe mass discharge rate. I therefore suggest a model in which the arrival of a small batch of magma, which intruded at shallow level and underwent rapid cooling, outgassing and crystallization,caused a prolonge d activity of scattered ash emission and progressive conduit opening, with eruption of large wall rock blocks. The fast rise of this magma (represented by vesicle-enri ched juvenile clasts) triggered the highly explosive eruption phases (pulsatory,collapsin g, and climactic). The increasing content of vesicle-rich juvenile clasts from bottom to top facies (Inw/Iw) and a progressive increase in eruption magnitude and intensity suggest that felsic magma had a primary role in controlling eruption explosivity.

6.3. Basin-fill architecture

Hamid Basin is unique composite volcano and its eruptive sequence is more typical of those eruptions associated with large-volum e silicic compositi on volcanoes with significant inter- eruptive periods. This basin is developed on a granite–gneissicand meta-vol cano-sediment ary succession bearing basement. The total thickness of the HVSS is �600 m. This thick succession registered successive extension al tectonics. The seven lithofacies



associations described above record abrupt lithofacies changes,suggesting significant changes in eruption, transport, and depositional processes . Such frequent changes are expected inbasins that remain tectonically hyperactive during sedimentation (cf. Martins-Neto, 2000 ). These depositional systems were in part contemporane ous and in part succeeded each other during the history of the basin evolution. Facies architecture in FA1 (Fig. 6)reveals initial coarsening and thickening (progradational), and later fining and thinning (retrogradational) successions . Facies architecture in FA-2, 3, 4, and 5 is progradational , whereas in FA- 6 and 7 the architectures are retrogradation al. The progradational and retrogradation al successions are the manifestatio ns of tectonic control on basin evolution. Prograda tional successions are the result of infilling of basin under stable tectonic regime, on the contrary retrogradation al successions are syntectoni c and developed when downsagging of the basin floor is much faster than sediment supply (Bhattachar ya and Bull, 2010 ).

The NED is dominated by metamorph ic and intrusive complexes which formed the substrate on which volcano-sed imentary successions like the Hamid successions evolved. Thus the most likely scenarios for basin formation in such terranes would be deep erosional dissection, rifting, transtension, pull-apart or caldera systems. A compressive scenario such as piggy back basins typical for active fold-and-thrus t belts is not probable, also since NED lacks indicative structures such as thrust faults and related folding.The studied lithofacies associations represent development offive-phase sequences that were different enough to produce deposits of contrasting lithofacies types and deposit shape (Fig. 14).

6.3.1. Phase 1: Alluvial fan period Alluvial fan sediments were develope d at the initial phase of

fault controlle d basin opening, resting on older granitoids with erosion unconformity (Fig. 7A). The sedimentar y succession ofthe Hamid basin thus records developmen t of multiple tectonism-induc ed sedimentar y cycles in an intracontinent al rift setting. Abrupt change in facies types and their thickness across faults suggests that synsedimentar y faulting separated sedimentary environments of greatly different water depths and subsidence rates during active rifting (Kwon et al., 2011 ). In the firststage of evolution (Fig. 14 1A/B), The Hamid Basin commences with massive polymict ic conglomerates (Cm) based on its feeding by, different sources, showing very high detrital concentrations comprising clasts-crys tal rich sandy matrix. The unstratified,coarse-grained , poorly-sorted and heterolithic distribution ofblocks and gravel fragments in a sandy- to conglomerate- domi- nated matrix support their rapid deposition from non-cohesive volcanogeni c debris flows (Manville et al., 2009 ). The introduct ions of volcanic debris and plutonic intrusions to surrounding areas are mainly gravity-d riven mass flows due to the characterist ic high relief topography. This scenario requires the presence of high and ex- tended mountain ranges surrounding the Hamid Basin exposing metasediment ary, metavolcan ic, plutonic and volcanic rocks ofNeoproteroz oic age. Therefore, post-eruptive sedimentation typically begins with a debris-and hypercon centrated-flow phase associated with braiding of the river system that terminates with the resumption of background sedimentation (i.e., inter-erupti ve sedimentation sensu Smith (1991). The scarcity of fine-grained deposits suggests that the sediment supply was much greater than the accommodati on generation (Shanley and McCabe, 1994; Wright and Marriott, 1993 ), so that most of the fine-grained particles passed to more distal areas. The provenance was derived from abroad catchment area south of the basin, probably reflecting anearly capture of a major fluvial system to the basin during the firststages of its history. The only record of deposition during quiescent intervals is the unconfined channel-fill conglomerates. Any flood-

