Origin and emplacement of Archean ophiolites of the central orogenic belt, North China craton

Precambrian Ophiolites and Related RocksEdited by Timothy M. KuskyDevelopments in Precambrian Geology, Vol. 13(K.C. Condie, Series Editor) 223© 2004 Elsevier B.V. All rights reserved.

Chapter 7

ORIGIN AND EMPLACEMENT OF ARCHEAN OPHIOLITES OFTHE CENTRAL OROGENIC BELT, NORTH CHINA CRATON

TIMOTHY M. KUSKYa, JAINGHAI LIb, ADAM GLASSc ANDX.N. HUANGb

aDepartment of Earth and Atmospheric Sciences, St. Louis University,St. Louis, MO 63103, USAbDepartment of Geological Sciences, Peking University, Beijing, ChinacDepartment of Geology, Cardiff University, Scotland, UK

Understanding Archean crustal and mantle evolution hinges upon identification andcharacterization of oceanic lithosphere. We have discovered and reported a complete, al-beit dismembered and metamorphosed Archean ophiolite sequence in the North ChinaCraton. Banded iron formation structurally overlies several tens of meters of variably de-formed pillow lavas and mafic flows. These are in structural contact with a 2 km thickmixed gabbro and sheeted dike complex with gabbro screens, exposed discontinuouslyalong strike for more than 20 km. The dikes consist of metamorphosed diabase, basalt,hb-cpx-gabbro, and pyroxenite. Many have chilled margins developed on their NE sides,indicating one-way chilling. The dike/gabbro complex is underlain by several kilometers ofmixed isotropic and foliated gabbro, which develop compositional layering approximately2 km below the sheeted dikes, and then over several hundred meters merge into stronglycompositionally layered gabbro and olivine-gabbro. The layered gabbro becomes mixedwith layered pyroxenite/gabbro marking a transition zone into cumulate ultramafic rocksincluding serpentinized dunite, pyroxenite and wehrlite, and finally into strongly deformedand serpentinized olivine and orthopyroxene-bearing ultramafic rocks interpreted as de-pleted mantle harzburgite tectonites. A U/Pb zircon age of 2.505 Ga from gabbro of theDongwanzi ophiolite makes it the world’s oldest recognized, laterally-extensive completeophiolite sequence. Characteristics of this remarkable ophiolite may provide the best con-straints yet on the nature of the Archean oceanic crust and mantle, and offer insights to thestyle of Archean plate tectonics and global heat loss mechanisms.

The Dongwanzi ophiolite is one of the largest well-preserved greenstone belts in theCentral Orogenic belt that divides the North China craton into eastern and western blocks.More than 1000 other fragments of gabbro, pillow lava, sheeted dikes, harzburgite, andpodiform-chromite bearing dunite occur as tectonic blocks (tens to hundreds of meterslong) in a biotite-gneiss and BIF matrix, intruded by tonalite and granodiorite, in theZunhua structural belt. Blocks in this metamorphosed Archean ophiolitic mélange pre-serve deeper levels of oceanic mantle than the Dongwanzi ophiolite. The ophiolite-related

DOI: 10.1016/S0166-2635(04)13007-7

224 Chapter 7: Origin and Emplacement of Archean Ophiolites of the Central Orogenic Belt

mélange marks a suture zone across the North China Craton, traced for more than 1600 kmalong the Central orogenic belt. Many of the chromitite bodies are localized in dunite en-velopes within harzburgite tectonite, and have characteristic nodular and orbicular chromitetextures, known elsewhere only from ophiolites. The chromites have variable but highchrome numbers (Cr/Cr + Al = 0.74–0.93) and elevated P, also characteristic of supra-subduction zone ophiolites. The high chrome numbers, coupled with TiO2 wt% < 0.2 andV2O5 wt% < 0.1 indicate high degrees of partial melting from a very depleted mantlesource and primitive melt for the chromite. A Re-Os model age from the chromites indi-cates an age of 2547 ± 10 Ma (Kusky et al., 2004), showing that they are the same ageas the Dongwanzi ophiolite. The ultramafic and ophiolitic blocks in the Zunhua mélangeare therefore interpreted as dismembered and strongly deformed parts of the Dongwanziophiolite. We suggest the name “Dongwanzi ophiolite belt” for these rocks.

Neoarchean (2.50 Ga) high-pressure granulites form a belt more than 700 km long alongthe western side of the Central Orogenic Belt. Several Neoarchean sedimentary basinsconsisting of conglomerate, graywacke and shale are located along the eastern side of theCentral Orogenic Belt, and are interpreted as remnants of a foreland basin. The three beltsrecord the Neoarchean subduction and collision between the western and eastern blocks ofthe North China Craton.

1. INTRODUCTION

Understanding early Earth evolution and obtaining reliable geologic markers of tem-poral changes in specific tectonic environments has proven elusive. Remarkably few con-straints exist on the nature of Archean oceanic crust and mantle, and even whether or notEarth lost heat in the Archean by plate tectonic mechanisms such as creating and cool-ing oceanic lithosphere, as it does today. Although many arguments have been made tosupport the operation of plate tectonics in the Archean, particularly in the Slave, Superior,Yilgarn, Kaapvaal, and Zimbabwe cratons, no well-documented and complete ophioliteswere known from the geological record until recently. We reported a complete, thoughdismembered and metamorphosed Archean ophiolite sequence in a Neoarchean orogenicbelt (2.60–2.50 Ga) of the North China Craton (Kusky et al., 2001). This remarkable ophi-olite may offer the best constraints yet on the nature of the late Archean oceanic crustand mantle. The ophiolite is associated with an intensely sheared linear belt of ophioliticmélange in which blocks of pillow lava, diabase dikes, gabbro, dunite, harzburgite, andpodiform chromitite occur in a biotite-gneiss and banded iron formation (BIF) matrix (Liet al., 2002). Many parts of the mélange belt were intruded by tonalitic and granodioriticmagmas, then deformed and metamorphosed again. In this contribution we document thefield, structural, geochronological, and mineralogical characteristics of this ophiolite andrelated mélange. This contributes to a better understanding of Archean crustal and mantleevolution, which is helpful to document the very way in which the Earth lost heat duringearly times of high heat production.

1. Introduction 225

Ophiolites are a distinctive association of allochthonous rocks interpreted to form ina variety of plate tectonic settings, including oceanic spreading centers, back-arc basins,forearcs, and arcs (Anonymous, 1972; Moores, 1982, 2002; Nicolas, 1989; Parson et al.,1992; Sylvester et al., 1997; Dilek et al., 2000; Karson, 2001). A typical, complete ophi-olite, grades downward from pelagic sediments into a mafic volcanic complex generallymade of mostly pillow basalts, underlain by a sheeted dike complex. These are under-lain by isotropic and layered gabbro exhibiting cumulate textures, then tectonized peri-dotite, resting above a thrust fault marking the contact with underlying rock sequences.As described below, the 2505 Ma Dongwanzi ophiolite contains all of the rock asso-ciations of a typical ophiolite, and is thus a complete ophiolite sensu stricto. Thereare many variations in the rock sequence preserved in ophiolites, some related to ini-tial variations between the crustal sequences formed in different tectonic settings (e.g.,Moores, 2002), and other variations related to emplacement-related deformation (e.g.,Kusky and Vearncombe, 1997). In fact, most ophiolites are not typical (Karson, 2001;Moores, 2002), and lack one or more of the above units, or have additional unusual rockunits such as silicic intrusives, hornblende-gabbros, or high-magnesium lavas. Ophiolitesare one of the hallmarks of collisional mountain belts interpreted to mark the sites alongwhich oceanic basins have closed, and therefore to demonstrate lateral motion betweenplates.

The next oldest nearly-complete ophiolites recognized in the geological record arethe 1960 Ma Jourma Complex, Finland (Kontinen, 1987; Peltonen and Kontinen, 2004),and the 1998 Ma Purtuniq ophiolite, Cape Smith Belt (Scott et al., 1991, 1992), and the1730 Ma Payson ophiolite, Arizona (Dann, 1991, 1997a, 1997b, 2004). The apparent lackof complete or nearly complete ophiolites in the older geological record has promptedmany theories that plate tectonics may have operated in fundamentally different waysin Earth’s early evolution (Karson, 2001; Sleep and Windley, 1982; Bickle et al., 1994;Abbott, 1996; Moores, 2002). However, numerous Archean greenstone belts contain twoor more parts of the full ophiolite sequence (Harper, 1985; de Wit et al., 1987, 1992;Kusky, 1987, 1990, 1991; Brandl and de Wit, 1997). This led others to theorize that partsof many greenstone belts may be dismembered ophiolites formed in a manner analogous toyounger dismembered ophiolites (Kusky and Polat, 1999). Similarly, most Proterozoic andPhanerozoic ophiolites including the Jourma, Payson, and Purtuniq complexes are dismem-bered, or partial sequences (Kontinen, 1987; Scott et al., 1991; Kröner, 1985; Berhe, 1990;Johnson et al., 2004; Stern et al., 2004). The documentation of a complete and laterally-extensive Archean ophiolite sequence 500 Ma older than the best Proterozoic examples isimportant to our understanding of Earth evolution and how the planet may have lost heatduring early times of high heat production. This discovery also has implications for thedevelopment of Earth’s early biosphere, since some of the chert and BIF probably formednear sea-floor hydrothermal vents, and may host early life forms (e.g., Rasmussen, 2000;Reed, 2002).

The main obstacle to recognizing Archean ophiolitic rocks is polyphase deformationand metamorphism, which may rework them into belts of mafic schist or gneiss. It isimportant to delineate sites of preservation of ophiolites during regional tectonic analy-

226 Chapter 7: Origin and Emplacement of Archean Ophiolites of the Central Orogenic Belt

sis, and to identify ancient ophiolites by their characteristic rock assemblages, includingpillow lavas, sheeted dikes, layered gabbro, and podiform chromitite. Detailed structuralmapping is important for documenting relationships between ancient ophiolites and coun-try rocks. Geochemistry, petrology, and magmatic stratigraphy can further resolve the earlytectonic setting of the ophiolite, whether arc, suprasubduction zone (SSZ), or mid-oceanridge (MOR).

In this paper, we review our current state of knowledge of the field, structural,geochronological, mineralogical, and chromite chemistry of this ophiolite and relatedmélange. We assess what these data mean in terms of Archean crustal and mantle evo-lution, thermal state of the early Earth, and ideas about Precambrian plate tectonics. Wediscuss the implications these observations have for our understanding of fundamentalproperties of Archean crust and mantle, and planetary evolution. Understanding Archeanoceanic processes in the North China craton provides a valuable contrast with similarprocesses recorded in younger ophiolites, indicating how mid-ocean ridge processes mayhave evolved from a period of high heat production to one of lower heat production.

The current review presented in this contribution represents the results of the first-phasereconnaissance and 1:100,000 scale mapping, and is presented as our team embarks on anew effort to complete more detailed mapping over the course of the next several years.Therefore, some of the details and outstanding questions will hopefully be answered byfuture work, and some interpretations are likely to change as more data becomes available.

2. REGIONAL GEOLOGY AND TECTONIC DIVISIONS OF THE NORTH CHINACRATON

The North China Craton (Fig. 1) occupies about 1.7 million square kilometers in north-eastern China, Inner Mongolia, the Yellow Sea, and North Korea. It is bounded by theQinling-Dabie Shan orogen to the south, the Yinshan-Yanshan orogen to the north, theLongshoushan belt to the west and the Qinglong-Luznxian and Jiao-Liao belts to theeast (Bai and Dai, 1996, 1998). The North China Craton (NCC) includes a large areaof intermittently-exposed Archean crust (Fig. 1), including ca. 3.8–2.5 Ga gneiss, TTG,granite, migmatite, amphibolite, ultramafite, mica schist and dolomitic marble, graphiticand sillimanititc gneiss (khondalites), banded iron formation (BIF), and metaarkose (Jahnand Zhang, 1984a, 1984b; Bai et al., 1992; Wu et al., 1998; Zhao, 1993; Jahn et al., 1987;Bai, 1996; He et al., 1991, 1992; Shen et al., 1992; Wang et al., 1997). The Archean rocksare overlain by the 1.85–1.40 Ga Mesoproterozoic Changcheng (Great Wall) system (Liet al., 2000a, 2000b). In some areas in the central part of the North China Craton, 2.40–1.90 Ga Paleoproterozoic sequences deposited in cratonic graben are preserved (Kusky andLi, 2003).

We divide the North China craton into two major blocks (Fig. 2) separated by theNeoarchean Central Orogenic Belt in which virtually all U-Pb zircon ages (upper in-tercepts) fall between 2.55 and 2.50 Ga (Kröner et al., 1998, 2002; Li et al., 2000b;Wilde et al., 1998; Zhao, 2001; Kusky et al., 2001). The Western Block, also known as the

2. Regional Geology and Tectonic Divisions of the North China Craton 227

Fig. 1. Map of the North China craton, showing the distribution of Archean, Proterozoic, andyounger rocks. Compiled from numerous sources.

Ordos Block (Bai and Dai, 1998; Li et al., 1998), is a stable craton with a thick mantle root,no earthquakes, low heat flow, and a lack of internal deformation since the Precambrian.In contrast, the Eastern Block is atypical for a craton in that it has numerous earthquakes,high heat flow, and a thin lithosphere reflecting the lack of a thick mantle root. The NorthChina Craton is one of the world’s most unusual cratons in that it had a thick tectosphere(subcontinental lithospheric mantle) developed in the Archean, which was present throughthe Ordovician as shown by deep xenoliths preserved in Ordovician kimberlites (Gao etal., 2002). However, the eastern half of the root appears to have delaminated or otherwisedisappeared during Paleozoic, Mesozoic, or Cenozoic tectonism. This is demonstrated byTertiary basalts that bring up mantle xenoliths of normal “Tertiary mantle” with no evi-dence of a thick root (e.g., Menzies et al., 1993; Griffin et al., 1998; Zheng et al., 1998;Gao et al., 2002). The processes responsible for the loss of this root are enigmatic butare probably related to the present day high-heat flow, Phanerozoic basin dynamics andorogenic evolution.

228 Chapter 7: Origin and Emplacement of Archean Ophiolites of the Central Orogenic Belt

Fig. 2. Major tectonic divisions of the North China craton, into the Western and Eastern blocks,Central orogenic belt, and Paleoproterozoic orogenic belt (after Li et al., 2000a).

The Central Orogenic belt (COB) includes belts of TTG, granite, and supracrustal se-quences metamorphosed from granulite to greenschist-facies. It can be traced for about1,600 km from west Liaoning to west Henan (Fig. 2). Widespread high-grade regionalmetamorphism including migmatization occurred throughout the Central Orogenic beltbetween 2.60 and 2.50 Ga, with final uplift of the metamorphic terrain at ca. 1.9–1.80 Gaassociated with extensional tectonism (Li et al., 2000a) or a collision on the northernmargin of the NCC (Kusky and Li, 2003). Amphibolite to greenschist grade metamor-phism predominates in the southeastern part of the COB (Fig. 2), but the northwest-ern part of the COB is dominated by granulite-facies to amphibolite facies rocks, in-cluding some high-pressure assemblages (10–13 kbars at 850 ± 50 ◦C; Li et al., 2000b;Zhao et al., 2001; see additional references in Kröner et al., 2002). The high-pressure as-semblages can be traced for more than 700 km along a linear belt trending ENE. Inter-nal (western) parts of the orogen are characterized by thrust-related horizontal foliations,flat-dipping shear zones, recumbent folds, and tectonically interleaved high-pressure gran-ulite migmatite and metasediments. It is widely overlain by sediments deposited in grabenand continental shelf environments, and intruded by several dike swarms (2.40–2.50 Ga,1.80–1.90 Ga). Several large anorogenic granites with ages of 2.20–2.00 Ga are identifiedwithin the belt. Recently, two linear units have been documented within the belt, includinga high-pressure granulite belt in the west (Li et al., 2000a) and a foreland-thrust fold beltin the east (Li et al., 2002). The high-pressure granulite belt is separated by normal-senseshear zones from the western block, which is overlain by thick metasedimentary sequences

3. Ophiolites of the Central Orogenic Belt, North China Craton 229

(khondalite, younger than 2.40 Ga, and metamorphosed at 1,862.7 ± 0.4; A. Kröner, per-sonal communication, 2003).

The Hengshan high-pressure granulite belt is about 700 km long, consisting of sev-eral metamorphic terrains, including the Hengshan, Huaian, Chengde, and west Liaoningcomplexes (Fig. 2). The HPG commonly occurs as inclusions within intensely shearedTTG (2.60–2.50 Ga) and granitic gneiss (2.50 Ga), and are widely intruded by K-granite (2.20–1.90 Ga) and mafic dike swarms (2.40–2.45 Ga, 1.77 Ga) (Li et al., 2000b;Kröner et al., 2002). Locally, khondalite and turbiditic slices are interleaved with the high-pressure granulite rocks, suggesting thrusting. The main rock type is garnet-bearing maficgranulite with characteristic Pl-OPX corona around the garnet (Figs. 3a–c), which showrapid exhumation-related decompression. The isothermal decompressive P-T-t path can bedocumented within the rocks, the peak PT is in the range of 1.2–1.0 GPa, at 700–800 ◦C. Atleast three types of REE patterns are shown by mafic rocks of the high-pressure granulites,from flat to LREE-moderately enriched, indicating a tectonic setting of active continentalmargin or island arc (Li et al., 2002). The high-pressure granulites were formed throughsubduction-collision, followed by rapid rebound-extension, recorded by mafic dike swarmsof 2.50–2.40 Ga and graben-related sedimentary sequences in the Wutai Mountain-TaihangMountain areas (Kusky and Li, 2003).

