Methane-rich fluid evolution of the Baogutu porphyry Cu–Mo ... · Methane-rich ﬂuid evolution of the Baogutu porphyry Cu–Mo–Au deposit, Xinjiang, NW China Ping Shena,⁎,

Chemical Geology 275 (2010) 78–98

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Chemical Geology

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Methane-rich fluid evolution of the Baogutu porphyry Cu–Mo–Au deposit, Xinjiang,NW China

Ping Shen a,⁎, Yuanchao Shen a, Jingbin Wang a,b, Heping Zhu a, Lijuan Wang a,b, Lei Meng a,b

a Key Laboratory of Mineral Resources, Institute of Geology & Geophysics, Chinese Academy of Sciences, P.O. Box 9825, Beijing 100029, Chinab China Non-ferrous Metals Resource Geological Survey, Beijing 100012, China

⁎ Corresponding author. Tel.: +86 10 82998189; fax:E-mail address: [email protected] (P. Shen).

0009-2541/$ – see front matter © 2010 Elsevier B.V. Adoi:10.1016/j.chemgeo.2010.04.016

a b s t r a c t
a r t i c l e i n f o
Article history:Received 20 February 2010Received in revised form 26 April 2010Accepted 28 April 2010

Editor: D.B. Dingwell

Keywords:Methane-rich fluid inclusionsThe Baogutu porphyry Cu–Mo–Au depositWestern JunggarXinjiang

Baogutu is the first porphyry Cu–Mo–Au deposit discovered in Western Junggar, Xinjiang, NW China. Theore-bearing intrusion is a dioritic intrusive complex that includes stage 1 diorites and minor stage 2 dioriteporphyries. The stage 1 diorites have produced concentric potassic and propylitic alteration zone andoverprinted phyllic alteration and host much of the Cu–Mo–Au mineralization at Baogutu. The Baogutuporphyry Cu–Mo–Au deposit consists of unusually disseminated and minor vein-style mineralization.The fluid evolution occurred in stage 1 diorite from late magmatic stage to hydrothermal stage (stages 1Band 1C) is constructed based on alteration and fluid inclusion analysis by microthermometry, RamanSpectroscopy and Quadrupole Mass Spectrometer.Five fluid types are distinguished and they are L1 (liquid-rich 2-phase), L2 (liquid-vapor 2-phase), V1(vapor-rich 2-phase), V2 (mono-phase vapor), and H (multi-phase solid). The dominant inclusions rich inreductive CH4 plus H2O while minor inclusions rich in CH4 and CO2. The inclusions in quartz from the latemagmatic stage contain the most brine inclusions and less other inclusions with CH4-rich and H2O-richcompositions. The inclusions in quartz from stage 1B contain all fluid types and have CH4-rich and CH4+H2Ocompositions with rare CO2. The inclusions in quartz from stage 1C are characterized by more vaporinclusions without brine inclusions and a clear CO2 concentration in a few assemblages with CH4+CO2

composition.Using microthermometry, we estimate fluid trapping conditions at TN400 °C and P=1500 to 3100 bar in latemagmatic stage, T=200 to 400 °C and P=50 to 320 bar in disseminated quartz in stage 1B, T=180 to 400 °Cand P=20 to 260 bar in vein quartz in stage 1B, and T=170 to 400 °C and P=20 to 230 bar in stage 1C. Suchlate magmatic conditions are compatible with the fluid evolution as a result of CH4-rich inclusions derivedfrom mantle magma. Such hydrothermal conditions in stages 1B and 1C with small P–T fluctuations arecompatible with the dominant disseminated mineralization at Baogutu that indicates a weak fracturing. Theoverall fluid path at Baogutu is toward lower pressure and small changed temperature, with compositiontransit of fluid from a high halite CH4-rich system to low halite CH4+CO2 system. It may lead to theformation of this Cu–Mo–Au deposit.

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1. Introduction

The Central Asian Orogenic Belt (Altaids or Altaid collage, Xiao etal., 2009) lies between the Siberian and Russian cratons to the north,and Tarim and North China cratons to the south (Fig. 1a). The latePaleozoic tectonics of the Central Asian Orogenic Belt (CAOB) wascharacterized by continuous accretion along the southern activemargin of Siberia and the northern active margin of the North Chinacraton (Fig. 1a, Xiao et al., 2009, 2010). It led to the formation ofthe giant metal deposits (Fig. 1). The fifteen most important porphyryCu (Mo–Au) deposits are distributed in the CAOB (Rui et al., 1984;

Kudryavtsev, 1996; Rui et al., 2002; Qin et al., 2002; Nie et al., 2004;Cooke et al., 2004, 2005). They are Baogutu (in this study), Tuwu-yandong, Wunuhetushan, and Duobaoshan in China, Erdenet, Tsa-gaan–Suvarga, and Oyu Tolgoi in Mongolia, Boshekul, Samarsk,Kounrad, Aktogai, Borly, Sayak, and Koksai in Kazakhatan, andKal'makyr in Uzbekistan (Fig. 1a). Most are porphyry Cu–Mo andCu deposits exception for Oyu Tolgoi, Samarsk, and Kal'makyrwhich are porphyry Cu–Au deposits (Kudryavtsev, 1996; Zhukovet al., 1997; Heinhorst et al., 2000; Rui et al., 2002; Nie et al., 2004;Sillitoe, 2010).

Baogutu is the first porphyry Cu–Mo–Au deposit discovered inWest Junggar, Xinjiang, China (Fig. 1b). Four diamond drill holestotaling 1200 mcompleted by the Institute of Geology andGeophysics,Chinese Academy of Sciences in 1985–1990, one of them encounteredcopper mineralization (N0.1% Cu) over a 30 m interval (Shen et al.,

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79P. Shen et al. / Chemical Geology 275 (2010) 78–98

1993; Shen and Jin, 1993). Since 2002, 62 diamond drill holes totaling34,000 m have been completed by the Institute of Geology, XinjiangGeoexploration Bureau for Non-ferrous Metals, in Baogutu. Dissemi-nated and vein-style mineralization was discovered at the Baogutuporphyry Cu–Mo–Au deposit (Zhang et al., 2005, 2006b). Anevaluation in 2009 showed that Baogutu complex contained metal63×104 tonnes averaging 0.28% Cu, 1.8×104 tonnes averaging 0.011%Mo, 14 tonnes averaging 0.1 ppm Au, and 390 tonnes averaging1.8 ppm Ag (Zhang, unpublished data). Based on these data, Baogutuis Xinjiang's second largest porphyry copper deposit, after Tuwu–Yandong located in Eastern Tianshan, NW China (Fig. 1a). It consists ofunusually dominant dissemination (N80%) relative to most porphyrydeposits.

Methane rich fluid inclusions have been widely reported indifferent types of metallic ore deposits and petroleum basins (e.g.,Dubessy et al., 2001; Charlou et al., 2002; Hurai et al., 2002). Methanerich fluid inclusions also are reported in some rocks, such as the skarnnear the giant REE–Nb–Fe deposit at Bayan Obo, Northern China (Fanet al., 2004) and the mafic–ultramafic intrusion in Cu–Ni sulfidesdeposit in Eastern TianshanMountain, NW China (Liu and Pan, 2006).However, methane rich fluid inclusions are typically not detected inmost porphyry Cu deposits (Rusk and Reed, 2002), CO2 has beenidentified in inclusions from most porphyry Cu deposits (e.g., Butte:Rusk and Reed, 2002; Bajo de la Alumbrera: Ulrich et al., 2002;Bingham: Redmond et al., 2004; Landtwing et al., 2005; and ElTeniente: Klemmet al., 2007; Rusk et al., 2008; Landtwing et al., 2010).By contrast, fluid inclusions in quartz from Baogutu porphyry Cu–Mo–Au deposit trapmethane rich fluids based on the analysis results of theRaman Spectroscopy and Quadrupole Mass Spectrometer. The pres-ence of CH4 influids fromdiorite at Baogutu is an interesting feature. Inthis paper, we attempt to elucidate the evolutionary history ofmethane rich fluids that is responsible for alteration and mineraliza-tion of the Baogutu porphyry Cu–Mo–Au system associated withintermediate intrusive rocks (e.g., diorite).