basin deposits that might have accumulated during this period were apparent ly removed by erosion. Preservation of only the channel part of the quiescent record is favoured in settings with relatively low subsidence rates (Smith, 1991 ). Thus, slow subsidence characterized the basin during accumulati on of the alluvial fan deposits. As discussed above, FA 1 has been interpreted as analluvial fan setting that is restricted to the vicinity of the basin-border faults following models from other areas of the Egyptian East- ern Desert (Dardir and Abu Zied, 1972; Grothaus et al., 1979; Willis et al., 1988 ). Rounding of boulders took place during fluvial transport in mountain rivers. The basin is closest to the inferred source and hence experienced thick braided river-style aggradation with abundan t debris flows during phase 1.

6.3.2. Phase 2: Sandy fluvial channel and lacustrine period The Hamid Basin experienced a drastic change in depositional

environm ent at the beginning of this period. This latter phase was characterized by the coexistence of a mixed-load sandy braided rivers and an ephemeral fluvial system (Fig. 14, 2A/B).The transition to the sandstone-mu dstone assemblage occurred when sediment yield began to decline and suspended sediment became an important part of the total sediment load of the river. The response was developmen t of a more stable channel and flood-ba-sin system. In the course of Phase 2, the transitions from lacustrine to sandy braided rivers and again to lacustrine plains are documented (Fig. 6) indicating a strong fluctuation in water depth during the flood events. The ephemeral nature of the streams issuggested by the recurrence of these cycles, marked by episodic peak discharge followed by waning flow, with no evidence oflong-lasti ng active channels. The abundance of facies Ml in FA3 suggests a calm-water environment, where settling of fine-grainedparticles took place. The rhythmic intercalation of FA 2 and FA3 with the sandston e facies (Fig. 6) indicates cycles of sand input,both from traction and suspension, followed by water stagnation and subaerial exposure. According to Paim (1994), the lake-level rises would be simultaneou s with increased sand supply, and the lake-level falls with the deposition of fine-grained particles from suspensi on. These two stages are recorded as well-defined distinct successions, here interpreted in terms of the external controls onthe basin architectur e.

The shift from the high energy environment of Phase 1 to the medium to low energy system of Phase 2 can theoretically be explained by strong denudation of the surrounding mountain ranges.However this shift was related to associated fault movement in the study area. Local faulting could have created a low area in the Ha- mid area, and disrupted both the groundwate r and surface-w ater systems. Lakes respond rapidly to fault-related changes in basin physiogr aphy, and juxtaposition of lacustrine deposits and subaerial deposits is a common record of basin faulting (e.g., Blair, 1987;Blair and Bilodeau, 1988; Blair and McPherson, 1994 ). Lack of adeep-mar ine basin-fill, the abrupt lateral/vertical lithofacies variations, coarse-grai ned alluvial fan sedimentati on, intrabasinal unconfor mities, uni-modal volcanism, and basin margin faults strongly suggest strike-slip influence for the Hamid Basin (cf.Mueller and Corcoran, 1998; Nilsen and McLaughlin , 1986 ). Thus,it is possible that the occurrence of abundant flood flow cycle sediments and hyperconcentrate d and debris flows deposits indicate the influence of topographic highs during the evolution of the basin margin strike-slip system. It should be concluded that a combination of extensional (transtensional regime) and strike-slip tectonic phases (transpressional regime) in the Eastern Desert has been happened at interval 600–530 Ma (Stern, 1985 ). The formation of the Dokhan Volcanics and the Hammam at Groups may confirm the association of these tectonics (Akaad et al., 1993;Abd El-Wahed , 2009 ; El Kalioubi and Osman, 1996 ).