The Qinglong foreland basin and fold-thrust belt (Fig. 2) is north to northeast-trending,and is now preserved as several relict folded sequences (Qinglong, Fuping, Hutuo, andDengfeng). Its general sequence from bottom to top can be further divided into three sub-groups of quartzite-mudstone-marble, turbidite, and molasse, respectively (Fig. 4). Thelower subgroup of quartzite-mudstone-marble is well preserved in central sections of theQinglong foreland basin (Taihang Mountain), with flat-dipping structures, interpreted as apassive margin developed prior to 2.5 Ga on the eastern block. It is overlain by lower-gradeturbidite and molasse type sediments (Figs. 3d–h). The western margin of Qinglong fore-land basin is intensely reworked by thrusting and folding, and is overthrust by the overlyingorogenic complex (TTG gneiss, ophiolitic, accretionary sediment). To the east its defor-mation becomes weaker in intensity (Fig. 5). The Qinglong foreland basin is intruded by agabbroic dike complex consisting of 2.40 Ga diorite, and overlain by graben-related sed-iments and flood basalts. In the Wutai and North Taihang basins, many ophiolitic blocksare recognized along the western margin of the foreland thrust-fold belt. These consistsof pillow lava, gabbroic cumulates, and harzburgite. The largest ophiolitic thrust compleximbricated with foreland basin sedimentary rocks is up to ten kilometers long, preservedin the Wutai-Taihang Mountains (Wan et al., 1998).

3. OPHIOLITES OF THE CENTRAL OROGENIC BELT, NORTH CHINA CRATON

We have identified several dismembered Archean ophiolites in the Central orogenicbelt, including some in Lioning Province, at Dongwanzi, north of Zunhua, and at WutaiMountain (Fig. 2). The best studied of these are the Dongwanzi and Zunhua ophioliticterranes, which are the main focus of this review. The Zunhua Structural belt (ZSB) of

230 Chapter 7: Origin and Emplacement of Archean Ophiolites of the Central Orogenic Belt

Fig. 3. Photographs of rocks in the Central Orogenic Belt. (a) High-grade gneiss with large gar-net porphyroblasts exhibiting retrograde high-P granulite textures, Hengshan Complex; (b) maficboudin in granulite gneiss, Hengshan Complex; (c) garnet-rich granulite gneiss, Hengshan Complex;(d) foreland basin flysch from the Qinglong basin; (e) conglomerate of the Hutuo Group, WutaiMountains; (f) graywacke/shale beds in Hutuo Group, interpreted as flysch; (g) west-dipping flyschof Hutuo Group, Wutai Mountains (small temple for scale is 1.5 m tall); (h) thrust slice of arkose(Lower Wutai Group), interpreted as stable continental margin imbricated with foreland basin flysch,Wutai Mountains.

3. Ophiolites of the Central Orogenic Belt, North China Craton 231

Fig. 4. Stratigraphic column showing Group and Formation names for different rock assemblagesin the Wutai Mountain area, along with their tectonic interpretation. Modified after Tian (1991).Abbreviations for formation names: BYK, Banyukou Formation; JGK, Jingangku Formation; ZW,Zhuangwang Formation; WX, Wenxi Formation; BZY, Baizhiyan Formation; HMY, HongmenyanFormation.

the Eastern Hebei Province (Fig. 6) preserves a cross section through most of the north-eastern part of the Central Orogenic belt. This belt is characterized by highly strainedgneiss, banded iron formation (BIF), 2.60–2.50 Ga greenstone belts and mafic to ultra-mafic complexes in what Li et al. (2002) interpret as a high-grade ophiolitic mélange.The belt is intruded by widespread 2.60–2.50 Ga tonalite-trondhjemite gneiss (TTG),granites (2.50 Ga) and is cut by ductile shear zones (Li et al., 2000a; Wu et al., 1998;Kusky et al., 2001). The Neoarchean high-pressure granulite belt (Chengde-HengshanHPG) strikes through the northwest part of the belt. The Zunhua structural belt is thrustover the Neoarchean Qianxi-Taipingzhai granulite-facies terrain (2.50 Ga), consisting ofenderbitic to charnockitic gneiss forming several small dome-like structures southeast ofthe Zunhua belt. The Zunhua structural belt clearly cuts across the dome-like Qian’an-

232 Chapter 7: Origin and Emplacement of Archean Ophiolites of the Central Orogenic Belt

Fig. 5. Schematic structural cross section through the Hengshan Complex, Wutai Complex, andFuping Complex, showing Hengshan thrust over Wutai, and Wutai thrust over the Fuping Complex.Modified after Tian (1991).

Qianxi structural patterns to the east. The Qian’an granulite-gneiss dome (3.80–2.50 Ga)forms a large circular dome in the southern part of the area (Fig. 6), and is composed oftonalitic-trondhjemitic gneiss, and biotite-granite. Mesoarchean (2.80–3.00 Ga) and Pale-oarchean (3.50–3.85 Ga) supracrustal sequences (Shen et al., 1992) outcrop in the easternpart of the region (Fig. 6). The Qinglong Neoarchean (2.70–2.50) amphibolite to green-schist facies supracrustal sequence strikes through the center of the area, and is interpretedhere to be a foreland fold-thrust belt, intruded by large volumes of 2.40 Ga diorite in theeast. The entire North China craton is widely cut by at least two Paleoproterozoic maficdike swarms (2.50–2.40 Ga, 1.80–1.70 Ga), associated with regional extension (Li et al.,2000a). Mesozoic-Cenozoic granite, diorite, gabbro and ultramafic plugs occur throughoutthe NCC, and form small intrusions in some of the belts.

The largest well-preserved sections of the Dongwanzi ophiolite are located approxi-mately 200 km NE of Beijing in the northeastern part of the Zunhua structural belt, near thevillages of Shangyin and Dongwanzi (Figs. 6 and 7). It consists of prominent amphibolite-facies mafic-ultramafic complexes (Fig. 8) in the northeast sector of Zunhua structural belt,previously mapped as belts of amphibolite-facies layered intrusion and associated rocks(Shen et al., 1992; Zhang et al., 1986, 1991). The southern end of the Dongwanzi ophiolitebelt near Shangyin is complexly faulted against granulite-facies gneiss, with both thrustfaults and younger normal faults present. The main section of the ophiolite dips steeply

3. Ophiolites of the Central Orogenic Belt, North China Craton 233

Fig. 6. Map of part of the Eastern Hebei Province, North China Craton, showing the location andtectonic setting of the Dongwanzi ophiolite. Note abundant similar belts of amphibolite-granulitegrade metabasites in the Zunhua structural belt. The Dongwanzi ophiolite is intruded by the (de-formed) ca. 2.4 Ga Cuizhangzi gneiss (Wang and Zhang, 1995) and has a Sm-Nd whole rock isochronmodel age of 2756 ± 177 Ma (Wu and Geng, 1991). The ophiolite is approximately the same ageas the 2.7–2.5 Ga Qinglong supracrustal sequence, consisting of interbedded metagraywackes andshales.

NW, is approximately 50 km long, and is 5–10 km wide (Figs. 6 and 7). We have obtaineda U/Pb-zircon age of 2505 ± 2 Ma for two gabbro samples from the Dongwanzi ophiolite(Kusky et al., 2001). Additional preliminary U/Pb zircon geochronology from dikes cuttingthe Central Belt of the Dongwanzi ophiolite has revealed the presence of a fraction of cleareuhedral zircons with ages falling between 300 and 200 Ma (R. Tucker, personal commu-nication, 2002). Zhai et al. (personal communication, 2002) report an 40Ar/39Ar plateau(metamorphic) age of 1.8 Ga from amphiboles from the same outcrop that the 300 MaU/Pb ages come from, which led us to re-examine the central belt of the Dongwanzi ophi-olite. We have determined that parts of the central belt are intruded by a mafic/ultramaficMesozoic pluton with related dikes (Fig. 9), and that this pluton is slightly larger than we

234 Chapter 7: Origin and Emplacement of Archean Ophiolites of the Central Orogenic Belt

Fig. 7. Modified reconnaissance map of the Dongwanzi ophiolite, showing belts of pillow lavas,sheeted dikes, gabbro complex, cumulate ultramafics, and harzburgite tectonite. Modified after Kuskyet al. (2001).

originally suggested (Kusky et al., 2001; Huson et al., 2004). More work is needed to de-termine if the central belt is intruded by even more younger rocks related to the Dushancomposite plutonic complex.

The base of the ophiolite is strongly deformed, and intruded by the 2391 ± 50 MaCuizhangzi diorite-tonalite complex (Zhang et al., 1986). The Dongwanzi ophiolite is as-sociated with a number of other amphibolite-facies belts of mafic plutonic and extrusiveigneous rocks (Fig. 6) in the Zunhua structural belt. These mafic to ultramafic slices andblocks can be traced regionally over a large area from Zunhua to West Liaoning (about200 km). Much of the ZSB is interpreted as a high-grade ophiolitic mélange (Li et al.,

3. Ophiolites of the Central Orogenic Belt, North China Craton 235

Fig. 8. Photographs of rocks of the Dongwanzi ophiolite. (a) Pillow lavas, (b) interpillow chert andhyaloclastite, (c) pillow lava, (d) high-level sheeted dikes, (e) layered gabbro, (f) thinly layered gab-bro.

2002), with numerous tectonic blocks of pillow lava, BIF, dike complex, gabbro, dunite,serpentinized harzburgite, and podiform chromitite in a biotite-gneiss matrix (Fig. 6), in-truded extensively by tonalite and granodiorite. Cross-cutting granite has yielded an ageof 2400 ± 15 Ma (Wu et al., 1998). We suggest that the blocks in the mélange correlatewith the Dongwanzi and other ophiolitic fragments in the ZSB. This correlation is sup-ported by recent Re-Os age determinations on several of these blocks, revealing that theyare 2.54 Ga old. Existing age constraints need to be supplemented by new U/Pb ages on

236 Chapter 7: Origin and Emplacement of Archean Ophiolites of the Central Orogenic Belt

other units of the ophiolite, cross-cutting granites, and other mafic belts in mélange thatmay be correlative with the Dongwanzi ophiolite.

4. FIELD DESCRIPTION OF OPHIOLITIC UNITS

The upper part of the Dongwanzi ophiolite consists of pillow lavas, pillow breccias,and interflow sedimentary rocks including rare chert, banded iron formation (with localsulfides), and metapelites (Figs. 7 and 8). There are other extensive thick fault-boundedunits of banded iron formation located near the top of the ophiolite. The pillow lavas,pillow breccias, and interpillow sediments are rarely well-preserved, extensively altered tochlorite and hornblende schists, and most are significantly deformed. Pillows are in manycases difficult to recognize (Kusky and Li, 2002), and are also interbedded with moremassive flows, and cut by sills (Fig. 8). Where well preserved, the pillows show typicalcuspate lower boundaries and lobate upper contacts, defining stratigraphic younging. Insome places pillows are delineated by 2–3 cm thick epidote-rich selvages surroundingfine-grained, 0.5–1.0 m wide pillow cores. In some locations metamorphism has causedthe pillow cores to become coarsely-crystalline with large hornblende crystals overgrowingearlier magmatic textures. Elongate pillow lava tubes are preserved in a few places, such as1 km west of the village of Shang Yin (Fig. 7). There, the pillow lavas are faulted againstgranulite facies tonalitic gneiss (ca. 2.50 Ga), and locally occur as inclusions in a youngertonalite (Fig. 8). Interpillow chert is also present but rare at Shang Yin. The pillow lavashave flat MORB-like to suprasubduction zone types of REE signatures (Kusky and Li,2002; Kusky et al., 2004).

The dike complex of the Dongwanzi ophiolite is at least 2 km thick and extends discon-tinuously along strike for more than 20 km (Figs. 8 and 9). Several, hundred meter longoutcrops exhibit 100% dikes, but in most places the dikes intrude gabbro, and exhibit mutu-ally cross-cutting relationships with the gabbro. In contrast, the largest-known sheeted dikecomplex in a Proterozoic ophiolite is the 700 m long by 150 m wide dike complex in theca. 1998 Ma Purtuniq ophiolite (Scott et al., 1991). The Dongwanzi dike complex exhibitspredominant one-way chilling of diabase dikes that are chilled on their NE sides, but areintruded by other dikes along their SW margins. More than 70% of the dikes in the sheetedcomplex exhibit one-way chilling, with hundreds of dikes measured. Gabbro screens arecommon in the dike complex, with diabase dikes exhibiting double and single-chill mar-gins chilled against the gabbro. The number and thickness of gabbro screens generallyincreases downward in the ophiolite, marking a gradual transition from the dike complex

Fig. 9. Geological map of the lower part of the transition zone from layered gabbro to cumulateultramafic rocks in the Dongwanzi ophiolite. Note also the lenses of harzburgite tectonite, gabbro,and the intrusion of two sets of mafic dikes that appear related to the sheeted dike complex that isstratigraphically above this unit. Large diorite intrusion on SW side of map area also intrudes theChangcheng system (1.85–1.40 Ga). Modified after Kusky et al. (2001).

4. Field Description of Ophiolitic Units 237

238 Chapter 7: Origin and Emplacement of Archean Ophiolites of the Central Orogenic Belt

Fig. 10. Photographs of mutually cross cutting gabbro and diabase dikes from the central belt.

into the gabbroic fossil magma chamber. In many locations the gabbro also exhibits inter-nal chill margins, with coarse-grained gabbro becoming finer-grained toward chill marginsthat also show a preferential one-way chilling along their NE margins, and intrusion byother gabbroic dikes along their SW margins (Fig. 8). This transition from dikes to gabbrois reminiscent of parts of the Semail ophiolite, where dikes grade into and were probablyfed from magmatically layered gabbro (Nicolas and Boudier, 2000). The gabbro is variablein mineralogy and texture, with some phases being very feldspar rich, while others showabundant, several cm long crystals of clinopyroxene largely altered to hornblende. Themineralogical composition of the gabbroic rocks ranges from diorite through true gabbro,hornblende gabbro, hornblende-pyroxene gabbro, to pyroxenite. Olivine is extremely rare.Basaltic-diabasic dikes commonly cut the gabbro, but in places gabbro also cuts and in-cludes xenoliths of diabasic dikes suggesting that the gabbro and diabase are synchronousor comagmatic phases, and both are related to magmatic extension. In places, diabase dikescan be shown to root in the gabbro complex, as they form from the coalescence of manysmall melt veinlets that are disseminated through the gabbro (Fig. 10). The gabbro and dia-base dike complex is cut by abundant epidote-clinozoisite veins, and massive sulfides (Cu,Fe) similar to sea-floor metamorphism veins in other ophiolites (Sylvester et al., 1997;Harper, 1999). The presence of these veins in the gabbro gives us an idea of the depthto which the sea floor hydrothermal system extended in this ophiolite. Other pyroxenite

4. Field Description of Ophiolitic Units 239

Fig. 11. Photograph of cross-cutting pyroxene-apatite dike that has yielded a preliminary date of ca.300 Ma (R. Tucker, personal communication, 2002), from the central belt.

through granodiorite veins and dikes also intrude the ophiolite and dike complex (Fig. 11),but these have yielded a preliminary U/Pb age of 300 Ma (R. Tucker, personal communica-tion) and are therefore probably related to the nearby Paleozoic-Dushan plutonic complex,and not the ophiolite as originally proposed (Kusky et al., 2001). Accordingly, we haverevised our map of the central (Fig. 9) belt to show as accurately as known the distributionof younger intrusive rocks.

Some normal faults have developed along the sheeted dike margins, which tilted thegabbroic screens between dikes. To the east, metagabbro and pyroxenite locally occurs asinclusions within Neoarchean gneiss. Sheeted dikes and gabbroic rocks are locally inter-leaved with the Mesoproterozoic Chang Cheng Series quartzite (1.85–1.40) and repeatedby Mesozoic thrusts.

The angular discordance between two sets of gabbro dikes observed in many outcropsis similar to dike relationships observed in many younger ophiolites. The two sets of dikescould be formed before and after rotation of normal fault blocks during sea-floor exten-sional tectonics. Alternatively, the second set of dikes might be related to a younger episodeof arc or other type of volcanism. A third possibility is that some of the dikes are Meso-zoic, as confirmed by the 300 Ma preliminary age obtained from a dike at Dongwanzivillage. We recognize so far that the Dongwanzi ophiolite is cut by at least four deformeddike swarms (three mafic dike swarms, including pyroxenite dikes, the two sets of shearedgabbroic dikes, and diabase dikes), followed by one granodiorite-pyroxenite suite (dated at∼ 300 Ma by R. Tucker).

240 Chapter 7: Origin and Emplacement of Archean Ophiolites of the Central Orogenic Belt

Since we now recognize that parts of the ophiolite are cut by Mesozoic granodiorite,gabbro, and pyroxenite, and one undeformed mafic to ultramafic dike swarm also intrudesthe Mesoproterozoic sequence, additional field efforts need to be aimed at differentiatingMesozoic from older Archean units through detailed mapping and geochronology. Whenadditional work aimed at more-clearly identifying the ages and characteristics of specificdike swarms is complete, some fundamental questions about the Archean oceanic thermalgradient may be answered. For instance, how do the thickness and textures of the Archeandikes vary with depth, and are these changes related to the thermal profile in the Archeanoceanic environment? This is especially critical because temperature gradients change dra-matically from the base of the dike complex to the gabbro unit in ophiolites (Nicolas,1989).