2. Regional geology

TheWestern Junggar terrain is located in the central section of theCAOB (Fig. 1a). It consists mainly of Palaeozoic island arc and back-arcbasin rocks (Shen and Jin, 1993; Chen and Jahn, 2004; Xiao et al.,2008; Shen et al., 2009). They were accreted onto the Kazakhstanplate as the Tarim, Kazakhstan and Siberian plates converged (Chenand Jahn, 2004; Chen and Arakawa, 2005; Xiao et al., 2008). Thisgeodynamic process led to the formation of volcanic-hosted andintrusion-related gold deposits (Shen and Jin, 1993; Shen et al., 1996,2007, 2008) and porphyry copper deposits (Shen and Jin, 1993; Zhanget al., 2006a,b; Wang and Xu, 2006; Shen et al., 2008, 2009). Theyconstitute the Hatu gold and the Baogutu copper belts, respectively.The two metallogenic belts are separated by the Darbut fault (Fig. 1b).

The Western Junggar terrain consists mainly of Lower Carbonif-erous volcanic rocks which are widespread throughout southeasternWestern Junggar, particularly near the Darbut fault (Fig. 1b). Shen andJin (1993) studied two Early Carboniferous volcanic belts (Anqi andDarbut volcanic belts) in the Western Junggar terrain which areseparated by the Anqi fault (Fig. 1b). The volcanic rocks in the Anqivolcanic belt contain tholeiitic rocks which are inferred to reflect themagmatism under conditions of regional extension and back-arcbasin (Shen and Jin, 1993). The volcanic rocks in the Darbut volcanicbelt consist of tholeiitic and calc-alkaline assemblages, as well as felsicvolcaniclastic sequences, indicating a transitional setting from back-arc basin to arc (Shen and Jin, 1993; Shen et al., 2009) or an immaturearc setting.

Three Early Carboniferous stratigraphic units in Anqi and Darbutvolcanic belts have been identified. From oldest to youngest, these arethe Tailegula, Baogutu, and Xibeikulasi group (Shen and Jin, 1993).The Tailegula group consists of basaltic–(andesitic) flows and pillow

lavas and felsic tuff with intercalations of silica rock. Pillow basalt andchert of the Tailegula group in the Anqi volcanic belt yielded Rb–Srages of 328±31 Ma (Shen et al., 1993; Li and Chen, 2004) and 323±22 Ma (Li and Chen, 2004), respectively. The felsic tuff of the Tailegulagroup has a U–Pb zircon LA-ICP-MS age of 357.5±5.4 Ma (Guo et al.,2010). Gold mineralization is stratabound within the Tailegula groupand defines the Hatu goldmetallogenic belt in the northern part of theDarbut fault (Fig. 1b).The Baogutu group includes volcaniclasticsiltstone and sandstone, as well as lithic–vitric felsic tuff withintercalations of pebbly graywacke and the lens of limestone, marland bioclastic limestone. The felsic tuff of the Baogutu group in theDarbut volcanic belt have a U–Pb zircon SHRIMP age of 328.1±1.8 Ma(Wang and Zhu, 2007) and a U–Pb zircon LA-ICP-MS age of 332.1±3.0 Ma (Guo et al., 2010). The Xibeikulasi group consists of greywackewith graded bedding, volcaniclastic siltstone and mudstone withformed bedding.

These Early Carboniferous sequences are intruded by ore-bearingdiorite stocks at about ∼325 Ma (Fig. 2; Tang et al., 2009; Shen,unpublished data) and barren granite batholiths at ∼300 Ma (Chenand Jahn, 2004; Wang et al., 2004; Su et al., 2006; Han et al., 2006;Wang and Xu, 2006). The diorite stocks and its adjacent wall rocks arespatially and temporally related to copper mineralization, and definethe Baogutu metallogenic belt in the southern part of the Darbut fault(Fig. 1b). Shen et al. (2009) used whole rock geochemical analyses toconfirm that the intrusions are diorites and quartz diorites with atransitional character from tholeiite to calc-alkaline. Mantle-normal-ized trace element patterns depict three types including distinctlyfractionated REE pattern, enriched LREE and large ion lithophileelements (LILE) pattern, and little enrichment of LREE and LILEpattern. We therefore further suggest that the ore-bearing dioritestocks still formed in a Darbut immature arc or initial arc.

The regional structure is characterized by faults that displaydominant northeast trends including the Darbut, Anqi, and Hatufaults, although NS- and WE-trending structures are also present(Fig. 1b). Of these, the Darbut fault is the most important because itwas the locus of intense magmatism and associated mineralization.The main structures in the south of the Darbut fault are a series ofapproximately N-trending faults and folds which are almost orthog-onal to those to the north of the Darbut fault (Fig. 1b). The Xibeikulasisyncline is dominant structure which is traversed by several major N-and NE-striking faults (Fig. 2). The Xibeikulasi syncline consists ofXibeikulasi Group in core and Baogutu Group and Tailegula Group intwo limbs. The ore-bearing diorite stocks intruded the core and twolimbs of the Xibeikulasi syncline (Fig. 2).

3. Local geology

3.1. Host wall-rocks

The local geology comprises Early Carboniferous volcaniclastic,sediment, and volcanic rocks belonging to the Baogutu Group andXibeikulasi Group (Fig. 3). The Baogutu group, which is dominant hostwall-rocks around the mineralized complex, is composed of volcani-clastic and volcanic rocks. They contain the dominant felsitic lithic–vitric tuff, silty tuff and volcaniclastic siltstone and minor andesite.The disseminated and vein mineralization occurred mainly in thevolcaniclastic siltstone, silty tuff and lithic–vitric tuff. The barrenXibeikulasi group, occurs in the western part of the mineralizedcomplex, consists of greywacke, tuffaceous mudstone and tuffaceoussiltstone.

3.2. Intermediate complex

The mineralized complex intruded the eastern limb of theXibeikulasi syncline into the Lower Carboniferous Baogutu andXibeikulasi Groups and occupies localized dilatant sites which provided

80 P. Shen et al. / Chemical Geology 275 (2010) 78–98

Fig. 2. Geological map of the Baogutu porphyry copper belt in the West Junggar (modified after Shen and Jin, 1993).


by a structural intersection of the N- and ENE-trending faults (Figs. 2and 3).

The host rocks at Baogutu have been described in different ways bypreviousworkers,mainly as granodiorite and quartz diorite (Shen andJin, 1993) or granodiorite porphyries (Zhang et al., 2005, 2006b;Cheng and Zhang, 2006; Zhang et al., 2006a; Song et al., 2007)respectively. We have identified the ore-bearing intrusion as a dioriticintrusive complex that includes equigranular diorite, weakly porphy-ritic diorite, and diorite porphyry, with rare granodiorite (Table 1).Two intrusive phases recognized by us at Baogutu based on the cross-cutting relationships are the main-stage (stage 1) equigranular toweakly porphyritic diorites and minor late-stage (stage 2) dioriteporphyries. Moreover, two breccias variants (mineralized hydrother-mal breccias and very weak mineralized matrix-rich breccias) havebeen also identified. Alteration and mineralization were closelyrelated to the main-stage diorites. Spatial relationships between thediscrete intrusive phases are shown in Figs. 3 and 4.

3.3. Ore bodies

The orebodies define an area that is 1100 m×800 m and extendfor more than 800 m downwards. They are hosted in the dioritecomplex and adjacent wall rocks based on the 62 drill holes (Figs. 3and 4). The orebodies are covered by Quaternary overburden. The0.2% Cu contour defines a wedge-shaped zone approximately 800 mwide and more than 700 m deep in a cross-sectional view (Fig. 4).Mineralization is more strongly developed in the northern andeastern parts of the complex, and Cu and Mo grades graduallyincrease with increasing depth. Highest-grade zones (N0.4% Cu or

Fig. 1. a. Schematic map of the Central Asian Orogenic Belt (modified after Xiao et al., 2009) shA are the following: 1 — Baogutu; 2 — Tuwu-yandong; 3 —Wunuhetushan; 4 — DuobaoshanKounrad; 11— Aktogai; 12— Borly; 13— Sayak; 14— Koksai; 15— Kal'makyr. b. Regional geJunggar showing the distribution of the Anqi volcanic belt and the Darbut volcanic belt andfault separates the Anqi volcanic belt and the Darbut volcanic belt and the Darbut fault sepaShen and Jin, 1993).