Fig. 14. Composite column and schematic cartoons showing the diverse facies associations and the detailed description of the five phases for the basin evolution in Hamid area, Northeastern Desert, Egypt. (Phase 1) Intense rilling and debris flow initiation occurs on steep slopes blanketed by volcanic and granitic materials. (Phase 2) Sediment–water ratios decline as yields from hillslopes decrease through depletion of the reservoir of easily remobilised material, completion and stabilization of rill systems, recovery of infiltration rates, and increased discharges resulting from drainage network integration. Channels form and begin to incise the underlying deposits vertically and headwards, resulting in increasingly distal transport and aggradation of volcaniclastic sediments (sandy fluvial and lacustrine rock associations). (Phase 3) Coherent lava floweruption combined with ejection of ignimbritic rocks (facies Il). (Phase 4) Plinian column-feeding phase, resulting in vitric tuffs and volcanic breccias in deflection zone.(Phase 5) Plinian column-collapsing phase, generating high temperature dilute currents and depositing ignimbritic rocks (facies Iiw/dw). High temperature deposition during this phase is characterized by agglutination and dense welding in proximal areas.



6.3.3. Phase 3: effusive and explosive volcanism During Phase 3, the lake perished or shrank. Again extremely

violent eruption of coherent volcanics from a vent or fissure produced lava flows and pyroclast ic fall out deposits (FA4 and FA5).The chemical composition of the lava flow indicates its intermediate composition and fits into the andesitic intraplate lineage (Fig. 12A/B). In the central basin area, a phreatomagm atic vent situated in unconsolida ted sediments (cf. Németh and Martin,2007) led to the formatio n of volcanog enic mass flow deposits (Cv), hyaloclastic deposits (Hy). As the eruption proceeded, less volatile-rich magma was encountered and the structural instabilities caused collapse of the conducit walls, abruptly shutting off the magma supply. As the result of falling gas content, eruption velocity decrease and increase of vent radius, collapse event started and resulted in forming of lithoysae-rich ignimbrites (fa-cies Il). The abrupt switch from fall to flow forming activity (FA7) is though to have been caused by a drastic increase in discharge rate. Co-genetically , various volcanic centers, both of silica-rich and silica-poor compositi on, were active (Fig. 14, 3A/ B). The lava flows coexist with pyroclastic and epiclastic deposits in the same accumulation space. The presence of juvenile magma clasts, jig-saw fit textures that were mainly defined by development of hairline cracks in crystals (Fig. 8E), and devitrificationof glassy fragments (Fig. 8F) indicate a major role of magma- water interactio n driven phreatomagm atic explosive events aswell for the FA6.

6.3.4. Phase 4: Explosive volcanism and base surge deposits Initial stages of the phase 4 involved plinian style activity. This

plinian column-feed ing stage generated vitric tuff deposits (Tv)over the northern area (profiles III and IV in Fig. 6) suggestin g apredomin antly sustained eruption column that experienced fluctu-ations in eruption intensity over time. Soon after partial collapse ofthe eruption column, vent widening in response to vent-wall rock instabilit ies and increased magma discharge rate overloaded the eruption column with dense, lithic debris resulting in southward collapse and deposition of a thick co-ignimbr ite lag breccia in the deflation zone (Fig. 14, 4A/B). They were too coarse and bulky tohave been transported far by the ensuing pyroclastic flow. Estimat- ing the height of the eruption column is difficult due to a lack ofoutcrop and poor preservation conditions. The recognition ofbase-surge deposits in such systems is particularly important because of their significance as an indicator of near vent volcanic processes (Bull and Cas, 2000 ). Pyroclastic vitric tuffs are cross- bedded, stratified to well bedded, fine grained and poorly sorted,characteri stics that are typical of deposition from laterally moving pyroclast ic density currents (Sohn, 1996 ). In addition, it exhibits many planar and gullied erosive basal contacts within the framework. The presence of some ambiguous rounded fine ash balls (likely accretionar y lapilli) indicate the presence of condensed water in the depositional system that is typical for phreatom agmatic explosive eruptions. Edmonds et al. (2006) and Seghedi(2011) indicate that accretionary lapilli in fallout deposits closely



associated with pyroclast ic flows in a proximal environment are more commonly associated with phreatomagmati c eruption s. In- deed the decrease in grain size can be explained by a decrease inhydrostatic pressure perhaps due to shoaling caused by the growth of the volcanic pile (Kokelaar , 1986 ), which allows for more violent phreatomagm atic explosions.

