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Geology, Fluid Inclusions, and Oxygen Isotope Geochemistry of the Baiyinchang Pipe-Style Volcanic-Hosted Massive Sulfide Cu Deposit

in Gansu Province, Northwestern China

HOU ZENGQIAN,†

Institute of Geology, Chinese Academy of Geological Sciences (CAGS), Beijing 100037, People’s Republic of China

KHIN ZAW, CODES ARC Centre of Excellence in Ore Deposits, University of Tasmania, Private Bag 79, Hobart, Tasmania, Australia

PETER RONA, Institute of Marine and Coastal Sciences and Department of Geological Sciences, Rutgers University,

71 Dudley Road, New Brunswick, New Jersey

LI YINQING, QU XIAOMING, SONG SHUHE,Institute of Mineral Resources, Chinese Academy of Geological Sciences (CAGS), Beijing 100037, People’s Republic of China

PENG LIGUI, Xi’an Institute of Geology and Mineral Resources, Xi’an 710054, People’s Republic of China

AND HUANG JIANJUN

Xi’an Institute of Metallurgical Geology, Xi’an 710054, People’s Republic of China

AbstractThe Baiyinchang massive sulfide Cu deposit (Zheyaoshan and Huoyanshan mines) is hosted by an early

Cambrian, submarine, felsic volcanic succession within an extrusive cryptodome associated with an overlyingbasaltic flow, in a Late Proterozoic-early Paleozoic submarine volcanic belt in the north Qilian orogen, north-western China. The deposit is comprised of two mineralized zones: a 30-cm-thick, strata-bound Zn-rich sulfidelens associated with hematitic Fe-Mn cherts, and an underlying, discordant massive ore-dominated sulfidezone enveloped by a hydrothermal alteration pipe that is zoned from chlorite in the center to quartz-sericite atthe margin. The discordant sulfide zone accounts for 90 percent of the Cu reserves of the Zheyaoshan mine.It consists of four main ore types: (1) pipelike pyrrhotite-pyrite ± chalcopyrite ore, (2) massive sulfide ore, (3)a disseminated ore halo, and (4) footwall stringer ore. The pyrrhotite-pyrite ± chalcopyrite pipe has an ellipti-cal shape in plan and is 30 × 50 m across. The pipe partially replaces the overlying massive pyrite lens andextends downward at least 150 m, to be gradually replaced by chalcopyrite-rich stringer veins and chalcopyrite-bearing quartz veins surrounded by a discordant hydrothermal alteration envelope. Massive chalcopyrite-pyritelenses discordant to volcanic bedding, containing relict patches of felsic volcanic host rocks, are commonly en-veloped by a disseminated sulfide halo within a chloritized volcanic unit. These features suggest thatZheyaoshan is a pipe-style deposit that formed mainly by subsea-floor replacement of volcanic host rocks.

Studies of fluid inclusions indicate that there are four types: (1) type I two-phase, aqueous fluid inclusions,(2) type II daughter mineral-bearing, multiphase fluid inclusions, (3) type III CO2-rich fluid inclusions, and(4) type IV CH4-rich fluid inclusions. Type II inclusions have high homogenization temperatures (Th) rangingfrom 320° to 430°C, contain high salinity fluids (31–38 wt % NaCl equiv), and coexist with CO2-rich fluidsfound in vapor-rich, high-Th (up to 487°C), moderate salinity (10–16 wt % NaCl equiv) inclusions in the dis-cordant sulfide zone and associated altered rocks, suggesting a possible contribution of a magmatic fluid to thehydrothermal system. The coexistence of vapor-rich, high-Th (>300°C) and aqueous, low-Th (<220°C) type Ifluid inclusions in the stringer zone suggests that heated seawater mixed with magmatic fluid (gas) in the feederzone. Most type I fluid inclusions in the massive chalcopyrite-pyrite body and in the strongly chloritized pipehave a low Th (62°–225°C) and high salinities (15.0–23.0 wt % NaCl equiv), suggesting that a dense brine zonedeveloped in fractures in the subsea floor where sulfides accumulated by open-space filling and replacementof host volcanic rocks.

Eleven quartz samples from the overlying discordant sulfide zone yielded a restricted range of δ18O valuesbetween 8.8 and 11.1 per mil, from which we calculate that the corresponding hydrothermal fluids had δ18Ovalues ranging from –5.3 to +3.1 per mil over a temperature range of 160° to 278°C. Whole-rock δ18O values

†Corresponding author: e-mail, houzengqian@126.com

©2008 Society of Economic Geologists, Inc.Economic Geology, v. 103, pp. 269–292

IntroductionMOST VOLCANIC-HOSTED massive sulfide (VHMS) depositsare comprised of massive sulfide lenses or sheets and an un-derlying discordant stringer zone, hosted by volcanic rocks(e.g., Franklin et al., 1981; Large et al., 2001; Gifkins et al.,2005). Economically significant stringer zones and pyrite-chalcopyrite pipes are present only in some deposits, such asthe Mount Lyell deposit in the Cambrian Mount Read Vol-canics of western Tasmania (Cox, 1981), the OrdovicianHighway and Reward deposits in the Mount Windsor sub-province, Queensland (Doyle and Huston, 1999), and the De-vonian Mount Morgan deposit, Queensland (Taube, 1986;Ulrich et al., 2002). Pipelike polymetallic sulfides are also pre-sent in the early Paleozoic Bawdwin deposit in northeasternMyanmar (Khin Zaw, 2003, 2004). The Baiyinchang Cu de-posit in the Late Proterozoic-early Paleozoic submarine vol-canic belt of northwestern China may be another example ofa pipe-style VHMS deposit. Bian (1989) argued that it is aporphyry-type deposit, based on the lack of exhalative-sedi-mentary massive sulfides and the close relationship of the sul-fide orebody with felsic porphyries in the district. The depositwas discovered in the 1950s and is being exploited from theZheyaoshan and Huoyanshan mines (Song, 1955, 1982).

The Zheyaoshan mine, with current copper reserves of~0.89 Mt remaining after exploitation of 51 Mt of ore, pro-duced ~0.6 Mt of Cu, whereas 15 Mt of ore was exploitedfrom the Huoyanshan deposit to produce ~0.26 Mt of Cu(Baiyin Company, unpub. data). The copper grade of theBaiyinchang deposits ranges from 0.4 to 4.6 percent, with anaverage of 1.17 percent Cu. Three significant observations inthe district indicate that the Baiyinchang deposit is atypical incomparison to many other submarine hydrothermal systems.

1. The hydrothermal alteration pipe and the associated dis-cordant sulfide zone dips steeply and extends 500 to 700 mdownward into the host felsic volcanic rocks below the pale-osea floor (Song, 1982; Peng et al., 1995). This provides aunique opportunity to study the deep structure of the subma-rine hydrothermal system and the sulfide deposit.

2. The sulfide orebody, hosted in felsic volcanic rocks, isspatially associated with mineralized, shallow-level felsic por-phyry bodies (Cheng, 1980; Song, 1982; Bian, 1989; Peng et

al., 1995) and this provides an opportunity to evaluate the roleof synvolcanic intrusions in the genesis of massive sulfides. (3)The stringer ores and the chalcopyrite-bearing quartz veinsare well developed in the root of the discordant sulfide pipe,and preliminary fluid inclusion studies from this zone foundhigh-temperature (>350°C), high-salinity hydrothermal fluids(Liu, 1982). Herein we describe the geology and the deepstructure of the Baiyinchang deposit, and discuss its genesis,based on detailed fluid inclusion and oxygen isotope studies.

Regional GeologyThe Baiyinchang Cu deposit is the largest VHMS deposit in

the Late Proterozoic to early Paleozoic marine volcanic beltof the north Qilian orogen, northwestern China (Fig. 1A).This belt is 1,200 km long and more than 50 km wide, strikeswest-northwest, and extends from Gansu, Qinghai, andShaanxi to Henan provinces in northwestern China. Thewestern segment of the belt is bounded by the large-scaleAlthy strike-slip fault, whereas the eastern segment is con-nected with the north Qinling early Paleozoic orogenic belt.The marine volcanic belt underwent a complex tectonic his-tory during the Qilian orogeny, beginning with rifting andcontinental break-up in the Late Proterozoic-Early Cam-brian, and followed by the development of a trench-arc-basinsystem in the Ordovician, an arc-continent collision in the lateOrdovician to Devonian, and finally, accretion to the southernmargin of the China-Korea platform (Fig. 1A; Xiang and Dai,1985; Xia et al., 1991, 1996; Wu et al., 1994). The marine vol-canic rocks of the belt host 54 known VHMS deposits (Fig.1A, B) with a total metal content of 5.8 Mt of combined Cu,Pb, and Zn (Wu et al., 1994; Hou et al., 1999).

Three major VHMS-bearing volcanic belts have been rec-ognized in the Qilian orogen based on their spatial-temporaldistribution and the tectonomagmatic associations of thesevolcanic rocks (Fig. 1A). The Late Proterozoic to Early Cam-brian belt (III2) is the most important and hosts a bimodalvolcanic sequence with a range of Sm-Nd ages from 522 to606 Ma (Xia et al., 1996); it is unconformably overlain by arctholeiitic basalt during the Late Cambrian and Ordovician(Fig. 1B; Xia et al., 1998). The bimodal volcanism, associatedwith rifting of the Proterozoic basement, formed at least fourvolcanic domes, termed the Baiyin and Heishishan domes in

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for the altered volcanic rocks in the pipe yielded a much wider range, from 1.6 per mil in the chlorite core to8.7 per mil in the outer sericite-chlorite zone, suggesting that a low-δ18O seawater-dominated hydrothermalfluid interacted with the footwall volcanic rocks. Oxygen isotope data for quartz from both the stringer zoneand the altered host volcanic rocks also record a contribution of the magmatic fluid to the Zheyaoshan subma-rine hydrothermal system. Assuming that the analyzed quartz precipitated from a hot (300°–430°C) hy-drothermal fluid, as suggested by the high-temperature and high-salinity fluid inclusion data, the δ18O valuesof the hydrothermal fluid in equilibrium with quartz (δ18O values of 9.0–11.1‰) range from 2.0 to 8.0 per mil,which corresponds to the δ18O range between magmatic fluid and seawater.

The ore-forming fluids responsible for the Cu-Zn mineralization at Baiyinchang belonged to the H2O-NaCl-CO2-CH4 system. A felsic magma chamber is thought to have been situated 1 to 1.5 km below the sea floor,and this likely supplied the necessary heat to drive seawater convection and introduced an H2O + CO2-domi-nated, high-temperature, high-salinity (magmatic) brine to the Baiyinchang hydrothermal system. A narrow,steeply dipping, funnel-shaped brine zone was present in the margin of the fracture zone above the magmachamber, and this was trapped within the porous felsic volcaniclastic rocks. This brine zone is considered tohave been the key factor for the formation of the Zheyaoshan pipe-shaped sulfide orebodies. The Cu-bearing,high-temperature (>300°C) fluids are considered to be the product of mixing of magmatic brine with seawaterand the replacement of the surrounding volcanic rocks resulted in the formation of the discordant, massive ore-dominated sulfide orebodies at Zheyaoshan.

Gansu Province (Fig. 1B) and the Qingshuigou-Bailiugou andShitougou-Xiangzigou domes in Qinghai Province. Twenty-two VHMS deposits, including Baiyinchang and Xiaotieshan,cluster in the felsic volcanic rocks associated with three of thefour domes, excluding the Heishishan dome.

The Late Cambrian to Early Ordovician volcanic belt (III1)(Fig. 1A) is the second most important belt for VHMS de-posits. Here, submarine volcanic rocks consist mainly of maficvolcanics with a range of Sm-Nd ages between 469 and 454Ma (Xia et al., 1996, 1998). The volcanic rocks are part of aback-arc type ophiolite suite (Xia et al., 1998) and host 23 de-posits, including the Yindonggou, Jiugequan, and Cuogou de-posits (Fig. 1A; Wu et al., 1994; Hou et al., 1999). The Or-dovician-age volcanic belt (III3) developed along the northernmargin of the middle Qilian continental block (Fig. 1A) and iscomprised of mafic and felsic volcanic rocks, with Sm-Ndages from 427.5 to 444.9 Ma (Xia et al., 1998). The volcanismin this belt is bimodal, and the mafic end member shows en-richments in K2O, Rb, Ba, and high field strength elements

(Nb, Ta, P, and Ti) (Xia et al., 1998), suggesting a postcolli-sional, crustal extensional setting. The volcanic rocks hostnine deposits, including the Honggou and Jiaolongzhang Cudeposits (Wu et al., 1994; Hou et al., 1999).

District GeologyThe Baiyin volcanic dome occurs in a fault block bounded

by north-northwest–, north-northeast–, and east-northeast–di-rected fault systems within the Late Proterozoic to Early Cam-brian volcanic belt (III2). This area hosts the four major VHMSdeposits of the Bayin district (Fig. 1B). The Cambrian to Or-dovician volcanosedimentary sequences away from the Baiyindome are strongly folded, whereas the volcanic district centeredon the dome mainly underwent deformation by thrust (F1) andstrike-slip faulting. The elliptical dome is about 7 km long and5.6 km wide and consists of a 700-m-thick pile of felsic volcanicrocks overlain by mafic volcanic rocks (Figs. 1B, 2). Generally,the volcanosedimentary sequences in the volcanic dome havebeen metamorphosed to the lower greenschist facies.

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A

B

Unconformity

Sedimentary rocks - Lower Silurian

Quartz phyric rhyolite - Middle Cambrian

Thrust fault

Fault

Dolerite

Granite dike

Sedimentary sequence - Middle Cambrian

Limestone and marble - Middle Cambrian

Mixed volcanics - Middle Cambrian

Mafic volcanics - Middle Cambrian

Basalt - Lower Ordovican

Sedimentary rocks - Upper Triassic

Main boundary faults(dashed line: inferred)Boundary faults VHMS deposits

City, town

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40 80 120 km0

Q QQ

N

Q

C2

O11

2

34

5

T3

82

85

0 1 2 km

43

36°40'

96°

Subei

YuimenJouquan

Sunan Zhangye

Wuwei

Baiyin

Lanzhou

Fig.1b

III2

Menyuan

100° 104°

36°

38°

I

III1

I

II

IIII2

III1

III3IV

IV

104°15'

C2

C2

C2

O1

T3

DH

DB

F1

S1

Fig.1aXi'an

Beijing

FIG. 1. Simplified geologic map of the marine volcanic belts and volcanic-hosted massive sulfide deposits in the early Pa-leozoic North Qilian orogen, northwestern China (modified after Xiang and Dai, 1985; Xia et al., 1996; Hou et al., 1999). A)Regional-scale tectonic elements and distribution of VHMS deposits in the North Qilian orogen. B) Geologic map andVHMS deposits of the Baiyin district. Q = Quaternary sedimentary rocks, I = Tarim-Sino-Korean platform, II = Hexi corri-dor transitional belt, III = Proterozoic-early Paleozoic marine volcanic belt in the North Qilian (III1: Late Cambrian-EarlyOrdovician northern volcanic belt; III2: Cambrian central volcanic belt; III3: Ordovician southern volcanic belt), IV = Mid-dle Qilian continental block, DB = Baiyin volcanic dome, DH = Heishishan volcanic dome, 1 = Zheyaoshan mine, 2 = Huoy-anshan mine, 3 = Xiaotieshan deposit, 4 = Tongchanggou deposit, 5 = Sigequan deposit.

The four VHMS deposits, each with different metal associ-ations, are (1) the Baiyinchang Cu deposit (Zheyaoshan andHuoyanshan mines), (2) the Xiaotieshan Zn-Pb-Cu deposit,(3) the Tongchanggou Cu-Zn, and (4) the Sigequan Zn-Pb-Cudeposits (Figs. 1, 2). Although these VHMS deposits occur indifferent parts of the dome, they are all hosted by felsic vol-caniclastic rocks (Figs. 2, 3).

Regional mapping indicates that there are at least four hy-drothermal alteration pipes of various sizes developed inthese four deposits. These alteration pipes are zoned fromchlorite in the core to quartz-sericite at the margins (Fig. 2).The pipes have undergone regional deformation and faultingand are aligned along a west-northwest–striking trend, sug-gesting that a west-northwest–directed fault system may havecontrolled the location of the discharging conduits of the sub-marine hydrothermal system.