N0.03% Mo) occur at depths of 300–700 m depth below the surface atthe northwestern and eastern parts of the complex associated withhigher fracture intensities and/or intensively developed hydrother-mal breccias (e.g., Fig. 4). The orebody dips northeast at an angle ofabout 45 to 55° implying that significant tilting may have occurredafter ore formation, given that porphyry deposits typically have sub-vertically oriented ore shells (e.g., Hedenquist et al., 1998; Wilsonet al., 2003).

The mineralized complex well-develops the disseminated, minorvein-style Cu–Mo–Au mineralization. Cu–Au mineralization is local-ized in the main parts of the complex and at its contact with theadjacent wall rocks. Au mineralization, which gold grades range from0.03 to 0.4 g/t Au with average 0.1 g/t, is by-product. Economic Momineralization (common N0.01% with minor N0.06%) has been de-tected at depth in the complex. The Baogutu ore-bearing complexcould have the Cu–Mo–Au assemblage.

3.4. Ages of ores and their host porphyries

The ages of the ores and their porphyries at Baogutu are deter-mined. The molybdenites from two veins of the Baogutu complexhave been dated by the Re–Os method, yielded age of 310.1±3.6 Maand 310.4±3.6 Ma (Song et al., 2007). Main-stage quartz diorite inthe Baogutu complex yielded Rb–Sr ages of 322±30 Ma (Shen et al.,1993). We obtained a U–Pb zircon SHRIMP age of 325.1±4.2 Ma(Shen, unpublished data). Late-stage diorite porphyry of the Baogutucomplex yielded U–Pb zircon LA-ICP-MS age of 309.9±1.9 Ma (Tanget al., 2009). These data indicate that the Baogutu complex range inage from 325 to 309 Ma and the molybdenites ores 310 Ma.

owing principal porphyry copper deposits. The number of themineral deposits in panel; 5 — Erdenet; 6 — Tsagaan-Suvarga; 7 — Oyu Tolgoi; 8 — Boshekul; 9 — Samarsk; 10 —

ological map of the Baogutu porphyry copper belt and the Hatu gold belt in theWesternthe location of the gold deposits and porphyry copper deposits. The NE-trending Anqirates the Hatu gold belt and the Baogutu copper belt (modified from Shen et al., 1993;

Fig. 3. Geological map of the Baogutu porphyry Cu–Mo–Au deposit showing the intrusion complex. WE01 show location of used in section shown in Figs. 4 and 6, dots indicate theposition of drill holes, and pentagons show the location of the samples.


4. Materials and methods

4.1. Sample selection

Selected sample originates from 3 diamond drill holes (Fig. 3)within the hypogene alteration zones (phyllic and potassic) in thestage 1 diorites. Fluid inclusions were studied in quartz crystals fromstage 1 diorites and different vein stages, corresponding to the spatial

Table 1Summary description of the intrusion sequence and rock types of the Baogutu complex.

Intrusionhistory

Intrusioncomposition

Location Texture

Main stage(stage 1)

Equigranulardiorite

Center and depth ofthe complex

Equigranular hypidiomorphic, figrained (1–3 mm)

Weaklyporphyriticdiorite

Exposes to the outer ofthe equigranulardiorite

Weak porphyritic texture, 20–3(0.5–3 mm), and coarse ground

Hydrothermally-cementedbreccias

Center and bottem ofthe complex

50–60% clasts up to 3–5 cm widminerals cement

Late stage(stage 2)

Diorite porphyry Distributes around theporphyritic-like diorite

Porphyritic, 10–20% phenocrysmicrocrystalline (b0.01 mm) angroundmass

Matrix-richbreccias

Center and bottem ofthe complex

Composed of 70–80% clasts andhydrothermal minerals cement

and temporal evolution of the magmatic–hydrothermal system atBaogutu; the evolution was determined from geologic relationshipsand petrographic investigation of successive quartz generations.

About 100 core samples containing stage 1 diorites and variousvein types from various depths were collected for laboratory analyses.Over 70 samples were observed for inclusion type, abundance, spatialdistribution, and size. Approximately 30 of these samples werefurther analyzed by microthermometry. Over 800 inclusions were

Mineralogy assemblage

ne- to medium- Plagioclase (30–40%), hornblende (15–20%), biotite(10–15%), pyroxene (b5%), and quartz (b5–10%)

0% phenocrystmass (0.05–0.5 mm)

Plagioclase (30–40%), hornblende (15–20%), biotite(10–15%), and quartz (5–15%)

e with hydrothermal Contain complex and wall-rock clasts, with hydrothermalminerals cement without matrix; breccias can be wellmineralized

ts (2–3 mm) andd (0.02–0.05 mm)

Plagioclase, hornblende, and biotite phenocrysts in quartz(5–15%) plagioclase, and hornblende groundmass

matrix with minor As above hydrothermal breccias, but rich in matrix and lackthe hydrothermal minerals cement with weakmineralization

Fig. 4. Geologic cross section along WE01 showing the host rocks to the Baogutu porphyry Cu–Mo–Au deposit (logged by authors). The copper ore bodies are modified from Zhanget al. (2006b). See Fig. 3 for location.


analyzed by microthermometry. The inclusions range from less than 3to ∼8 μm in diameter, with the vast majority of inclusions being lessthan 6 μm. Most inclusions analyzed by microthermometry werebetween 4 and 6 μm in diameter. All analysis is carried out at theInstitute of Geology and Geophysics, Chinese Academy of Sciences.

4.2. Methods

4.2.1. MicrothermometryThemicrothermometric studyoffluid inclusionswas carriedoutwith

a Leitz microscope and a Linkam THMS 600 programmable heating-freezing stage (e.g. Shepherd et al., 1985; Luet al., 2004). The precision oftemperature measurements on cooling runs is about ±0.1 °C. Onheating runs, the precision is ±2 °C. Ice-melting temperatures wereobserved at a heating rate of no more than 0.1 °C/s. Homogenizationtemperatures were observed at a heating rate of 1 °C/s. Homogenizationof halite-bearing inclusions was obtained on a heating cycles of about5 °C.

4.2.2. Raman spectroscopyIn order to confirm the suggested fluid inclusion volatile species,

representative samples were analyzed using a Renishaw 1000 Ramanmicrospectrometer according to the method of Burke (2001). With thistechnique, inclusions are homogenized in a heating stagemounted on a

microscope attached to a Renishaw 1000 Raman microprobe. The laserbeamwith a wavelength of 514.5 nm and a spot size of about 1 μm,wasfocused on the bubble for each fluid inclusion through a light micro-scope. The Raman peaks of the CH4 and H2O are 2914–2916 cm−1 and3500 cm−1, respectively. The Raman peaks of the CO2 are about 1281and 1386 cm−1 in this study.

4.2.3. Quadrupole mass spectrometerIn order to further determine gaseous composition of fluid

inclusions and contents, four samples were analyzed using a PrismaTM QMS200 Quadrupole Mass Spectrometer according to the methodof Zhu and Wang (2000). The dried washed sample weighing 50 mgwas put in a clean quartz tube, and heated to 100°C. Then, valve wasturned on and the gas pipe was vacuumed. When the pressure in thequartz tube was less than 6×10−6 Pa, the burst stove was heated to550°C at the speed of 1°C/3 s. The vacuum valve was turned on fordetermination of gaseous composition.