The gabbro complex of the Dongwanzi ophiolite is up to 5 km thick and grades down-ward from isotropic gabbro, into faintly layered gabbro, into strongly layered gabbro, theninto a mafic/ultramafic transition zone of mixed layered gabbro and cumulate ultramaficrocks (Figs. 7–9). The gabbro complex is metamorphosed to amphibolite facies. Layeringin the gabbro ranges from several cm-thick discontinuous layers, through decameter thicklayers, to faint layers several meters thick. These include cm to meter thick alternationsbetween clinopyroxene and plagioclase rich layers, with individual layers varying in com-position between anorthosite and clinopyroxenite (Fig. 9). In some locations in the thinlylayered gabbro the gabbro exhibits folds that could be slumps resulting from layers thataccumulated along walls of the magma chamber sliding down to the chamber floor. Thin,relatively homogeneous gabbro dikes a few cm to decimeters thick intrude the layeredgabbro at small angles to the compositional layering, and pods of gabbro-pegmatite occurlocally.

The gabbro and mixed cumulate gabbro and ultramafic parts of the ophiolite gradegradually downward into a unit of about 50% coarse-grained gabbro and 50% pyroxenite(Fig. 9). This unit represents the “transition zone” between mafic rocks above, and ultra-mafic cumulates and depleted mantle below. These layers are locally cut by metabasalticdikes with unknown ages, but are mineralogically and texturally similar to mafic dikes athigher levels of the ophiolite. The leucogabbro is locally strongly foliated, and the gab-bro at this level contains many pods of ultramafic rocks. The lower part of the cumulatemafic/ultramafic complex consists almost entirely of ultramafic cumulates (Fig. 9). Or-thocumulate pyroxenite, rare dunite, harzburgite (Fig. 12), and other peridotites are inter-layered with olivine-pyroxene gabbro, and olivine gabbro layered cumulates. In a few lo-calities layers grade from a dunite base up into wehrlite and clinopyroxenite tops. Thinlayers and disseminated grains of chromite are present but rare (Fig. 9). However, wehave documented numerous podiform chromitites in 200–400 m large tectonic blocks inmélange underlying the ophiolite, traced to the southwest 60 km along strike (see Huanget al., 2004).

Small amounts of strongly tectonized and serpentinized ultramafic rocks have beenfound along the exposed base of the Dongwanzi ophiolite (Fig. 9). These rocks, show-ing field evidence for high-temperature deformation, include strongly foliated and lineateddunite and layered olivine-orthopyroxene harzburgite tectonite (Fig. 12), some containing

4. Field Description of Ophiolitic Units 241

Fig. 12. Photographs of ultramafic tectonite rocks from the southern part of the central belt.

individual septa of olivine and orthopyroxene crystals, and ultramafic mylonites. The py-roxene crystals in the peridotites are commonly aligned, and bands of chromite in duniteprobably formed during early, high-temperature deformation. The ultramafic tectonites areextensively serpentinized making identification of primary mineralogy difficult. Despitethe pervasive serpentinization, this unit is clearly more deformed than overlying units, itcontains depleted harzburgite (with orthopyroxene in contrast to clinopyroxene dominatedunits in the crustal section), and preserves evidence for early, high-temperature deforma-tion. We provisionally interpret this as the lower, residual or depleted mantle part of theophiolite, from which the overlying magmatic rocks were extracted.

242 Chapter 7: Origin and Emplacement of Archean Ophiolites of the Central Orogenic Belt

5. CONTACT RELATIONSHIPS

The base of the Dongwanzi ophiolite is an early shear zone, locally in contact withbiotite-gneiss of the Zunhua structural belt. It has been extensively intruded by the2391 ± 50 Ma Cuizhanghizi diorite-tonalite complex (Zhang et al., 1991) and othertonalitic gneiss to the east, then deformed again after intrusion of the diorite complex. Dior-ite forms dikes, layer-parallel sills, and is strongly deformed. The base of gabbro/layeredultramafic cumulates and the ultramafic tectonites has progressively more diorite intrudingit downward until the Dongwanzi ophiolite is only represented as ultramafic and garnet-amphibolite inclusions of tens of cm to several meters in scale in the dioritic to tonaliticgneiss to the east (Kusky et al., 2001). The lower contact relationships require additionaldetailed mapping, especially because the area of the Zunhua structural belt southwest ofthe Dushan granite contains a large number of structurally-bounded lozenges of gabbroand ultramafic rock that may be dismembered parts of the base of the ophiolite. Rocksof the Zunhua Structural belt therefore preserve a much more extensive record of mantleprocesses than originally estimated (Li et al., 2002).

The upper contact of the main thrust sheet containing the Dongwanzi ophiolite is alsoa fault, but it is at least in part Proterozoic or Mesozoic, as the ophiolite is imbricatedwith Mesoproterozoic quartzites. Regional relationships show that this is a Mesozoic thrustbelt that imbricates the Dongwanzi ophiolite with Mesoproterozoic sediments (Xu, 1990;Li et al., 2000a; Davis et al., 1996; Ziegler et al., 1996). Additional work is needed toseparate Mesozoic and Proterozoic structural elements from the older Archean record.

6. HIGH-GRADE OPHIOLITIC MÉLANGE IN THE ZUNHUA STRUCTURAL BELT

The Zunhua structural belt (ZSB) is mainly an amphibolite-facies terrain separated froman Archean granulite-gneiss dome (3.85–2.50 Ga) of the Eastern Block by a major thrust-sense shear zone (Figs. 6 and 13). It contains NE-striking, intensely strained gneiss and am-phibolites exhibiting tight composite folds and ductile shear zones. Various thrust slices,such as TTG gneiss, mafic plutonic rocks, supracrustal sequences, mafic volcanics, BIF,garnet-bearing gneiss, and granites are tectonically intercalated with each other (Kuskyet al., 2001; Wu et al., 1998; Wu and Zhong, 1998; Shen et al., 1992; Zhang et al., 1991;Fang et al., 1998). More than 1000 mafic to ultramafic boudins, ranging from several metersto several kilometers in length are recognized in the ZSB. Locally, layered gabbro and cu-mulate rocks form belts up to several kilometers long (Shen et al., 1992; Zhang et al., 1991;Zhang et al., 1986; Fang et al., 1998), that are intruded by ca. 2.56–2.50 Ga tonalitic gneiss(Wu and Geng, 1991; Wu et al., 1998). Mafic volcanics in the ZSB commonly have anoceanic affinity as shown by their flat REE patterns to LREE-depleted patterns, similar tobasalts from suprasubduction and mid-oceanic ridge settings (Zhao, 1993; Wu et al., 1998;Kusky and Li, 2002). These mafic and ultramafic boudins occur in a fine-grained biotite-gneiss matrix, and are interpreted as tectonic blocks in ophiolitic mélange (Li et al., 2002).Many parts of the mélange were intruded by tonalite and granodiorite, and deformed again.

6. High-Grade Ophiolitic Mélange in the Zunhua Structural Belt 243

Fig. 13. Detailed map of the western part of the Zunhua structural belt, showing the location ofpodiform chromite and dismembered ophiolitic ultramafic pods.

In the Paleoproterozoic, the ZSB was intruded by mafic dike swarms that are now meta-morphosed, and ca. 2400 ± 15 Ma granite (Wu and Geng, 1991). Finally, it was overlainunconformably by unmetamorphosed Mesoproterozoic sequences (younger than 1.85 Ga),demonstrating that major tectonothermal episodes associated with the Zunhua structuralbelt occurred before 1.85 Ga.

Mapping of ultramafic to mafic blocks within the Zunhua structural belt has resultedin the recognition of various rock types typical of oceanic crust and mantle. We interpretsome large boudins as strongly dismembered fragments of an originally more continu-ous sequence, tectonically transposed with Neoarchean biotite-gneiss, BIF, and tonalite.The boudins are preserved as blocks in a high-grade metasedimentary and tonalitic gneiss(Li et al., 2002). The Dongwanzi ophiolite may represent one such unusually large andcomplete block, particularly the main part that contains sheeted dikes, gabbro, and cumu-late ultramafic rocks. The smaller tectonic blocks and pods display typical ophiolitic rocktypes, including partly serpentinized harzburgite, peridotite tectonite, dunite, serpentinite,podiform chromitite, hornblendite, wehrlite, pyroxenite, metagabbro, cumulates and pillowlavas, massive metabasalt, and greenschist. Locally, well-preserved sheeted dikes of severalmeters width are recognized. Garnet-amphibolite blocks are also quite common. Graniteintruding the mafic to ultramafic rocks have yielded ages of 2400 ± 15 Ma age (Wu andGeng, 1991). The mafic to ultramafic blocks preserve delicate textures and structures thatare characteristic of ophiolites (Nicolas, 1989). In particular, the podiform chromitites pro-vide very important information on the nature of the Archean oceanic mantle. Commonly,

244 Chapter 7: Origin and Emplacement of Archean Ophiolites of the Central Orogenic Belt

it is generated at high-temperatures (1200–1300 ◦C), and can escape late metamorphismand intense deformation.

Ultramafic to mafic pods and tectonic blocks are stretched and occur within a stronglydeformed matrix consisting of foliated and sheared, fine-grained biotite-gneiss andhornblende-gneiss with some layers of amphibolite and BIF (Figs. 13 and 14). Theseblocks are intensely sheared and tectonically transposed along their margins. In con-trast, internal structures of the blocks commonly show distinct foliation, layering, andfold patterns, discordant to the external foliations outside the blocks in the surround-ing shear zones. Gabbro and pyroxenite boudins exhibit well-preserved relict cumulatetextures (Fig. 15). Metamorphic tectonite fabrics are well-preserved within the cores ofblocks and pods, which are defined by oriented orthopyroxene porphyroclasts, strings ofchromite, and elongated ribbons of olivine. Polyphase and isolated foliation patterns arerecognized within larger peridotite blocks. The early tectonic fabrics defined by composi-tional layering include chromite seams, disseminated chromite, oriented nodular and orbic-ular chromite, and flattened antinodular chromite. In younger ophiolite complexes, thesetextures are commonly interpreted to form during high-temperature plastic deformation inthe upper mantle associated with oceanic spreading (Nicolas and Azri, 1991; Zhao, 1993;LeBlanc and Nicholas, 1992; Leblanc, 1997; Matsumoto and Arai, 1999).

The early high-temperature tectonite fabrics in the peridotite are cut by late steeply-dipping shear zones, which are commonly parallel to tectonic contacts with country rocks.Serpentinization is concentrated along late shear zones and fractures cutting across the ear-lier foliation. Within these shear zones, ultramafic protoliths are separated into numeroussmall-scale pods and lenses of tens of centimeters to several meters scale, which are flat-tened and stretched. Many preserve asymmetric shapes. The presence of more than two setsof tectonic fabrics suggests that the harzburgitic mantle tectonites have been overprintedby late obduction-related deformation.

7. DESCRIPTION OF ROCK TYPES IN THE BLOCKS OF MÉLANGE

Peridotitesare mainly composed of serpentinized olivine, relict orthopyroxene, and mi-nor chromite and magnetite (Fig. 16). Stretched orthopyroxene grains form augen up to2–3 mm in diameter enclosed in a serpentinite matrix, and display ribbon-shaped tails.Some orthopyroxene porphyroclasts preserve embayed outlines associated with corrosionby melt. Relict olivine aggregates show elongated geometry (Fig. 16). Minor subhedral toeuhedral chromites are present. Some peridotite tectonites show penetrative foliations andstretching lineations. High-temperature metamorphic tectonite fabrics defined by alignedand stretched olivine are well-preserved in cores of blocks (Fig. 16).

Serpentinized dunitein the Zunhua structural belt typically exhibits well-defined lay-ering defined by needle-like chromite grains. Relict olivine forms extended ribbons withasymmetrical geometry. Augen of olivine exhibit deformation kink bands (Fig. 16). Smallamounts of chromite (< 5%) are euhedral to subhedral. These fabrics are attributed to man-

7.

Descrip

tion

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ock

Types

inth

eB

locks

ofM

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245

Fig. 14. Outcrop map of Zunhua structural belt podiform chromite area, showing locations of rock types in metasedimentary and metatonaliticgneiss. Modified from Li et al. (2002).

246 Chapter 7: Origin and Emplacement of Archean Ophiolites of the Central Orogenic Belt

Fig. 15. Podiform chromite photos. (a) Flattened lens of serpentinized harzburgite; (b) Serpen-tinized harzburgite block in mélange, surrounded by weathered biotite gneiss; (c) Gabbroic lens(coarse-grained, bottom of photo) within sheared biotite gneiss; (d) Ophiolitic mafic to ultramaficboudins as inclusion with granitic gneiss; (e) Pillow lava.

tle flow at high temperatures (Li et al., 2002). Olivine crystals are commonly serpentinized,with magnetite distributed along the foliation planes.

Boudins and tectonic blocks of various types of gabbrorange in size from 20 × 15 cmto 10 × 100 m. They occur within intensely sheared garnet-bearing gneiss (Fig. 15).A 400 × 60 cm block of cumulate gabbro and layered gabbro is identified within augengneiss. Rarely, pods of dunite are recognized within olivine gabbro. These rocks are gen-

7. Description of Rock Types in the Blocks of Mélange 247

Fig. 16. Details of peridotite mineralogy and textures. (a, b) Polished surfaces of hand specimensillustrating principal microstructures and textures of flattened harzburgite core with rings of serpenti-nite. Early high-temperature foliations are preserved in cores of pods; (c) Harzburgite showing Opx,chromite, and olivine crystals; (d) Asymmetrical olivine porphyroclast with recrystallized tail withinharzburgite tectonite; (e, f) Kink bands within relict olivine in serpentinized harzburgite.

erally less-deformed than the ultramafic complex. Cumulate textures including alternatingpyroxene-rich and plagioclase-rich banding are locally preserved. The layers are generallyabout 2 to 5 cm thick, characterized by alternating internal texture and mineral composi-tion, representing primary magmatic layering. They clearly underwent amphibolite-faciesmetamorphism as pyroxenite layers are commonly transformed into foliated hornblenditealong their margins.

248 Chapter 7: Origin and Emplacement of Archean Ophiolites of the Central Orogenic Belt

Sheeted and multiple dike complexesare recognized from several places in the Zunhuastructural belt. The width of individual half dikes are generally tens of centimeters. Manydikes have only one chilled margin, which is a consequence of repeated intrusions in thecenter of a single opening fissure. The chilled margins are recognized as 2–3 cm thickboundaries defined by strong alignment of fine-grained hornblende (Fig. 14). One dikecomplex can be traced for more than 100 m along strike. The mineral assemblage is madeup of plag + cpx + hb, characteristic of amphibolite facies conditions. In one outcrop,more than ten dikes have been recognized over a distance of 4 m, with individual dikes be-ing about 5–30 cm wide. They preserve differences in grain size, and several have slightlydifferent mineralogy than adjacent dikes, suggesting a complex or polyphase history.

Spectacular pillow lava structures are preserved locally in weakly deformed mafic vol-canic domains (Figs. 14 and 15). Pillows vary in size from tens of centimeters to one meter,and are interbedded with amygdular massive basalt. The pillow lavas are pervasively al-tered to albite and chlorite assemblages, associated with hydrothermal alteration. Rarely,the pillows preserve dark cryptocrystalline margins, representing original glassy selvages.Younging toward the northwest is indicated by the shapes of pillows. Layers of pillow brec-cia and volcanoclastic sediment are intercalated with the pillow basalt, and these units aremetamorphosed into plagioclase-biotite schist and biotite schist. Some ultramafic lensesare intercalated with pillow lavas, suggestive of large amounts of shearing.

At least six large and numerous smaller chromite-rich peridotite massifsare recognizedwithin the eastern Zunhua Structural Belt (Figs. 13 and 14). The chromitites are concordantwith the foliation of the enclosing harzburgites, are commonly lens-shaped, and hosted indunite envelopes within intensely serpentinized harzburgite. Serpentinized pods and lensesshow concentric rings with systematic variations in mineral composition and textures.Outer rings are commonly 2–10 cm thick and composed of serpentine, whereas inner corespreserve dunite or massive harzburgite with clear tectonic foliations. Narrow deformedpyroxenite and gabbroic dikes (1–10 cm wide) with branches are recognized within ser-pentinized harzburgites, which are interpreted as melt channels or trapped parental meltsto oceanic basalt (Fig. 17). Tectonic fabrics defined by chromite are well-preserved inthe cores of the serpentinized pods. Open folds, tight folds and rootless folds are iden-tified with the chromitite bands (Fig. 17). Dunite envelopes are commonly one to threecentimeters wide, and separate chromite from harzburgite. They are a common feature ofpodiform chromites, and they are known to form almost exclusively in the mantle or crust-mantle transition zone of suprasubduction zone (harzburgite type) ophiolites of differentages (Nicolas and Azri, 1991; Zhou et al., 1996; Bai et al., 1992; Edwards et al., 2000;Li et al., 2002).

Most of the chromitites are strongly deformed by plastic flow, although nodular chromi-tites are locally preserved, especially in discordant pods. Nodular, orbicular, banded, mas-sive, antinodular, and disseminated chromitite textures are all present (Fig. 17), and inmany places grade into each other. Discordant structures are preserved within weakly-strained domains. Nodular textures consist of small balls of chromite in a dunite matrix,whereas orbicular chromitites consist of thin-rings of chromite surrounding cores of dunite.Nodular and orbicular chromite texturesare only known to form in ophiolitic settings.