Volcanic Architecture and Host Rocks of the Baiyinchang Deposit

Volcanic architecture

At least three volcanic systems (centers) have been recog-nized by mapping of the volcanic facies in the Baiyin districtwithin the Baiyin volcanic dome (Peng et al., 1995). The Baiy-inchang volcanic center was likely located on the northernside of the Zheyaoshan mine. An elliptical felsic breccia body,250 m in diameter and 560 m deep, was developed near thevolcanic center (Peng et al., 1995), around which was de-posited a gently dipping volcaniclastic sequence consisting oftuffaceous breccia, tuffaceous lava, and tuffaceous sedimen-tary rocks (Fig. 4). The host volcanic sequence in the centerof the volcanic dome is a mixture of intrusive and volcaniclas-tic rocks. The quartz-bearing rhyolitic rocks partly define anextrusive cryptodome, whereas the associated volcaniclastic

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TongchanggouXiaotieshan

Baiyinchangdeposit

Sigequan

Lapaigou

Zheyaoshanmine

N

0 300 600m

Massive sulfide orebody

Unconformity

Phyllite-limestone Fe-Mn chert

Clastic Sequence (S1)

Quartz-sericite zone

Chlorite zone

Normal Fault

Basaltic lava(C2)

Pillow basalt flows (O1)

Rhyolitic volcanic complex(C2)

Huoyanshan

104°15'

36°40'59"

36°40'57"

Quaternary sediments (Q)

FIG. 2. Geologic map of the Baiyinchang district showing VHMS-bearing felsic volcanic rocks and associated hydrother-mal alteration (modified from Peng et al., 1995).

BAIYINCHANGMiddle Cambrian

980

m22

60 m

8310

m

Mixed basaltic and rhyolitictuffaceous rocks and phyllite

Rhy

oliti

c Vo

lcan

ic C

ompl

ex

Quartz phyric rhyolitic tuffaceousvolcanics, quartz-albite porphyry,and Fe-Mn chert and ore zone

Rhyolitic flow, breccia,and tuffaceous rocks

Rhyolitic agglomerate andtuffaceous breccia

Pillow basalt, massive basaltictuffaceous rock and clasticsedimentary rock

Rhyolitic agglomerateand tuffaceous breccia

Rhyolitic flow, breccia,and tuffaceous rocks

Quartz phyric rhyoliticvolcanicsClastic sedimentaryrockQuartz-albite porphyry

Fe-Mn chert and orezoneRhyolitic tuffaceous rock

PhylliteBasaltic rock

Pillow basalt

FIG. 3. Stratigraphic column of the Baiyinchang deposit, China.

rocks consist of quartz rhyolitic tuffaceous and breccia units.On the western side of the volcanic center, a sedimentary se-quence consisting of phyllite (pyritic shale), limestone, andFe-Mn cherts is developed in a volcanic depression (Fig. 4).A small quartz-albite stock also intruded along a synvolcanicfault in the south and southwest.

Host rocks

The host-rock succession at the Zheyaoshan mine consistsof a volcaniclastic-dominated felsic dome in which por-phyritic volcanic and tuffaceous rocks of quartz rhyolite com-position host most of the discordant massive sulfide orebod-ies. The felsic volcanic rocks are porphyritic in texture andcharacterized by the presence of fractured phenocrysts of al-bite and quartz. These fractured phenocrysts account for 30to 70 vol percent of all the phenocrysts and occur in a ground-mass of albite, quartz, and the secondary hydrothermal alter-ation minerals, sericite, quartz, and chlorite. The quartz rhy-olitic tuffaceous unit is more than 200 m thick, and commonlycontains fragments of quartz and albite crystals and lesserfelsic rock fragments in a tuffaceous matrix consisting of amicrocrystalline assemblage of chlorite, sericite, quartz, andalbite.

A quartz-albite porphyry body, the last product of the felsicmagmatism, intruded along faults surrounding the felsic vol-canic center in the Baiyin district (Fig. 4). These porphyries

have subsequently been chloritized, silicified, and sericitized.At the Huoyanshan mine, the porphyry body is partly miner-alized and forms the no. 194 orebody with porphyry clasts setin a matrix of chalcopyrite and pyrite (Cheng, 1980).

Exhalite

Fe-Mn exhalite is common at Zheyaoshan, where it formsan ore-equivalent horizon near the sulfide orebodies. The ex-halite also occurs as lenses in the western segment of themain orebody (Fig. 4). In the southern wall of theZheyaoshan pit, an exhalite lens is about 30 to 50 m thick andmound shaped. It overlies the strongly altered volcanic hostrocks and is, in turn, conformably overlain by a ~50-m-thicksequence of black shale and limestone. The exhalite unit is>80 m thick and includes hematitic chert, Fe-Mn chert, andsiliceous Mn-Fe nodules. The hematitic chert is red or red-brown and banded or crudely layered. It consists of quartz(75 vol %), hematite (19 vol %) and minor amounts ofsericite and albite (Peng et al., 1995). This exhalite is similarto the hematitic chert or jasper of some Kuroko-type de-posits (Kalogeropoulos and Scott, 1983; Hou et al., 2001).The Fe-Mn chert is dark red and occurs as crudely layeredor irregular lenses that commonly overlie the hematiticchert. The Fe-Mn chert is composed of quartz, hematite,and minor chlorite, albite, sericite, and calcite and has amuch higher Mn content (up to 1.25 wt % MnO) than the

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Zheyaoshan Mine

N

A

A’

Hematitic chert

Phyllite and limestone

Fe-Mn chert and nodule

Tuffaceous breccia Quartz-albite porphyry Massive ore

Disseminated ore

Rhyolitic tuffaceous rock

Quartz phyric rhyolite

Pyrrhotite-pyrite pipePillow basalt Fault

Breccia, tuffaceous lavaRhyolitic breccia

0 100 200 300 m

FIG. 4. Simplified geologic map of the Baiyinchang deposit, showing the volcanic architecture of the ore-bearing rhy-olitic complex, China (modified from Peng et al., 1995; Xia et al., 1998). The highly generalized cross section A-A’ is shownin Figure 6.

hematitic chert (Peng et al., 1995). The siliceous Mn-Fe nod-ules are brown-gray in color and occur as elliptical bodies, 5to 20 cm in length and 4 to 9 cm in width. These nodulescommonly occur in the sequence of black pyritic shale andlimestone. The Mn-Fe nodules are comprised of goethiteand hematite (30 vol %), quartz (27 vol %), hydroxybraunite(21 vol %), and minor anhydrite and chlorite, and are char-acterized by high content of Mn (up to 16 wt % MnO) andFe (up to 24 wt % Fe2O3). Peng et al. (1995) indicated thatthese three types of exhalites have similar REE patterns andstrong negative Eu anomalies, suggesting formation fromlow-temperature hydrothermal fluids (Peter, 2003) and acommon exhalative origin similar to that of the exhalites ofother VHMS belts (Franklin et al., 1981; Kalogeropoulos andScott, 1983; Eastoe et al., 1987; Large, 1992; Leistel et al.,1998).

MineralizationThe Baiyinchang deposit is exploited from two mines,

Zheyaoshan and Huoyanshan, that are located on opposite

sides of the F1 fault (Fig. 2). The deposit is dominated bymassive sulfide and disseminated orebodies with minor stock-work or stringer orebodies (Figs. 5, 6). It belongs to the Cu orCu-Zn type (Large, 1992), and the massive sulfides are pyriterich (Song, 1955, 1982; Wu et al., 1994; Peng et al., 1995).Two mineralized zones have been recognized, based on pet-rography of the host rocks, occurrence and shape of the sul-fide orebodies, and type and structure of the sulfide ores.These main ore zones are comprised of strata-bound sulfidesand exhalites that are interpreted to have been deposited onthe ancient sea floor, and others are discordant massive sul-fide ore hosted in a felsic volcanic rock. The Zheyaoshan de-posit has the morphology of a strata-bound polymetallic mas-sive sulfide lens and underlying discordant massive-stringervein zone. However, the size of the strata-bound lens issmaller than that of the underlying discordant massive andstringer veins enveloped by the alteration pipe. Therefore,the deposit is interpreted to be a typical pipe-style deposit ofLarge (1992), mainly consisting of massive, crosscuttingpyrite-chalcopyrite orebodies.

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FIG. 5. A. Geologic map showing planar distribution of the sulfide orebodies at Zheyaoshan (modified from Wu et al.,1994). The numbered lines are prospecting lines at Zheyaoshan. B. Variation of Cu contents in the sulfide orebodies at the1,799-m level of the Zheyaoshan deposit (after Wu et al., 1994). The disseminated ores predominate between prospectinglines 7 and 11, whereas the massive ores predominate between prospecting lines 2 and 7.

Strata-bound sulfide zone

The strata-bound sulfide zone is small in size but economi-cally significant. This zone is commonly about 30 cm thickand consists of discrete, small sulfide lenses that conformablyoverlie the quartz-phyric rhyolitic tuffaceous rock. In theeastern part of the Zheyaoshan deposit, a banded Zn-rich lensabout 50 m long and 30 cm thick overlies the quartz rhyolitictuffaceous rock and is, in turn, overlain by a 1-m-thickhematitic chert, which in turn is overlain by a 5- to 10-m-thicktuffaceous slate and phyllite unit. The Zn-rich ores are mas-sive, mainly banded or crudely layered (Fig. 7A) and exhibitlaminated and bedded sedimentary structures with associatedcolloform textures (Wu et al., 1994; Peng et al., 1995). In thewestern part of the Zheyaoshan deposit, a smaller Zn-richlens (now mined out) was closely associated with thehematitic chert and Fe-Mn chert (Song, 1982; Peng et al.,1995). At Huoyanshan, similar strata-bound lenses have notbeen observed (Peng et al., 1995).

Massive sulfide zone

This sulfide zone accounts for 90 percent of the total Cu re-serves in the Baiyinchang deposit. Massive sulfide orebodiesand associated hydrothermal alteration zones at Zheyaoshanare discordant to local bedding and are partly enclosed by a

succession of felsic tuffaceous rock and extrusive facies of acryptodome. The zone is about 600 m long and 50 m wide andconsists of a series of steeply dipping sulfide lenses (Figs. 5,6). The ore can be divided into four principal types based onmineralogy, texture, and the relationship to the host rocks: (1)massive pyrrhotite-pyrite pipe, (2) massive sulfide lens, (3)disseminated ore halo, and (4) footwall stringer zone.

Massive pyrrhotite-pyrite pipe: The pyrrhotite-pyrite pipedips steeply to the southwest and consists of a body of mas-sive pyrite, pyrrhotite, and chalcopyrite (Cheng, 1980; Yan,1983; Peng et al., 1995). The pipe is elliptical, 30 × 50 m inplan (Figs. 4, 5), and occurs predominantly beneath themain, massive no. 1 orebody. The pipe extends 150 m down-ward into the hydrothermally altered footwall rocks. Thecontact between the pipe-shaped body and the massive ore-body is abrupt, and there is no evidence to suggest that thispipe-shaped body is related to structural deformation in thedistrict.

The massive pyrrhotite-pyrite pipe has a relatively simplemineralogy of pyrite, pyrrhotite, and chalcopyrite, with minorsphalerite, galena, and magnetite. The pipe shows distinct,concentric ore zoning in plan view. The marginal zone is com-posed of 5 to 10 m of porous, honeycomb-textured pyrite; themiddle zone consists of massive pyrrhotite and pyrite, and theinner zone consists of dense stringer veinlets of chalcopyrite.

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NE25°

Faul

t1715 m

1655 m

1535 m

1475 m

1415 m

Chl

Ser+Qtz

ChlSer+Qtz

A'

A

0 100 200 m

Massive chalcopyrite-pyrrhotite-pyrite ore

Massive pyrite-sphalerite-galena ore

Massive pyrite oreDisseminated ore

Stockwork or stringer ore

Rhyolitic breccia

Hematitic chert

Tuffaceous breccia

Limestone

Tuffaceous rockQuartz-albite stock

Fe-Mn chertRhyolitic volcanic rock

Sericite-quartz zoneSer+Qtz

Chlorite zoneChl

Supergene enrichment zone

Oxidized zone

FIG. 6. Generalized cross section of the Zheyaoshan mine showing vertical ore zoning and alteration zoning of the pipein the felsic volcanic complex (modified from Song, 1982).

The pyrrhotite-pyrite sulfides also form steeply dipping orvertical, pillar-shaped ore lenses (e.g., the tabular no. 41 ore-body), located on the northern side of the no. 1 main massiveorebody at Zheyaoshan (Peng et al., 1995). The upper part ofthe pillar-shaped ore lens is in direct contact with the overly-ing no. 1 main massive orebody. Its base splits into severalsmaller pillars and extends steeply, 500 m stratigraphicallydownward, into the footwall rocks. The ore pillars display nomineral zoning but have a mineral assemblage similar to thepyrrhotite-pyrite pipe. Similar ore pillars also occur at Huoy-anshan, where they directly underlie the massive pyrite lens,and give way downward to a stringer (vein) zone (Cheng,1980; Peng et al., 1995).

Massive sulfide lenses: Massive sulfide lenses dominate inthe discordant sulfide zone hosted by the felsic volcanic fa-cies. The larger massive lenses are 350 to 600 m long and ex-tend downdip 450 to 600 m, whereas the smaller ones are 100to 250 m long and extend downdip 50 to 200 m (Fig. 6). Map-ping of the volcanic facies indicates that these massive lensesare discordant to the bedding of the quartz tuffaceous rock atan oblique angle up to 70° (Cheng, 1980).

Inclusions of wall rocks are generally absent within themassive sulfide lenses. However, relict fragments of chlori-tized felsic host rocks are commonly observed along the mar-gins of the lenses (Fig. 7B). They are generally discontinuousalong strike and irregular in shape. These relict patches of

chloritized rocks with disseminated sulfides commonly com-prise 20 to 30 vol percent of the ore within a 5- to 20-m-widezone at the margin of the massive sulfide lenses. Toward themargins of the massive sulfide lens, an interfingered bound-ary between the wall rock and the massive sulfides can be ob-served locally (Fig. 7B).

These massive sulfide lenses do not show typical sedimen-tary textures, such as lamination, banding, bedding, cross-bedding, or graded bedding, but chemical-mineralogical zon-ing has been observed. Three ore types can be described,based on mineralogy and composition of the massive ores: (1)massive pyrite ± chalcopyrite ore (Fig. 7C), (2) massive spha-lerite-rich pyrite ore (Fig. 7D), and (3) massive chalcopyrite-rich pyrite ores (Fig. 7E). Massive pyrite ore is fine grainedand occurs in the uppermost part of massive sulfide lenses;massive sphalerite-rich pyrite ore is located stratigraphicallybelow the massive pyrite ore and is characterized by a 10 to20 vol percent sphalerite. Massive chalcopyrite-rich pyrite oreoccurs predominantly in the stratigraphically lower part of themassive sulfide lenses and is characterized by a high Cu grade(>1.0 %). The massive chalcopyrite-rich pyrite ore is replacedby a chalcopyrite-rich stringer zone (Fig. 7F) stratigraphicallybeneath it (Fig. 6). Massive sulfide lenses grade laterally out-ward to disseminated sulfide ores (Fig. 5A).