5. Alteration and veins

5.1. Alteration zonation and alteration stages

We carried out alteration mapping and document the alterationcharacteristics of the study Baogutu complex. The intrusive complex


has been subjected to intense hydrothermal alteration, especiallywithin and adjacent to the ore-bearing intrusion (Fig. 5). Based onobservations from outcrops (Fig. 3) and 10 drill holes (Fig. 4), detailedexamination of alteration overprinting relationships was carried outon 766 polished thin sections using transmitted and reflected lightmicroscopy. Figs. 5 and 6 show the distribution of alteration mineralsin detail plan and across an east–west vertical cross section at theBaogutu deposit, respectively. We divided the alteration into twostages and five sub-stages at Baogutu and identified disseminationsand/or vein generations within each of the paragenetic stages(Table 2). The alteration and mineralization occurred in the stage 1diorites are emphasized only here.

Within the stage 1 diorites overprinted by the hydrothermalbreccias, Ca–Na silicate alteration assemblage (stage 1A), potassicalteration assemblage (stage 1B) and phyllic alteration assemblage(stage 1C) have been recognized; the Ca–Na silicate alteration zonewas not clearly outlined; the potassic zone have been subdividedinto the inner potassic subzone in the complex (a range of 1200 m×1400 m) and outer potassic subzone in the wall rocks (100 m to500 m in width); the distal propylitic zone is 50 m to 500 m in width;the phyllic zone (700 to 1000 m in area) overprinted in the potassiczone rather than the concentric ring around the potassic zone asmapped previously (Cheng and Zhang, 2006; Zhang et al., 2006b). Cuand Mo sulfides are associated spatially and temporally with potassicand phyllic assemblages. The distal propylitic and early inner Ca–Nasilicate alteration assemblages lack significant sulfides. Within thestage 2 diorite porphyries, pervasive potassic alteration (stage 2A)occurred in the groundmass of porphyries with mineralization. Thematrix-rich breccias (stage 2B) have weak alteration in cement withvery weak mineralization (Table 2).

Fig. 5. Alteration zonation at Ba

5.2. Quartz generations and vein sequences

5.2.1. Quartz in late magmatic stageConstraints on the spatial and temporal evolution of the magmatic–

hydrothermal systemat Baogutuwere based on thegeologic constraintsof quartz-bearing samples, both late magmatic and hydrothermal. Theearliest fluids in the magmatic–hydrothermal system are preserved ininterstitial quartz (I-quartz) hosted in the diorites (Fig. 7a). Most ofthese quartzes have anhedral forms due to latest growth within thecrystallizing magma. Commonly, they show no evidence for recrystal-lization. Therefore, they are thought to represent the late magmaticstage of the magmatic–hydrothermal system.

5.2.2. Quartz in hydrothermal stage 1BHydrothermal processes in the Baogutu were deduced largely

from the distribution of the different hydrothermal quartz in dioritesand veins. The earliest quartz type, the disseminated quartz (D-quartz), is represented by anhedral granular quartz with variableamounts of biotite, magnetite, chlorite, and rutile (Fig. 7b). As a resultof the hydrothermal metasomatism, D-quartzes occur in the possiticdiorites andwall-rocks of the Baogutu Group, in close proximity to theintrusive rocks. Petrographically later quartz type, the hydrothermallycemented quartz (H-quartz), is subhedral granular quartz withvariable amounts of biotite and magnetite (Fig. 7c). The H-quartz(now matrix) occurs in the mineralized hydrothermal breccias(Fig. 8b) which located in the centre and deepest parts of Baogutucomplex (Fig. 4) and is associated with potassic alteration.

The stage 1B corresponds to local quartz veinlets or veins without adistinct alteration halo. The vein type includes the early quartz–chalcopyrite–pyrite–biotite ± pyrrhotite veinlets (B1-veinlets, Fig. 8c),

ogutu Cu–Mo–Au deposit.

Table 2Major characteristics of alteration and veins in the Baogutu complex (in relative age sequence, oldest at top).

Stage Alterationtype

Alteration assemblage Mineralization type Vein type Vein distribution Texture Veinthickness

This study

Stage1A

Ca–Na silicatealteration

Act–Alb–Mt ± Ep Rare in diorites Disseminated Not used

Stage1B

Potassicalteration

Bio–Qrz–Mt ± Rut ± Chlin diorites; Bio–Qtz–Mt–Ksp–Apa in wall rocks

Intense disseminated D-quartz Abundant in diorites Disseminated UsedModerate breccias overprintwith Qtz–Bio–Cp–Pyassemblage

H-quartz Moderately abundant indiorites

Cement-richhydrothermal breccias

Used

Qtz–Cp–Py veinlets B1-veinlets Moderately abundant indiorites

Veins with irregularparallel walls

b5 mm Used

Qtz–Cp–Py–Mo veinlets B2-veinlets Moderately abundant indiorites

Veins with irregularparallel walls

b5 mm Used

Qtz–Cp–Py veins B3-veins Moderately abundant indiorites and wall rocks

Veins with parallel walls 1–2 cm, upto 10 cm

Used

Barren Bio–Qtz veinlets B4-veinlets Common in diorites andwall rocks

Veins with wavy walls 0.1–0.2 mm Not used

Barren Ksp–Mt veinlets B5-veinlets Rare in wall rocks Veins with wavy walls 0.1–0.5 mm Not usedBarren Apa veinlets B6-veinlets Rare in wall rocks Veins with wavy walls 0.1–0.5 mm Not used

Propylitialteration

Chl–Ep–Py–Ser–Cal Qtz–Cp–Py veins B3–veins Peripheral wall rocks Veins with parallel walls 1–2 cm Not used

Stage1C

Phyllicalteration

Ser–Qtz–Py ± Chl ± Cal Qtz–Mo–Cp veins withphyllic alteration envelopes

C1-veins In diorites Veins with wavy walls 1–2 cm Used

Qtz–Cp–Py veins withphyllic alteration envelopes

C2-veins In diorites Veins with irregularwalls

b1 cm Used

Pyr–(Qtz–Cal) veinlets C3-veinlets In diorites Massive sulfide veinletswith regular walls

b1 cm Not used

Gpy - (Qtz–Cal) veinlets C4-veinlets In stage 1 diorites Veins with regular walls b1 cm Not usedStage2A

Potassicalteration

Bio–Qtz–Mt Moderate disseminated Abundant throughoutdiorite porphyry

Disseminated ingroundmass

Not used

Stage2B

Potassicalteration

Gyp–Qtz –Cal–Bio Weak breccias overprint withGyp–Qtz–Cal–Bio–Cp–Pyassemblage

In complex and theircontact zone

Matrix-richhydrothermal breccias

Not used

Mineral abbreviations: act = actinolite, alb = albite, ap = apatite, bio = biotite, bn = bornite, cal = calcite, cc = chalcocite, chl = chlorite, cp = chalcopyrite, ep = epidote, gyp =gypsum, ksp = K-feldspar, mo = molybdenite, mt = magnetite, prl = pyrophyllite, pyrrhotite = pyr, py = pyrite, qtz = quartz, rut = rutile, sl = sphalerite, ser = sericite.

Fig. 6. Cross section WE01 showing domains of secondary biotite and propylitic alteration in stage 1B and sericite alteration in stage 1C.


Fig. 7. Photomicrographs showing the features of the different quartzs from Baogutu Cu–Mo–Au deposit. (a) I-quartz (interstitial quartz) hosted in weakly porphyritic biotite dioritein late magmatic stage; (b) D-quartz (disseminated quartz) in biotite diorite in stage 1B; (c) H-quartz (hydrothermally cemented quartz) occurred in matrix of the hydrothermalbreccias in stage 1B; (d) B1-veinlets in stage 1B with internal symmetry. All photomicrographs were taken under cross polarized light exception for (b). Abbreviations: Pl:plagioclase; P-Bio: primary biotite; S-Bio: secondary biotite; Mt: magnetite; Ser- sericite; Q: quartz.


quartz–chalcopyrite–pyrite–molybdenite–biotite veinlets (B2-veinlets),and the late quartz–chalcopyrite ± pyrite veins (B3-veins, Fig. 8d). Theyare similar to A-type veins occurred in porphyry deposits (Sillitoe, 2010).Of these, B1-veinlets are predominant. B1- and B2-veinlets (b1 cmwidth) are filled with subhedral quartz, anhedral biotite, chalcopyrite,pyrite and pyrrhotite. Subhedral quartz is the predominant ganguemineral and contains abundant fluid inclusions. The internal symmetryor open vugs (Fig. 7d) and the irregularwalls suggest that they formed inthe brittle rocks. They occur throughout the inner potassic alterationsubzone and also extend into the adjacent volcanic–sedimentary wallrocks of the Baogutu Group. B3-veins are planar and continuous veins.The veins aremore than 5 cmwide and cut thewall rocks of the BaogutuGroup.