7. Description of Rock Types in the Blocks of Mélange 249

Fig. 17. Polished surfaces of handspecimens illustrating principal microstructures and textures ofchromitites ores. Scale bar is 1 cm in all photos. (a) Banded chromitites with dunite envelop; (b) Root-less fold showing thickening of chromitite band; (c) Chromite vein within serpentinized dunite;(d) Nodular chromites in weakly-deformed domain; (e) Nodular and orbicular chromitites; (f, g)Antinodular chromite; (h) Layered antinodular chromites.

250 Chapter 7: Origin and Emplacement of Archean Ophiolites of the Central Orogenic Belt

Nodular and orbicular chromites, with diameters of 2–10 mm and occasionally larger than10 mm, are generally flattened into elliptical shapes, and some nodules form flattened rings.They occur together or separately. Some nodular and orbicular chromites grade into mas-sive chromites, veins, or disseminated ores. Nodular and orbicular textures are the mosttypical magmatic structures of ophiolitic chromitites (Nicolas, 1989; Nicolas and Azri,1991). Pull-apart textures are common in the massive, banded, and antinodular chromititedeposits, and similar textures have been attributed to high-temperature (> 1000 ◦C) mantledeformation in other ophiolites (e.g., Holtzman, 2000; Nicolas, 1989). Abundant olivineoccurs as rounded inclusions in the chromitite, although they are widely altered into ser-pentinites. Orthopyroxene porphyroclasts show asymmetrical recrystallized tails indicatinghigh-temperature shearing.

Rounded to flattened chromite balls are characteristic of nodular chromititein the Zun-hua structural belt, with the diameter of the chrome nodules ranging from 0.2 to 1 cm.Commonly, the nodules impinge into orbicules, deforming the initially spherical rings.The nodules and orbicules show patterns of flattening and mutual impression along theircontacts with each other (Figs. 17 and 18), suggesting that they settled while they were stillsoft. Most nodules with flattened geometry are oriented parallel to the foliation. Locally,nodules are sorted into layers by their sizes. These features are interpreted to be a resultof rapid deposition of chromite nodules while they were still plastic (Lago et al., 1982).The nodules and orbicules commonly exhibit stretching fabrics interpreted to have formedwhen they were still in liquid form (Li et al., 2002). They are elongated by plastic strain andshow a preferred orientation, forming easily recognized lineations in many samples. Theouter boundary of single nodules are typically smooth and rounded (Figs. 17 and 18). Incontrast, their inner parts display individual chromite grains that grew inward. The devel-opment of nodular chromites record dynamic magmatic flow or partial melting conditions,needed to keep chromite suspended and growing concentrically into the magma (Edwardset al., 2000). The delicate magmatic structures preserved show that they have not beensignificantly deformed after their formation, which is attributed to their rigid nature.

In some cases, nodules grade into antinodules in the same hand-specimen. They recordmagmatic growth and settling in the upper mantle (e.g., Edwards et al., 2000). These origi-nal magmatic structures are commonly destroyed or become incomplete as the shear strainincreases. For example, nodular and orbicular textures are strongly stretched and trans-posed into layering or antinodular chains. Compared with nodules, orbicules are morestrongly stretched, their ratio of X/Z are up to 5:1. Minor orthopyroxene occurs as porphy-roclasts. Rounded inclusions of olivine are recognized within chromite grains, and someinclusions of olivine show kink bands (Fig. 16), recording plastic deformation before orduring growth of the chromite. We attribute this deformation to upper mantle flow in theoceanic mantle.

Antinodular chromititescontain 30–50% chromite. They consist of rounded or flatteneddunite aggregates surrounded by chromite chains or networks (Fig. 17). Olivine composesthe cores surrounded by chromite net-like bands, suggestive of strong flattening or ex-tension. Flattened antinodular texture is typical of plastic deformation in oceanic man-tle, which is a result of straining of weaker olivine inclusions in a chromite-rich matrix

7. Description of Rock Types in the Blocks of Mélange 251

Fig. 18. Microscopic textures of chromite ores. (a) Flattened orbicule of chromite (reflected light);(b) Flattened nodules of chromite and igneous contact between two nodules (reflected light); (c) Flat-tening of contact between orbicule and nodule of chromite; (d) Flattening of chrome nodule withasymmetrical tail; (e) Inclusion of orthopyroxene (OPX) and olivene (OL) showing kink bands withinchromites; (f) Deformation textures of chromitite ore showing stretched olivine crystals in dunite core(strain ellipse outlined in black) with antinodular chromite; (g) Pulled apart chromites; (h) Cumulatelayer of chromite nodules.

252 Chapter 7: Origin and Emplacement of Archean Ophiolites of the Central Orogenic Belt

(Nicolas, 1989). Alignment of needle-like chromite also indicates strong shearing. Dis-seminated chromititescontain 10–30% chromite, with grain sizes of chromite being largerthan 0.1 mm. The disseminated chromites locally grade into nodular chromite, and also lo-cally have nodules with diameters of 2–5 mm scattered throughout the disseminated ores.In disseminated chromitites, chromite grains are uniformly scattered in an olivine ma-trix. Layered and banded textures consist of anastomosing chains of chromite surroundingovoids of olivine, which were generated by shearing, rather than crystallization and accu-mulation. They are characterized by intense shearing. Locally, they are cross-cut by late,less than one cm wide veins of chromite, suggesting that at least two phases of chromiteswere originally present. Banded and massive chromititesshow characteristic banding orlayering, with 2–4 cm thick bands of chromite (Fig. 17). In a few places they are cross-cut by pyroxenite and dunite dikes. Some chromite layers occur as rootless to tight foldsor asymmetric lensoidal boudins, and other layers and lenses consist of nodules. Minorgrains of orthopyroxene are preserved. Olivine is commonly serpentinized. Massive orescontain at least 70% fine-grained chromite distributed homogeneously in a serpentinitematrix. Massive chromitites may grade into disseminated chromites, whereas other layersdisplay podiform geometry with indicators of pre-full crystallization shearing. Pull-aparttextures are filled by olivine (Li et al., 2002), forming indicators of magmatic shearing withthe massive chromites.

8. CHEMISTRY OF CHROMITITE DEPOSITS

Chromite from two samples of chromitite in the Zunhua ophiolitic mélange (dn, dn2),and one from dunite with disseminated chromite from the Dongwanzi ophiolite (dd) wereanalyzed by WDS on a Cameca S × 100 at the Manchester Electron Microprobe Facility(MEMF), UK. The operating conditions for WDS analysis were an acceleration voltage of20 kV, a counting time of 20 s, and a 20 nA beam for Cr, Al, Mg and Fe with a countingtime of 80 s with a 100 nA beam for Zn, Co, Ti, V, Mn, Ni and Si. Natural oxide andsynthetic standards were used, with the manufacturers ZAF correction program. Iron wasdetermined as FeOt and FeO and Fe2O3 were calculated by stoichiometry using the methodof Droop (1987). Detection limits are between 10 and 20 ppm. The data is shown in Table 1.Twenty three additional samples from the Zunhua chromitites were analyzed using theWDS techniques on a Cameca-50 at the Geological Institute Academia Sinica in Beijing.These data are shown in Table 2.

The chromite from the ultramafic fragments in the mélange beneath Dongwanzi ophio-lite has very refractory composition, with Cr#s Cr/(Cr + Al) of 0.74 to 0.93, TiO2 at lessthan 0.3 wt%, and V2O5 wt% at less than 0.1 for the samples analyzed; with the exceptionof sample 1525 which has a Cr# of 0.42. Excluding sample 1525 there are two distinctgroups; massive chromitite and nodular chromite in dunite (samples dn and dn2) and dis-seminated chromite in dunite (dd). These groups have almost identical TiO2 wt% but thedisseminated chromite has a lower Cr# and higher V2O5 content. The Fe2O3 wt% in thesamples analyzed is variable, with a range from 3.2 to 37.2, though with the exception of

8. Chemistry of Chromitite Deposits 253

Table 1a. Major element concentrations of chromite measured on the Manchester electron micro-probe facility

Sample Oxide wt% TotalAl2O3 Cr2O3 MgO MnO TiO2 ZnO SiO2 V2O5 NiO Fe2O3 FeO

dn2 5.05 59.67 6.25 0.49 0.14 0.05 0.00 0.03 0.07 5.89 22.99 100.63dn2 5.53 60.52 7.73 0.45 0.14 0.05 0.00 0.04 0.07 5.18 20.83 100.54dn2 5.18 59.15 4.93 0.48 0.15 0.06 0.00 0.04 0.06 5.36 25.01 100.41dn2 5.19 59.00 4.79 0.49 0.12 0.05 0.00 0.04 0.06 5.31 25.14 100.17dn2 5.32 59.65 5.95 0.49 0.18 0.04 0.00 0.04 0.07 5.41 23.57 100.73dn2 4.91 58.67 4.28 0.51 0.13 0.05 0.00 0.03 0.07 5.91 25.91 100.47dn2 5.24 58.41 3.54 0.51 0.13 0.06 0.00 0.03 0.06 5.25 27.04 100.27dn2 5.19 58.18 4.06 0.51 0.14 0.06 0.00 0.03 0.06 5.87 26.28 100.37dn2 5.29 56.92 3.36 0.53 0.10 0.06 0.00 0.03 0.05 7.02 27.43 100.80dn2 4.50 58.14 4.51 0.51 0.10 0.06 0.00 0.03 0.07 7.24 25.50 100.64dn2 5.36 59.35 5.41 0.50 0.15 0.05 0.15 0.03 0.07 4.91 24.50 100.48dn2 5.15 60.59 6.29 0.49 0.10 0.05 0.00 0.03 0.07 5.01 22.98 100.77dn2 5.11 58.97 4.41 0.52 0.15 0.05 0.00 0.03 0.07 5.49 25.81 100.62dn2 4.84 59.61 4.73 0.51 0.09 0.06 0.00 0.03 0.05 5.24 25.15 100.32

dn 4.88 61.32 8.48 0.48 0.16 0.05 0.00 0.03 0.13 5.52 19.52 100.55dn 4.68 62.05 8.39 0.48 0.15 0.05 0.00 0.03 0.11 4.65 19.48 100.06dn 5.06 60.94 8.49 0.48 0.17 0.05 0.00 0.03 0.13 5.80 19.59 100.74dn 4.98 60.78 8.38 0.48 0.16 0.05 0.00 0.03 0.13 5.49 19.49 99.98dn 4.94 60.94 8.51 0.48 0.17 0.05 0.00 0.03 0.12 5.59 19.38 100.20dn 4.97 61.39 8.64 0.48 0.17 0.05 0.00 0.03 0.11 4.68 18.97 99.49dn 4.81 62.09 8.72 0.47 0.16 0.05 0.00 0.03 0.13 4.49 18.96 99.88dn 4.92 61.76 8.98 0.47 0.16 0.05 0.15 0.03 0.12 4.27 18.68 99.60dn 4.93 61.94 8.71 0.47 0.17 0.04 0.00 0.03 0.10 4.32 18.95 99.67dn 4.91 61.45 8.72 0.47 0.16 0.05 0.00 0.03 0.12 5.10 19.00 100.02dn 4.95 62.32 9.02 0.47 0.16 0.05 0.00 0.03 0.11 3.94 18.40 99.46dn 4.92 61.86 8.83 0.48 0.14 0.05 0.00 0.03 0.11 4.93 18.89 100.23dn 4.80 62.72 10.17 0.42 0.16 0.05 0.11 0.03 0.12 4.13 16.81 99.50dn 5.01 62.30 9.06 0.46 0.17 0.05 0.08 0.03 0.12 3.98 18.56 99.80dn 4.91 61.87 8.82 0.46 0.16 0.06 0.00 0.03 0.12 4.31 18.66 99.39

dd 10.75 46.20 4.97 0.56 0.19 0.22 0.03 0.08 0.13 11.45 25.45 100.03dd 10.72 46.29 5.20 0.54 0.20 0.20 0.01 0.08 0.13 11.81 25.23 100.40dd 10.83 46.17 5.38 0.53 0.20 0.20 0.00 0.09 0.13 11.83 24.95 100.30dd 10.06 47.23 4.83 0.57 0.17 0.22 0.00 0.08 0.11 11.61 25.69 100.55dd 10.33 46.87 4.56 0.58 0.18 0.23 0.00 0.08 0.11 12.07 26.39 101.40dd 10.77 44.99 4.40 0.58 0.19 0.23 0.02 0.08 0.13 12.43 26.33 100.15dd 9.88 46.16 4.56 0.58 0.11 0.21 0.03 0.08 0.12 12.74 26.00 100.45dd 10.35 45.63 4.35 0.58 0.20 0.22 0.02 0.09 0.13 12.38 26.40 100.34dd 10.41 46.40 5.25 0.55 0.18 0.19 0.02 0.08 0.13 11.99 25.05 100.25dd 10.63 44.16 3.75 0.59 0.15 0.23 0.01 0.08 0.11 13.33 27.29 100.32dd 10.20 44.88 3.51 0.60 0.11 0.22 0.00 0.08 0.07 12.72 27.42 99.80dd 10.91 46.45 5.31 0.55 0.20 0.20 0.01 0.09 0.13 11.64 25.19 100.66dd 10.78 47.01 5.77 0.53 0.19 0.18 0.01 0.08 0.12 11.54 24.50 100.71dd 10.19 47.00 4.63 0.58 0.17 0.22 0.02 0.08 0.11 11.65 26.07 100.72

254 Chapter 7: Origin and Emplacement of Archean Ophiolites of the Central Orogenic Belt

Table 1b. Trace element concentrations of chromite measured on the Manchester electron micro-probe facility

Sample Molar concentration in ions Total RatiosAl Cr Mg Mn Ti Zn Si V Ni Fe3+ Fe2+ O Cr# Mg# Fe3+#

dn2 0.21 1.66 0.33 0.01 0.00 0.00 0.00 0.00 0.00 0.15 0.66 4.00 3.04 0.89 0.33 0.076dn2 0.23 1.66 0.40 0.01 0.00 0.00 0.00 0.00 0.00 0.13 0.60 4.00 3.04 0.88 0.40 0.066dn2 0.22 1.66 0.26 0.01 0.00 0.00 0.00 0.00 0.00 0.14 0.73 4.00 3.04 0.88 0.26 0.070dn2 0.22 1.67 0.25 0.01 0.00 0.00 0.00 0.00 0.00 0.14 0.74 4.00 3.04 0.88 0.26 0.069dn2 0.22 1.66 0.31 0.01 0.00 0.00 0.00 0.00 0.00 0.14 0.68 4.00 3.04 0.88 0.31 0.070dn2 0.21 1.66 0.23 0.02 0.00 0.00 0.00 0.00 0.00 0.16 0.76 4.00 3.04 0.89 0.23 0.077dn2 0.22 1.66 0.19 0.02 0.00 0.00 0.00 0.00 0.00 0.14 0.80 4.00 3.04 0.88 0.19 0.069dn2 0.22 1.65 0.22 0.02 0.00 0.00 0.00 0.00 0.00 0.16 0.77 4.00 3.04 0.88 0.22 0.077dn2 0.23 1.62 0.18 0.02 0.00 0.00 0.00 0.00 0.00 0.19 0.81 4.00 3.05 0.88 0.18 0.092dn2 0.19 1.65 0.24 0.02 0.00 0.00 0.00 0.00 0.00 0.19 0.75 4.00 3.05 0.90 0.24 0.094dn2 0.22 1.66 0.28 0.01 0.00 0.00 0.01 0.00 0.00 0.13 0.71 4.00 3.04 0.88 0.29 0.064dn2 0.21 1.68 0.33 0.01 0.00 0.00 0.00 0.00 0.00 0.13 0.66 4.00 3.04 0.89 0.33 0.064dn2 0.22 1.66 0.23 0.02 0.00 0.00 0.00 0.00 0.00 0.14 0.76 4.00 3.04 0.89 0.24 0.071dn2 0.20 1.68 0.25 0.02 0.00 0.00 0.00 0.00 0.00 0.14 0.74 4.00 3.04 0.89 0.25 0.068