Footwall stringer ore zone: This zone occurs stratigraphi-cally below the massive sulfide lenses and the massive pyrite

276 ZENGQIAN ET AL.

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Cpy

E F

QtzQtz

Cpy+Py

CA

Qtz

D

Cpy+PyB

Altered rock

FIG. 7. Photographs of different ore types from Bayinchang. A. Zn-rich strata-bound sulfide ore, exhibiting crudely lay-ered structures (sample no. BC-26). B. Relics of chloritized rhyolitic volcanic rock from the margin of the massive sulfidelenses, showing interfingered boundary between the wall rock and the massive sulfides (sample no. GBY-071). C. Massivepyrite ± chalcopyrite ore, from the upper part of massive sulfide lenses, interlayered with quartz veinlets (sample no. BZ19).D. Massive Zn(-Pb)-rich pyrite. E. Massive chalcopyrite-rich pyrite ore (sample no. GBY-24), characterized by a high Cugrade, from the stratigraphic lower part of the massive sulfide lenses. F. Chalcopyrite-bearing quartz vein, from stratigraph-ically below the massive sulfide lenses and the massive pyrite pipe, consisting of quartz and minor chalcopyrite. Quartz fromthis material was used for fluid inclusions study (sample no. BZ-09). Cpy = chalcopyrite, Ga = galena, Py = pyrite, Qtz =quartz, Sp = sphalerite.

pipe, and consists of chalcopyrite-rich stringer veins and chal-copyrite-bearing quartz vein swarms (Fig. 7F). The contactbetween the stringer veins and the overlying massive lensand/or pipe is sharp. Although the main contact is located ap-proximately at the 1,550-m level (Fig. 6), chalcopyrite-quartzveins crosscutting the massive Cu-rich pyrite lenses havebeen observed at the 1,573-m level. Chalcopyrite-richstringer veins commonly occur in the center of the stringerzone, and the veins vary in width from 0.2 to 10 cm, with veinspacing ranging between 10 and 30 cm. Chalcopyrite-bearingquartz vein swarms replaced the chalcopyrite-rich stringerveins downward into the altered footwall volcanic rocks. Withdepth, in the quartz vein swarms, chalcopyrite content dra-matically decreases and quartz veins of variable width be-come dominant (Fig. 7F).

Disseminated ore halo: A disseminated ore halo envelopesthe massive pyrite body laterally and vertically. It is 600 to 800m long in an east-west direction and 30 to 50 m wide in anorth-south direction (Figs. 4–6). In general, the intenselydisseminated ore halo is mainly developed in the western partof the Zheyaoshan deposit, whereas a narrow disseminatedore halo occurs in the eastern part (Figs. 4, 5).

Mineralogical zoning

Lateral mineral zoning is characterized by a core of Cu-richmassive sulfides surrounded by a halo of Cu-poor dissemi-nated ore (Fig. 5). In the western part of the Zheyaoshanmine (lines 7–11, Fig. 5B), there are at least two Cu-rich cen-ters, where intensely disseminated and semimassive Cu-richsulfides were developed. In the eastern part (Lines 2–7, Fig.5B), there is an east-west–trending Cu-rich zone, corre-sponding to a large, massive orebody and associated pyrite-pyrrhotite pipe.

Vertical mineral zoning is characterized by a stratigraphicdownward change from a strata-bound Zn-rich zone througha massive sulfide zone into a stringer zone. This progressioncomprises fine-grained massive pyrite at the top, throughmassive sphalerite-pyrite, massive chalcopyrite-pyrite, andchalcopyrite-rich stringer and chalcopyrite-bearing quartzvein swarms (Fig. 6). Corresponding chemical zoning variesfrom a Zn-Pb–rich zone at the top to a Cu-rich zone strati-graphically downward.

Fluid Inclusion StudiesFluid inclusions in quartz from altered host rocks, mineral-

ized stocks, and sulfide ores have been investigated by Liu(1982), Xia et al. (1985), and Peng et al. (1995). However,owing to lack of description of the sample locations and rela-tionships to different sulfide ore types, their data are not suf-ficient to establish the evolution of a submarine hydrothermalsystem at Baiyinchang. In this study, a detailed investigationof fluid inclusions in quartz from the stringer ore, massive sul-fide ore, altered host volcanic rocks, and associated quartz-al-bite porphyry were conducted on over 50 doubly polishedsections. From these, 25 sections were chosen for microther-mometric measurements. The microthermometric data werecollected using Chaixmeca and Linkam heating-freezingstages, with a measured temperature range from –180° to+600°C. Accuracy of the measurements was ensured by cali-bration against the triple point of CO2 (–56.6°C) and the

freezing point (0.0°C) of pure water. The precision of themeasurements is reproducible to within 0.1°C for freezingand 1°C for heating. Tests for the presence of volatile species(CO2, N2) and hydrocarbons (CH4, C2H4, C2H6) were madeon 20 fluid inclusions at the Institute of Mineral Resources,Beijing, using a Microdil 28 laser Raman microprobe.

Petrography of fluid inclusions

Classification of inclusions as primary, pseudosecondary,and secondary was made according to the criteria given byRoedder (1984). The secondary fluid inclusions are generallyvery small (<3 µm) and commonly occur along healed mi-crofractures crosscutting mineral grain boundaries. Primaryinclusions mainly occur as scattered and isolated inclusions inquartz and have regular forms, such as negative crystal andpolygonal shapes. The primary inclusions also occur as clus-ters that are commonly parallel to the hexagonal morphologyof the host quartz crystals. The primary fluid inclusions fromZheyaoshan can be divided into four compositional types (Ta-bles 1, 2), based on the number of phases at room tempera-ture and estimates of fluid compositions based on the ob-served phase transitions during freezing-heating runs.

Type I two-phase aqueous inclusions: Type I inclusionsoccur in quartz from all ore types and in hydrothermally al-tered host rocks (Table 1). Irregular, negative-crystal or elon-gated shapes are common, and two visible phases, liquid andvapor, are present. These inclusions range in size from 5 to 25µm and have a vapor volume varying from 10 to 85 percent.At least three subtypes of fluid inclusions can be distin-guished from the degree of filling and the geometry of fluidinclusions: (Ia) liquid-rich, with a vapor volume of 5 to 20 per-cent (Fig. 8A), (Ib) liquid-vapor, with a vapor volume of 25 to55 percent, and (Ic) vapor-rich, with a vapor volume of 60 to90 percent (Fig. 8B). Generally, type Ia inclusions occurthroughout both the massive and stringer ores, whereas typesIb and Ic inclusions are mainly restricted to the stringer ores.Laser Raman spectroscopic analyses indicate that liquidphases of the type I inclusions are almost solely H2O, and thevapor phases are dominated by H2O with minor amounts ofCH4 (Fig. 9A, B).

Type II daughter mineral-bearing multiphase inclusions:These inclusions were previously observed by other workersin the discordant stringer zone and in the altered volcanicrocks of the Baiyinchang district (Liu, 1982; Xia et al., 1985).The inclusions consist of one or more solid daughter miner-als, an aqueous phase, and a vapor bubble. They have an ir-regular to negative crystal form and are commonly 4 to 14 µmin diameter. In the stringer zone, most of these inclusions areisolated, and some occur as clusters in the interior of thequartz crystal. In the quartz from the altered host rhyolitictuffaceous rock, these type II inclusions commonly occuralong healed fissures of the quartz phenocrysts. The mor-phology and optical properties of the solids in the inclusionsindicate that halite cubes are the most abundant daughterminerals. Other daughter crystals include a KCl mineral (Liu,1982) and an unidentified opaque mineral, which does notdissolve even at temperatures above 400°C.

Type III CO2-rich inclusions: These inclusions contain oneto three phases at room temperature and can be further di-vided into monophase CO2 (IIIa; Fig. 8C), liquid CO2-vapor

GEOLOGY, FLUID INCLUSIONS, & OXYGEN ISOTOPES OF THE BAIYINCHANG VHMS DEPOSIT, GANSU, CHINA 277

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278 ZENGQIAN ET AL.

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TAB

LE

1. S

umm

ary

of M

icro

ther

mom

etri

c D

ata

for

Prim

ary

Aqu

eous

Flu

id I

nclu

sion

s at

Zhe

yaos

han

Min

e, C

hina

Salin

ity

Sam

ple

Roc

k–O

reH

ost

FI

Vapo

rT

h (°

C )

Fin

al m

eltin

g T

of i

ce (

°C )

(wt %

NaC

l equ

iv)

no.

Loc

atio

nty

pem

iner

alty

pe(v

ol %

)R

ange

A

vera

geR

ange

Ave

rage

Ran

geA

vera

ge

Type

I: T

wo-

phas

e aq

ueou

s in

clus

ions

FI

in th

e st

rata

-bou

nd s

ulfid

e zo

ne

BC

-26

Stra

ta-b

ound

ore

Py-s

p-ga

ore

Qua

rtz

Ia15

175

(1)

–3.9

(1)

26.

16B

C-2

6St

rata

-bou

nd o

rePy

-sp-

ga o

reQ

uart

zIa

1528

3 ~

283

(2)

283

–3.8

(1)

6.30

BC

-26

Stra

ta-b

ound

ore

Py-s

p-ga

ore

Qua

rtz

Ia20

245

~ 25

7 (3

)25

1

FI

in th

e m

assi

ve s

ulfid

e zo

ne

BZ-

19N

o.3;

L10

; 157

3mM

assi

ve o

reQ

uart

zIa

1021

4 ~

225

(5)

218

–5.1

~ –

5.7

(5)

–5.4

8.0

~ 8.

88.

5B

Z-19

No.

3; L

10; 1

573m

Mas

sive

ore

Qua

rtz

Ia10

112

~ 19

6 (1

2)13

8–5

.1 ~

–5.

7 (5

) –5

.48.

0 ~

8.8

8.5

BZ-

19N

o.3;

L10

; 157

3mM

assi

ve o

reQ

uart

zIa

511

5 ~

132

(2)

124

–25.

9 ~

–26.

0 (2

)–2

6.0

23.0

~ 2

3.0

23.0

BZ-

17N

o.1;

L5-

6; 1

573m

Mas

sive

ore

Qua

rtz

Ia5

62 ~

131

(8)

109

–10.

6 ~

–14.

1 (8

)–1

2.7

14.6

~ 1

7.9

16.8

BZ-

02N

o.1;

L3-

4; 1

573m

Mas

sive

ore

Qua

rtz

Ia5

127

(1)

–6.9

(1)

10.4

BZ-

03N

o.1;

L3-

4; 1

573m

Mas

sive

ore

Qua

rtz

Ia10

165

~ 20

8 (9

)18

2–3

.8 ~

–5.

9 (9

)–4

.46.

2 ~

9.1

7.7

FI

in th

e st

ring

er z

one

BZ-

05N

o.1;

L3-

4; 1

573m

Cp-

qtz

vein

Qua

rtz

Ia10

147

~ 17

3 (7

)16

2–1

.6 ~

–3.

2 (7

)–2

.52.

7 ~

5.3

4.2

BZ-

09N

o.1;

L3-

4; 1

573m

Cp-

py v

ein

Qua

rtz

Ia7

136

~ 14

6 (4

)14

0–0

.9 ~

–1.

3 (4

) –1

.11.

6 ~

2.2

2.2

BZ-

09N

o.1;

L3-

4; 1

573m

Cp-

py v

ein

Qua

rtz

Ia7

150

~ 22

0 (6

)17

6–5

.2 ~

–7.

1 (6

)–6

.38.

1 ~

10.6

9.5

BZ-

10N

o.1;

L3-

4; 1

573m

Cp-

py v

ein

Qua

rtz

Ia5

142

~ 22

0 (5

)16

8–3

.3 ~

–3.

6 (5

)–3

.45.

4 ~

5.9

5.6

BC

-3N

o.1,

142

5mC

p-qt

z ve

inQ

uart

zIa

1015

7 ~

241

(5)

193

–4.5

~ –

6.3

(3)

–5.1

7.2

~ 9.

68.

0B

C-3

No.

1, 1

425m

Cp-

qtz

vein

Qua

rtz

Ia10

200

~ 40

5 (8

)27

7–1

2.1

~ –1

6.4

(5)

–14.

816

.1 ~

19.8

18.4

BC

-5N

o.1,

142

5mC

p-qt

z ve

inQ

uart

zIa

511

6 ~

185

(7)

158

–1.8

~ –

2.5

(5)

–2.3

3.1

~ 4.

33.

9B

C-5

No.

1, 1

425m

Cp-

qtz

vein

Qua

rtz

Ia5

311

~ 39

8 (4

)35

4–1

2.1

~ –1

2.1(

3)–1

2.1

16.1

~ 1

6.1

16.1

BC

-7N

o.1,

142

5mC

p-qt

z ve

inQ

uart

zIa

1011

8 ~

258

(7)

173

–1.1

~ –

5.7

(6)

–3.2

1.9

~ 8.

84.

3B

C-9

No.

1, 1

425m

Cp-

qtz

vein

Qua

rtz

Ia15

192

~ 22

0 (5

)20

4–2

.7 ~

–5.2

(2)

–4.0

4.5

~ 8.

16.

3B

C-1

1N

o.1,

142

5mC

p-qt

z ve

inQ

uart

zIa

1091

~ 1

51 (

4)11

1B

C-1

1N

o.1,

142

5mC

p-qt

z ve

inQ

uart

zIa

1222

2 ~

231

(3)

226

–2.3

~–2

.6 (

2)3.

9 ~

4.3

4.1

BC

-11

No.

1, 1

425m

Cp-

qtz

vein

Qua

rtz

Ia15

318

~ 36

2 (4

)33

4B

Z-03

No.

1; L

3-4;

157

3mQ

uart

z ve

inQ

uart

zIb

3032

2 (1

)B

Z-09

No.

1; L

3-4;

157

3mC

p-py

vei

nQ

uart

zIb

3538

3 ~

416

(2)

400

–6.7

~ –

7.0

(2)

–6.9

10.1

~ 1

0.5

10.3

BC

-3N

o.1,

142

5mC

p-qt

z ve

inQ

uart

zIb

5539

0 ~

405

(2)

398

BC

-5N

o.1,

142

5mC

p-qt

z ve

inQ

uart

zIb

3023

7 ~

248

(2)

243

BC

-7N

o.1,

142

5mC

p-qt

z ve

inQ

uart

zIb

3031

7 ~

413

(7)

375

BC

-9N

o.1,

142

5mC

p-qt

z ve

inQ

uart

zIb

3035

9 ~

387

(3)

369

BC

-12

No.

1, 1

425m

Cp-

qtz

vein

Qua

rtz

Ib25

186

~ 23

9 (8

)21

1–3

.2 ~

–3.6

(5)

–3.3

5.3

~ 5.

95.

5B

Z-03

No.

1; L

3-4;

157

3mQ

uart

z ve

inQ

uart

zIc

7537

8 ~

455

(6)

v41

5B

Z-09

No.

1; L

3-4;

157

3mC

p-py

vei

nQ

uart

zIc

6042

0 ~

480

(5)

v45

2–6

.7 ~

–8.

2 (5

)–7

.910

.1 ~

11.

9 11

.5B

C-5

No.

1, 1

425m

Cp-

qtz

vein

Qua

rtz

Ic60

445

~ 48

5 (3

) v

472

–12.

1 (1

)16

.1 ~

16.

116

.1

FI

in th

e al

tere

d ho

st v

olca

nic

rock

s

BC

-34

10–2

0m n

orth

to N

o.1

Chl

zon

e, A

PQ

uart

zIa

1068

~ 1

63 (

16)

125

–12.

7 ~

–16.

6 (1

1)–1

5.1

16.6

~ 2

0.8

18.6

BC

-31

20–3

0m n

orth

to N

o.1

Chl

+ser

zon

e, A

PQ

uart

zIa

1013

0 ~

169

(3)

150

–2.5

~ –

2.5

(3)

–2.5

4.2

~ 4.

24.

2B

C-3

020

–30m

nor

th to

No.

1C

hl+s

er z

one,

AP

Qua

rtz

Ia10

130

~ 16

0 (5

)14

8–2

.9 ~

–3.

3 (5

)–3

.24.

8 ~

5.6

5.2

BC

-32

30–5

0m n

orth

to N

o.1

Ser+

qtz

zone

Qua

rtz

Ia12

182

~ 26

9 (8

)20

5–2

.6 ~

–2.

7 (8

)–2

.74.

3 ~

4.5

4.4

BC

-18

L1–

2; e

ast t

o N

o.1

Ser+

qtz

zone

Qua

rtz

Ia12

148

~ 18

5 (4

)16

6–1

.3 ~

–1.

7 (4

)–1

.62.

2 ~

2.9

2.7

BC

-19

L1–

2; e

ast t

o N

o.1

Ser+

qtz

zone

Qua

rtz

Ia13

134

~ 23

7 (1

1)19

3–1

.6 ~

–1.

7 (6

)–1

.72.