5.2.3. Quartz in hydrothermal stage 1CThe late vein type includes C1-veins, C2-, C3- and C4-veinlets. C1-

veins have a quartz–molybdenite–chalcopyrite–pyrite assemblage,and are similar to the B-type veins occurred in porphyry deposits(Sillitoe, 2010). C1-veins are typically thick (1–5 cm), molybdenite-rich, and have been observed within diorites (Fig. 8e). Vein walls areparallel and slightly wavy. C2-veinlets have a quartz–chalcopyrite–pyrite–biotite±pyrrhotite assemblagewith obvious phyllic alterationhalos (including quartz and sericite) from 1 mm to 2 cm thick andformed within diorites as well as in the adjacent wall rocks (Fig. 8f, g).C2-veins are the same as the D-type veins occurred in porphyrydeposits (Sillitoe, 2010). C3-veinlets are characterized by a pyrrhotite–

Fig. 8. Photographs of mineralization from Baogutu Cu–Mo–Au deposit. (a) intensive dissemchalcopyrite andmagnetite, note the abundant of chalcopyrite andmagnetite; (b) breccia orein the disseminated mineralization in diorite; (d) vein ore in stage 1B, quartz–chalcopyrite–with phyllic alteration halos; (f) vein ore in stage 1C, quartz–chalcopyrite–pyrite veinlets wdiorite porphyry with abundant quartz–chalcopyrite–pyrite veinlets; (h) vein ore in stage 1Abbreviations: Q: quartz; Cp: chalcopyrite; Py: pyrite; Mo: molybdenite; Pyr: pyrrhotite; G

calcite ± quartz assemblage (Fig. 8h) and vary from b3 mm to 1 cmwide and occur only within the strong sericite alteration domains. Thelatest C4-veinlets have a gypsum–calcite ± quartz assemblage(Fig. 8h) and vary from 1 to 5 mm wide and cut the phyllic zone andrelated veins. Quartz in stage 1C is mostly granular and orientedperpendicular to the vein walls, suggesting that veins formed in thebrittle rocks. Quartz veins in stage 1C cut the potassic zone and relatedveins which indicate that they generally postdate veins in stage 1B.

6. Fluid inclusion petrography and microthermometry

6.1. Fluid inclusion types

The classification of fluid inclusion types observed in this study isprimarily based on phase proportions at room temperature. Alldescriptions strictly refer to fluid inclusion assemblages (Wilkinson,2001). Samples examined for fluid inclusion study are representativeof the different quartz stages described above and include I-quartz, D-quartz, H-quartz, vein quartz (B1-veinlets, B2-veinlets, B3-veins, C1-veins, C2-veins). Five fluid types are distinguished in this study basedon petrographic and microthermometric criteria. They are L1 (liquid-rich 2-phase), L2 (liquid-vapor 2-phase), V1 (vapor-rich 2-phase), V2(mono-phase vapor), and H (multi-phase solid). The fluid inclusiontypes distinguished in this study are presented in Table 3 and shown inFig. 9.

inated ore in stage 1B, medium-grained diorite with disseminated hydrothermal biotite,in stage 1B; (c) veinlet ore in stage 1B, quartz–chalcopyrite–pyrite veinlets overprintedpyrite veins in diorite; (e) vein ore in stage 1C, quartz–molybdenite–chalcopyrite veinsith phyllic alteration halos occurred in possitic alteration zone; (g) vein ore in stage 1C,C, pyrrhotite–calcite ± quartz veinlets is cut by and gypsum–calcite ± quartz veinlets.yp: gypsum.


Table 3Microthermometry data and pressure estimates of the fluid inclusions from the quartzes at Baogutu.

Sample no. Drill hole no. Depth(m)

Host quartz(stage)

Alterationzone

FI type Th (°C) Tm (ice) or Ts(halite) °C

Salinity(wt.% NaCl equiv)

P (bars) Basis for P estimate

ZK102-528 ZK102 528 I-quartz(late magmatic)

Diorite H 265–281 (23) 400–530 (13) 47.4–63.9 1542–3128 Halite disappearance

ZK102-530 ZK102 530 I-quartz(late magmatic)

Diorite H 209–284 (15) 420–482 (15) 47.4–57.1 1565–3103 Halite disappearance

ZK203-260 ZK203 260 D-quartz(stage 1B)

Potassic L2 309–310 (2)

H 279–282 (9) 345–395 (9) 41.5–46.4 99–1020 Halite disappearanceZK211-783 ZK211 783 D-quartz

(stage 1B)Potassic L1 271–274 (3)

L2 271–399 (19) −2.2∼−5.4 (8) 3.71–8.41 54–160 LV curveV1 361–378 (2)H 279–281 (2) 315–375 (2) 39.1–44.4 136–320 Halite disappearance


Potassic L1 218–246 (3) −5.8 (1) 8.98 28 LV curve

L2 243–365 (10) −0.9∼−10.9 (6) 1.57–14.87 63–194 LV curveV1 357–358 (2)


Potassic L1 207–278 (8)

L2 201–348 (13) −2.8∼−4.0 (3) 4.64–6.45 72–112 LV curveV1 281–358 (2)H 189–191 (2) 383–448 (2) 45.33 249–307 Halite disappearance


Potassic L1 215–333 (2) −3.5 (1) 5.71 20 LV curve

L2 270–379 (10) −2.9∼−4.0 (4) 4.80–6.45 56–96 LV curveV1 339H1 180 371 (1) 44.32 863 Halite disappearance

ZK211-582 ZK211 582 H-quartz(stage 1B)

Potassic L2 201–378 (2) −5.3∼−6.4 (2) 8.28–9.73 32–108 LV curve

V1 231–369 (24)V2 395–422 (3)

ZK211-155 ZK211 155 B1-veinlets(stage 1B)


L2 259–386 (19) −6.4∼−8.9 (5) 9.73–12.73 41–97 LV curveV1 270–441 (6) −6.8∼−12.8 (2) 10.24–16.71 161 LV curve


Potassic L1 215–221 (3)

L2 234–299 (10) −1.8 (1) 3.06 68 LV curveV1 324–336 −3.7 (1) 6.01 138 LV curve


Potassic L1 162–230 (6) −0.1∼−2.8 (4) 0.2–4.7

L2 165–300 (13) −2∼−4 (7) 3.4–6.5 21–82V1 330 (1)


Potassic L1 200–321 (8) −3∼−10 (5) 5–14 6–20 LV curve

L2 205–355 (23) −1.8∼−11 (11) 3–15 38–39 LV curveV1 200 (1)



L2 221–368 (23) −1.5∼−5 (17) 2.67–15.5 41–129 LV curveV1 273–330 (7) −0.8∼−3.8 (6) 1.4–6.2 84–200 LV curve


Potassic L1 201–269 (3) −1.2∼−9.3 (2) 2.07–13.18

L2 302–381 (19) −6.1∼−13.6 (7) 9.34–17.43 77–127 LV curveV1 362–418 (8)H 234–261 (3) 344–550 (3) 41.45–66.75 593–612 Halite disappearance


Potassic L1 205–257 (2)

L2 170–278 (21) −1.2∼−17.7 (9) 2.07–20.75 15–31 LV curveV1 183–260 (12) −7.8∼−20 (5) 11.46–22.71V2 345 (1)H 332 (1) 39.76 87 Halite disappearance


Potassic L1 178–194 (13)

L2 221–309 (27) −5∼−9 (9) 7.9–12.8 20–52 LV curveV1 300–440 (8) −8∼−9 (4) 11.7–12.8 272–263 LV curve