The significant time break between the vitric tuffs and capping ignimbritic volcanic units (Inw/Iw) as inferred from the presence ofa thick discontinuous channelized gravelly volcanic lithic conglomerate beds (Cb) suggests a long-lasting inter-erup tive period prior to resumption of volcanic activity, forming small-volume, low as- pect ratio welded pyroclastic density current deposits (Fig. 4). This interpretation is favoured over fluvial origin, due to the general high elevation where this cover can be located and its volcanic flank-mantling nature. The braided rivers sampled volcanic debris from the surroundi ng volcanic terrains, likely from the nearby Ha- mid volcanic massif. As a result a braided river network formed with volcanic lithic-rich gravelly bars and fine-grained fluvial volcanic sands. All these rocks form the upper succession of the Ha- mid volcano. Generally the eruptive products of the vitric tuffs are laterally restricted but there are exceptions that represent eruptive phases capable of producing sub-Plinian to Plinian styles that form laterally extensive and thick pyroclast ic fall and flowdeposits (Schumac her and Schmincke, 1990; Van der Bogaard and Schmincke, 1984 ). A relationshi p between surges and flowshas been predicted (e.g., Wohletz and Sheridan, 1979 ) and the deposits from some pyroclastic density currents suggest that there is a gradation in transport and depositional processes between surges and flows (e.g., Mount St. Helens blast deposit, Druitt,1992); member B of the Neapolitan Yellow Tuff, Cole and Scarpati (1993). Such eruption s are commonl y associated with magma from complex sources (Woerner and Schmincke, 1984 ; Woerner and Wright, 1984 ) and their eruption can result in significant environmental effects similar to those from long-lived, composite volcanoes (Schminck e et al., 1999 ).

6.3.5. Phase 5: Caldera eruption and ignimbrite-f orming deposits As the surges progressive ly dried up during emplaceme nt,

and lost more and more of the carried load due to deposition,the dynamics of the flows also changed. The initial large-scale plinian phase was progressive ly followed by a column-collap sing phase (Phase 5). Eruptive activity in the north likely ended with deposition of ignimbrites (Iiw/dw) when partial collapse of the magma chamber roof in response to significant magma with- drawal resulted in further decompress ion of the magma system and blocking of the vent (Fig. 14, 5A/B). Thick pyroclastic flowdeposits (ignimbrites) are present in the northern part of the Ha- mid Basin. This evidence may indicate that this basin presents acomplex succession of eruptive phases, gravitational or explosive collapse, and erosional intervals during the growth of the volcanic edifice (Freundt et al., 2000; Sato et al., 1992; Rotolo et al.,2013). These deposits could also be associated with lava domes and produced by dry explosive and collapse processes initiated directly by a volcanic eruption (e.g., in Merapi and Unzen volcanoes, Newhall and Melson, 1983 ) or alternativel y produced less commonly where magma encounters an aquifer or surface water and dynamical ly interacts with it (e.g., Takayama Volcano, Kanoand Takarada, 2007 ). Chemical analyses of pumice lapilli indicate their dacitic to rhyolitic composition (Table 2 and Fig. 12). This pumiceous pyroclastic succession is stratifie, massive and has erosional contacts with the underlying vitric tuffs which suggests depositio n from dilute pyroclastic density currents (Figs. 4and 9A). The high temperature emplaceme nt of the pyroclastic flows is reflected by the different zones of welding, the pink col- or, and columnar jointing associated with the post- emplaceme ntcompaction of the ignimbrites (Cas, 1983; Hoblitt and Kellogg,

1979). High magmatic flux evidenced by the increasing signs ofhigh temperat ure emplacement of pyroclast s in the upper part of the pyroclastic section indicates a gradual increase of magmatic fragmentati on of magma in the course of the eruption and growth of the felsic volcanics. Vertical transition from lithic-rich pumiceo us beds to welded pyroclastic units (Fig. 11)indicates that these units are formed from pumiceous pyroclastic density currents, e.g., small volume ignimbri tes (Allen, 2004;Freundt et al., 2000 ). The origin of vertical grading of lithic and pumice clasts is related to mechanis ms operating during the transport process within the high particle concentration basal avalanche of pyroclastic flows. Vertical grading results from the balance between gravitational and dispersive forces, and istransferred to a lateral grading by vertical velocity gradients within a nonturbulen t flow zone of pyroclastic flows. The pyroclastic flows are modeled as Bingham-type fluids (Danilo et al.,1995). The gradual increase in the abundan ce of pumiceo usclasts together with the increased number of large elongated fiamme indicates an increase of hot magma feeding the eruption.It is concluded that compactional load, local variations in accumulation rate and clast sizes, alongside high eruptive temperature are all partially responsible for welding in the ignimbri tic deposits especially in the upper horizons (Figs. 9A and 11, Boyceand Gertisser, 2012 ). The large number of volcanic lithics in the basal portion of the ignimbritic sections suggests that their likely origin is from the underlying braided river gravel beds that the eruption has punctured (Fig. 4). Pumiceous clasts with moderate vesiculari ty and thick glass shard walls (Fig. 10E and F) indicate a minor role of magma-water interaction driven phreatom agmatic explosive events, enhancin g its explosive eruption, which subsequent ly quickly turned into a purely magmatic fragmentation-domi nated eruption forming an eruption column providing pyroclast s. This coincides well with the presence of volcanic con- glomerat es sandwiched between the vitric tuffs and ignimbri tic deposits suggesting some resetting of the braided river systems into which the Hamid volcano erupted (Vespermann and Schminck e, 2000 ). The role of phreatomagmati c explosive eruptions in the vent opening and clearing stage of small-volume silicic dome-form ing eruptions are well-documen ted from both field studies and experimental volcanology (Austin-Er ickson et al., 2008 ).