7 ~

2.9

2.8

CO2 (IIIb; Fig. 8D), and liquid-vapor CO2-H2O inclusions(IIIc; Fig. 8E) (Fig. 9C-E; Table 2). These inclusions occurmainly in the stringer zone and altered host rocks and havenot been observed in the massive ore zone. Type III inclu-sions occur in quartz as clusters or isolated inclusions andshow subhedral to euhedral or negative crystal shapes. Theyare generally between 5 and 20 µm in diameter, althoughmuch larger inclusions (up to 50 µm across) are observed in afew chalcopyrite-quartz vein samples. Type IIIa inclusionsconsist of a single CO2-rich fluid at room temperature,whereas type IIIb inclusions have a wide range of vaporCO2/liquid CO2 ratios, varying from 10 to 45 vol percent(Table 2). Type IIIc inclusions contain a carbonic- and anH2O-rich phase at room temperature. Visual estimates of thevolumetric proportions of the carbonic and aqueous phases inthese inclusions vary widely, from 10 to 90 percent (Table 2).Laser Raman spectroscopic analyses of type III inclusions in-dicate that CO2-rich inclusions contain various amounts of hy-drocarbon (i.e., CH4) (Fig. 9C-E).

Type IV CH4 -rich inclusions: Laser Raman spectroscopicanalyses show that type IV inclusions are dominated by CH4,

GEOLOGY, FLUID INCLUSIONS, & OXYGEN ISOTOPES OF THE BAIYINCHANG VHMS DEPOSIT, GANSU, CHINA 279

0361-0128/98/000/000-00 $6.00 279

FI

in th

e qu

artz

-alb

ite p

orph

yrie

sB

C-2

0L

1–2;

eas

t to

No.

1Q

tz–a

lb p

orph

yry

Qua

rtz

Ia10

167

~ 27

0 (1

1)19

7–3

.7 ~

–4.

7 (4

)–4

.56.

0 ~

7.5

7.1

BC

-21

L1–

2; e

ast t

o N

o.1

Qtz

–alb

por

phyr

yQ

uart

zIa

1214

9 ~

174

(3)

164

–2.1

~ –

2.9

(3)

–2.4

3.6

~ 4.

84.

1B

C-2

2L

1–2;

eas

t to

No.

1Q

tz–a

lb p

orph

yry

Qua

rtz

Ia8

140

~ 17

6 (3

)15

9–2

.4 ~

–2.

7 (3

)–2

.24.

0 ~

4.7

4.4

BC

-2L

5-6;

sou

th to

No.

1Q

tz–a

lb p

orph

yry

Qua

rtz

Ia15

218

~ 24

9 (5

)22

9–1

.9 ~

–4.

2 (5

)–3

.03.

3 ~

6.7

5.0

Type

II:

Dau

ghte

r m

iner

al-b

eari

ng m

ultip

hase

incl

usio

nsZD

–71

No.

1, L

1A

ltere

d qu

artz

Q

uart

zII

10 ~

15

400

38ph

yric

rhy

olite

frag

men

tZ–

31N

o.1,

L4

Alte

red

quar

tz

Hyd

ro.

II10

~ 1

530

031

phyr

ic r

hyol

itequ

artz

Z–7

1N

o.1,

L4

Alte

red

quar

tz

Hyd

ro.

II10

~ 1

532

032

phyr

ic r

hyol

itequ

artz

Z4–3

1N

o.1,

L4

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FIG. 8. Photomicrographs of fluid inclusions in quartz from the discordantsulfide zone, the strata-bound Zn-rich lens, and the host-rock unit atZheyaoshan. A. H2O-dominated (type Ia, liquid to vapor ratio L/V<20) and(type Ib, L/V = 25–55) two-phase inclusions in chalcopyrite-quartz veins. B.Vapor-rich (type Ic, L/V= 60–90) two-phase inclusions in chalcopyrite-pyrite-quartz veins. C. Monophase CO2 inclusions in chalcopyrite-pyrite veins. D.CO2-H2O inclusions (CCO2 + LH2O) in chalcopyrite-quartz veins. E. CO2-H2Oinclusions (VCO2 + L CO2 + LH2O) in chalcopyrite-quartz veins. F. CH4-rich in-clusions in chalcopyrite-pyrite veins.

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TABLE 2. Microthermometric Data of Primary CO2 and CH4-rich Fluid Inclusions in Quartz at Zheyaoshan, China

Salinity InitialSample Inclusion Vapor Partial Th Density Hydrate (wt % NaCl Complete melting Tno. Location Ore type Size (µm) sub-type (vol %) (V→L) (g/cm3) Th (°C) equiv) Th (°C) (°C)

Type III: CO2-rich fluid inclusionsType IIIa (monophase CO2 inclusions)BC-5 No.1, 1425m Qtz-cp vein 14 × 7 CO2 15.0 0.81 –58.6BC-5 No.1, 1425m Qtz-cp vein 7 × 7 CO2 11.0 0.848 –58.6BC-5 No.1, 1425m Qtz-cp vein 7 × 7 CO2 11.0 0.848 –58.6BC-5 No.1, 1425m Qtz-cp vein 10 × 8 CO2 19.7 0.769BC-5 No.1, 1425m Qtz-cp vein 6 × 3 CO2 12.8 0.833BC-5 No.1, 1425m Qtz-cp vein 13 × 10 CO2 16.8 0.798 –58.4BC-5 No.1, 1425m Qtz-cp vein 14 × 7 CO2 23.8 0.720 –58.6BC-5 No.1, 1425m Qtz-cp vein 12 × 6 CO2 17.2 0.749BC-5 No.1, 1425m Qtz-cp vein 15 × 7 CO2 19.6 0.770BC-5 No.1, 1425m Qtz-cp vein 7 × 5 CO2 21.5 0.749BC-11 No.1, 1425m Qtz-cp vein 9 × 5 CO2 28.2 0.649 –57.3BC-11 No.1, 1425m Qtz-cp vein 6 × 4 CO2 29.0 0.632 –57.3BC-11 No.1, 1425m Qtz-cp vein 6 × 4 CO2 28.6 0.641 –57.3BC-21 No.1, 1425m Qtz-cp vein 9 × 7 CO2 14.9 0.815 –57.6BZ-05 No. ×1: 1573m Qtz-cp vein 18 × 14 CO2 10.4 0.853 –60.6BZ-05 No. 1: 1573m Qtz-cp vein 20 × 16 CO2 9.9 0.857 –60.1BZ-09 No. 1: 1573m Qtz-cp vein 6 × 3 CO2 15.6 0.808BZ-10 No. 1: 1573m Qtz-cp vein 28 × 7 CO2 15.9 0.805 –59.7BZ-10 No. 1: 1573m Cp-py vein 7 × 5 CO2 11.3 0.846BZ-10 No. 1: 1573m Cp-py vein 5 × 5 CO2 13.0 0.831 –59.7BZ-10 No. 1: 1573m Cp-py vein 7 × 7 CO2 14.4 0.819 –59.7BZ-10 No. 1: 1573m Cp-py vein 7 × 4 CO2 22.0 0.743 –59.7BZ-10 No. 1: 1573m Cp-py vein 18 × 12 CO2 12.3 0.837 –59.8BZ-10 No. 1: 1573m Cp-py vein 12 × 9 CO2 11.0 0.848 –59.9BZ-10 No. 1: 1573m Cp-py vein 14 × 9 CO2 16.1 0.803 –59.9BZ-10 No. 1: 1573m Cp-py vein 5 × 5 CO2 15.8 0.806 –59.9BZ-10 No. 1: 1573m Cp-py vein 7 × 6 CO2 17.6 0.790 –59.9BZ-10 No. 1: 1573m Cp-py vein 7 × 7 CO2 18.6 0.780 –59.9BZ-13 No. 1: 1573m Cp-py vein 10 × 6 CO2 7.8 0.870 –58.7BZ-13 No. 1: 1573m Cp-py vein 18 × 6 CO2 15.3 0.811 –59.2BZ-13 No. 1: 1573m Cp-py vein 54 × 14 CO2 22.2 0.741 –59.1BZ-13 No. 1: 1573m Cp-py vein 72 × 15 CO2 20.6(v) 0.196 –59.1BZ-13 No. 1: 1573m Cp-py vein 10 × 12 CO2 10.9 0.894 –61.6BZ-13 No. 1: 1573m Cp-py vein 8 × 5 CO2 5.3 0.886BZ-13 No. 1: 1573m Cp-py vein 13 × 10 CO2 2.5 0.903BZ-13 No. 1: 1573m Cp-py vein 10 × 5 CO2 3.3 0.898 –58.5BZ-13 No. 1: 1573m Cp-py vein 11 × 5 CO2 21.3(v) 0.202 –58.1

Type IIIb (Two-phase CO2 inclusions)BC-5 No.1, 1425m Qtz-cp vein 14 × 9 CO2-CO2 45 21.8 0.745BC-5 No.1, 1425m Qtz-cp vein 12 × 12 CO2-CO2 15 21.0 0.755 –58.6BC-5 No.1, 1425m Qtz-cp vein 6 × 6 CO2-CO2 10 20.7 0.758BC-5 No.1, 1425m Qtz-cp vein 16 × 7 CO2-CO2 15 26.4 0.681BC-21 No.1, 1425m Qtz-cp vein 22 × 7 CO2-CO2 15 21.1 0.753 –57.6

Type IIIc (CO2-H2O inclusions)BC-5 No.1, 1425m Qtz-cp vein 40 × 50 CO2-H2O CO2=60 27.5 0.663 234 (H2O)BC-5 No.1, 1425m Qtz-cp vein 11 × 11 CO2-H2O CO2=95 20.5 0.761 348 (CO2) –58.4BC-21 No.1, 1425m Qtz-cp vein 13 × 7 CO2-H2O CO2=55 30.0 0.596 6.8 6.12 335 (CO2) –57.6BC-21 No.1, 1425m Qtz-cp vein 22 × 7 CO2-H2O CO2=35 24.2 0.714 6.7 6.29 270 (H2O) –57.6BC-21 No.1, 1425m Qtz-cp vein 15 × 4 CO2-H2O CO2=55 18.1 0.785 266 (CO2) –57.6BC-21 No.1, 1425m Qtz-cp vein 7 × 7 CO2-H2O CO2=75 25.1 0.702 325 (CO2) –57.6BZ-05 No. 1: 1573m Qtz-cp vein 14 × 7 CO2-H2O CO2=15 15.5 0.809 7.6 4.69 228 (H2O)BZ-10 No. 1: 1573m Cp-py vein 20 × 18 CO2-H2O CO2=35 16.8 0.798 8.6 2.81 270 (H2O)BZ-10 No. 1: 1573m Cp-py vein 9 × 7 CO2-H2O CO2=18 7.2 0.875 226 (H2O)BZ-13 No. 1: 1573m Cp-py vein 18 × 5 CO2-H2O CO2=40 13.1 0.830 8.9 2.22 245 (H2O) –59.1BZ-13 No. 1: 1573m Cp-py vein 14 × 6 CO2-H2O CO2=15 9.8 0.857 5.6 8.13 –58.7

Type IV: CH4-rich fluid inclusionsBZ-17 No. 1: 1573m Massive ore 20 × 14 CH4 –106.8 0.32100BZ-17 No. 1: 1573m Massive ore 16 × 16 CH4 –100.0 0.30137BZ-17 No. 1: 1573m Massive ore 14 × 5 CH4 –102.8 0.30982BZ-17 No. 1: 1573m Massive ore 13 × 13 CH4 –100.5 0.30296BZ-17 No. 1: 1573m Massive ore 36 × 9 CH4 –109.2 0.32720BZ-17 No. 1: 1573m Massive ore 18 × 10 CH4 –109.1 0.32695

with minor amounts of other volatile species (Fig. 9F). TheCH4-rich inclusions occur widely in quartz from both the dis-cordant massive ores and the chalcopyrite-pyrite stringerzones (Table 2). They are approximately 5 to 36 µm in diam-eter and have subhedral to euhedral or negative crystalshapes (Fig. 8F). These inclusions are present in isolation oras clusters and commonly coexist with CO2-rich inclusions inindividual quartz grains. The coexistence of CH4-rich inclu-sions with vapor-rich two-phase inclusions (type Ic) has alsobeen observed in chalcopyrite-pyrite-quartz veins.

Microthermometric results

Type I liquid-vapor two-phase aqueous inclusions: Eutectictemperatures were determined for some larger inclusions(about 20-µm diam) and found to vary from –32° to –21°C,indicating that some NaCl, KCl, FeCl2, and CaCl2 may bepresent as dissolved salts (Crawford, 1981; Vanko et al.,2004). Final-melting temperatures of ice range from –26.0° to–1.0°C, corresponding to salinities of 2.6 to 23.0 wt percentNaCl equiv, respectively (Table 1; Potter et al., 1978; Bodnar,1993). Type Ia fluid inclusions in the discordant massive sul-fide zone yield a salinity range of 6.2 to 23.0 wt percent NaClequiv (Fig. 10A) and a homogenization temperature range of62° and 225°C (Fig. 11A).

Variable salinities from 3.1 to 19.8 wt percent NaCl equivwere recorded for type Ia inclusions in the stringer zone (Fig.10B), and these type Ia inclusions also have a wide range ofhomogenization temperatures from 90° to 405°C (Fig. 11B).Type Ib inclusions in the stringer zone have a salinity range of5.5 to 10.5 wt percent NaCl equiv and homogenization tem-peratures of 235° to 415°C (Figs. 10B, 11B). The vapor-richtype Ic inclusions are restricted to the stringer zone and havea moderate salinity (11.9–16.1 wt % NaCl equiv) and thehighest homogenization temperatures, from 378° to 485°C(Figs. 10B, 11B; Table 1).

Type Ia inclusions in the chlorite-altered volcanic host rockyield salinities of 4.2 to 20.8 wt percent NaCl equiv (Fig. 10C)and a homogenization temperature range of 75° to 175°C(Fig. 11C). In comparison, type Ia inclusions in the outerquartz-sericite–altered host volcanics and porphyry stockhave a lower salinity range of 2 to 8 wt percent NaCl equivand a homogenization temperature range of 125° to 250°C(Figs. 10C, 11C).

Type II daughter mineral-bearing multiphase inclusions:These inclusions commonly occur in quartz from the stringerzone and in quartz phenocrysts in the host rocks. Homoge-nization of these inclusions to liquid generally occurs in therange of 280° to 430°C (Liu, 1982). Heating runs on these in-clusions indicate dissolution temperatures between 300° and400°C for the halite cubes and 55° to 100°C for KCl (Liu,1982), based on the dissolution temperatures of the daughtersolids using the H2O-NaCl-KCl system (Roedder, 1971; Hallet al., 1988). The salinity of the inclusions is therefore esti-mated to range from 31 to 38 wt percent NaCl equiv, andsalinities are similar for inclusions in both the stringer zoneand the hydrothermally altered host volcanic rocks (Fig. 10B,C). A similar homogenization temperature range of 350° to400°C was also recorded for type II inclusions in the stringerzone and the hydrothermally altered volcanic rock (Fig. 11B,C).

Type III CO2-rich inclusions: Monophase CO2-rich inclu-sions occur in quartz from the stringer zone and associated al-tered host rocks. Inclusions show complete solid-CO2 melting(Tm, carbonic) between –60.6° and –57.6°C, which is lower thanthe CO2 triple point (–56.6°C). The homogenization temper-ature of CO2 (Th, carbonic) varies from 2.5° to 29.0°C (Fig. 12A),which is lower than the critical temperature of pure CO2

(31°C). Both Tm, carbonic and Th, carbonic indicate the probablepresence of minor amounts of additional gas species, such asCH4 and/or N2 (Burruss, 1981), as confirmed by laser Raman

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BZ-17 No. 1: 1573m Massive ore 11 × 3 CH4 –104.7 0.31528BC-3 No.1, 1425m Qtz-cp vein 15 × 7 CH4 –85.9 0.23683BC-3 No.1, 1425m Qtz-cp vein 8 × 7 CH4 –84.5 0.22004BC-3 No.1, 1425m Qtz-cp vein 8 × 7 CH4 –84.0 0.21404BZ-05 No. 1: 1573m Cp-py vein 18 × 11 CH4 –83.3 0.20570BZ-13 No. 1: 1573m Cp-py vein 29 × 6 CH4 –95.8 0.28710BZ-13 No. 1: 1573m Cp-py vein 18 × 6 CH4 –85.7 0.23443BZ-13 No. 1: 1573m Cp-py vein 20 × 4 CH4 –92.3 0.27314BZ-13 No. 1: 1573m Cp-py vein 25 × 8 CH4 –89.4 0.25925BZ-13 No. 1: 1573m Cp-py vein 18 × 4 CH4 –89.5 0.25979BZ-13 No. 1: 1573m Cp-py vein 29 × 9 CH4 –83.0 0.20205BZ-13 No. 1: 1573m Cp-py vein 14 × 5 CH4 –97.3 0.29245BZ-13 No. 1: 1573m Cp-py vein 17 × 6 CH4 –91.2 0.26816BZ-13 No. 1: 1573m Cp-py vein 18 × 9 CH4 –93.7 0.27889BZ-13 No. 1: 1573m Cp-py vein 13 × 14 CH4 –96.5 0.28852BZ-13 No. 1: 1573m Cp-py vein 14 × 14 CH4 –95.1 0.28451BZ-13 No. 1: 1573m Cp-py vein 29 × 14 CH4 –95.3 0.28524BZ-13 No. 1: 1573m Cp-py vein 54 × 22 CH4 –95.0 0.28415

Abbreviations: Cp = chalcopyrite, py = pyrite, Qtz = quartz, Qtz-cp = quartz-chalcopyrite vein

TABLE 2. (Cont.)