ZK211-527 ZK211 526.5 B3-veins(stage 1B)

Potassic L1 221–276 (8) −1.8∼−2.8 (3) 3.06–4.65

L2 234–358 (21) −5.1 (1)ZK211-209 ZK211 209 B3-vein

(stage 1B)Potassic L1 156–391 (11) −8.4∼−18.4 (3) 12.16–22.26 11–42 LV curve

L2 234-368 (19) −3.8∼−17 (2) 6.16–20.22 148–174 LV curveV1 335–360 (7)V2 321 (1)

ZK102-272 ZK102 272 C2-veinlet(stage 1B+1C)

phyllic L2 250–385 (27) −3.1∼−15.2 (6) 3.4–16 32–218 LV curve

V1 210–460 (16) −2.1∼−21.1 (7) 2.7–23 18–237 LV curve


Table 3 (continued)

Sample no. Drill hole no. Depth(m)

Host quartz(stage)

Alterationzone

FI type Th (°C) Tm (ice) or Ts(halite) °C

Salinity(wt.% NaCl equiv)

P (bars) Basis for P estimate

V2 320–421 (3) −4∼−11 (2) 6.5–15 101–133 LV curveZK203-44 ZK203 44 C2-vein

(stage 1C)phyllic L1 164–298 (6) −2.2∼−8.2 (4) 3.71–11.93 19–21 LV curve

L2 240–316 (12) −2.6∼−10 (5) 4.34–13.94 10–81 LV curveV1 267–400 (7) −1∼−8.8 (3) 1.7–12.62 62–125 LV curveV2 354–391 (3) −0.7 (1) 1.23

ZK211-363 ZK211 363 C2-vein(stage 1C)

phyllic L1 151–251 (7) −5.3∼−6.7 (4) 8.28–10.11 4–38 LV curve

L2 191–276 (13) −3.1∼−7.6 (7) 5.11–11.22 35–100 LV curveV1 236–369 (6) −3.3∼−5.4 (2) 5.41–8.41


phyllic L1 214–273 (19) −0.5∼−7.2 (9) 0.88–10.73 24–49 LV curve

L2 226–386 (17) −2.1∼−5.0 (11) 3.55–7.86 26–33 LV curveV1 293–321 (17)


phyllic L1 173 (1) −5.4 (1) 8.41 8 LV curve

L2 197–333 (12) −1.6∼−8.6 (4) 2.74–12.39 25–110 LV curveV1 345 (1)

Abbreviations: FI = fluid inclusions, Th = vapor bubble disappearance temperature, Ts = halite dissolution temperature, Tm = final ice melting temperature.Pressure for H inclusions calculated according to Bodnar (1994) for halite homogenizing inclusions and the assumed pressure, respectively.


H inclusion is halite saturated at room temperature, generally hasnegative crystal shapes, and is typically between 4 and 6 μm in size. Itcontains liquid plus solid phases plus 10 to 20vol.% vapor at roomtemperature (Fig. 9a). In contrast to most porphyry Cu deposits (e.g.,Alumbrera, Bingham, Santa Rita, Red Mountain); most H inclusions atBaogutu contain no other daughterminerals in addition to halite, minorH inclusions contain several solid phases (e.g. halite, sylvite, opaque).Halite crystals are larger than the other solids and can be readilydistinguishedby their cubic shape. Sylvite (whenpresent) is smaller andhas a clear rounded shape. The very rare opaque may be sulfides. Hinclusions are commonly scattered variably in quartz (Fig. 9a).

Vapor-rich inclusions were divided, according to their filling ratio,into V1 inclusions (60–80% vapor) and V2 inclusions (N85% vapor).Both types commonly have rounded isometric or negative crystalshapes and range between 4 and 8 μm in size. No daughter crystals andCO2 were observed. They are scattered in individually quartz grains(Fig. 9b–f).

Aqueous inclusions contains liquid and b30–40vol.% vapor andlacks halite daughter minerals at room temperature (Fig. 9b–f). Theyoccur as negative crystal shape, rounded or irregular shaped inclusionsand divided into two groups: L1 and L2. The L1 inclusions have a smallvapor bubble of 10 to 20% inclusion volume. The L2 inclusions have alarge vapor bubble that occupies about 30–40% of the inclusionvolume. No daughter crystals and CO2 were observed. Aqueousinclusions are generally widespread in veins and range in size fromaround 3 up to 8 μm. There is a continuous gradient in bubble sizebetween L1 and L2 inclusions; therefore, distinguishing between L1and L2 inclusions at room temperature is sometimes difficult. They arescattered or occur along secondary or pseudosecondary trails.

6.2. Fluid inclusion microthermometry

We observe that most inclusions are randomly scattered andminor are aligned along cross-cutting healed micro-fractures atBaogutu. It differs from the individual quartz crystals and veinsections with continuous growth zonation and no significant crackfilling or dissolution discontinuities at the El Teniente porphyry Cu–Mo deposit, Chile (Klemm et al., 2007). It also differs from typical veinreactivation textures documented at Butte, Montana (Rusk and Reed,2002), Bingham, Utah (Redmond et al., 2004; Landtwing et al., 2005),Copper Canyon, Nevada (Nash, 1976), Bajo de la Alumbrera, andArgentina (Ulrich et al., 2001). Careful attention was required duringsampling to distinguish between potassic alteration features and thelater alteration overprint and veins occurred in stages 1B and 1C in

order to identify the superimposition of some generations ofinclusions in samples. Our descriptions of potassic assemblages andrelated veins rely heavily on petrographic specimens determined tohave undergone only limited later alteration. Based on these, someinclusions which are relatively large (N4 μm) and preferably polyg-onal-shaped scattered in late magmatic–hydrothermal stage atBaogutu are measured in this study (Table 3).

Two common thermometric procedures, freezing and heating, wereemployed to determine the approximate salinity (wt.%NaCl equivalent)and homogenization temperature (Th), respectively. Freezing wascarried out mainly for non-halite-bearing liquid-rich inclusions, andthe temperature of the melting point of ice was recorded. The collecteddata were converted into corresponding salinity values by usingdiagrams for NaCl–H2O system (Shepherd et al., 1985). The heatingstage was used for all types of inclusion. For non-halite bearing inclu-sions the homogenization temperature of liquid and vapor (predom-inant L+V→L and rare L+V→V) was recorded. In the halite-bearinginclusions, two points: (1) Th (temperature of vapor and liquidhomogenization) and (2) Ts (NaCl) (the temperature at which halitedissolves) were recorded. The solid phase is predominantly halite,occasionally accompanied by sylvite; therefore, almost all inclusionsmeasured are halite-bearing inclusions and Ts (KCl) (the temperature atwhich sylvite dissolves) was not recorded. The only Ts (NaCl) valueswere converted into wt.% NaCl by using phase diagrams of binarysystem NaCl–H2O (Shepherd et al., 1985). The unidentified opaqueminerals seldom form more than 1% of the volume of an inclusion andtherefore, do not significantly affect its homogenization temperature. Inthe case of V2 inclusions, the final melting temperature of ice wasdifficult to determine, because of the high vapor/liquid ratios.

6.2.1. I-quartz in late magmatic stageFluid inclusions in the interstitial quartz occurred in the diorites

contains the predominant halite saturated inclusions andminor vaporinclusions. They are thought to represent the late magmatic stage ofthe magmatic–hydrothermal system. All hyper-saline fluid inclusionsare homogenized by halite dissolution. Final homogenization tem-peratures by halite dissolution ranged from 401° to 550 °C (Fig. 10a,Table 3). Salinities calculated from these halite dissolution tempera-tures range between 47 and 65wt.% NaCl equiv (Fig. 10a).

6.2.2. D-quartz in stage 1BFluid inclusions in the disseminated quartz occurred in the potassic

zone homogenize mainly between 200° and 400 °C (Fig. 10b, Table 3).The hyper-salinefluid inclusions are homogenized by halite dissolution.


Fig. 10. Histograms of homogenization temperatures and salinities (wt.% NaCl equivalent) from microthermometric data for all fluid inclusions from different quartzes.