The differenc es between the lithofacies Inw/Iw and other phre- atomagmat ic pyroclastic density current deposits are primarily the result of the relatively high particle concentratio n and momentum of the current and its unsteady nature. These characterist ics were largely inherited from the collapsin g phreatoplini an eruption column, especially the presence of steam within the column, the pulsatory nature of the discharge and the great height from which the eruption column collapsed. Facies Tv and Inw/Iw are interpreted tobe individual pulses of a discontinuously depositin g current that record changes in componentry over time as the deposit progressively aggraded (Branney and Kokelaar , 1992, 2002 ). The latter two facies successions are intruded by co-genetic subvolcanic bodies, which are considered the last volcanic activity in the studied area. The deposits from pyroclast ic density currents reflect processes occurring only at the final stage of deposition (within the depositio nal boundary layer) and do not reflect the long- distance transport. This is a characteristic well established in pyroclast icsurges (e.g., Sigurdsson et al., 1987 ), but also attributed to high particle concentration pyroclastic flows (Branney and Kokelaar, 1992 ;Fisher et al., 1993; Wilson and Walker, 1982 ). Evidence for a floworigin includes topographic control on thickness variation, the presence of a poorly sorted coarse-ash matrix, the local imbrication of juvenile clasts and the presence of a reverse graded upper layer,common at the base of most ignimbrites (Mellors and Sparks,1991; Sparks et al., 1973 ).



6.4. Comparison with volcanic units in adjacent basins

Although high-precision geochronology is lacking for much ofthe Dokhan Volcanics in the Eastern Desert, stratigraphi crelationship s and similarities in eruptive style suggest that corre- lations between other volcanic units can be made. Neoproterozo- ic crustal components of the NED include metavolcanic complexes, older granite (OG), Dokhan Volcanic (DV), Hamma- mat sediments (HS), younger granite (YG), and abundant dykes (Fig. 1; Khalaf, 2004; Mohamed and El-Sayed, 2008 ). The DVand HS interfinger in volcanosedime ntary successions that accumulated in Ediacaran extensional basins (Breitkreuz et al., 2010 ).The intermon tane HS consist of coarse alluvial conglomer ates,fluvial sandstones and lacustrine pelites, which are partially vol- canoclastic in compositi on (Eliwa et al., 2010; Grothaus et al.,1979; Khalaf, 2004 ). These successions unconformably overlie OG and metavolcan ics and have been intruded by YG with sharp contacts. The studied areas for stratigrap hic comparison in the NED stretch between Ras Gharib city in the north and Safaga city southward (Fig. 1). A comparis on with these areas shows good stratigraphi c similarity (Fig. 15). The Wadi Bali across Esh ElMellaha, 30 km north of Hurghada, exposes a complex volcano- sedimentary succession (Fig. 15A). At the base of the succession,andesitic lavas are directly overlain by a thick layer of welded ignimbrite and capping volcano-s edimentary sequence (Khalaf,1986). The steep hill east of Gabal El Kharaza exposes a section that starts with alluvial sediments , unconforma bly overlying OG(Fig. 15B). Gabal El Urf, located 12 km to the south of Gabal Kharaza, represents a volcano-s edimentary succession with athickness of 2000 m (Fig. 15D; Eliwa et al., 2010 ). This succession starts with thick Hammamat-ty pe conglomerates and lacustrine sediments, overlying OG with an erosional unconformity.The upper part of both Gabal Kharaza and El Urf Successions isdominated by two thick ignimbrite units. The preliminar y profile