Salinity InitialSample Inclusion Vapor Partial Th Density Hydrate (wt % NaCl Complete melting Tno. Location Ore type Size (µm) sub-type (vol %) (V→L) (g/cm3) Th (°C) equiv) Th (°C) (°C)

spectroscopic analyses (Fig. 9C-E). Two-phase CO2 inclu-sions have melting temperatures (Tm, carbonic) that range be-tween –58.6° and –57.6°C, and partial homogenization tem-peratures (Th, carbonic) from 20.7° to 26.4°C (Table 2; Fig. 12A).The estimated XCH4 in the inclusions is less than 0.07, basedon Thiery et al. (1994), Brown and Hagemann (1995), andFan et al. (2000). H2O-CO2 inclusions have final-meltingtemperatures of solid CO2 from –59.1° to –58.1°C, suggestingthe presence of various amounts of components such as CH4.Further warming of these inclusions within the interval of7.2° to 27.5°C resulted in partial homogenization (Th, carbonic)to liquid (Table 2; Fig. 12A). Clathrate-melting (Tm, clathrate) in

these inclusions occurred at –5.6° to +8.4°C (Table 2), imply-ing a salinity of 2.3 to 8.1 wt percent NaCl equiv, using theequation of Diamond (1992) and the method of Collins(1979). Complete homogenization (Th, total) of type III inclu-sions either to liquid or to vapor occurred between 226° and348°C (Table 2).

Type IV CH4 -rich inclusions: These inclusions nucleated avapor bubble upon cooling below –100°C. No solid CO2 orCO2 clathrate was formed before all phases froze to form asolid at approx. –182.3°C. These inclusions homogenized toliquid, and partially homogenized (Th, carbonic) to liquid be-tween –109.2 and –100.0°C for the massive ores, and between

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FIG. 9. Laser Raman spectra for types I, III and IV inclusions in quartz from the discordant stringer and massive ore zonesat Zheyaoshan. A and B. Liquid and vapor phases of type I inclusions (see Fig. 9B; quartz, BZ-09), respectively. C and D.Liquid and vapor phases of type III inclusions (see Fig. 9D; quartz, BZ-05), respectively. E. Type III monophase CO2 in-clusion (see Fig. 9C; quartz, BZ-05). F. type IV CH4-rich inclusion (see Fig. 9F; quartz, BZ-13).

A B

C D

E F

–97.3 and –83.0°C for the stringer zone (Table 2; Fig. 12B).Compared with the critical point temperature of CH4

(–82.6°C) and the triple point temperature of pure H2S(–85.6°C), the slightly lower Th, carbonic values for the type IVinclusions indicate that the fluid is dominated by CH4 (Bur-russ, 1981; Ramboz et al., 1985; Fan et al., 2000; Jia et al.,2000) but may also contain minor amounts of other gasspecies (e.g., H2S).

Oxygen Isotope DataSamples analyzed for oxygen isotopes were collected from

the discordant massive and stringer zones and the altered fel-sic host rocks, as well. Whole-rock samples were prepared by

pulverizing, and quartz separates were obtained by handpick-ing and separation using heavy liquids. The δ18O analyseswere performed at the Institute of Mineral Resources, Chi-nese Academy of Geological Science (Beijing), using the con-ventional method developed by Clayton and Mayeda (1963).The δ18O data are reported relative to the V-SMOW standard(Table 3) with an analytical precision of ±0.2 per mil (2σ).

δ18O in quartz from sulfide ores

Eleven samples of quartz separates from the Zheyaoshandeposit were analyzed. These samples yielded a limited rangeof δ18O values (8.8–11.1‰; Table 3), which is similar to pre-vious analytical results (9.6–10.4‰; Peng et al., 1995), exceptfor one sample with a higher δ18O value. This range of δ18Ovalues is slightly higher than that of quartz from the stringerzones of some Japanese Kuroko deposits (avg +9.1‰:Pisutha-Arnond and Ohmoto, 1983; Huston, 1999) but much

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0 2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 32 34 36 38 40

22

20

18

16

14

12

10

8

6

4

2

0

Stringer sulfide zone

Type II

Type Ia

0 2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 32 34 36 38 40

22

20

18

16

14

12

10

8

6

4

2

0

Hydrothermally altered host rock

Salinity (wt. % NaCl equiv.)

16

14

12

10

8

6

4

2

00 2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 32 34 36 38 40

Discordant massive sulfide ore

Type Ia

C

B

A

Num

ber

Num

ber

Num

ber

Salinity (wt. % NaCl equiv.)

Salinity (wt. % NaCl equiv.)

Chlorite zone (Volcanics)Qtz+Ser zone (Volcanics)Qtz+Ser zone (Stock)

Altered host volcanicsType II

Type Ia

Type IbType Ic

SW

SW

SW

FIG. 10. Histograms of salinities of fluid inclusions in quartz from A) dis-cordant massive sulfide ores, B) stringer-sulfide zone, and C) altered hostrocks in the Zheyaoshan deposit. SW = seawater. Data is for type II inclusionsfrom Liu (1982) and Xia et al. (1985).

0

2

4

6

8

10

0 50 100 150 200 250 300 350 400 450 500

0 50 100 150 200 250 300 350 400 450 500

0 50 100 150 200 250 300 350 400 450 500

Stringer sulfide zone

Hydrothermally altered host rock

Homogenization Temperature (oC)

Discordant massive sulfide ore

0

2

4

6

8

10

12

14

16

18

20

0

2

4

6

8

10

12

14

16

18

12A

B

C

Num

ber

Num

ber

Num

ber

Homogenization Temperature (oC)

Homogenization Temperature (oC)

Type II

Chlorite zone (Volcanics)Qtz+Ser zone (Volcanics)Qtz+Ser zone (Stock)

Altered host volcanics

Type Ia

Type II

Type Ia

Type Ia

Type IbType Ic

FIG. 11. Histograms of homogenization temperatures of fluid inclusions inquartz from the A) discordant massive sulfide ores, B) stringer-sulfide zone,and C) altered host rocks in the Zheyaoshan deposit. Data for type II inclu-sions from Liu (1982) and Xia et al. (1985).

lower than that of quartz from the Triassic Gacun Kuroko-type deposit (13.7–16.4‰: Hou et al., 2001). Peng et al.(1995) reported a high δ18O value of 14.5 per mil for one sam-ple of quartz from the massive ore at Zheyaoshan; however,not enough information was provided for us to evaluate thishigh δ18O value.

According to oxygen isotope exchange coefficients betweenquartz and water (Matsuhisa et al., 1979), and from homoge-nization temperatures of fluid inclusions (types Ia and Ib) inquartz of 160° to 278°C (Table 1), the hydrothermal fluidsthat precipitated quartz had a range of δ18O values from –0.16to +3.91 per mil for the massive ores and from –5.27 to +3.10per mil for the stringer zone (Table 3). These values suggestan average δ18OH2O value close to 0.0 per mil. However, thefluid inclusions in quartz analyzed for δ18O have a wide Th

range (100° to 413°C: Table 1), and therefore the calculatedδ18OH2O has a wide range for the equilibrium hydrothermalfluid.

δ18O in altered felsic volcanic host rocks

Six whole-rock samples from the altered pipe at Zheyaoshanhave a wide range of δ18O values (1.6 to 10.5‰) (Table 3).Felsic rocks in the marginal zone of the altered pipe have alimited δ18O range from 9.3 to 10.5 per mil, close to that ofthe least-altered volcanic rocks and quartz-albite porphyry(9.4–10.2‰: Xia et al., 1991), suggesting that no significantisotopic exchange of oxygen took place in these rocks duringthe district-scale low-T alteration (<200°C). The strongly al-tered felsic rocks yield a much wider range of δ18O values,from 1.6 per mil in the chlorite core to 8.7 per mil in the outersericite + chlorite zone. Such a lateral variation in δ18O valuesis similar to that described in many VHMS deposits (Hattoriand Muehlenbachs, 1980; Costa et al., 1983; Green et al.,

1983; Urabe et al., 1983; Barrett and MacLean, 1991; Hoy,1993; Huston, 1999). The δ18O values decrease markedly to-ward the altered core, coincident with a feeder zone hydro-thermal fluid having a much lower δ18O value (e.g., ~0‰).Assuming that the analyzed quartz precipitated from a hot(300°–430°C) hydrothermal fluid, as suggested by the fluidinclusion data, the δ18O values of the hydrothermal fluid inequilibrium with quartz having δ18O values from 9.0 to 11.1 permil (Table 3) would have a range of δ18O of 2.0 to 8.0 per mil.

DiscussionA seawater origin for mineralizing fluids is widely accepted

for the formation of most VHMS deposits. For most ancientVHMS deposits, available fluid inclusion data indicate lowsalinities (3.0–10 wt % NaCl equiv), close to that of seawater(3.2 wt % NaCl equiv: Bischoff and Rosenbauer, 1985), and awide temperature range (100°–350°C). However, high-salin-ity fluid inclusions have also been observed in some ancientand modern VHMS deposits (Khin Zaw et al., 1996, 2003;Lecuyer et al., 1999; Hou et al., 2001; Solomon et al., 2004a,b; Vanko et al., 2004), indicating that some fluids were con-siderably more dense than seawater, with the capacity to forma brine pool under the right sea-floor topographic conditions(Pottorf and Barnes, 1983; Solomon and Khin Zaw, 1997;Hou et al., 2001; Solomon et al., 2004a, b). Some authorshave suggested that these high-salinity fluid inclusions recordmagmatic fluid input into the submarine hydrothermal sys-tem (e.g., Urabe et al., 1995; Yang and Scott, 1996, 2002; Houet al., 2001, 2005), whereas others argue that alternativemechanisms (e.g., two-phase separation of seawater) can re-sult in a dense saline fluid coexisting with low-density vapor(Kelley and Delaney, 1987; Nehlig, 1991; Lecuyer et al.,1999; Vanko et al., 2004).

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14

12

10

8

6

4

2

0 5 10 15 20 25 30

5

4

3

2

1

-80 -85 -90 -95 -100 -105 -110

IIIa

IIIb

IIIc

Type III: CO2-rich inclusionsA Type IV: CH4-rich inclusionsB

Massive ore

Stringer ore

ThCH4(°C)

Num

ber

Num

ber

ThCO2(°C)

FIG. 12. Histograms of partial homogenization temperatures of type III CO2-rich (A) and IV CH4-rich (B) inclusions insulfide ores at Zheyaoshan.

Fluid inclusions in quartz from the discordant sulfide zonesand hydrothermally altered host rocks at Zheyaoshan show awide variety of inclusion types, ranging from apparently pureCO2 and CH4 to H2O-CO2-NaCl and H2O-NaCl fluids. Mostfluid inclusions that occur widely in quartz from the stringerzone and from the altered host rocks (quartz-sericite zone)have a low Th (116°–205°C) and low salinity (2.2–5.6 wt %NaCl equiv: Table 1 and Figs. 10, 11). The similarity of thesesalinities to that of normal seawater (3.2 wt % NaCl equiv)suggests a large input of seawater into the Zheyaoshan sub-marine hydrothermal system. This suggestion is supported byoxygen isotope data for the hydrothermally altered host rocksand Cu-bearing quartz veins (i.e., calculated δ18OH2O values~0‰: Table 3). In addition to the low-temperature and low-salinity fluids, three groups of high-salinity fluids have beenrecognized: (1) a high-temperature, high-salinity fluid, (2) ahigh-temperature, vapor-rich fluid of moderate salinity, and(3) a low-temperature brine.

The high-temperature, high-salinity fluids were foundmainly in the chalcopyrite-rich stringer zone and in the hy-drothermal quartz from the host rocks (Table 1; type II daugh-ter mineral-bearing multiphase inclusions and type I two-phase aqueous inclusions). These high-salinity fluids areinterpreted to reflect the contribution of magmatic fluid to thesubmarine hydrothermal system at Zheyaoshan (see below).

High-temperature, vapor-rich, moderate salinity fluidsoccur only in gangue quartz in the stringer zone. This fluid isrepresented by type Ic inclusions, which homogenize to vapor

and mainly coexist with CO2-rich fluid inclusions in thestringer zone, suggesting that the fluid was trapped as a vapor.Similar gas-rich, high-temperature fluid inclusions also havebeen observed in the Gacun VHMS deposit, southwesternChina (Hou et al., 2001). Researchers concluded that thistrapped fluid was magmatic fluid (gas) escaping from the fel-sic magma chamber.

A low-temperature brine, represented by type Ia inclu-sions, is found in the discordant massive sulfide ores of theno. 1 orebody and in strongly altered rock of the chlorite coresurrounding the sulfide orebodies. From the center to themargin of the main no. 1 orebody in the discordant massivesulfide zone, the salinities vary from a range of 14 to 23 to arange of 6 to 10 wt percent NaCl equiv. Similarly, surround-ing the sulfide orebody, the salinity of the brine decreasesmarkedly from 16.6 to 20.8 wt percent NaCl equiv in porous,strongly chloritized volcanic breccia to 2.2 to 7.5 wt percentNaCl equiv in weakly chloritized and sericitized tuffaceousrocks (Table 1). This zonation suggests that mixing between ahigh-salinity (magmatic) fluid and low-salinity seawater oc-curred within and around the discordant massive sulfide pipe.