dn 0.20 1.68 0.44 0.01 0.00 0.00 0.00 0.00 0.00 0.14 0.56 4.00 3.04 0.89 0.44 0.070dn 0.19 1.71 0.44 0.01 0.00 0.00 0.00 0.00 0.00 0.12 0.56 4.00 3.04 0.90 0.44 0.059dn 0.21 1.67 0.44 0.01 0.00 0.00 0.00 0.00 0.00 0.15 0.56 4.00 3.04 0.89 0.44 0.073dn 0.20 1.68 0.44 0.01 0.00 0.00 0.00 0.00 0.00 0.14 0.56 4.00 3.04 0.89 0.44 0.070dn 0.20 1.68 0.44 0.01 0.00 0.00 0.00 0.00 0.00 0.14 0.55 4.00 3.04 0.89 0.44 0.071dn 0.20 1.69 0.45 0.01 0.00 0.00 0.00 0.00 0.00 0.12 0.55 4.00 3.04 0.89 0.45 0.060dn 0.20 1.70 0.45 0.01 0.00 0.00 0.00 0.00 0.00 0.12 0.54 4.00 3.03 0.90 0.45 0.057dn 0.20 1.69 0.46 0.01 0.00 0.00 0.01 0.00 0.00 0.11 0.53 4.00 3.03 0.89 0.46 0.055dn 0.20 1.70 0.45 0.01 0.00 0.00 0.00 0.00 0.00 0.11 0.54 4.00 3.03 0.89 0.45 0.055dn 0.20 1.69 0.45 0.01 0.00 0.00 0.00 0.00 0.00 0.13 0.54 4.00 3.04 0.89 0.45 0.065dn 0.20 1.71 0.47 0.01 0.00 0.00 0.00 0.00 0.00 0.10 0.53 4.00 3.03 0.89 0.47 0.050dn 0.20 1.69 0.46 0.01 0.00 0.00 0.00 0.00 0.00 0.13 0.54 4.00 3.04 0.89 0.46 0.063dn 0.19 1.71 0.52 0.01 0.00 0.00 0.00 0.00 0.00 0.11 0.48 4.00 3.03 0.90 0.52 0.053dn 0.20 1.70 0.47 0.01 0.00 0.00 0.00 0.00 0.00 0.10 0.53 4.00 3.03 0.89 0.47 0.051dn 0.20 1.70 0.46 0.01 0.00 0.00 0.00 0.00 0.00 0.11 0.54 4.00 3.03 0.89 0.46 0.055

dd 0.45 1.30 0.26 0.02 0.01 0.01 0.00 0.00 0.00 0.30 0.73 4.00 3.08 0.74 0.27 0.144dd 0.45 1.30 0.28 0.02 0.01 0.01 0.00 0.00 0.00 0.30 0.72 4.00 3.08 0.74 0.28 0.148dd 0.45 1.29 0.28 0.02 0.01 0.01 0.00 0.00 0.00 0.30 0.71 4.00 3.08 0.74 0.29 0.148dd 0.42 1.33 0.26 0.02 0.00 0.01 0.00 0.00 0.00 0.30 0.74 4.00 3.08 0.76 0.26 0.146dd 0.43 1.31 0.24 0.02 0.00 0.01 0.00 0.00 0.00 0.31 0.75 4.00 3.08 0.75 0.24 0.151dd 0.46 1.28 0.24 0.02 0.01 0.01 0.00 0.00 0.00 0.32 0.76 4.00 3.08 0.74 0.24 0.157dd 0.42 1.31 0.24 0.02 0.00 0.01 0.00 0.00 0.00 0.33 0.75 4.00 3.08 0.76 0.25 0.160dd 0.44 1.29 0.23 0.02 0.01 0.01 0.00 0.00 0.00 0.32 0.76 4.00 3.08 0.75 0.23 0.156dd 0.44 1.31 0.28 0.02 0.00 0.01 0.00 0.00 0.00 0.31 0.72 4.00 3.08 0.75 0.28 0.151dd 0.45 1.26 0.20 0.02 0.00 0.01 0.00 0.00 0.00 0.35 0.79 4.00 3.08 0.74 0.20 0.168dd 0.44 1.29 0.19 0.02 0.00 0.01 0.00 0.00 0.00 0.33 0.80 4.00 3.08 0.75 0.19 0.162dd 0.45 1.30 0.28 0.02 0.01 0.01 0.00 0.00 0.00 0.30 0.72 4.00 3.08 0.74 0.28 0.145dd 0.45 1.31 0.30 0.02 0.00 0.00 0.00 0.00 0.00 0.29 0.69 4.00 3.08 0.75 0.30 0.144dd 0.43 1.32 0.25 0.02 0.00 0.01 0.00 0.00 0.00 0.30 0.75 4.00 3.08 0.76 0.25 0.146

8.

Chem

istryofC

hro

mitite

Dep

osits

255

Table 2a. Major element concentrations of chromite measured on the Academia Sinica electron microprobe facility

SampleNo.

15z12 z2 15z15 1535 1552 1553 1555 1566 1572 1577 1590 1515 1589 1578 1579 1584

SiO2 (%) 39.4989 39.2334 39.8596 40.3738 36.1557 38.56 33.1993 42.0188 39.418 34.6402 36.1302 32.0265 35.718 38.439 33.6712 39.9353TiO2 0.002 0.0049 0.0048 0.0278 0.0111 0.0063 0.012 0.0236 0.0083 0.0221 0.0167 0.05 0.006 0.0084 0.0259 0.0077Al2O3 0.0733 0.0766 0.1023 0.8517 0.5285 0.0732 0.5382 0.5628 0.2647 0.7825 0.509 1.3361 0.0932 0.2498 0.8663 0.1652Fe2O3 (T) 6.9663 5.2075 6.3428 10.8189 4.9546 5.6639 4.6792 13.0095 5.7205 5.3911 8.295 8.959 10.1639 5.3177 6.1065 6.3332MnO 0.0522 0.0611 0.0552 0.1035 0.0544 0.0616 0.0758 0.1903 0.0302 0.0482 0.043 0.1505 0.0656 0.1038 0.0664 0.0346MgO 37.8122 42.2416 37.8945 35.314 37.6019 38.5175 37.6144 33.5305 38.7221 35.5799 35.6311 31.7702 35.251 37.7137 34.6841 37.5135CaO 0.0854 0.4468 0.0583 0.0664 0.7603 0.3381 0.1986 0.1212 0.0472 0.0991 0.0734 1.3642 1.7351 0.9868 0.2764 0.076Na2O 0.0182 0.0201 0.0671 0.015 0.0354 0.0031 0.0147 0.0673 0.0122 0.0472 0.0503 0.0399 0.0516 0.0054 0.0306 0.0071K2O 0.0071 0.0084 0.0092 0.0148 0.008 0.007 0.0062 0.0091 0.0093 0.0077 0.0098 0.0077 0.0062 0.0074 0.0057 0.008P2O5 0.0121 0.0103 0.0123 0.0159 0.0107 0.013 0.0107 0.0203 0.0092 0.0093 0.0078 0.0071 0.0087 0.0089 0.007 0.0063Cr2O3 0.7135 0.7032 0.7862 1.1852 5.9187 0.7744 7.4619 0.529 1.1029 10.0148 5.8001 13.4612 0.8231 0.8352 11.2609 1.3413NiO (ppm) 5798 3917 5430 2785 3842 4116 4117 3622 3787 3693 3795 3076 4546 3623 3189 3589FeO (%) 1.6 1.48 0.52 4.3 0.72 0.55 0.65 6.15 1 0.9 1.6 2.89 2.3 1.33 0.95 1.35L.O.I (%) 13.61 11.98 13.76 10.66 12.87 15.09 16.16 9.69 13.59 12.33 12.36 10.22 14.55 15.32 12.01 13.5Ba (ppm) 0 0 4.5636 6.3761 13.1976 4.7537 3.2319 385.001 5.8972 8.61 8.5551 38.8209 11.3105 7.3739 35.7902 8.7355Co 107.101 95.8663 128.9148 78.8935 96.266 101.535 96.5728 91.9813 73.7451 95.8065 88.6797 108.7291 80.3125 83.9053 93.1489 75.9036Cu 0.45 0.3309 2.1103 4.1154 2.0759 1.9214 0.6746 13.9635 1.419 0 0 0 0 1.9759 2.2691 2.3433Ga 1.0319 5.0339 4.2656 4.2826 3.0222 4.7669 6.1056 5.5092 0.9256 4.7213 2.5772 6.4718 4.0258 2.625 5.9127 5.2211V 18.0361 14.0914 14.2615 29.5437 24.9442 12.8521 25.1216 25.4958 21.9646 26.1656 22.6376 49.0554 14.6973 12.5445 32.4333 16.2245Zn 74.0762 29 49.5923 93.9131 30.822 40.7071 36.6401 164.5144 48.7478 44.3615 59.5131 112.784 30.7199 43.555 59.6983 60.8765

256C

hapte

r7:O

rigin

and

Em

pla

cem

ent

ofA

rchean

Ophio

lites

ofth

eC

entra

lOrogen

icB

elt

Table 2b. Trace element concentrations of chromite measured on the Academia Sinica electron microprobe facility

Sample No. z2 15z12 15z15 1515 1535 1544 1552 1553 1555 1566 1572 1577 1578 1579 1584 1589 1590

Sc 1.936 1.404 1.228 2.344 3.943 1.494 1.804 1.964 1.679 5.012 3.596 1.7 1.775 1.66 2.792 2.007 2.134V 3.285 7.423 11.114 54.139 31.714 140.925 21.58 4.742 14.263 23.874 18.567 24.451 4.554 30.424 15.573 5.202 24.286Co 91.961 112.673 126.154 125.001 77.773 227.269 101.468 94.487 100.47 83.484 85.176 123.277 78.029 124.14 71.899 79.679 86.951Rb 0.596 0.6 1.266 0.484 2.651 5.599 2.321 0.56 3.443 10.584 13.925 1.071 0.434 0.98 2.803 1.225 2.841Sr 23.979 4.08 2.799 43.455 3.74 1.194 20.631 7.319 23.001 37.549 17.788 1.669 21.213 3.159 3.376 31.353 2.499Y 0.263 0.84 0.945 0.952 4.567 1.036 0.263 0.471 2.693 6.784 4.751 0.305 0.556 0.386 0.629 1.139 0.734Zr 0.273 0.312 0.322 1.689 1.992 11.454 0.597 0.234 0.315 3.402 1.632 0.295 0.245 0.31 0.414 0.33 1.202Nb 0.079 0.029 0.126 0.271 2.152 1.004 0.209 0.061 0.077 0.669 0.213 0.009 0.036 0.012 0.067 0.055 0.162Cs 0.014 0.12 0.216 0.015 0.042 0.134 0.137 0.052 0.064 0.188 0.318 0.047 0.134 0.015 0.049 0.029 0.048La 0.48 0.57 0.773 0.688 1.954 0.849 0.348 0.27 4.918 5.767 3.082 0.276 0.228 0.248 0.546 1.141 0.607Ce 0.84 0.829 0.983 1.622 4.464 1.32 0.751 0.459 10.846 12.774 7.056 0.421 0.365 0.596 1.082 4.02 1.343Pr 0.088 0.138 0.218 0.217 0.574 0.189 0.065 0.061 0.718 1.306 0.754 0.056 0.068 0.086 0.138 0.267 0.161Nd 0.273 0.557 0.903 0.931 2.468 0.727 0.224 0.204 2.38 4.671 2.688 0.225 0.286 0.354 0.526 0.947 0.617Sm 0.04 0.105 0.168 0.214 0.689 0.138 0.034 0.04 0.458 0.961 0.666 0.043 0.065 0.075 0.095 0.178 0.123Eu 0.012 0.031 0.038 0.237 0.184 0.03 0.025 0.013 0.064 0.306 0.113 0.037 0.037 0.03 0.053 0.037 0.03Gd 0.048 0.118 0.172 0.21 0.82 0.144 0.044 0.044 0.462 1.015 0.689 0.047 0.077 0.056 0.118 0.209 0.14Tb 0.004 0.014 0.018 0.025 0.129 0.023 0.006 0.007 0.062 0.147 0.106 0.007 0.011 0.008 0.016 0.03 0.021Dy 0.028 0.085 0.106 0.145 0.769 0.166 0.03 0.043 0.35 0.895 0.603 0.046 0.076 0.055 0.109 0.187 0.127Ho 0.006 0.02 0.025 0.025 0.155 0.033 0.006 0.009 0.07 0.192 0.123 0.009 0.018 0.013 0.023 0.04 0.028Er 0.019 0.073 0.077 0.082 0.482 0.114 0.017 0.033 0.218 0.622 0.378 0.031 0.063 0.043 0.068 0.13 0.09Tm 0.004 0.012 0.011 0.009 0.074 0.015 0.003 0.007 0.036 0.107 0.06 0.006 0.01 0.008 0.012 0.022 0.015Yb 0.026 0.115 0.092 0.065 0.522 0.11 0.021 0.053 0.246 0.817 0.42 0.033 0.076 0.052 0.086 0.15 0.107Lu 0.008 0.023 0.016 0.008 0.079 0.015 0.003 0.012 0.043 0.128 0.064 0.004 0.015 0.008 0.017 0.025 0.02Hf 0.007 0.001 0.003 0.033 0.112 0.481 0.049 0.005 0.007 0.117 0.076 0.009 0.002 0.011 0.01 0.006 0.042Ta 0.058 0.019 0.018 0.038 0.862 0.371 0.096 0.035 0.05 0.096 0.033 0.006 0.033 0.006 0.036 0.036 0.104Pb 2.212 3.691 1.123 48.536 16.165 6.651 1.253 0.98 6.68 190.611 5.241 0.804 1.856 3.554 13.349 1.583 0.423Th 0.148 0.079 0.09 0.092 0.733 0.426 0.16 0.146 3.088 3.241 5.623 0.018 0.018 0.071 0.024 0.258 0.058U 0.203 0.935 1.481 0.165 0.177 0.291 0.145 0.172 0.221 0.244 4.165 0.1 0.024 0.619 0.062 0.17 0.079

9. Interpretation of Chrome Chemistry 257

sample 1525f it is within the range 3.2–13.2. The high Fe2O3 content in sample 1525f iscaused by alteration from chromite to ferrit-chromite, as described in Ulmer (1974), andas such is not considered to a represent crystallization composition.

9. INTERPRETATION OF CHROME CHEMISTRY

Chromite analysis has a long history in the interpretation of ophiolitic rocks (Irvine,1965, 1967; Dick and Bullen, 1984; Kamenetsky et al., 2001; Matveev and Balhaus,2002) and for Archean sequences (Stowe, 1994; Cotteril, 1969; Chadwick and Crewe,1986). Chromite is used as it is often the only unaltered mineral and the composition isbelieved to represent crystallization compositions (Irvine, 1967; Dick and Bullen, 1984;Arai, 1992; Ahmed et al., 2001), particularly for chromite from lava flows, and withinmassive chromitite, and dunite. The low diffusivity of Cr, Al, Ti and V, within olivine,makes these elements ideal for indicating crystallization compositions (Arai and Yu-rimoto, 1995). The Cr# indicates the degree of depletion of the mantle source fromwhich the melt was derived, and to a lesser extent, the degree of melt fractionation(Irvine, 1965, 1967; Dick and Bullen, 1984). The TiO2 wt% indicates the degree ofdepletion of the mantle source and the degree of melt fractionation (Shervais, 1982;Dick and Bullen, 1984; Arai, 1992; Kamenetsky et al., 2001). Vanadium content indi-cates the degree of depletion of the mantle source, the degree of melt fractionation, and thefO2 during partial melting, melt fractionation and chromite crystallization (Shervais, 1982;Canil, 1999). Fe2O3 content of spinel has been cited as indicating the fO2 of the man-tle source, magmatic system and crystallization environment (Wood and Virgo, 1989;Arai, 1992; Kepezhinskas et al., 1993).

Interpretation of chromite compositions, through geochemical modeling of the chromiteforming magmatic system, and comparison with both Archean greenstone belts andPhanerozoic SSZ ophiolite complexes allow inferences to magma type and tectonic settingto be made. Power et al. (2000) contest that the use of chrome spinel chemistry for the inter-pretation of the tectonic setting of dismembered and deformed ophiolitic and other igneouscomplexes is inherently unreliable. They suggest that changes in chromite composition arecaused by sub-solidus equilibrium, serpentinization of the host rock, and metasomatism,all of which result in an increase in Fe and Ti contents. However, with care during analysis,altered chromite can easily be identified. BSE imaging can identify alteration, and scruti-nization of the chemical results can reveal which samples are altered (e.g., sample 1525f isobviously ferrit-chromite).

The style of chromite mineralization, chromitite grading into dunite and eventuallyharzburgite, is consistent with formation through reaction between an exotic melt anda harzburgitic mantle (Kelemen, 1990; Kelemen et al., 1990; Zhou et al., 1994, 1996;Oberger et al., 1995; Zhou and Robinson, 1997; Arai, 1997a, 1997b). Cr# is often sim-ilar between mantle-hosted and crustal cumulate chromitites (e.g., Dick and Bullen, 1984;Oberger et al., 1995).

258 Chapter 7: Origin and Emplacement of Archean Ophiolites of the Central Orogenic Belt

The high Cr# (0.74 to 0.93) of the chromite (Figs. 19 and 20), from the mélange be-neath the Dongwanzi ophiolite, is indicative of a very high degree of mantle partial melt-ing, or partial melting of a previously depleted mantle source. In modern and Phanero-zoic ophiolitic environments, Cr#s this high are associated with boninitic or arc picritemagmatism in SSZ spreading center and island arc tectonic environments, with slab fluidaided melting of previously depleted mantle that has previously been depleted by meltextraction (Dick and Bullen, 1984; Jan and Windley, 1990; Kamenetsky et al., 2001;Parkinson and Pearce, 1998). However, higher geothermal gradients in the Archeanmay have produced a melt rich enough in Cr to form high Cr# chromite without pre-vious partial melting of the mantle. High Cr# chromite could also form at a MORenvironment after previous extraction of a more depleted melt such as a komatiite.Both modern and Archean komatiites have lower Cr#s, typically 0.5–0.6 (Echeverria,1980) and 0.66–0.77 (Zhou and Kerrich, 1992) respectively, than those associated withthe Dongwanzi ophiolite and Zunhua ophiolitic mélange blocks. However, these Cr#sare higher than those from MOR chromite, less than 0.0–0.7 (Dick and Bullen, 1984;Kamenetsky et al., 2001). It is likely that a melt produced from a source that has previ-ously been depleted by a melt, that was the result of a greater degree of partial meltingthan MORB, would produce chromite with such high Cr#s.