Fig. 15. Schematic lithological profiles of volcano-sedimentary successions in the Ras GhEl Maleha Range (Khalaf, 1986 ). (B) East of Gabal Kharaza (Breitkreuz et al., 2010 ). (C) GaStandard vertical bedding sequence in subaerial pyroclastic flow deposits (Sparks et al., 19have different scales. For location see Fig. 1, for legend see Fig. 6.

of Gabal Nuqara southwe st of Safaga (Fig. 15C) shows a succession of andesitic volcanics interbedded with fluvial sandston esand alluvial conglomerates overlain by non-welded to welded ignimbri tic rocks followed upward by rhyolitic lavas (Khalaf,2012). Furthmore, the upper volcanic facies of the Hamid area (Fig. 15E) generally conform to the standard sequence (Fig. 15F) of Sparks et al. (1973), where subaerial pyroclast icdensity current deposits laterally and vertically become pyroclastic surge deposit at the base, overlain by lithic-rich layers (2a),and capped by pumice-r ich layers (2b). Fine ash deposit (layer3) has not been recorded in the Hamid Basin. So, these basins which contain both Hammam at and Dokhan facies especiall y Ha- mid and El Urf basins (columns D and E, Fig. 15), comprise alluvial-fan deposits, fluvial braided flood-plain deposits, deep lacustrine deposits, andesitic phreatom agmatic volcanic deposits,explosive pyroclast ic deposits, and coherent bodies of lava flows,sills, and dikes. These deposits accumulated in a structurally controlle d intermontane basins, beginning with alluvial-fan con- glomerat es and sandston e eroded from flanking high mountains,followed by the developmen t of a deep lake, perhaps as a result of down faulting. Shrinkage of the lake is marked by a return ofhigh-ener gy arenaceo us and rudaceous sedimentation concurrent with the onset of silica-ric h and silica-poor volcanic centers leading to the formation of volcanog enic mass flow deposits,hyaloclasti c deposits, and lavas. The terminal history of the basins were marked by large ignimbri te-forming caldera eruptions.

Such multiple eruptive episodes are also common from silicic tuff ring and dome volcanoes such as Cerro Pinto in central Mexico (Zimmer et al., 2010 ). Cerro Pinto’s evolution is a relatively isolated, small volcano with multiple eruptive phases including explosive , effusive, dome- (cryptodome-) forming and cone destruction al episodes separated by long eruption repose periods allowing the accumulation of volcanicl astic aprons during inter- eruptive times (Carrasco-Nune z and Riggs, 2008 ). These eruption s

arib segment, North Eastern Desert, Egypt. (A) wadi Bali in the southern part of Eash bal Nugara (Khalaf, 2010). (D) Gabal El Urf (Eliwa et al., 2010 ). (E) Hamid Basin. (F)73 ). Ages for ignimbritic Dokhan Volcanics are from Breitkreuz et al. (2010). Profiles



behavior are typical for long lived, polygenetic composite volcanoes over the lifespan of such composite volcanoes (Fisher and Smith, 1991; Smith and Lowe, 1991; Manville et al., 2009 ). Hamid area however has not erupted through such multiple magma pulses as indicated by its very homogen eous geochemical signa- ture suggesting a small isolated magma chamber responsible for the formation of this volcano. Cerro Pinto shows strong evidence of being a single vent system by its morphology and stratigraphy which are similar to Hamid area in spite of the recorded chemical variations in its eruptive products indicating some degree of differentiation, crustal assimilation or their combination over the eruptive period of the volcano (Carrasco-Nune z and Riggs, 2008 ).

It is likely that Hamid area has either a single vent or closely spaced vents as suggested by the simple morphology of the volcano and these vents tapped magma from similar sources (Fig. 12). This behavior is typical for complex composite volcanoes that have erupted over long time. Hamid’s complex magmatic system and associated distinct eruptive phases recorded in its volcanic architecture make it a unique volcano. Hamid area also differs from Cerro Pinto, not having preserved lava domes in its edifice. There is also no evidence to support existence of lava dome forming events that were destroyed by dome collapses that could be recorded in block and ash flow deposits. Such deposits are lacking at Hamid area. The presence of welded pyroclastic units in the capping succession of Hamid area suggests, that its eruption has reached a high magma discharge rate stage in its final phase poten- tially in an eruption that can be looked a unimodal eruption style,also more common feature in long lasting eruptions of composite volcanoes.