Final homogenization temperatures by halite dissolution ranged from315° to 440 °C. Salinities calculated from these halite dissolutiontemperatures range between 39 and 46wt.% NaCl equiv (Fig. 10b).Aqueous inclusions have a salinity mainly range of 3.7 to 9.7wt.% NaClequiv andhomogenize from200° to390 °C.Minor vapor inclusionshavethe high apparent homogenization temperatures (395° to 442 °C) toaqueous inclusions (Table 3).

6.2.3. Vein quartz in stage 1BFluid inclusions in quartz occurred in B1-, B2-veinlets and B3-veins

have similar characterization (Fig. 10c). They are characterized bydominant L2 inclusions and abundant V1, L1 inclusions. V2 inclusionis rare. L1 inclusion homogenize between 180° and 310 °C, with onediscrete population occurring between 200° to 240 °C. L2 inclusionshave the high apparent homogenization temperatures (220° to380 °C). The temperature of homogenization of V1 and V2 fluidinclusions vary from 270° to 390 °C. Salinities of L1 and L2 inclusionsvary from 1.7 to 19.0 wt.% NaCl equiv. Salinities of V1 inclusions varyfrom 3.7 to 12.08 wt.% NaCl equiv.

In summary, fluid inclusions in vein quartz form stage 1B homoge-nize between 180 °C and 420 °C, with a bimodal distribution of 180 °Cto 260 °C and 260° to 420 °C. It indicates that the ore-forming fluids atBaogutu could have two different sources (magmatic fluid with aninflux of a low-salinity fluid of external origin).

Fig. 9. Fluid inclusion types and features from different stages at Baogutu. (a) Inclusions in isalinity inclusions which are interpreted to directly exsolved from the magma; (b) InclusInclusions in B1-veinlet quartz in stage 1B resemble the inclusions in D-quartz; (d) Inclusionhigh-salinity inclusions; (f) Inclusions in C2-vein quartz in stage 1C have the highest Cu gra

6.2.4. Vein quartz in stage 1CFluid inclusions in quartz gangue occurred in C1- and C2-veins

have similar characterizations. Veins are characterized by L2, V1 andV2 inclusions. L1 inclusion homogenize between 170° and 270 °C(Table 3, Fig. 10d). L2 inclusion homogenize between 170° and 390 °C(Table 3), with 85% of fluid inclusions homogenizing at temperatures210–330 °C. V1 inclusions have the high apparent homogenizationtemperatures (180° to 460 °C) and concentrated in between 250° and390 °C (Fig. 10d; Table 3). The temperature of homogenization of V2fluid inclusions (N320 °C) overlaps the homogenization temperaturesof coexisting V1 fluid inclusions (Table 3). Salinities of the L1 overlapthe salinities of coexisting V fluid inclusions and range from 1.3 to23wt.% NaCl equiv (Fig. 10d).

6.3. Results of Raman spectroscopy

In contrast to most porphyry Cu deposits (Butte, Alumbrera,Bingham, Santa Rita, Red Mountain), at Baogutu, clathrate formationwas not observed in any of the inclusions assemblages. Raman spec-troscopy is used in this study in order to detected gaseous com-position of fluid inclusions. The results of more than 70 individualfluid inclusions from one assemblage (I-quartz) formed in latemagmatic stage and three assemblages (D-quartz, B- and C-veinquartzes) formed in hydrothermal stage are plotted in Fig. 11. I-quartz

nterstitial quartz (I-quartz) in late magmatic stage are characterized by the most high-ions in disseminated quartz (D-quartz) in stage 1B contain all types in Baogutu; (c)s in B2-veinlet quartz in stage 1B; (e) Inclusions in C1-vein quartz in stage 1C without

de (N1%) and are characterized by the most vapor inclusions representing local boiling.

Fig. 11. Laser Raman spectra of fluid inclusions of the Baogutu deposit.


fluid inclusions include H2O-rich (Fig. 11a) and CH4-rich with micro-CO2 (Fig. 11b). So, fluids brought from late magmatic stage bythe quartz in diorite are CH4–H2O type. D-quartz fluid inclusions instage 1B include H2O-rich (Fig. 11c), CH4-rich (Fig. 11d), CH4 + H2O(Fig. 11e), and H2O + CH4 (Fig. 11f) with notable absence of CO2.Vein-quartz fluid inclusions in stage 1B have similar gaseous com-position which include CH4 + H2O (Fig. 11g) and CH4-rich (Fig. 11h),without CO2. Vein-quartz fluid inclusions in stage 1C are characterizedby CH4 + H2O (Fig. 11i) and CH4 + CO2 (Fig. 11j). CO2 with a similarabundance as CH4 is detected only in a few analyses.

In summary, fluids carried by these quartzes in latemagmatic stageand hydrothermal stage 1B are rich in reductive gaseous species CH4

and H2O, with absence of CO2; while fluids carried by these quartzes inhydrothermal stage 1C are rich in CH4 and H2O plus minor CO2.

6.4. Results of quadrupole mass spectrometer

In order to further determine gaseous composition and theircontent of fluid inclusions, Quadrupole Mass Spectrometer is used inthis study. The analytical results for four respective samples formedduring from stage 1B to stage 1C are listed in Table 4. This result showthat the fluid inclusions in stage 1B contain H2O (82.12%), CH4

(12.89%), CO2 (3.49%) and minor N2 (1.14%), C2H6 (0.22%), and H2S(0.04%), indicating that fluid is rich in CH4 and H2O plus minor CO2.However, the fluid inclusions in stage 1C contain H2O (75.73–77.53%),CH4 (8.041–11.46%), CO2 (10.43–11.29%) and minor N2 (1.89–2.54%),C2H6 (0.21–0.27%), and H2S (0.006–0.009%) indicating that fluid isrich in CH4, H2O and CO2. The fluid inclusions (ZK102-272) in stage 1Boverprinted by stage 1C have gaseous composition and content which

Table 4Gaseous composition (mol%) of fluid inclusions from quadrupole mass spectrometry.

Sample no. Veins Stage H2O N2 He Ar* O2 CO2 CH4 C2H6 H2S

ZK102-458 B1-veinlet Stage 1B 82.12 1.142 − 0.135 – 3.493 12.89 0.22 0.041ZK102-272 B1-veinlet Stage 1B+1C 81.84 1.612 – 0.274 – 5.875 9.951 0.439 0.076ZK211-509 C1-vein Stage 1C 75.73 1.896 – 0.258 – 10.43 11.46 0.215 0.009ZK211-526 C1-vein Stage 1C 77.53 2.548 – 0.306 – 11.29 8.041 0.278 0.006

– not detected; * values for reference.


vary between stage 1B and stage 1C. Therefore, fluids that are broughtby the Baogutu diorite are rich in reductive volatiles CH4 and H2O instage 1B and change gradually from reductive volatiles to oxidativevolatiles from stage 1B to stage 1C.

7. Discussion

7.1. Pressure estimation and deep magmatic system

Pressure estimation from fluid inclusions can be obtained by fluiddata in the binary system NaCl–H2O as an approximation (Hedenquistet al., 1998; Driesner andHeinrich, 2002, 2007) and outlined in Table 3.Pressures determined for nonboiling assemblages are derived from thehomogenization temperature and representminimum values (Rusk etal., 2008). Boiling assemblages will give absolute fluid entrapmenttemperatures (Roedder, 1984), although compositional deviations ofthe fluids from the binary model system introduce systematicuncertainties. At Baogutu, the lack of the significance boiling assem-blages shows that pressures determined only represent the minimumvalues.