The TiO2 wt% of less than 0.3 wt% (Tables 1 and 2, Figs. 19 and 20) is typical ofchromite from SSZ ophiolites (Dick and Bullen, 1984; LeBlanc and Nicholas, 1992) anddiffers from deposits associated with komatiites, which have greater than 0.3 wt% TiO2(Zhou and Kerrich, 1992) and those from modern MOR, OIB and LIP settings (Dick andBullen, 1984; Kamenetsky et al., 2001). A comparison of the Cr# with the TiO2 wt%(Fig. 19) indicates that within a deposit Cr# is roughly constant whilst TiO2 wt% is vari-able. The variable TiO2 wt% could be the product of alteration, or reflect differing magmaTi contents during chromite crystallization.

Chromite from the Dongwanzi and Zunhua fragments have much lower V2O5 (0.03to 0.085) than chromite in Phanerozoic SSZ complexes (e.g., Jijal (0.1–0.3; Glass, un-published); Border Ranges (0.07–0.25; Kusky et al., in preparation); Mayarí-Barcoa(0.07–0.18; Proenza et al., 1999) and komatiite associated chromite deposits; Belingwe(0.17–1.42; Zhou and Kerrich, 1992) and Shurugwi (0.1–0.27; Stowe, 1987). The low V(Tables 1 and 2, Fig. 19) could be caused by:(1) High fO2 during magmatic fractionation and chromite crystallization: High fO2 will

cause more vanadium to be in the V4+ and V5+ oxidation states (Shervais, 1982)and behave incompatibly with respect to the crystallizing chromite (Canil, 1999). LowfO2 will cause more V to be in the V3+ oxidation state and enter the chromite lat-tice more easily. Therefore the low V values associated with the Dongwanzi chromitecould be caused by high fO2 in the magmatic system during chromite crystalliza-tion.

(2) Low fO2 during mantle partial melting: High fO2 will cause more vanadium to be inthe V4+ and V5+ oxidation states and behave incompatibly with respect to the man-tle during partial melting (Shervais, 1982). Low fO2 will cause more V to be in theV3+ oxidation state. V3+ behaves compatibly with respect to the mantle and is re-

9. Interpretation of Chrome Chemistry 259

Fig. 19. Mineral chemistry plots of TiO2 vs Cr#, V2O5 vs Cr#, and V2O5 vs TiO2 for Zunhua andDongwanzi chromites.

260 Chapter 7: Origin and Emplacement of Archean Ophiolites of the Central Orogenic Belt

Fig. 20. Geochemical plots of TiO2 vs Fe3 + (Fe3 + Cr + Al) and Cr# vs Mg# for Zunhua andDongwanzi chromites. Mg# is Mg/(Mg + Fe); Cr# is Cr/(Cr + Al).

9. Interpretation of Chrome Chemistry 261

tained in spinel (Canil, 1999). Therefore, low V chromite could be caused by low fO2

during partial melting. In the case of the chromites associated with the Dongwanziophiolite and Zunhua mélange this is considered unlikely, as the Cr2O3 content is sohigh and V3+ behaves similarly to Cr (Canil, 1997, 1999). Archean komatiite lavashad a similar or higher fO2 than modern oceanic basalts (Canil, 1997). If this repre-sents the mantle source and magmatic system, it is logical to assume that an ArcheanSSZ spreading center magma would be highly oxidizing and thus vanadium wouldbe dominantly in the V4+ and V5+ oxidation state, and not enter the chromite crystallattice.

(3) Previous mantle partial melting: Chromite associated with Archean komatiitic magmashas high V2O5 contents (Canil, 1997; Zhou and Kerrich, 1992) and indicates substan-tial extraction of vanadium from the mantle. Subsequent melts produced are likely tobe low in vanadium. Any significant melt extraction under high fO2 will deplete themantle of vanadium, as more vanadium would have been in the V4+ and V5+ oxida-tion states and thus incompatible with respect to the mantle (Shervais, 1982; Canil,1997, 1999).

There is an inverse relationship between Cr# and V content (Fig. 19) for the chromitesfrom the Zunhua structural belt. This indicates either chromium occupying the octahedralsite, in the chrome spinel lattice, in preference to V3+, or a different magmatic history, asthe higher V values are from disseminated chromite in dunite.

Mg and Fe in chromite have also been used to indicate magmatic compositions. Mg andFe contents may be reliable indicators for rapidly cooled volcanic rocks (Arai, 1992) butin intrusive and mantle deposits Mg and Fe in chromite will undergo subsolidus diffusionwith olivine (Irvine, 1965; Roeder et al., 1979; Sack and Ghiorso, 1991). Fig. 20 shows awide Mg# in relation to the Cr#, indicating extensive re-equilibration of Mg and Fe whichprecludes the use of these elements as petrogenetic indicators.

The Fe2O3 content of chromite is dependent on the amount of fractionation (Irvine,1974; Arai, 1992), the oxidation state of the mantle during partial melting, melt evolu-tion (Arai, 1992; Wood and Virgo, 1989; Kepezhinskas et al., 1993), and alteration andmetamorphism (Ulmer, 1974; Evans and Frost, 1975). The Fe2O3 wt% of the chromitesanalyzed in this study (3.2–13.2) is greater than those reported for chromite from MORsettings (0–4 wt% Fe2O3; Arai, 1992) and most SSZ spreading center settings (e.g., Shet-land, 2.5–4 wt% Fe2O3, own data; Oman, 3–6 wt% Fe2O3; Auge, 1987) but similar tothose from highly oxidizing boninitic and island arc magmas (Irvine, 1974) and Archeankomatiites (5–13, Zhou and Kerich, 1992). The high Fe2O3 content is unlikely to be causedby fractionation as the increase in Fe2O3 content is not accompanied by an increase inTiO2 wt% (Irvine, 1974; Arai, 1992; Kepezhinskas et al., 1993). Though ferric-ferrousiron ratios may be a reliable indicator for chrome spinel in volcanic rocks (Arai, 1992) ser-pentinization and metamorphism can lead to an increase in the Fe2O3 wt% (Ulmer, 1974;Evans and Frost, 1975). This precludes the use of iron as an effective indicator of magmaoxygen fugacity in altered ultramafic rocks.

262 Chapter 7: Origin and Emplacement of Archean Ophiolites of the Central Orogenic Belt

When compared to Phanerozoic ophiolite complexes the field geology and chromitechemistry show more similarities with a depleted SSZ spreading center than with an islandarc or mid-ocean ridge spreading center environment.

10. WHAT DOES THE OPHIOLITIC ORIGIN FOR THE DONGWANZI BELT ANDMANTLE FRAGMENTS IN MÉLANGE REVEAL ABOUT ARCHEANOCEANIC MAGMATIC PROCESSES?

The origin and tectonic setting of Archean greenstone belts has been one mosthotly debated questions in Precambrian geology for much of the last century. Portionsof several Archean greenstone belts have been interpreted to contain dismembered orpartial ophiolites (Kusky and Polat, 1999; Abbott, 1996; de Wit et al., 1982, 1987,1992; Kusky, 1987, 1990, 1991; Kusky and Vearncombe, 1997; Helmstaedt et al., 1986;MacLaughlin and Helmstaedt, 1995; Isachsen et al., 1991; Isachsen and Bowring, 1997;Wilks and Harper, 1997), but none of these contain the complete ophiolite sequence. Sev-eral of the Archean greenstone belts that have been interpreted as dismembered Archeanophiolites have three or four of the main magmatic components of a full ophiolite, al-though all of these have been disputed (Bickle et al., 1994). The Dongwanzi ophioliteand related belts in the North China craton contain a complete, albeit dismembered andmetamorphosed ophiolite sequence, including pillow lavas, gabbro, dike complexes, man-tle tectonites, and podiform chromitites. The preservation of a complete Archean ophiolitesequence in the North China craton is therefore of great importance for understandingprocesses of Archean sea-floor spreading, as it is the most complete record of this processknown to exist.

The Dongwanzi ophiolite has a strongly-deformed and serpentinized mantle section.However, about 60 km to the southwest primary mantle minerals and textures are well-preserved in many tectonic blocks within the Zunhua mélange in the same structural belt.In this mélange, the lower part of the ophiolite is preferentially preserved, with the mainrock types including harzburgite, dunite, podiform chromitites, meta-gabbro, and mafic andultramafic cumulates. The oceanic mantle rocks show intense serpentinization, consistentwith the Central Orogenic belt representing the suture between the East and West blocksof the North China craton (Fig. 2). In a few places pillow lava and sheeted dike complexesare also preserved in the mélange. Some of these have flat REE signatures (Kusky and Li,2002) and are chemically similar to the Dongwanzi ophiolite. We suggest therefore thatthe Dongwanzi ophiolite and Zunhua mélange ophiolitic fragments preserved dismem-bered fragments and ophiolites from the closure of the same ocean basin. But what canthe minerals and textures preserved in these ophiolitic fragments tell us about Archean seafloor spreading processes?

Harzburgite blocks in the mélange host podiform chromitites with dunite envelopes.The blocks grade up-section into wehrlite, pyroxenite, olivine gabbro (troctolite), and gab-bro. Podiform chromitites are commonly hosted by alpine-type or ophiolitic peridotite andare a normal component of ophiolites of different ages. Podiform chromite deposits are

10. What Does the Ophiolitic Origin for the Dongwanzi Belt and Mantle Fragments Reveal? 263

located in the transition zone between layered gabbro and peridotite tectonite, and the lher-zolite/harzburgite (asthenosphere-lithosphere) transition in ophiolites (Nicolas and Azri,1991). They are commonly attributed to oceanic settings such as mid-oceanic ridges, in-traoceanic suprasubduction zones (back-arc basins or island arcs) (LeBlanc and Nicholas,1992; Zhou et al., 1996). Their geological occurrence is closely associated with oceanicspreading processes (Nicolas and Azri, 1991). Late Proterozoic podiform chromitites inophiolites have been described in several areas, including Ethiopia, Saudi Arabia, Mo-rocco, south China, and Egypt, and Phanerozoic examples are numerous (Li et al., 2002).The oldest relatively intact podiform chromitite previously recognized is that from the Out-okumpu ophiolite complex (2.0 Ga), Finland (Vuollo et al., 1995). The Zunhua chromiteores exhibit remarkable similarities to the podiform ores described from the examples men-tioned above.

Characteristic features of podiform chromites include: (1) lensoidal geometry, typicallyelongate along foliation although more rarely as discordant pods; (2) chromite bands withrootless to tight folds; (3) unique magmatic textures and structures, such as dunite en-velopes for the chromite ores, and nodular, orbicular, and high-temperature shear fabricsof antinodular chromites; and (4) strong plastic deformation associated with harzburgite,with metamorphic foliations and lineations. The origin of the podiform chromitites is at-tributed to melt-rock reaction, or dynamic magmatism within melt channels in the upperoceanic mantle (Nicolas and Azri, 1991; Zhou et al., 1996). The presence of water in themelt is thought to be important for the formation of podiform chromite (e.g., Edwards etal., 2000).

The Zunhua chromitites are typical podiform chromitites as classified by Thayer (1969).The chromites are variably deformed from strongly-stretched to weakly reworked. Thenodular and orbicular chromites, apparently first described by Johnston (1936) fromthe Josephene ophiolite of Northern California, are characteristic of alpine-type peri-dotites or ophiolitic chromite ores (Nicolas, 1989; Gass et al., 1984; Zhou et al., 1996;Lippard et al., 1986; Peters et al., 1991; Dilek et al., 2000). It is now believed that this typeof chromite accumulated below the transition between oceanic crust and mantle based onnumerous investigations in ophiolites. Inclusions within chromites, olivine, and orthopy-roxene of the host peridotites record high-temperature plastic deformation. The flatteningand elongation of chromite parallel to foliation and lineation are intensive high-temperatureshear strain. These textures probably record the plastic flow of the upper mantle, nowmainly preserved in the core of tectonic blocks. These early lineations defined by deformedmagmatic inclusions and the elongation of ore zones are not parallel to later lineations re-lated to the emplacement of the blocks along shear zones, supporting the idea that theyrepresent early mantle-deformation related fabrics. Podiform chromitites are remarkablyresilient to later deformation and metamorphism since they are generated at high tem-peratures (1200–1300 ◦C) and become very rigid when cooled, thus resisting later shear.These asthenospheric chromite pods are miniature time capsules preserving extraordinaryamounts of information about the Archean mantle that we have only begun to tap andunderstand.

264 Chapter 7: Origin and Emplacement of Archean Ophiolites of the Central Orogenic Belt

Chromite can be used as a petrogenetic indicator for tectonically uncertain ultramaficmassifs. The mineral chemistry of the Zunhua chromitites suggests that they formed in asuprasubduction zone ophiolitic environment. The Cr#, TiO2 wt% and V2O5 wt% indicatepartial melting of a depleted source under high fO2 conditions. Applied to the moderntectonic paradigm these compositions indicate formation in a SSZ setting and fluid aidedmelting of a previously depleted source. However, due to different geothermal gradients,and possibly magmatic systematics, analogy with modern and Phanerozoic systems maynot be directly applicable. Higher heat flow in the Archean could lead to great enoughdegrees of partial melting to form high Cr# number chromite without fluid aided meltingof previously depleted mantle.

The Zunhua chromitites occur within dunite envelopes, grading into mantle harzburgite.Such textures form by interaction between a melt and a depleted mantle source. This raisesthe possibility that mantle and magmatic dynamics were similar to modern systems andthat suprasubduction zone environments operated by similar magmatic systems to those oftoday.

Comparison between Zunhua chromite ores and younger examples reveals a sur-prising similarity in their textures and structures. Podiform chromitites are present al-most exclusively in ophiolites, being generated in the uppermost oceanic mantle beneathactive spreading ridges above intraoceanic suprasubduction zones (Lago et al., 1982;Leblanc, 1997; Zhou et al., 1996). Coupled with the presence of a full ophiolite sequencein the Dongwanzi complex, the documentation of the Zunhua chromitites provides con-vincing evidence for the operation of sea-floor spreading and plate tectonics during theArchean before 2.50 Ga. We prefer to ascribe a faster to moderate spreading rate to theformation of the Zunhua podiform chromitites, as podiform chromite is mainly associ-ated with harzburgite-type (HOT) ophiolites (Nicolas and Azri, 1991). Although the fieldand petrographic observations are consistent with the Neoarchean ophiolites of the CentralOrogenic belt preserving relatively hot mantle features, we do not have evidence that thismantle record was any hotter than the present day range of mantle temperatures. However,the hot Archean North China mantle is consistent with some of the higher heat productionduring the Archean being accommodated by faster creation of oceanic lithosphere from aslightly hotter oceanic asthenosphere.

11. EMPLACEMENT OF ARCHEAN OPHIOLITES AND MANTLE BOUDINS INTHE CENTRAL OROGENIC BELT; IMPLICATIONS FOR THE EVOLUTION OFTHE NORTH CHINA CRATON AND ARCHEAN CONTINENTAL GROWTH

The Central Orogenic belt is over 1,600 km long, and separates the distinctly differentEastern and Western Blocks of the North China craton. It contains a diverse suite of rocksincluding high-pressure granulites, strongly deformed metasedimentary and metavolcanicrocks, and a number of different intrusive plutonic suites and dike swarms. The CentralOrogenic belt is bounded on the east by a group of sedimentary basins that are filled withflysch-molasse like sequences of graywacke, shale, and conglomerate. We interpret the

11. Emplacement of Archean Ophiolites and Mantle Boudins in the Central Orogenic Belt 265

Central Orogenic belt as the ca. 2.5 Ga suture between the Eastern and Western Blocks ofthe North China craton.

We interpret the mafic and ultramafic blocks in the biotite gneiss matrix to representa strongly dismembered ophiolite in a metasedimentary and metavolcanic matrix. Rela-tionships are strongly reminiscent of younger ophiolitic mélange terranes, where blocksof ophiolite are preserved in a metasedimentary accretionary prism/trench complex (e.g.,Kusky et al., 1997; Kusky and Polat, 1999; Kusky and Young, 1999). The Zunhua ophi-olitic blocks in mélange do not preserve an overall younging direction, although a fewof the blocks show younging directions toward the west. Similarly, the Dongwanzi ophi-olite to the northeast preserves an overall westward-younging sequence. The foliation inthe Zunhua structural belt strikes north to northeast, and dips steeply west to northwest.The ophiolitic relicts clearly underwent intense tectonic transposition and amphibolite-facies metamorphism during or after their structural emplacement. The ultramafic rocksof the ophiolite are characterized by intense strain and structural dislocation associatedwith major thrusting and recumbent folding. As a result they are widely tectonically inter-leaved with country rock, and have mylonitic margins. Shear zones within gneiss separatedifferent parts of the ophiolite complex. These tectonic slices were dismembered duringobduction over an older gneissic terrane to the east.