7. Summary and conclusion

The Hamid Basin accumulate d a thick fluvio-lacustrine fill inwhich intermedi ate and felsic volcanic edifices and their deposits can overlap with each other and with the sediments produced bythe background sedimentation . This basin deposit is underlain byandesitic lava flows and volcanicl astics at the base (lower succession), changing to dacitic to rhyolitic pyroclast ics density current deposits at the top, suggestive of a magma chamber with a felsic roof zone and a more mafic lower portion (Duritt, 1983; Druitt et al., 1999; Spark and Wright, 1979 ). Geochem ical data show that these deposits represent a unimodal nature with eruptive products belonging to a typical calc-alkalin e andesite to rhyolite and a crustal influenced intra-contin ental felsic lineage. The stratigraphy ofthe HVSS can be divided into a lower and an upper succession,based on field characterist ics. The lower Hamid succession is composed of alluvial sediments , coherent lava flows, and a variety ofvolcaniclast ic rocks related to explosive volcanism. The different lithologies indicate that the early stage of volcanism was effusive,which later changed to explosive activity, producing porphyritic lava flows and pyroclastic rocks in a subaerial environment. The upper Hamid succession includes pyroclastic density current deposits and feeder-dy kes/sills that are more volumino us than the lava flows. Their deposits are more typical for caldera volcanoes such as welded pumiceous silicic pyroclastic density current deposits and therefore could be referred to as a ‘‘mini caldera’’.The facies analysis of the HVSS led to the subdivision of fourteen lithofacies types, grouped into seven lithofacies associations . Their depositional sequence shows that an increase in felsic volcanism occurred at the NW part, represented by a high rate of volcanicl astic sediments supply. Basin evolution started with a thick, massive and clast-supporte d conglomerate of alluvial fan facies (phase 1). Itcontinued with sandy sheet floods braided river and lacustrine deposits (phase 2), and followed by coherent lava flows, phreatomagmatic deposits and lithophys ae-rich ignimbri tes (phase 3).

The upper succession of the Hamid area is dominate d by pyroclastic density current deposits, starting with a thick package of co- ignimbri te breccias lag deposits and vitric tuffs of surge facies,the latter related to phreatomagm atic vents (Phase 4). Final caldera collapse occurred shortly following emplacement of a thick succession of incipiently to densely welded ignimbrites (phase 5), with intercala tion of coarse bedded conglomerates presumably in response to evacuation of the magma chamber and final collapse ofthe magma chamber roof. The last stage of silicic volcanism was mostly intrusive, producing the subvolcani c quartz-fel dspar porphyry facies that was less volumino us than the lava flow facies.

The Hamid eruptive products described here are accumulate ddiscordan tly as one of the Late Neoprote rozoic volcano-sed imentary succession of the Dokhan Volcanics and Hammamat Group in the NED. Hamid area is a unique volcano; however, its eruptive sequence is more typical of those eruptions associated with large-volume silicic composition volcanoes with significant inter- eruptive periods. The compositional variations recorded in each eruptive phase (basal and capping) indicate a complex magmatic feeding and tapping system, typical for unimodal style of polygenetic and polymagmati c volcanism, usually associated with the prolonge d and complex eruption history of a volcanic system. Such volcanic systems can either consist of a single volcanic edifice, ornested and/or overlapping volcanic edifices such as stratovol ca- noes (e.g., Kereszturi et al., 2010 ; Lexa et al., 2010 ). Therefore Hamid area is inferred to be an isolated polygene tic unimoda lvolcanics .

Acknowled gments

The author thanks Mr. H. Khamis for his assistance in the fieldand Drs. M. Khalaf and S. Afifi for their help in carrying out the chemical analyses by XRF, at the Nuclear Material Authority, Egypt.Two anonymons reviewers have improved the quality of the paper and are gratefully acknowledged .

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http://dx.doi.org/10.1029/2001GL014372

http://dx.doi.org/10.1029/2001GL014372

Other uses, including reproduction and distribution, …scholar.cu.edu.eg/?q=ezzeldinkhalaf/files/aes_1838.pdftailed structural study of the Eastern Desert of Egypt,Greiling et al.

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