Low-temperature brine inclusions were trapped coevallywith CH4-rich inclusions in the same massive sulfide samples,suggesting that the brine was a CH4-rich, H2O-NaCl fluid.CH4-rich inclusions are not uncommon in modern and an-cient VHMS deposits (e.g., Simoneit, 1988; Paull et al., 1997;Hou and Zhang, 1998; Hou et al., 2005). The presence ofCH4 is attributed either to a localized reaction between the

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TABLE 3. Oxygen Isotope Composition of Quartz from the Discordant Stringer Zone and the Surrounding Alteration Zone at the Zheyaoshan Mine and Calculated Isotopic Composition of Assumed Equilibrium with Fluid (H2O)

Sample δ18O Temperature δ18OH2O

no. Sample location Rock and ore Mineral SMOW (‰) (avg. °C) (‰) Description

BZ-04 No.1; L3-4; 1573m Cp-qtz vein Quartz 9.3 160 –5.27 Chalcopyrite-bearing quartz veinBZ-05 No.1; L3-4; 1573m Cp-qtz vein Quartz 11.1 160 –3.50 Chalcopyrite-bearing quartz veinBZ-10 No.1; L3-4; 1573m Cp-qtz vein Quartz 9.9 180 –3.21 Chalcopyrite-bearing quartz veinBZ-13 No.1; L3-4; 1573m Cp-py vein Quartz 9.8 180 –3.31 Chalcopyrite-pyrite stringer oreBC-3 No.1, 1425m Cp-qtz vein Quartz 10.9 278 +3.10 Chalcopyrite-bearing quartz vein BC-5 No.1, 1425m Cp-qtz vein Quartz 8.8 254 –0.01 Chalcopyrite-bearing quartz veinBC-6 No.1, 1425m Cp-qtz vein Quartz 10.3 267 +2.06 Chalcopyrite-bearing quartz veinBC-7 No.1, 1425m Cp-qtz vein Quartz 9.4 267 +1.17 Chalcopyrite-bearing quartz vein BC-9 No.1, 1425m Cp-qtz vein Quartz 9.0 254 +0.19 Chalcopyrite-bearing quartz vein BC-10 No.1, 1425m Cp-qtz vein Quartz 9.6 230 –0.43 Chalcopyrite-bearing quartz veinBC-11 No.1, 1425m Cp-qtz vein Quartz 9.0 230 –0.99 Chalcopyrite-bearing quartz vein 253a No.1, L5-6 Cp-qtz vein Quartz 10.3 220 –0.26 Chalcopyrite-bearing quartz veins252a No.1, L5-6 Massive Cu pyrite Quartz 10.4 220 –0.16 Massive chalcopyrite-rich pyrite ore256a No.1, L5-6 Massive Cu pyrite Quartz 9.6 220 –0.95 Massive chalcopyrite-rich pyrite ore257a No.1, L5-6 Massive Cu pyrite Quartz 14.5 220 +3.90 Massive chalcopyrite-rich pyrite ore279a No.1, L5-6 Sericitized Bulk-rock 10.5 160 m north of the main orebody,

tuffaceous breccia sericite-chlorite alteration zone280a No.1, L5-6 Chloritized Bulk-rock 8.7 20 m north of the main orebody;

tuffaceous rock chlorite alteration zone281a No.1, L5-6 Chloritized Bulk-rock 7.9 Northern wall rock of main orebody;

tuffaceous rock chlorite alteration zone282a No.1, L5-6 Chlorite core Bulk-rock 1.6 Strongly altered chlorite core283a No.1, L5-6 Chloritized Bulk-rock 7.8 Southern wall rock of main orebody;

tuffaceous breccia chlorite alteration zone284a No.1, L5-6 Sericitized quartz Bulk-rock 9.3 80 m south of the main orebody;

phyric rhyolite least altered porphyry

Notes: Data source (a) Peng et al. (1995); No. 1 = no. 1 orebody, L5-6 (3-4) = prospecting lines 5-6 (3-4), 1425m = elevation level, Cp = chalcopyrite, qtz= quartz; average temperature is from homogenization temperature of fluid inclusions; δ18OH2O is a calculated value for hydrothermal fluids that are assumedto be in equilibrium with gangue quartz (‰)

hydrothermal fluids and graphitic, carbonaceous wall rocks orserpentinization reactions (Bottrell et al., 1988; Sakai et al.,1990; Shepherd et al., 1991).

Possible sources of the high-salinity fluids

The absence of evaporites in the Baiyinchang district rulesout the possibility that the high-salinity fluids were derivedfrom leaching of evaporite deposits by circulating seawater.Two-phase separation of seawater under certain pressure andtemperature conditions also is considered unlikely becauselow-Th, low-salinity, vapor-rich fluid inclusions coexisting withlow-Th saline inclusions have not been observed atZheyaoshan (Table 1), although both saline, high-tempera-ture (>300°C) fluid and vapor-rich fluid inclusions occur to-gether in gangue quartz from the discordant sulfide zone. Wefavor the input of magmatic fluids to the Zheyaoshan hy-drothermal system to explain the presence of high-salinity in-clusion fluids. This suggestion is supported by the followingthree lines of evidence.

1. High-temperature and high-salinity type II inclusions arepresent not only in gangue quartz from the stringer ores, butalso in quartz phenocrysts and fragmental quartz from thequartz rhyolites that host the mineralization. Liu (1982) re-ported that the Th of such fluid inclusions in both phenocrystand gangue quartz varies from 300° to 400°C, and salinitiesrange between 31 and 38 wt percent NaCl equiv. (Fig. 13A).Xia et al. (1985) also observed similar inclusions in phenocrys-tic quartz from the least-altered rhyolites, and reported similarTh and salinity ranges varying from 300° to 430°C, and 31 to 38wt. percent NaCl equiv, respectively. These observations sug-gest that the felsic magma could have produced high-tempera-ture saline fluids in the late stages of magmatic differentiation.

2. As a group, high-temperature and high-salinity type Iainclusions mostly coexist in the discordant sulfide zone withthe high-temperature, vapor-rich inclusions (types Ib and Ic)and with CO2-rich fluid inclusions of type IIIc. These vapor-rich aqueous inclusions have a wide range of the homoge-nization temperatures (350°–485°C) that reach higher tem-peratures than in many ancient and modern submarinehydrothermal systems (cf. Pisutha-Arnond and Ohmoto,1983; Campbell et al., 1988; Hou et al., 2005). Type Ic inclu-sions (L/V = 60–90) coexist with CO2-rich fluid inclusions inindividual quartz grains and homogenize to vapor at >350°C.A significant feature of these vapor-rich inclusions is that theirlarge decrease in Th is not correlated with a change in salinity(11.9–16.1 wt % NaCl equiv). This suggests that this fluid isprobably a trapped magmatic vapor which was introducedinto the submarine hydrothermal system at Zheyaoshan. TheCO2-rich inclusions also have high homogenization tempera-tures (to vapor; up to 348°C: Table 2). Laser Raman spectradata for CO2-rich fluid inclusions in both phenocryst andgangue quartz reveal mixtures of CO2, H2O, CH4, and H2S.CO2 concentrations vary from 35 to 83 mol percent in thevapor and 17 to 95 mol percent in the liquid, whereas H2Oconcentrations vary from 8 to 22 mol percent in the vapor and3 to 68 mol percent in the liquid (Peng et al., 1995). Yang andScott (1996, 2002) demonstrated that the metal-rich mag-matic fluid component in submarine hydrothermal systemsthey studied is dominated by CO2 and a lesser amount of

H2O. This is consistent with the suggestion that enrichmentof CO2 in the H2O-CO2-NaCl fluid system is characteristic offluids derived from the last-stage felsic magmatic differentia-tion (Sakai et al., 1990; Yang and Scott, 1996, 2002; Hou andZhang, 1998; Hou et al., 2005).

3. Oxygen isotope data for quartz from both the stringerzone and the hydrothermally altered host volcanic rocks indi-cate a contribution of magmatic fluid in the Zheyaoshan sub-marine hydrothermal system. The δ18O values of the hy-drothermal fluid in equilibrium with this quartz range from2.0 to 8.0 per mil, which is intermediate between magmaticwater and seawater.

Fluid mixing

Mixing of at least two fluid end members occurred atZheyaoshan, as indicated by the bimodal distribution of Th

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B

A40

20

35

30

25

15

10

5

050 100 150 200 250 300 350 400

100 200 300 400 500

45

40

35

30

25

20

15

10

5

0

Path 1

Path 2

Mixing Array

Brine

Seawater

Homogenization Temperature (°C)

Discordant Sulfide Zone

Salin

ity (N

aCl w

t.% e

quiv

.)

Homogenization Temperature (°C)

Hydrothermally Altered Host Rocks

Liu (1982)

Chlorite core

Quartz-sericite halo

Altered quartz-albite stockAltered host volcanic rocks

Salin

ity (N

aCl w

t.% e

quiv

.)

Massive sulfide oreStringer sulfide zoneStringer zone and altered host rocks

FIG. 13. Bivariant plots of homogenization temperature versus salinity oftypes I and II inclusions in quartz from A) Hydrothermally altered host rocksand B) The massive ore-dominated sulfide zone at Zheyaoshan. Path 1 is amixing trend between magmatic fluids and cold seawater. Path 2 is a mixingtrend between seawater and vapor-rich, high-temperature fluids derivedfrom degassing of the felsic magma.

and salinity data of the fluid inclusions (Table 1; Figs. 10, 11)and by the systematic variation of salinity versus Th (Fig.13B). Two mixing arrays for these fluid inclusions have beenrecognized (paths 1 and 2). Path 1 represents a typical mix-ing trend between cold seawater and a hydrothermal brinederived from a magma (Hou et al., 2001). The high-temper-ature and high-salinity inclusions in path 1 are mainlytrapped in the underlying discordant stringer zone strati-graphically below the largest massive orebody (no. 1). Thelow-temperature saline fluid inclusions in path 1 mainlyoccur in the massive ore pipe zone, suggesting that mixing ofthe magmatic fluid with seawater continued vertically up-ward from the stringer zone through the discordant sulfidezone toward the sea floor. Low salinities (<10 NaCl wt %equiv) of fluid inclusions in massive ore samples collectedfrom the margins of the no. 1 main massive orebody suggestthe input of a large amount of seawater into the hydrother-mal brine occurred at the margin of the pipe. The chloritecore of the alteration pipe was formed by water-rock interac-tion under high (water/rock) ratios and moderate to low tem-peratures (Fig. 13A), perhaps due to prolonged interactionbetween the magmatic brine and entrained seawater drawndown from above.

Fluid inclusions in all other discordant stringer zone andsemimassive ores lie along path 2 (Fig. 13B). Although thereis a continuous variation in Th for these fluid inclusions, thereare two distinct groups. A high-Th group is characterized bytemperatures >300°C and moderate salinities (10–16 wt %NaCl equiv), whereas the low-Th fluid group is characterizedby temperatures <250°C and low salinities (1.6–10 wt %NaCl equiv). The bimodal distribution of Th values for thesefluid inclusions is recorded within individual quartz grains inthe discordant stringer zone, suggesting upward movement ofmagmatic vapor into the circulating heated seawater. We sug-gest that this modified seawater was heated up to 250°C bythe shallow-level emplacement of felsic magmas.

Fluid immiscibility

Unmixing of C-O-H fluid at Zheyaoshan could explain co-evally trapped CO2- and H2O-dominated fluid inclusions. Im-miscibility is further suggested by certain criteria of unmixedH2O-CO2-NaCl fluids (Ramboz et al., 1982), including theclose spatial association of type I and III fluid inclusions inthree-dimensional arrays with various proportions of carbonicand aqueous phases, variable Th (total), and diverse XCO2 in typeIII inclusions in the stringer zone (Table 2). The wide varia-tion of CO2/H2O ratios in type III inclusions from the stringerzone suggests intermittent vapor-liquid immiscibility duringthe formation of the stringer zone. Complete homogenizationto either a CO2 gas or H2O liquid phase occurs at tempera-tures up to 348°C, suggesting that immiscibility took place inthe early stage of fluid evolution. Higher CO2 contents in theearly-stage fluid suggest that the CO2 was derived from mag-matic degassing, and unmixing would have occurred as a re-sult of changes in pressure, temperature, and salinity (Jia etal., 2000). It is inferred that H2O-rich (type I) and CO2-rich(type III) inclusions represent a spectrum of coeval fluid in-clusions that were trapped simultaneously in different partsof the quartz growing in the unmixed H2O-CO2-NaCl hy-drothermal system.

Unmixing can be caused by phase separation, fluid-carbon-ate reaction (Shepherd et al., 1991), hydraulic fracturing(Kerrich and Allison, 1978; Robert and Brown, 1986; Mc-Cuaig and Kerrich, 1998) and pressure fluctuation (Khin Zawet al., 1994; Jia et al., 2000). At Zheyaoshan, the evidence forimmiscibility is restricted to the stringer zone and hydrother-mally altered rocks. Although immiscibility would have re-sulted in precipitation of sulfides in the stringer zone, there isno evidence for fluid immiscibility in the discordant massiveorebody.

Subsea-floor replacement

Recent studies of VHMS deposits indicate that syngeneticsubsea-floor replacement processes generate a proportion ofmassive sulfide ores in many deposits (e.g., Goodfellow andBlaise, 1988; Khin Zaw and Large, 1992; Large, 1992; Good-fellow and Franklin, 1993; Galley et al., 1995; Marani et al.,1997; Doyle and Huston, 1999). Several criteria proposed todistinguish subsea-floor replacement (Doyle and Allen,2003), include (1) relics of the host rock within the sulfide de-posit, (2) replacement fronts between the sulfide and hostunit, (3) discordance between sulfides and the bedding of en-closing host rocks, and (4) strong hanging-wall alteration.

The Baiyinchang VHMS deposit is hosted by a rhyolitictuffaceous unit and a cryptodome facies associated with auto-clastic breccias. Similar deposits interpreted to have formedby subsea-floor replacement, include the Maurliden deposit,Sweden (Montelius et al., 2000), Highway-Reward (Large,1992; Doyle and Huston, 1999), Mount Morgan deposit(Taube, 1986; Ulrich et al., 2002), and parts of the SulfurSpring deposit, Australia (Morant, 1995). The Baiyinchangdeposit also includes strata-bound Pb-Zn–rich lenses precipi-tated on the paleosea floor, but these lenses are small andeconomically insignificant. The main orebodies are the pipe-style massive sulfide zone formed within a felsic intrusion-vol-caniclastic succession. Although exotic rock fragments are ab-sent within the massive sulfide lenses, relics or relict patchesof the host volcanic rock were preserved and occur within a5- to 20-m-wide zone at the margin of the massive lenses.There is continuous variation in the size and proportion ofthese relics that range from weakly mineralized host rock tosemimassive sulfide. These relics have been strongly chlori-tized but are similar in character to the host rocks, suggestingincomplete replacement of a coherent host facies (Doyle andHuston, 1999; Doyle and Allen, 2003).

At Zheyaoshan, the position of the replacement fronts, asindicated by the presence of disseminated, spotty, semimas-sive sulfides (cf. Khin Zaw and Large, 1992), were developedaround massive sulfide orebodies. These fronts are commonlyirregular in shape, gradational, and are characterized by aprogressive increase in sulfide and chlorite contents towardthe massive orebody. In plan, replacement fronts arebounded by a disseminated ore halo that surrounds massivesulfide and contains discrete relict patches of the host rock. Incross section, replacement fronts are restricted to the chlori-tized zone and marked by steeply dipping, disseminatedzones on both sides of the pipe-shaped orebody (Fig. 6).Lithofacies mapping by Cheng (1980) showed the discor-dance (up to 70°) between the massive orebodies and thebedding in the quartz rhyolitic tuffaceous rock at Zheyaoshan

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and demonstrated that the contacts are not faulted or relatedto interfingering between the volcanic facies and the massivesulfide orebodies.

Figure 14 shows schematically the spatial variations of thetemperatures inferred for the hydrothermal fluids from thepreviously published fluid inclusion data of Peng et al. (1995)and Table 1. The high-temperature (>300°C) zone is mainlylocated in the stringer sulfides but extends upward along thesulfide pipe, whereas the low-temperature (<200°C) zone isconfined to the massive sulfides within the discordant sulfidezone hosted in the volcanic facies (Figs. 6, 14). The low-tem-perature center generally corresponds to massive Zn- and Cu-rich orebodies, whereas the high-temperature column(~300°C) corresponds to pipelike chalcopyrite-pyrrhotite-pyrite ores (Fig. 14; Peng et al., 1995). The replacementfronts marked by disseminated ore lenses formed at temper-atures of 150°–200°C, higher than that of the massive sulfideores (60°–130°C) (Fig. 14). Due to mixing and cooling of thedense brine with down drawn seawater in the subsurface,most of the base metals are interpreted to have deposited be-neath the sea floor, with very little stratiform ore developedon the sea floor.

The chalcopyrite-rich pyrrhotite-pyrite pipe, which ac-counts for 10 vol percent of the discordant massive sulfide ore-bodies, consists of sulfide precipitates localized in dischargeconduits along which hot hydrothermal fluids moved upwardand eventually completely replaced the quartz rhyolitic hostrock. The pipe extends upward into the overlying discordantmassive orebody and merges downward with the pyrite-chal-copyrite zone, suggesting that sulfide accumulation took placeby open-space filling of the fractured volcaniclastic rocks inthe subsurface and was fed by focused, discharging fluids. Asthe salinity of the fluid that deposited the massive ores is muchhigher than that of seawater, a large amount of such densesaline fluid could have settled in the subsea floor, filling thepore spaces and fractures in the volcanic rocks.

Comparison with modern analogs

The setting of the Baiyinchang deposit in a bimodal vol-canic sequence related to rifting of the Proterozoic basementin the Cambrian volcanic belt (III2) of the North Qilian oro-gen is similar to modern back-arc settings. The Manus basin,bordered by New Britain, New Ireland, and the Manus Is-lands, is a back-arc basin with respect to the New Britain sub-duction zone and the active New Britain volcanic arc (Rona,1984; Scott and Binns, 1995; Sinton et al., 2003).