These relationships suggest, although do not require, that the ophiolites were em-placed into the mélange during westward directed subduction, then thrust over the east-ern block during closure of the intervening ocean basin (Fig. 21). In this model the con-temporaneous arc would be located to the west of the Zunhua structural belt. We inter-pret a narrow belt of deeply eroded and strongly metamorphosed 2.55–2.50 Ga arc-typeTTG plutonic rocks and a greenstone belt in the Wutai-Hengshan-Taihang Mountains tothe southwest (Figs. 2 and 21) to represent the remnants of this arc (Li et al., 2000a;Wilde et al., 1998). The ca. 2.50–2.40 Ga Qinglong, Hutuo, and Dengfeng sedimentary se-quences and other similar basinal deposits east of the Central Orogenic belt (Figs. 2 and 21)may represent the foreland basin sequence resulting from the collision of the east and westblocks. These basin sequences consist of lower turbidite and upper molasse sequences,with more intense thrusting and folding in the west adjacent to the Central Orogenic belt.The 2.50–2.40 Ga granitoids that intrude the base of the ophiolite and much of the CentralOrogenic belt could represent collisional to post-collisional granites formed during crustalthickening during orogenesis. This model also explains the exhumation of ca. 2.50 Gahigh-pressure granulites and retrograde eclogites in the Hengshan belt to the west (Kuskyand Li, 2003) (Fig. 21). The Hengshan granulite belt is over 700 km long, and is associatedwith the Neoarchean arc rocks on the western side of the Central Orogenic belt. If the con-tinental margin of the eastern block were deeply subducted beneath the overriding arc ofthe Western block, it would rebound isostatically soon after collision. This rapid uplift andexhumation would remove most of the arc sequence, now preserved as the eclogites andgranulites of the prominent Hengshan belt on the western side of the Central Orogenic belt(Fig. 21). Deepest levels of exhumation are recorded in the north, whereas central sectorsof the Central Orogenic belt (Wutai Mountain) preserve lower-grade 2.50 Ga arc plutonicrocks and greenstone sequences. Other evidence for this late-stage exhumation and col-

266 Chapter 7: Origin and Emplacement of Archean Ophiolites of the Central Orogenic Belt

Fig. 21. Tectonic evolution of the Central Orogenic belt of the North China craton.

lapse of the Central Orogenic belt is provided by a swarm of 2.50–2.40 Ga extensionalmafic dikes and flood basalts which occur in a larger area of the Central Orogenic belt(Fig. 21). After 2.40 Ga, an extensional graben system to cratonic basin developed withdeposition of marine carbonate-mudstone, represented by the Middle to Upper Paleopro-terozoic sequence (2.40–1.90 Ga) of the North China craton (Kusky and Li, 2003).

The Dongwanzi ophiolite and Zunhua ophiolitic mélange are located in a previouslyidentified suture zone separating two different cratonic blocks, and there is a parallel high-

Acknowledgements 267

pressure granulite belt (2.5 Ga) almost distributed in the same trend. These two differentbelts record the subduction and final closure of a former oceanic basin (2.50 Ga). It hasbeen suggested that there were one or two supercontinents in the Neoarchean (Kenoraland;Heaman, 1997). The 2.5 Ga orogeny in the North China craton might record the finalassembly of this supercontinent.

The Central Orogenic Belt is cored by an internal zone with high-grade metamorphicnappes structurally overlying a lower-grade thrust imbricate zone containing many ophi-olitic fragments thrust over a passive margin sequence. A foreland basin sequence hassyn-orogenic clastic rocks in the west succeeded eastward by a thick flysch sequence. Theoverall structural zonation of the orogen is much like younger examples, showing thatcollisional plate processes similar to those of the Phanerozoic were in operation in theArchean.

ACKNOWLEDGEMENTS

This work was supported by the U.S. National Science Foundation grants 02-07886and 01-25925 (awarded to T. Kusky), China National Natural Science Foundation grant49832030 (awarded to J.H. Li), Peking University Project 985, and St. Louis University.This work has benefited greatly from discussions with numerous scientists, with specialemphasis by Brian Windley, Kevin Burke, Alfred Kröner, Mingo Zhai, Ali Polat, and KentCondie.

REFERENCES

Abbott, D., 1996. Plumes and hotspots as sources of greenstone belts. Lithos 37, 113–127.Ahmed, A.H., Arai, S., Attia, A.K., 2001. Petrological characteristics of podiform chromites and

associated peridotites of the Pan African Proterozoic ophiolite complexes of Egypt. MineraliumDeposita 36, 72–84.

Anonymous, 1972. Ophiolites. Geotimes 17, 24–25.Arai, S., 1992. Chemistry of chromian spinel in volcanic rocks as guide to potential magma chem-

istry. Mineralogical Magazine 56, 173–184.Arai, S., 1997a. Origin of podiform chromitites. Journal of Asian Earth Sciences 15, 303–310.Arai, S., 1997b. Control of wall-rock composition on the formation of podiform chromitites as a

result of magma/peridotite interaction. Resource Geology 47, 177–187.Arai, S., Yurimoto, H., 1995. Possible sub-arc origin of podiform chromitites. The Island Arc 4,

104–111.Auge, T., 1987. Chromite deposits in the northern Oman ophiolite: Mineralogical constraint. Miner-

alium Deposita 22, 1–10.Bai, J., Dai, F.Y., 1996. The early Precambrian crustal evolution of China. Journal of Southeast Asian

Earth Sciences 13, 205–214.Bai, J., Dai, F.Y., 1998. Archean crust of China. In: Ma, X.Y., Bai, J. (Eds.), Precambrian Crustal

Evolution of China. Springer Geological Publishing, Beijing, pp. 15–86.

268 Chapter 7: Origin and Emplacement of Archean Ophiolites of the Central Orogenic Belt

Bai, J., 1996. Precambrian Crustal Evolution of China. Geological Publishing House, Beijing, p. 165.Bai, J., Wang, R.Z., Guo, J.J., 1992. The Major Geological Events of Early Precambrian and Their

Dating in Wutaishan Region. Geological Publishing House, Beijing, pp. 1–60.Berhe, S.M., 1990. Ophiolites in north and east Africa: implications for Proterozoic crustal growth.

Journal of the Geological Society of London 147, 41–57.Bickle, M.J., Nisbet, E.G., Martin, A., 1994. Archean greenstone belts are not oceanic crust. Journal

of Geology 102, 121–138.Brandl, G., de Wit, M.J., 1997. The Kaapvaal Craton. In: de Wit, M.J., Ashwal, L.D. (Eds.), Green-

stone Belts. In: Oxford Monographs on Geology and Geophysics, vol. 35, pp. 581–607.Canil, D., 1997. Vanadium partitioning and the oxidation state of Archean komatiite magmas. Na-

ture 389, 842–844.Canil, D., 1999. Vanadium portioning between orthopyroxene, spinel and silicate melt and the redox

states of mantle source regions for primary magmas. Geochimica et Cosmochimica Acta 63, 557–572.

Chadwick, B., Crewe, M.A., 1986. Chromite in the early Archean Akilia association (ca 3800 m.y.)Ivisartoq region, inner Gothabsfjord, southern Greenland. Economic Geology 81, 184–191.

Cotterill, P., 1969. The chromite deposits of Selukwe, Rhodesia. Economic Geology Monograph 4,154–186.

Dann, J.C., 1991. Early Proterozoic ophiolite, central Arizona. Geology 19, 590–593.Dann, J.C., 1997a. Pseudostratigraphy and origin of the Early Proterozoic Payson ophiolite, central

Arizona. Bulletin of the Geological Society of America 109, 347–365.Dann, J.C., 1997b. Branching sheeted dikes and seafloor spreading within an Early Proterozoic intra-

arc basin. Journal of Geophysical Research 102, 24,917–24,929.Dann, J.C., 2004. The 1.73 Ga Payson Ophiolite, Arizona, USA. In: Kusky, T.M. (Ed.), Precam-

brian Ophiolites and Related Rocks. In: Developments in Precambrian Geology, vol. 13. Elsevier,Amsterdam, pp. 73–93.

Davis, G.A., Qian, X.G., Zheng, Y.D., Tong, H.M., Yu, H., Wang, C., Gehrils, G., Shafiquallah, M.,Fryxell, J., 1996. Mesozoic deformation and plutonism in the Yunmeng Shan: a metamorphiccore complex north of Beijing, China. In: Yin, A., Harrison, T.M. (Eds.), The Tectonic Evolutionof Asia. Cambridge Univ. Press, pp. 253–280.

de Wit, M.J., Hart, R.A., Martin, A., Abbott, P., 1982. Archean abiogenic and probable biogenicstructures associated with mineralized hydrothermal vent systems and regional metasomatism,with implications for greenstone belt studies. Economic Geology 77, 1783–1802.

de Wit, M.J., Hart, R.A., Hart, R.J., 1987. The Jamestown ophiolite complex, Barberton mountainbelt: a section through 3.5 Ga oceanic crust. Journal of African Earth Science 6, 681–730.

de Wit, M.J., Roering, C., Hart, R.J., Armstrong, R.A., de Ronde, C.E.J., Green, R.W.E., Tredoux,M., Peberdy, E., Hart, R.A., 1992. Formation of an Archean continent. Nature 357, 553–562.

Dick, H.J.B., Bullen, T., 1984. Chromian spinel as a petrogenetic indicator in abyssal and alpine-typeperidotites and spatially associated lavas. Contributions to Mineralogy and Petrology 86, 54–76.

Dilek, Y., Moores, E., Elthon, D., Nicolas, A. (Eds.), 2000. Ophiolites and Oceanic Crust, New In-sights from Field Studies and the Ocean Drilling Program. Geological Society of America SpecialPaper 349, 552.

Droop, G.T.R., 1987. A general equation for estimating Fe3+ concentrations in ferromagnesian sil-icates and oxides from microprobe analyses using stoichiometric criteria. Mineralogical Maga-zine 51, 431–435.

Echeverria, L.M., 1980. Tertiary or Mesozoic komatiites from Gorgona Island, Colombia field rela-tions and geochemistry. Contributions to Mineralogy and Petrology 73 (3), 253–266.

References 269

Edwards, S.J., Pearce, J.A., Freeman, J., 2000. New insights concerning the influence of water duringthe formation of podiform chromitite. In: Dilek, Y., Moores, E., Elthon, D., Nicolas, A. (Eds.),Ophiolites and Oceanic Crust, New Insights from Field Studies and the Ocean Drilling Program.Geological Society of America Special Paper 349, 139–148.

Evans, B.W., Frost, B.R., 1975. Chrome-spinels in progressive metamorphism—a preliminary analy-sis. Geochimica et Cosmochimica Acta 39, 959–972.

Fang, L., Friend, C.R.L., Li, Q., Li, S.Z., Liu, W., Powell, D., Thirwall, M.F., Yang, Z.S., Zhang,Q.H., 1998. Geology of the Santunying Area of Eastern Hebei Province. Geological PublishingHouse, Beijing, p. 134.

Gao, S., Rudnick, R., Carlson, R., McDonough, W., Liu, Y.S., 2002. Re-Os evidence for replace-ment of ancient mantle lithosphere beneath the North China craton. Earth and Planetary ScienceLetters 6135, 1–15.

Gass, I.G., Lippard, S.J., Shelton, A.W., 1984. Ophiolites and Oceanic Lithosphere. In: The Geolog-ical Society Special Publications, vol. 13. Blackwell Scientific, Oxford, p. 413.

Griffin, W.L., Zhang, A., O’Reilly, S.Y., Ryan, C.G., 1998. Phanerozoic evolution of the lithospherebeneath the Sino-Korean craton, in Mantle Dynamics and Plate Interactions in East Asia, A.G.U.Geodynamics 27, 107–126.

Harper, G.D., 1999. Structural styles of hydrothermal discharge in ophiolite/sea floor systems. Re-views in Economic Geology 8, 53–73.

Harper, G.D., 1985. A dismembered Archean ophiolite, Wind River Mountains, Wyoming (USA).Ophioliti 10, 297–306.

He, T.X., Lin, Q., Fang, Z.R., et al., 1992. The Petrogenesis of Granitic Rocks in Eastern Hebei.Jiling Science and Technology Press, Changchun. pp. 1–4 (in Chinese with English abstract).

He, G.P., Lu, L.Z., Ye, H.W., 1991. The Early Precambrian Metamorphic Evolution of the EasternHebei and the Southeastern Inner Mongolia. Jilin Univ. Press, Changchun, pp. 1–17 (in Chinesewith English abstract).

Heaman, L.M., 1997. Global mafic magmatism at 2.45 Ga; remnants of an ancient large igneousprovince? Geology 25, 299–302.

Helmstaedt, H., Padgham, W.A., Brophy, J.A., 1986. Multiple dikes in the lower Kam Group, Yel-lowknife greenstone belt: Evidence for sea floor spreading? Geology 14, 562–566.

Holtzman, B., 2000. Guaging stress from mantle chromitite pods in the Oman ophiolite. In: Dilek, Y.,Moores, E.M., Elthon, D., Nicolas, A. (Eds.), Ophiolites and Oceanic Crust: New Insights fromField Studies and the Ocean Drilling Program. Geological Society of America Special Paper 349,149–158.

Huang, X.N., Li, J.H., Kusky, T.M., Chen, Z., 2004. Microstructures of the Zunhua 2.50 Ga Podiformchromite, North China Craton and implications for the deformation and rheology of the Archeanoceanic lithosphere mantle. In: Kusky, T.M. (Ed.), Precambrian Ophiolites and Related Rocks.In: Developments in Precambrian Geology, vol. 13. Elsevier, Amsterdam, pp. 321–337.

Huson, R., Kusky, T.M., Li, J.H., 2004. Geochemical and petrographic characteristics of the centralbelt of the Archean Dongwanzi ophiolite complex. In: Kusky, T.M. (Ed.), Precambrian Ophiolitesand Related Rocks. In: Developments in Precambrian Geology, vol. 13. Elsevier, Amsterdam,pp. 283–320.

Irvine, T.N., 1965. Chromian spinel as a petrogenetic indicator: Part 1, Theory. Canadian Journal ofEarth Science 2, 648–671.

Irvine, T.N., 1967. Chromian spinel as a petrogenetic indicator: Part 2, Petrologic applications. Cana-dian Journal of Earth Science 4, 71–103.

270 Chapter 7: Origin and Emplacement of Archean Ophiolites of the Central Orogenic Belt

Irvine, T.N., 1974. Petrology of the Duke Island ulramafic complex, southeastern Alaska. GeologicalSociety of America Memoir 138, 240.

Isachsen, C.I., Bowring, S.A., Padgham, W.A., 1991. U-Pb zircon geochronology of the Yellowknifevolcanic belt, N.W.T.: New constraints on the timing and duration of greenstone belt magmatism.Journal of Geology 99, 55–67.

Isachsen, C.I., Bowring, S.A., 1997. The Bell Lake Group and Anton Complex—a basement coversequence beneath the Yellowknife greenstone belt revealed and implicated in greenstone beltformation. Canadian Journal of Earth Science 34, 169–189.

Jahn, B.M., Zhang, Z.Q., 1984a. Archean granulite gneisses from Eastern Hebei Province, China:Rare Earth geochemistry and tectonic implications. Contributions to Mineralogy and Petrol-ogy 85, 225–243.

Jahn, B.M., Zhang, Z.Q., 1984b. Radiometric ages (Rb-Sr, Sm-Nd, U-Pb) and REE geochemistryof Archean granulite gneisses from Eastern Hebei Province, China. In: Kröner, A., Hanson, G.,Goodwin, A. (Eds.), Archean Geochemistry. Springer-Verlag, Berlin, pp. 204–244.

Jahn, B.M., Auvray, B., Cornichet, J., Bai, Y.L., Shen, Q.H., Liu, D.Y., 1987. 3.5 Ga amphibolitesfrom Eastern Hebei Province, China field occurrence, petrography, Sm-Nd isochron age and REEgeochemistry. Precambrian Research 34, 311–346.

Jan, M.Q., Windley, B.F., 1990. Chromian spinel-silicate chemistry in ultramafics rocks of the JijalComplex, Northwest Pakistan. Journal of Petrology 31, 667–715.

Johnson, P.R., Kattan, F.H., Al-Saleh, A.M., 2004. Neoproterozoic ophiolites in the Arabian Shield:Field relations and structure. In: Kusky, T.M. (Ed.), Precambrian Ophiolites and Related Rocks.In: Developments in Precambrian Geology, vol. 13. Elsevier, Amsterdam, pp. 129–162.

Johnston Jr., W.D., 1936. Nodular, orbicular, and banded chromite in Northern California. EconomicGeology 31, 417–427.

Kamenetsky, V.S., Crawford, A.J., Meffre, S., 2001. Factors controlling chemistry of magmaticspinel: an empirical study of associated olivine, Cr spinel and melt inclusions from primitiverocks. Journal of Petrology 42.

Karson, J.A., 2001. Oceanic crust when Earth was young. Science 292, 1076–1077.Kelemen, P.B., 1990. Reaction between ultramafic rock and fractionating basaltic magma I. Phase

relations, the origin of cal-alkaline magma series and the formation of discordant dunite. Journalof Petrology 31, 51–98.

Kelemen, P.B., Joyce, D.B., Webster, J.D., Holloway, J.R., 1990. Reaction between ultramafic rockand fractionating basaltic magma II. Experimental investigation of reaction between olivinetholeiite and harzburgite at 1150–1050 and 5 kb. Journal of Petrology 31, 99–134.

Kepezhinskas, P.K., Taylor, R.N., Tanaka, H., 1993. Geochemistry of plutonic spinels from the NorthKamchatka Arc: comparisons with spinels from other tectonic settings. Mineralogical Maga-zine 57, 575–589.

Kontinen, A., 1987. An Early Proterozoic ophiolite—The Jourma Mafic-Ultramafic Complex, north-eastern Finland. Precambrian Research 35, 313–341.