The host succession of the Baiyinchang deposit atZheyaoshan is a volcaniclastic-dominated felsic complex inwhich porphyritic dome and tuffaceous rocks of quartz-phyricrhyolitic composition host the massive sulfide orebodies. Thefine-grained volcaniclastic rocks at Baiyinchang are thoughtto have been produced by explosive volcanism, which is dis-tinctly different from the felsic breccias that host the PAC-MANUS and known deposits in the Manus basin (Paulick etal., 2004). This difference may simply be related to hydrosta-tic pressure (water depth) and the magnitude and rate ofmagma vesiculation (McBirney, 1993). Cu + Zn-rich, fluid-filled cavities within glass melt inclusions in the phenocrystsof Pual Ridge andesites (Yang and Scott, 1996), as well as thegas compositions and high fluorine contents of vent fluids,provide direct evidence for input of magmatic fluids, similarto that suggested by fluid inclusion evidence in the Baiyin-chang deposit.

Ongoing replacement processes at PACMANUS are similarto processes inferred for the Baiyanchang massive sulfides.Pyrite (1–5%) is present as disseminated grains and veins inthe altered volcanic rocks and as centimeter-scale pods andveins of massive pyrite in a stockwork zone underlying areasof active high-temperature black smoker chimneys (Binns etal., 2002). Based on these observations and extensive loggingof the holes, Barriga et al. (2002) inferred that the originalrhyodacite is being altered by hydrothermal fluids and re-placed by sulfide-anhydrite-silica. In general, the high per-meabilities related to discrete networks of fractures in pillowbreccia zones measured in volcanic rocks of oceanic crust areconsidered favorable for subsea-floor replacement processesin ore formation, including pipe-style sulfide bodies such asthose at Baiyinchang.

A model for the Baiyinchang deposit

The Baiyinchang deposit formed within a submarine, syn-volcanic cryptodome volcanic center, similar to that hosting

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200

Ser+Qtz

Chl

Ser+Qtz

0 100 200 m

Chl

Elevation(m)

1775

1715

1655

1535

1475

1415

Oxidized zone

150

200

20010

015

0

100

200

250

250

350

400

200

300

300

250

Isotherm (°C)Rhyolitic volcanic rock

Massive and disseminated ores Rhyolitic breccia

Stockwork or stringer ore Hematitic chert

Ser+Qtz Sericite-quartz zone Fe-Mn chert

Chlorite zone Quartz-albite stockChl

Rhyo

litic

tuffa

ceou

sro

ck

FIG. 14. Schematic cross section of the Baiyinchang hydrothermal systemshowing spatial variation in the temperature of the submarine hydrothermalfluids estimated from the fluid inclusion data previously reported and deter-mined in this study. End-member temperatures of hydrothermal fluids re-sponsible for precipitation of the stringer sulfide zone are estimated from thehigh-Th fluid inclusions (Table 1), whereas temperatures for massive sulfideore are estimated from the average Th of the fluid inclusions (Peng et al.,1995; Table 1). Temperature data for hydrothermally altered rocks are fromPeng et al. (1995) and this study. The isotherms are highly schematic.

the Highway-Reward VHMS deposit in Mount Windsor sub-province (Doyle and Huston, 1999) and that hosting someVHMS deposits in the Skellefte district (Allen et al., 1996a,b). Our analyses of fluid inclusions and oxygen isotopes con-tribute to the genetic model for the formation of the Baiyin-chang deposit (Zheyaoshan mine). Three principal factorswhich controlled the location and the formation of the de-posit are as follows: (1) a high geothermal gradient developedin the submarine volcanic center, (2) intrusion of a high-levelcryptodome, accompanied by expulsion of high-salinity mag-matic fluids along fractures toward the sea floor, and (3) mix-ing of the magmatic fluids and seawater in fractured rhyoliticvolcanic rocks below the sea floor (Fig. 15).

Peng et al. (1995) concluded that the partial extrusion of afelsic cryptodome within the felsic volcanic tuffaceous unitand the shallow-level emplacement of felsic porphyry bodiesalong ring-shaped faults were an indication of a felsic magmachamber located as shallow as 1 to 1.5 km below the paleosea-floor. The magma chamber may have served as a heat engine,driving the hydrothermal convection of seawater, causingleaching of metals from the volcanic pile and their contribu-tion to the ore-forming fluid (Fig. 15A). Degassing of the fel-sic magma released a CO2-dominated metal-bearing mag-matic fluid, as evidenced by high-temperature (up to 485°C),high-salinity (31–38 wt % NaCl equiv), and CO2-rich fluid in-clusions in the chalcopyrite-rich stringer veins and quartzphenocrysts from the altered host rocks.

The dense nature of the saline hydrothermal fluids mayhave allowed them to accumulate in the pore spaces and frac-tures of the rhyolitic breccia, resulting in a “brine zone” in thesubsea floor at Zheyaoshan. The morphology of the brinezone remains unconstrained, but it was most likely a narrow,steeply dipping funnel, possibly covered by overlying pyriticshale and limestone and fed by hydrothermal fluids that werefocused through a synvolcanic fault during hydraulic fractur-ing or pressure fluctuations. Within the brine-saturated fun-nel, precipitation of sulfides may have been caused by influxof cold seawater into the hot brine (Fig. 14B). Massive sulfideores formed in the zone of maximum fluid mixing within thebrine funnel. In general, chalcopyrite-rich massive pyrite oresmainly accumulated in the hottest and deepest part of thefunnel, with sphalerite-rich ores above, due to vertical de-crease in the temperature of the brine. Within the feederzone the mixing of hot, Cu-bearing fluids with convectingheated seawater promoted the precipitation of semimassivechalcopyrite-pyrite ore and the formation of chalcopyrite-richstringer veins (Fig. 15B). Away from the feeder zone, lateralinfiltration of hydrothermal fluids and replacement of hostvolcanic facies resulted in the formation of a halo of dissemi-nated ore. Locally, the mixed hydrothermal fluids at the topof the funnel were expelled into the water column andformed vents on the sea floor, which resulted in the formationof the dispersed strata-bound Zn-rich lenses and hematiticcherts (Fig. 15B).

AcknowledgmentsThe authors thank staff members of the Institute of Min-

eral Resources, CAGS, for their assistance in laboratory workand constructive discussions. We also thank all past and pre-sent mine geologists from the Baiyin mine and the Baiyin

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FIG. 15. A possible genetic model for the formation of the Baiyinchang hy-drothermal system, China. A. Emplacement of a felsic magma chamber actedas a heat engine, driving the hydrothermal convection of seawater, and thefelsic magma expelled saline hydrothermal fluids resulting in a “brine zone”in the subsea-floor environment. B. Discordant massive sulfide ores formedwithin the brine-saturated funnel due to mixing of magmatic brine and influxof cold seawater. Lateral infiltration of hydrothermal fluids and replacementof host volcanic facies resulted in the formation of a halo of disseminated ore.Locally, the mixed hydrothermal fluids at the top of the funnel were expelledinto the water column to form vents on the sea floor, which resulted in theformation of the dispersed, strata-bound Zn-rich lenses and exhalites.

Metal Company for assisting us at the mine sites. This studywas supported by the National Science Foundation of China(49773177) and the key project (95-Pre-39) of the Ministry ofScience and Technology, China. The authors thank journal re-viewers David Vanko, Bruce Taylor, and an anonymous re-viewer for their valuable and thoughtful comments. Co-guest-editor Jan Peter is thanked for his incisive review andsuggestions which have substantially improved the paper.Special thanks are due to Ross Large for reading the manu-script and making constructive comments. The authors alsowould like to thank Mark Hannington for his careful andthorough review of the final version of the manuscript andhelpful comments.

REFERENCESAllen, R.L., Weihed, P., and Svenson, S.A., 1996a, Setting of Zn-Cu-Au-Ag

massive sulfide deposits in the evolution and facies architecture of a 1.9 Gamarine volcanic arc, Skellefte, Sweden: ECONOMIC GEOLOGY, v. 91, p.1022–1053.

Allen, R.L., Lundstrom, I., Ripa, M., Simeonov, A., and Christofferson, H.,1996b, Facies analysis of a 1.9 Ga, continental margin, back-arc, felsiccaldera province with diverse Zn-Pb-Ag- (Cu-Au) sulfide and Fe oxide de-posits, Bergslagen region, Sweden: ECONOMIC GEOLOGY, v. 91, p.979–1008.

Barrett, T.J., and MacLean, W.H., 1991, Chemical, mass and oxygen isotopechanges during extreme hydrothermal alteration of an Archean rhyolite,Noranda: ECONOMIC GEOLOGY, v. 86, p. 406–414.

Barriga, F.J.A.S., Dias, A., Marques, A.F.A., Costa, I.A., Relvas, J., Pinto, A.,Fouquet, Y., and Costa, I.R., 2002, Replacement processes in volcanogenicmassive sulfide deposits: The key to giant orebodies [abs.]: Geological So-ciety of America Abstracts, v. 34, no. 6, p. 444.

Bian, Q-T., 1989, Structure of the Baiyin ore field and genetic model of theBaiyinchang deposit: Beijing, Earthquake Press, 152 p. (in Chinese).

Binns, R.A., Barriga, F.J.A.S., Miller, D.J., et al., 2002, Proceedings of theOcean Drilling Program [ODP] Initial Reports: Ocean Drilling Program,College Station, Texas, Texas A&M University, 193 p. (CD-ROM)

Bischoff, J.L., and Rosenbauer, R.J., 1985, An empirical equation of state forhydrothermal seawater (3.2 percent NaCl): American Journal of Science, v.285, p. 725–763.

Bodnar, R.J., 1993, Revised equation and table for determining the freezingpoint depression of H2O-NaCl solutions: Geochimica et CosmochimicaActa, v. 57, p. 683–684.

Bottrell, S.H., Shepherd, T.J., Yardley, B.W.D., and Dubessy, J., 1988, A fluidinclusion model for the genesis of the ore of the Dolgellau gold belt, NorthWales: Journal of the Geological Society [London], v. 145, p. 139–145.

Brown, P.E., and Hagemann, S.G., 1995, MacFlincor and its application tofluids in Archean lode-gold deposits: Geochimica et Cosmochimica Acta, v.59, p. 3943–3952.

Burruss, R.C., 1981, Analysis of phase equilibria in C-O-H-S fluid inclusions:Mineralogical Association of Canada Short Course Handbook, v. 6, p.39–74.

Campbell, A.C., Bower, T.S., Measures, C.I., Falkner, K.K., Khadem, M. andEdmond, J.M., 1988, A time series of vent fluid compositions from 21°NEast Pacific Rise (1979, 1981, 1985), and the Guaymas basin, Gulf of Cali-fornia (1982, 1985): Journal of Geophysical Research, v. 93, p. 4537–4549.

Cheng, G., 1980, Some geological characteristics of the Baiyinchang pyrite-type deposit: Geology and Exploration, no. 9, p. 12–20 (Chinese with Eng-lish abstract).

Clayton, R.N., and Mayeda, T.K., 1963, The use of bromine pentafluoride inthe extraction of oxygen from oxides and silicates for isotopic analysis:Geochimica et Cosmochimica Acta, v. 27, p. 43–52.

Collins, P.L.F., 1979, Gas hydrates in CO2-bearing fluid inclusions and use offreezing data for estimation of salinity: ECONOMIC GEOLOGY, v. 74, p.1435–1444.

Costa, U.R., Barnett, R.L., and Kerrich, R., 1983, The Mattagami Lake mineArchean Zn-Cu sulfide deposit, Quebec: Hydrothermal precipitation oftalc and sulfides in a sea-floor brine pool: Evidence from geochemistry,18O/16O, and mineral chemistry: ECONOMIC GEOLOGY, v. 78, p. 1144–1203.

Cox, S.F., 1981, The stratigraphic and structural setting of the Mt. Lyell vol-canic-hosted sulfide deposits: ECONOMIC GEOLOGY, v. 76, p. 231–245.

Crawford, M.L., 1981, Phase equilibria in aqueous fluid inclusions: Miner-alogical Association of Canada Short Course Handbook 6, p. 75–100.

Diamond, L.W., 1992, Stability of CO2-clathrate-hydrate + CO2-liquid +CO2-vapor + aqueous KCl-NaCl solutions: Experimental determinationand application to salinity estimates of fluid inclusions: Geochimica et Cos-mochimica Acta, v. 54, p. 545–552.

Doyle, M.G., and Huston, D.L., 1999, The subsea-floor replacement originof the Ordovician Highway-Reward volcanic-associated massive sulfide de-posit, Mount Windsor subprovince, Australia: ECONOMIC GEOLOGY, v. 94,p. 825–844.

Doyle, M.G., and Allen, R.L., 2003, Subsea-floor replacement in volcanic-hosted massive sulfide deposits: Ore Geology Reviews, v. 23, p. 183–222.

Eastoe, C.J., Solomon, M., and Walshe, J.L., 1987, District-scale alterationassociated with massive sulfide deposits in the Mount Read Volcanics, west-ern Tasmania: ECONOMIC GEOLOGY, v. 82, p. 1239–1258.

Fan, H-R., Groves, D.I., Mikucki, E.J., and McNaughton, N.J., 2000, Con-trasting fluid inclusion types at the Nevoria gold deposit in the SouthernCross greenstone belt, Western Australia: Implications of auriferous fluidsdepositing ores within an Archean banded iron-formation: ECONOMIC GE-OLOGY, v. 95, p. 1527–1536.

Franklin, J.M., Sangster, D.F., and Lydon, J.W., 1981, Volcanic-associatedmassive sulfide deposits: ECONOMIC GEOLOGY 75TH ANNIVERSARY VOLUME,p. 485–627.

Galley, A.G., Watkinson, D.H., Jonasson, I.R., and Riverin, G., 1995, Thesubsea-floor formation of volcanic-hosted massive sulfide: Evidence fromthe Ansil deposit, Rouyn-Noranda: ECONOMIC GEOLOGY, v. 90, p.2006–2017.

Gifkins, C., Herrmann, W., and Large, R.R., 2005, Altered volcanic rocks: Aguide to description and interpretation: Centre for Ore Deposit Studies[CODES], Hobart, University of Tasmania, 275 p.

Goodfellow, W.D., and Blaise, B., 1988, Sulfide formation and hydrothermalalteration of hemipelagic sediments in Middle Valley, Northern Juan deFuca Ridge: Canadian Mineralogist, v. 26, p. 675–696.

Goodfellow, W.D., and Franklin, J.M., 1993, Geology, mineralogy, and chem-istry of sediment-hosted clastic massive sulfides in shallow cores, MiddleValley, Northern Juan de Fuca Ridge: ECONOMIC GEOLOGY, v. 88, p.2037–2068.

Green, G.R., Ohmoto, H., Date, J., and Takahashi, T., 1983, Whole-rock oxy-gen isotope distribution in the Fukazawa-Kosaka area, Hokuroko district,Japan, and its potential application to mineral exploration: ECONOMIC GE-OLOGY MONOGRAPH 5, p. 395–411.

Hall, D.L., Sterner, S.M., and Bodnar, R.J., 1988, Freezing point depressionof NaCl-KCl-H2O solutions: ECONOMIC GEOLOGY, v. 83, p. 197–202.

Hattori, K., and Muehlenbachs, K., 1980, Marine hydrothermal alteration ata Kuroko ore deposit, Kosaka, Japan: Contributions to Mineralogy andPetrology, v. 74, p. 285–292.

Hou, Z.-Q., and Zhang, Q.-L., 1998, CO2 and hydrocarbon-rich fluids in sub-marine active hydrothermal system in the Okinawa trough: Evidence fromfluid inclusions: Sciences in China, v. 18, p. 142–148.

Hou, Z.-Q., Deng, J., Sun, H.-T., and Song, S.-H., 1999, Volcanogenic mas-sive sulfide deposits in China: Setting, feature, and style: Exploration andMining Geology, v. 8, p. 149–175.

Hou, Z.-Q., Khin Zaw, Qu, X.-M., Ye, Q.-T., Yu, J.-J., Xu, M.-J., Fu, D.-M.,and Yin, X.-K., 2001, Origin of the Gacun volcanic-hosted massive sulfidedeposit in Sichuan, China: Fluid inclusion and oxygen isotope evidence:ECONOMIC GEOLOGY, v. 96, p. 1491–1512.

Hou, Z.-Q., Khin Zaw, Li, Y.-H., Zhang, Q.-L., Zeng, Z.-G., and Urabe, T.,2005, Contribution of magmatic fluid to the active hydrothermal system inthe JADE field, Okinawa trough: Evidence from fluid inclusions, and oxy-gen and helium isotopes: International Geology Reviews, v. 47, p. 420-437.