Kröner, A., 1985. Ophiolites and the evolution of tectonic boundaries in the late Proterozoic Arabian-Nubian Shield of north east Africa and Arabia. Precambrian Research 27, 277–300.

Kröner, A., Cui, W.Y., Wang, S.Q., et al., 1998. Single zircon ages from high-grade rocks of theJianping complex, Liaoning province, NE China. Journal of Asia Earth Sciences 16 (5–6), 519–532.

Kröner, A., Wildes, S., Wang, K., Zhao, G.C., 2002. Age and evolution of a late Archean to earlyProterozoic upper to lower crustal section in the Wutaishan/Hengshan/Fuping terrain of northernChina, A Field Guide. GSA Penrose Conference, Beijing, China, September 2002.

References 271

Kusky, T.M., 1987. Comment on “Multiple dikes in the lower Kam Group, Yellowknife GreenstoneBelt: Evidence for Archean sea-floor spreading?” Geology 15, 280–282.

Kusky, T.M., 1990. Evidence for Archean ocean opening and closing in the southern Slave Province.Tectonics 9, 1533–1563.

Kusky, T.M., 1991. Structural development of an Archean orogen, western Point Lake, NorthwestTerritories. Tectonics 10, 820–841.

Kusky, T.M., Li, J.H., 2002. Is the Dongwanzi complex an Archean Ophiolite? Response to Zhai,M., Zhao, G., Zhang, Q. Science 295, A923.

Kusky, T.M., Li, J.H., Tucker, R.T., 2001. The Archean Dongwanzi ophiolite complex, North ChinaCraton: 2.505 Billion Year Old Oceanic Crust and Mantle. Science 292, 1142–1145.

Kusky, T.M., Li, J.H., 2003. Paleoproterozoic tectonic evolution of the North China Craton. Journalof Asian Earth Sciences 22, 383–397.

Kusky, T.M., Polat, A., 1999. Growth of Granite-Greenstone Terranes at Convergent Margins andStabilization of Archean Cratons. In: Marshak, S., van der Pluijm, B. (Eds.), Special Issue onTectonics of Continental Interiors. Tectonophysics 305, 43–73.

Kusky, T.M., Vearncombe, J., 1997. Structure of Archean Greenstone Belts. In: de Wit, M.J., Ashwal,L.D. (Eds.), Tectonic Evolution of Greenstone Belts. In: Oxford Monographs on Geology andGeophysics, vol. 35, pp. 95–128.

Kusky, T.M., Bradley, D.C., Haeussler, P.J., Karl, S., 1997. Controls on accretion of flysch andmélange belts at convergent margins: Evidence from the Chugach Bay thrust and Icewormmélange, Chugach Terrane, Alaska. Tectonics 16, 855–878.

Kusky, T.M., Li, J.H., Raharimahefa, T., Carlson, R.W., 2004. Re-Os isotope chemistry andgeochronology of chromite from mantle podiform chromites from the Zunhua ophiolitic mélangebelt, N. China: Correlation with the Dongwanzi ophiolite. In: Kusky, T.M. (Ed.), PrecambrianOphiolites and Related Rocks. In: Developments in Precambrian Geology, vol. 13. Elsevier, Am-sterdam, pp. 275–282.

Kusky, T., Young, C., 1999. Emplacement of the Resurrection Peninsula ophiolite in the southernAlaska Forearc During a Ridge-Trench Encounter. Journal of Geophysical Research 104 (B12),29,025–29,054.

Kusky, T., Glass, A., Tucker, R., Harper, G., Bradley, D., Ozdogan, M., in preparation. Structure,Chemistry, Age and Origin of the Border Ranges Ultramafic/Mafic Complex: A SuprasubductionZone Ophiolite Complex. Journal of Geophysical Research.

Lago, B.L., Rabinowicz, M., Nicolas, A., 1982. Podiform chromitite ore bodies: a genetic model.Journal of Petrology 23, 103–125.

Leblanc, M., 1997. Chromitite and ultramafic rock compositional zoning through a paleotransformfault Poum, New Caledonia—A reply. Economic Geology 92, 503–504.

LeBlanc, M., Nicholas, A., 1992. Ophiolitic chromitites. International Geology Review 34, 653–686.Li, J.H., Qian, X.L., Gu, Y.C., 1998. Outline of Paleoproterozoic tectonic division and plate tectonic

evolution of North China Craton. Earth Science 23, 230–235.Li, J.H., Qian, X.L., Huang, X.N., Liu, S.W., 2000a. The tectonic framework of the basement of

North China Craton and its implication for the early Precambrian cratonization. Acta PetrologicaSinica 16 (1), 1–10.

Li, J.H., Kröner, A., Qian, X.L., O’Brien, P., 2000b. The tectonic evolution of early Precambrianhigh-pressure granulite belt, North China Craton (NCC). Acta Geological Sinica 274 (2), 246–256.

272 Chapter 7: Origin and Emplacement of Archean Ophiolites of the Central Orogenic Belt

Li, J.H., Kusky, T.M., Huang, X., 2002. Neoarchean podiform chromitites and harzburgite tec-tonite in ophiolitic mélange, North China Craton, Remnants of Archean oceanic mantle. GSAToday 12 (7), 4–11.

Lippard, S.J., Shelton, A.W., Gass, I.G., 1986. The ophiolite of northern Oman. Geological Societyof London Memoir 11, 178.

MacLaughlin, K., Helmstaedt, H., 1995. Geology and geochemistry of an Archean mafic dike com-plex in the Chan Formation: basis of a revised plate tectonic model of the Yellowknife greenstonebelt. Canadian Journal of Earth Science 32, 614–630.

Matsumoto, I., Arai, S., 1999. Morphological variations of chromian spinel in dunite and harzburgitefrom the Sangun Zone, Southwest Japan, as a marker of melt/peridotite reaction. Science Reportsof the Kanazawa University 44 (1–2), 11–24.

Matveev, S., Balhaus, C., 2002. Role of water in the origin of podiform chromite deposits. Earth andPlanetary Science Letters 203, 235–243.

Menzies, M.A., Fan, W., Zhang, M., 1993. Paleozoic and Cenozoic lithoprobes and the loss of> 120 km of Archean lithosphere, Sino-Korean craton, China. In: Prichard, H.M., Alabaster, T.,Harris, N.B., Neary, C.R. (Eds.), Magmatic Processes and Plate Tectonics. Geological SocietySpecial Publication 76, 71–81.

Moores, E.M., 1982. Origin and emplacement of ophiolites. Reviews in Geophysics 20, 735–750.Moores, E.M., 2002. Pre-1 Ga (pre-Rodinian) ophiolites: their tectonic and environmental implica-

tions. Geological Society of America Bulletin 114, 80–95.Nicolas, A., 1989. Structures of Ophiolites and Dynamics of Oceanic Lithosphere. Kluwer Acad-

emic, Boston, p. 367.Nicolas, A., Azri, H.A., 1991. Chromite-rich and chromite-poor ophiolites: the Oman case. In: Pe-

ters, T.J., Nicolas, A., Coleman, R.G. (Eds.), Ophiolite Genesis and Evolution of the OceanicLithosphere. Kluwer Academic, Boston, pp. 261–274.

Nicolas, A., Boudier, F., 2000. Large mantle upwellings and related variations in crustal thicknessin the Oman ophiolite. In: Dilek, Y., Moores, E.M., Elthon, D., Nicolas, A. (Eds.), Ophiolitesand Oceanic Crust: New Insights from Field Studies and the Ocean Drilling Program. GeologicalSociety of America Special Paper 349, 67–73.

Oberger, B., Loran, J.P., Girardeau, J., Mercier, J.C.C., Pitragool, S., 1995. Petrogenesis of ultramaficrocks and associated chromitites in the Nan Uttaradit ophiolite, Northern Thailand. Lithos 35,153–182.

Parkinson, I.J., Pearce, J.A., 1998. Peridotites from the Izu-Bonin-Mariana Forearc (ODP Leg 125):Evidence for mantle melting and melt-mantle interaction in a supra-subduction zone setting. Jour-nal of Petrology 39, 1577–1618.

Parson, L.M., Murton, B.J., Browning, P. (Eds.), 1992. Ophiolites and Their Modern OceanicAnalogs. Geological Society Special Publication 60, 330.

Peltonen, P., Kontinen, A., 2004. The Jormua ophiolite: A mafic-ultramafic complex from an an-cient ocean-continent transition zone. In: Kusky, T.M. (Ed.), Precambrian Ophiolites and RelatedRocks. In: Developments in Precambrian Geology, vol. 13. Elsevier, Amsterdam, pp. 35–71.

Peters, Tj., Nicolas, A., Coleman, R.G., 1991. Ophiolite Genesis and Evolution of the OceanicLithosphere. Kluwer Academic, Boston, p. 903.

Power, M.R., Pirrie, D., Anderson, J.C., Wheeler, P.D., 2000. Testing the validity of chrome spinelchemistry as a provenance and petrogenetic indicator. Geology 28, 1027–1030.

Proenza, J., Gervilla, F., Melgarejo, J.C., Bodinier, J.L., 1999. Al- and Cr-rich chromitites fromthe Mayarí-Baracoa Ophiolitic Belt (Eastern Cuba): Consequence of interaction between volatile-rich melts and peridotites in suprasubduction mantle. Economic Geology 94, 547–566.

References 273

Rasmussen, B., 2000. Filamentous microfossils in a 3.235 million-year-old volcanogenic massivesulphide deposit. Nature 405, 676–679.

Reed, C., 2002. Chimneys from an Ancient Ocean. Geotimes 23.Roeder, P.L., Campbell, I.H., Jamieson, E., 1979. A re-evaluation of the olivine-spinel geothermome-

ter. Contributions to Mineralogy and Petrology 68, 325–334.Sack, R.O., Ghiorso, M.S., 1991. Chromian spinels as petrogenetic indicators: thermodynamics and

petrological applications. American Mineralogist 76, 827–847.Scott, D.J., St. Onge, M.R., Lucas, S.B., Helmstaedt, H., 1991. Geology and chemistry of the Early

Proterozoic Purtuniq ophiolite, Cape Smith Belt, Northern Quebec, Canada. In: Peters, T.J. (Ed.),Ophiolite Genesis and Evolution of the Oceanic Lithosphere. Kluwer Academic, Boston, pp. 817–849.

Scott, D.J., Helmstaedt, H., Bickle, M.J., 1992. Purtuniq ophiolite, Cape Smith belt, northern Quebec,Canada: A reconstructed section of Early Proterozoic oceanic crust. Geology 20, 173–176.

Shen, Q.H., Xu, H.F., Zhang, Z.Q., 1992. The Early Precambrian Granulites in China. GeologicalPublishing House, Beijing, pp. 134–140 (in Chinese).

Shervais, J.W., 1982. Ti-V plots and the petrogenesis of modern and ophiolitic lavas. Earth andPlanetary Science Letters 59, 101–118.

Sleep, N.H., Windley, B.F., 1982. Archean plate tectonics: constraints and inferences. Journal ofGeology 90, 363–379.

Stern, R.J., Johnson, P.R., Kröner, A., Yibas, B., 2004. Neoproterozoic ophiolites of the Arabian-Nubian Shield. In: Kusky, T.M. (Ed.), Precambrian Ophiolites and Related Rocks. In: Develop-ments in Precambrian Geology, vol. 13. Elsevier, Amsterdam, pp. 95–128.

Stowe, C.W., 1987. Chromite deposits of the Shurugwi Greenstone Belt, Zimbabwe. In: Stowe, C.W.(Ed.), Evolution of Chromium Ore Fields. Van Nostrand-Reinhold, New York, pp. 71–88.

Stowe, C.W., 1994. Compositions and tectonic setting of deposits through time. Economic Geol-ogy 89, 528–546.

Sylvester, P.J., Harper, G.D., Byerly, G.R., Thurston, P.C., 1997. Volcanic Aspects. In: de Wit, M.J.,Ashwal, L.D. (Eds.), Greenstone Belts. In: Oxford Monographs on Geology and Geophysics,vol. 35, pp. 55–90.

Thayer, T.P., 1969. Gravity differentiation and magmatic re-emplacement of podiform chromite de-posits. In: Magmatic Ore Deposits. In: Economic Geology Monographs, vol. 4, pp. 132–146.

Tian, Y.Q., 1991. Geology and Mineralization of the Wutai-Hengshan Greenstone Belt. Shanxi Sci-ence and Technology Press, Taiyuan, pp. 137–152 (in Chinese).

Ulmer, C.G., 1974. Alteration of chromite during serpentinisation in the Pennsylvania-Marylanddistrict. American Mineralogist 59, 1236–1241.

Vuollo, J., Liipo, J., Nykanen, V., Piirainen, T., 1995. An early Proterozoic podiform chromitite inthe Outokumpu ophiolite complex, Finland. Economic Geology 90, 445–452.

Wan, Y.S., Geng, Y.S., Wu, J.S., 1998. The geochemistry of early Precambrian metabasaltic rocks ofNorth China Craton. In: Cheng, Y.Q. (Ed.), Proceeding of Precambrian Geology of North ChinaCraton. Geological Publishing House, Beijing, pp. 39–59.

Wang, Q.C., Zhang, S.Q., 1995. The age of the Hongqiyingzi Group: a further discussion. RegionalGeology of China 2, 173–180.

Wang, K., Li, J.H., Hao, J., 1997. Late Archean mafic-ultramafic rocks from the Wutaishan, ShanxiProvince; a possible ophiolitic mélange. Acta Petrologica Sinica 132, 139–151.

Wilde, S.A., Cawood, P.A., Wang, K.Y., Nemchin, A., 1998. SHRIMP U-Pb zircon dating of granitesand gneisses in the Taihangshan-Wutaishan area: implications for the timing of crustal growth inthe North China craton. Chinese Science Bulletin 43, 144–145.

274 Chapter 7: Origin and Emplacement of Archean Ophiolites of the Central Orogenic Belt

Wilks, M.E., Harper, G.D., 1997. Wind River Range, Wyoming Craton. In: de Wit, M.J., Ashwal,L.D. (Eds.), Greenstone Belts. In: Oxford Monographs on Geology and Geophysics, vol. 35,pp. 508–516.

Wood, B.J., Virgo, D., 1989. Upper mantle oxidation state: Ferric iron contents of lherzolite spinelsby 57Fe Mössbauer spectroscopy and resultant oxygen fugacity. Geochimica et CosmochimicaActa 53, 1277–1291.

Wu, J.S., Geng, Y.S., 1991. Major Geological Events of Early Precambrian in North China Platform.Geological Publishing House, Beijing, pp. 1–11 (in Chinese).

Wu, J., Geng, Y.S., Shen, Q.H., 1998. Archean Geology Characteristics and Tectonic Evolution ofSino-Korea Paleocontinent. Geological Publishing House, Beijing, pp. 1–104.

Wu, C.H., Zhong, C.T., 1998. The Paleoproterozoic SW-NE collision model for the central NorthChina Craton. Progress of Precambrian Research 21, 28–50 (in Chinese).

Xu, Z.G., 1990. Mesozoic volcanism and volcanogenic iron ore deposits in eastern China. GeologicalSociety of America Special Paper 237, 46.

Zhang, Q.S., Yang, Z.S., Gao, D.Y., 1991. The Archean High-Grade Metamorphic Geology and GoldDeposits in Jinchangyu Area of Eastern Hebei. Geological Publishing House, Beijing, pp. 1–5 (inChinese).

Zhang, Y.X., Ye, T.S., Yang, H.Q., 1986. The Archean Geology and Banded Iron Formation of Ji-dong, Hebei Province. Geological Publishing House, Beijing, pp. 1–22 (in Chinese with Englishabstract).

Zhao, G., 2001. Paleoproterozoic assembly of the North China Craton. Geological Magazine 138,87–91.

Zhao, G.C., Wilde, S.A., Cawood, P.A., Sun, M., 2001. Archean blocks and their boundaries in theNorth China Craton: lithological, geochemical, structural and P-T path constraints and tectonicevolution. Precambrian Research 107, 45–73.

Zhao, Z.P., 1993. The Precambrian Geological Evolution of Sino-Korean Paraplatform. Science Pub-lishing House, Beijing, pp. 1–83 (in Chinese).

Zheng, Z., O’Reilly, S.Y., Griffin, W.L., Lu, F., Zhang, M., 1998. International Geology Review 40,471–499.

Zhou, M.F., Kerrich, R., 1992. Morphology and composition of chromite in komatiites from theBelingwe Greenstone Belt, Zimbabwe. Canadian Mineralogist 30, 303–317.

Zhou, M.F., Robinson, P.T., Bai, W.J., 1994. Formation of podiform chromitites by melt/rock inter-action in the upper mantle. Mineralium Deposita 29, 98–101.

Zhou, M.F., Robinson, P.T., Malpas, J., Zijin, L., 1996. Podiform chromitites in the Luobusa Ophio-lite (southern Tibet): Implications for melt-rock interaction and chromite segregation in the uppermantle. Journal of Petrology 37, 3–21.

Zhou, M.F., Robinson, P.T., 1997. Origin and tectonic environment of podiform chromite deposits.Economic Geology 92, 259–262.

Ziegler, A.M., Rees, P.M., Rowley, D.B., Bekker, A., Qing, Li, Hulver, M., 1996. Mesozoic assemblyof Asia: constraints from fossil floras, tectonics, and paleomagnetism. In: Yin, A., Harrison, T.M.(Eds.), The Tectonic Evolution of Asia. Cambridge Univ. Press, pp. 371–400.