Hoy, L.D., 1993, Regional evolution of hydrothermal fluids in the Norandadistrict, Quebec: Evidence from δ18O values from volcanogenic massivesulfide deposits: ECONOMIC GEOLOGY, v. 88, p. 1526–1541.

Huston, D.L., 1999, Stable isotopes and their significance for understandingthe genesis of volcanic-hosted massive sulfide deposits: A review: Reviewsin Economic Geology, v. 8, p. 157–181.

Jia, Y., Li, Y., and Kerrich, R., 2000, A fluid inclusion study of Au-bearingquartz vein systems in the Central and North Deborah deposits of theBendigo gold field, central Victoria, Australia: ECONOMIC GEOLOGY, v. 95,p. 467–494.

Kalogeropoulos, S.I., and Scott, S.D., 1983, Mineralogy and geochemistry oftuffaceous exhalite (Tetsusekiei) of the Fukazawa mine, Hokuroko district,Japan: ECONOMIC GEOLOGY MONOGRAPH 5, p. 412–432.

290 ZENGQIAN ET AL.

0361-0128/98/000/000-00 $6.00 290

Kelley, D.S., and Delaney, J.R., 1987, Two-phase separation and fracturing inmid-ocean ridge gabbros at temperatures greater than 700°C: Earth andPlanetary Science Letters, v. 83, p. 53–66.

Kerrich, R., and Allison, I., 1978, Vein geometry and hydrostatic pressureduring Yellowknife mineralization: Canadian Journal of Earth Sciences, v.15, p. 1653–1660.

Khin Zaw, 2003, Geology and sulphur isotope implications of Bawdwin de-posit, Northern Shan State, Myanmar: An Ag-rich, volcanic-hosted, poly-metallic massive sulphide deposit, in Eliopoulos, D.G., eds., Mineral ex-ploration and sustainable development: Rotterdam, Millpress, 7th Societyfor Geology Applied to Mineral Deposits Conference, Athens, Greece, Au-gust 24–28, 2003, p. 217–220.

——2004, Geological setting and formation of Bawdwin deposit, NorthernShan State, Myanmar: A silver-rich, volcanic-hosted polymetallic massivesulphide deposit [abs.]: 17th Australian Geological Convention, Hobart,Australia, February, Geological Society of Australia Abstracts, no. 73, p. 94.

Khin Zaw, and Large, R.R., 1992, The precious metal-rich South Herculesmineralization, western Tasmania: A possible subsea-floor replacement vol-canic-hosted massive sulfide deposit: ECONOMIC GEOLOGY, v. 87, p.931–952.

Khin Zaw, Huston, D.L., Large, R.R., Mernagh, T., and Hoffmann, C.F.,1994, Microthermometry and geochemistry of fluid inclusions from theTennant Creek gold-copper deposits: Implications for ore deposition andexploration: Mineralium Deposita, v. 29, p. 288–300.

Khin Zaw, Gemmell, J.B., Large, R.R., Mernagh, T.P., and Ryan, C.G., 1996,Evolution and source of ore fluids in the footwall stringer zone, HellyerVHMS deposit, western Tasmania: Evidence from fluid inclusion mi-crothermometry and geochemistry: Ore Geology Reviews, v. 10, p.251–278.

Khin Zaw, Hunns, S.R., Large, R.R., Gemmell, B.J., Ryan, C.G., and Mer-nagh, T.P., 2003, Microthermometry and chemical composition of fluid in-clusions from the Mt Chalmers volocanic-hosted massive sulphide deposits,central Queensland, Australia: Implications for ore genesis: Chemical Ge-ology, v. 194, p. 225–244.

Large, R.R., 1992, Australian volcanic-hosted massive sulfide deposits: Fea-tures, styles and genetic models: ECONOMIC GEOLOGY, v. 87, p. 469–470.

Large, R.R., McPhie, J., Gemmell, J.B., Herrmann, W., and Davidson, G.J.,2001, The spectrum of ore deposit types, volcanic environments, alterationhalos, and related exploration vectors in submarine volcanic successions—some examples from Australia: ECONOMIC GEOLOGY, v. 96, p. 913–938.

Lecuyer, C., Dubois, M., Marignac, C., Gruau, G., Fouquet, Y., Ramboz, C.,1999, Phase separation and fluid mixing in subseafloor back arc hydrother-mal systems: A microthermometric and oxygen isotope study of fluid inclu-sions in barite-sulfide chimneys of the Lau basin: Journal of GeophysicalResearch, v. 104, p. B17911-B17927.

Leistel, J.J., Marcoux, E., and Deschamps, Y., 1998, Chert in the Iberianpyrite belt: Mineralium Deposita, v. 33, p. 59–81.

Liu, B., 1982, Studies of solid and fluid inclusions on the copper deposits inBaiyinchang: Acta Geologica Sinica, v. 56, p. 165–173 (Chinese with Eng-lish abstract).

Marani, M.P., Gamberi, F., and Savelli, C., 1997, Shallow-water polymetallicsulfide deposits in the Aeolian island arc: Geology, v. 25, p. 815–818.

Matsuhisa, Y., Goldsmith, J.R., and Clyton, R.N., 1979, Oxygen isotopic frac-tionation in the system quartz-albite-anorthite-water: Geochimica et Cos-mochimica Acta, v. 43, p. 1131–1140.

McBirney, A.R., 1993, Igneous petrology, 2nd ed.: Boston, Jones and Bartlett,508 p.

McCuaig, T.C., and Kerrich, R., 1998, P-T-deformation-fluid characteristicsof lode gold deposits: Evidence from alteration systematics: Ore GeologyReviews, v. 12, p. 381–453.

Montelius, C.B., Allen, R.L., Svenson, S.A., and Weihed, P., 2000, The Mau-rliden volcanogenic Zn-Cu-Ag massive and network sulfide deposits,Skellefte district, Sweden [abs.]: Volcanic Environments and Massive Sul-fide Deposits Conference, Hobart, Australia, CODES, Abstract Volume, p.135–136.

Morant, P., 1995, The Panorama Zn-Cu VMS deposits, Western Australia:Australian Institute of Geoscientists Bulletin, v. 16, p. 75–84.

Nehlig, P., 1991, Salinity of oceanic hydrothermal fluids: A fluid inclusionstudy: Earth and Planetary Science Letters, v. 102, p. 10–325.

Paulick, H., Vanko, D.A., and Yeats, C.J., 2004, Drill core-based facies re-construction of a deep-marine, felsic volcano hosting an active hydrother-mal system (Pual Ridge, Papua, New Guinea, ODP Leg 193): Journal ofVolcanological and Geothermal Research, v. 130, p. 31–50.

Paull, C., Dickens, G.R., and Borowski, W., 1997, ODP drill gas hydrate, theworld’s largest source of fossil fuel: ODP’s Greatest Hits, p. 22–23.

Peng, L-G., Ren, Y-X., Li, P-Z., and Liu, X-M., 1995, A metallogenic modelfor copper-polymetallic deposits of Baiyinchang, Gansu Province: Beijing,Geological Publishing House, 211 p. (Chinese with English abstract).

Peter, J.M., 2003, Ancient iron-rich metalliferous sediments (iron forma-tions): Their genesis and use in the exploration for stratiform base metalsulphide deposits, with examples from the Bathurst Mining Camp, inLentz, D.R., ed., Geochemistry of sediments and sedimentary rocks: Secu-lar evolutionary considerations to mineral deposit-forming environments:Geological Association of Canada, Geotext v. 4, p. 145–173.

Pisutha-Arnond, V., and Ohmoto, H., 1983, Thermal history, and chemicaland isotopic compositions of the ore-forming fluids responsible for theKuroko massive sulfide deposits in the Hokuroko district of Japan: ECO-NOMIC GEOLOGY MONOGRAPH 5, p. 523–558.

Potter, R.W., II., Clynne, M.A., and Brown, D.L., 1978, Freezing point de-pression of aqueous sodium chloride solutions: ECONOMIC GEOLOGY, v. 73,p. 284–285.

Pottorf, R.J., and Barnes, H.L., 1983, Mineralogy, geochemistry and ore gen-esis of hydrothermal sediments from the Atlantis II Deep, Red Sea: ECO-NOMIC GEOLOGY MONOGRAPH 5, p. 198–223.

Ramboz, C., Pichavant, M., and Weisbrod, A., 1982, Fluid immiscibility innatural processes. Use and misuse of fluid inclusion data. II: Interpretationof fluid inclusion data in terms of fluid immiscibility: Chemical Geology, v.37, p. 29–48.

Ramboz, C., Schnapper, D., and Dubessy, J., 1985, The P-V-T-X-fO2 evolu-tion of H2O-CH4-CO2-bearing fluid in a wolframite vein: Reconstructionfrom fluid inclusion studies: Geochimica et Cosmochimica Acta, v. 49, p.205–219.

Robert, F., and Brown, A.C., 1986, Archean gold-bearing quartz veins at theSigma mine, Abitibi greenstone belt, Quebec: Part I. Genetic relations andformation of the vein system: ECONOMIC GEOLOGY, v. 81, p. 578–592.

Roedder, E., 1971, Fluid inclusion studies on porphyry type ore deposit atBingham, Utah, Butte, Montana and Climax, Colorado: ECONOMIC GEOL-OGY, v. 66, p. 110–114.

Roedder, E., 1984, Fluid inclusions: Reviews in Mineralogy, v. 12, 644 p.Rona, P.A., 1984, Hydrothermal mineralization at seafloor spreading centers:

Earth Science Reviews, v. 20, p. 1–104. Sakai, H., Gamo, T., Kim, E.S., Tsutsumi, M., Tanka, T., Ishibashi, J., Wakita,

H., Yamano, M., and Oomori, T., 1990, Venting of carbon dioxide-rich fluidand hydrate formation in mid-Okinawa trough backarc basin: Science, v.248, p. 1093–1096.

Scott, S.D., and Binns, R.A., 1995, Hydrothermal processes and contrastingstyles of mineralization in the western Woodlark and eastern Manus basinsof the western Pacific, in Parson, L.M., Walker, C.L., and Dixon, D.R.,eds., Hydrothermal vents and processes: Special Publication of the Geo-logical Society [London], v. 87, p. 191–205.

Shepherd, T.J., Bottrell, S.H., and Miller, M.F., 1991, Fluid inclusionvolatiles as an exploration guide to black shale-hosted gold deposits, Dol-gellau gold belt, North Wales, U.K: Journal of Geochemical Exploration, v.42, p. 5–24.

Simoneit, B.R.T., 1988, Petroleum generation in submarine hydrothermalsystems: An update: Canadian Mineralogist, v. 26, p. 827–840.

Sinton, J.M., Ford, L.L., Chappell, B., and McCullough, M.T., 2003, Magmagenesis and mantle heterogeneity in the Manus back-arc basin, Papua NewGuinea: Journal of Petrology, v. 44, p. 159–195.

Solomon M., and Khin Zaw, 1997, Formation on the sea floor of the Hellyervolcanogenic massive sulfide deposit: ECONOMIC GEOLOGY, v. 92, p.689–695.

Solomon, M., Gemmell, J.B., and Khin Zaw, 2004a, The nature and origin ofthe fluids responsible for forming the Heyller volcanic-hosted massive sul-phide deposit, as indicated by fluid inclusions, and stable and radiogenicisotopes of sulphides, sulphates, silicates and carbonates: Ore Geology Re-views, v. 25, p. 89–124.

Solomon, M., Tornos, F., Large, R.R., Badham, J.N.P., Both, R.A., and KhinZaw, 2004b, Zn-Pb-Cu volcanic-hosted massive sulphide deposits: Criteriafor distinguishing brine pool-type from black smoker-type sulphide deposi-tion: Ore Geology Reviews, v. 25, p. 259–283.

Song, S-H., 1955, Massive Cu-bearing pyrite deposits and regional metal-logenic features in Qilian area: Acta Geologica Sinica, v. 36, p. 1–22(Chinese).

——1982, Pyrite-type deposits: Comparison study of the genetic types ofsome significant deposits in the world: Chinese Academy of Geological

GEOLOGY, FLUID INCLUSIONS, & OXYGEN ISOTOPES OF THE BAIYINCHANG VHMS DEPOSIT, GANSU, CHINA 291

0361-0128/98/000/000-00 $6.00 291

Sciences, Bulletin of the Institute of Mineral Deposits, no. 3, p. 1–12 (Chi-nese with English abstract).

Taube, A., 1986, The Mount Morgan gold-copper mine and environment,Queensland: A volcanogenic massive sulfide deposit associated withpenecontemporaneous faulting: ECONOMIC GEOLOGY, v. 81, p. 1322–1340.

Thiery, R., Van der Kerkhof, A.M., and Dubessy, F., 1994, VX properties ofCH4-CO2 and CO2-N2 fluid inclusions: Modeling for T<30°C and P<400bar: European Journal of Mineralogy, v. 6, p. 753–771.

Ulrich, T., Golding, S.D., Kamber, B.S., Khin Zaw, and Taube, A., 2002, Dif-ferent mineralization styles in a volcanic-hosted ore deposit: The fluid andisotopic signatures of the Mt. Morgan Au-Cu deposit, Australia: Ore Geol-ogy Reviews, v. 22, p. 61–90.

Urabe, T., Scott, S.D., and Hattori, K., 1983, A comparison of footwall-rockalteration and geothermal systems beneath some Japanese and Canadianvolcanogenic massive sulfide deposits: ECONOMIC GEOLOGY MONOGRAPH5, p. 345–364.

Urabe, T., Baker, E.T., Ishibash, J., Feely, R.A., and Marumo, K., 1995, Theeffect of magmatic activity on hydrothermal venting along the superfast-spreading East Pacific Rise: Science, v. 269, p. 1092–1095.

Vanko, D.A., Bach, W., Roberts, S., Yeats, C.J., and Scott, S.D., 2004, Fluidinclusion evidence for subsurface phase separation and variable fluid mix-ing regimes beneath the deep-sea PACMANUS hydrothermal field, Manusbasin back arc rift, Papua New Guinea: Journal of Geophysical Research, v.109, p. B03201–B03201.

Wu, J.-R., Ren, B.-C., Huang, Y., Zhu, M., Zhao, T., and Yu, P.-S., 1994, Mas-sive sulfide deposits in marine volcanic rock area in northwest China:Wuhan, China University of Geosciences Press, 271 p. (Chinese with Eng-lish abstract).

Xia L.-Q., Peng, L.-G., and Liu, B., 1985, Melt inclusions in quartz kerato-phyre at Baiyinchang: Bulletin of Xi’an Institute of Geology and MineralResources, no. 11, p. 1–8 (Chinese with English abstract).

Xia, L.-Q., Xia, Z.-C., and Ren, Y.-X., 1991, Marine volcanic rocks in the Qil-ian-Qinling Mountains: Wuhan, China University of Geosciences Press,258 p. (Chinese).

Xia, L.-Q., Xia, Z.-C., and Xu, X.-C., 1996, Petrogenesis of marine volcanicrocks in north Qilian: Beijing, Geological Publishing House, 153 p. (Chi-nese).

Xia, L.-Q., Xia, Z.-C., and Ren, Y.-X., 1998, Volcanism and mineralization inQilian Mountain and surrounding area: Beijing, Geological PublishingHouse, 169 p. (Chinese).

Xiang, D.-F., and Dai, T.-F., 1985, Regional metallogenic characteristics ofthe volcanogenic massive sulfide deposits in Northwestern Qilian Moun-tain: Mineral Deposits, v. 4, p. 64–74 (Chinese with English abstract).

Yan, J.-N., 1983, Origin of the Baiyinchang pyrite-type deposits in the QilianMountain: Mineral Deposits, v. 2, p. 15–23 (Chinese with English abstract).

Yang, K., and Scott, S.D., 1996, Possible contribution of metal-rich magmaticfluid to a seafloor hydrothermal system: Nature, v. 383, p. 420–423.

——2002, Magmatic degassing of volatiles and ore metals into a hydrother-mal system on the modern sea floor of the eastern Manus back-arc basin,western Pacific: ECONOMIC GEOLOGY, v. 97, p. 1079–1100.

292 ZENGQIAN ET AL.

0361-0128/98/000/000-00 $6.00 292