Almost brine inclusions in quartz at Baogutu consistently showhomogenization by halite dissolution after bubble disappearance. Suchhomogenization behavior is relatively common in halite saturatedinclusions in magmatic–hydrothermal ore deposits (Ulrich et al., 2001;Bouzari and Clark, 2006). They have been interpreted to indicate en-trapment of a halite-saturated hydrothermal brine (Cloke and Kesler,1979; Wilson et al., 1980), homogeneous trapping at high pressures(Bodnar, 1994; Rusk et al., 2008), or post-entrapment H2O loss orvolume shrinkage (Sterner et al., 1988).Halite-homogenizing inclusionsat Baogutu do not occur widely on boiling trails, which would be clearevidence for post-entrapment modification. The analyzed fluid inclu-sion assemblages showed nopetrographic evidence of post-entrapmentdisturbance. Moreover, halite homogenization at temperatures wellabove homogenization of the bubble in brine inclusions do not yieldunreasonably high trapping temperatures, suggesting that significantpost-entrapmentmodificationmaynot occur.We therefore suggest thathomogenization by halite dissolution at Baogutu could record theentrapment at high pressures.

The depth of formation of porphyry Cu–Mo deposits is typicallydifficult to determine and in many deposits it is poorly known(Seedorff et al., 2005). Fluid inclusions may provide estimates oftrapping pressure, but the resultant depth estimates are notstraightforward because pressures may be either lithostatic orhydrostatic or somewhere in between (Rusk et al., 2008). In somecases depth can be reliably inferred from geologic relationships. AtBaogutu, we have indirect geologic estimates of depth. Baogutucomplex includes dominant equigranular diorite and weakly porphy-ritic diorite, indicating crystallization formed in depth. Baogutucomplex is emplaced into folded rocks of the Xibeikulasi fold (Fig. 2).The carapace of cogenetic Carboniferous Darbut Volcanics is missingover the Baogutu complex in the Baogutu district (Fig. 3), supportingdeep erosion.

Halite saturated inclusions in I-quartz were trapped at tempera-tures of about 401–550 °C, then pressures of I-quartz formation werecalculated from the homogenization temperature of brine inclusionsand were in the range of 1500 to 3100 Pa (Table 3). Under lithostatic

pressures, where rock density is 3 gm/cm3, I-quartz within Baogutucomplex formed at depths of 5 to 10 km. This depth estimate coincideswith that for the crystallization of the host Baogutu complex. TheBaogutu complex was, however, emplaced prior to porphyry Cu–Mo–Au mineralization at Baogutu, so this estimate yields crystallizationdepths rather than directly mineralization depths.

D-quartz from stage 1B precipitated from moderate-temperaturefluids (200–400 °C), probably during transition from a lithostatic to ahydrostatic regime or a hydrostatic regime. Here, pressures have beencalculated from the homogenization temperature of brine inclusions,yielding values of b320 bars. The estimated depth may be b3.2 km.

Petrographically later inclusions in stage 1B quartz have similarcharacterizes as D-quartz but with minor brine inclusions. Theuncertainty in the values derived from vapor-rich inclusions isgreatest owing to the very poor visibility of the liquid phase. Pressureestimates for V1 inclusions commonly exceed 84 bars, up to 260 bars.Aqueous L1 and L2 Inclusions were trapped at temperatures of about180–390 °C and have an estimated pressure from 10 to 150 bars.Geologic evidence with tension internal symmetry in quartz veinsindicates that fluid pressures of these low-moderate temperatureinclusions (180–390 °C) are in hydrostatic conditions during fluidentrapment; hence, a depth of b2.6 km is estimated.

The internal symmetry or open vugs and the irregular walls instage 1C suggest that they formed in still brittle rocks. Vein quartz instage 1C, therefore, precipitated from low-moderate temperaturefluids (170–400 °C) in a hydrostatic regime. Similarly, estimates forthe vapor-rich inclusions (V1 and V2) yield values of b230 bars,whereas minimum pressures for L1 and L2 aqueous liquids rangebetween 10 and 210 bars. The estimated depth may be b2.3 km.

7.2. Unusual methane-rich magmatic–hydrothermal fluid system and itsevolution

7.2.1. Unusual methane-rich magmatic–hydrothermal fluid systemOur unpublished data for lead isotope composition of pyrites at

Bagutu span a narrow range (206Pb/204Pb=17.86 to 18.50, 207Pb/204Pb=15.44 to 15.60, 208Pb/204Pb=37.49 to 38.28) and plot in themantle evolution curve. Shen and Jin (1993), Song et al. (2007) andour unpublished data for δ34S from the fluid systems at Baogutu closeto zero, indicating that they derived from the mantle source region(Gemmell and Large, 1992). Therefore, fluids at Baogutu are derivedfrom mantle.

In most porphyry Cu–Mo–Au deposit, CO2 has been identified ininclusions, but other gases such as CH4, H2S, and N2 are typically notdetected (Rusk and Reed, 2002). The fluids involved in the latemagmatic diorite are predominance of NaCl brines and CH4. The fluidsinvolved in hydrothermal quartz are CH4-rich in stage 1B and CH4 +CO2 in stage 1C. The fluids at Baogutu belong to a NaCl–H2O–CH4–

(CO2) system.Lithospheric upper mantle is relatively oxidized (Lécuyer and

Ricard, 1999; Catling et al., 2001), with volatile componentsdominated by CO2 (Lowenstern, 2001). CO2-rich fluids equilibratewith peridotites on the depths of less than 130 km or up to 200 km inthe mantle, respectively, whereas CH4-rich fluids equilibrate withmetal–silicatemelts on the depths of greater than 200 km in themuchreduced asthenospheric mantle (Simakov, 1998). CO2 fluid instead of


CH4 fluid is abundant in the upper mantle; it is likely that the parentmagma and inclusion fluids of the Baogutu diorite derive fromthe transition zone of the mantle or asthensphere underneath thePaleozoic orogens of the Western Junggar.

7.2.2. Methane-rich magmatic–hydrothermal fluid system evolutionMagmatic volatiles separating from calc-alkaline magmas pre-

dominantly consist of water and chloride salts (Heinrich, 2005), andmagmatic fluids are therefore commonly discussed with reference tothe experimentally well-studied phase relations in the binary systemNaCl–H2O (Hedenquist et al., 1998; Driesner and Heinrich, 2002,2007), even though quantitative pressure–temperature–composition(P–T–X) relationships are shifted by the common presence of othercolatiles (Heinrich, 2005). Our interpretation of the P–T–X fluidevolution is illustrated in Fig. 12 based on the binary system NaCl–H2O, as described by Heinrich (2005).

The fluid inclusions in I-quartz found in diorite at Baogutu arecharacterized as CH4-rich, H2O-rich and high-saline inclusionsand they belong to a NaCl–H2O–CH4 system by the results of RamanSpectroscopy (Fig. 11). Almost hyper-saline fluid inclusions homog-enize by halite dissolution suggesting a very high salinity in primaryfluid and homogeneous trapping at high pressures (Bodnar, 1994;Rusk et al., 2008). The significant post-entrapment modificationmay not occur at Baogutu. We therefore suggest that the halite–homogenizing inclusion assemblages could record the entrapment ofa single-phase brine at pressures well above the liquid-vaporphase boundary (Bodnar, 1994). There is a trend from nonboilingbrines with ∼60wt.% NaCl equiv toward lower salinity, essentiallyfollowing a high-temperature isotherm of the two-phase boundary(Fig. 12).

Significant salinity and density variations of fluids present duringpotassic alteration with biotite–quartz–magnetite assemblages andprobably the initial stages of copper deposition can be explained by aP–T path of general cooling and decompression of onemagmatic fluid.In D-quartz occurred in the potassic zone, fluid inclusions contain the

Fig. 12. Phasediagramof the systemNaCl–H2O(Hedenquist et al., 1998;Driesner andHeinrich,represent stage 1B and stage 1C samples, respectively.

dominant aqueous liquids of type L1 and L2 (N60%) and common V1inclusions (b30%) with minor V2 and H inclusions at Baogutu. Thefluid inclusions consist of dominant low-salinity (b23 wt.% NaClequiv) and minor high-salinity (N38 wt.% NaCl equiv) magmaticfluids. The fluids belong to a NaCl–H2O–CH4 system (Fig. 11). The lowpressure (up to 320 bar) fluids trapped above 250 °C are interpretedas the earliest inclusion assemblages in hydrothermal stage. Ascentand cooling in the single-phase field would cause this magmatic fluidto first intersect the two-phase surface on the liquid side of the criticalcurve (Fig. 12) and produces a low density vapor. Therefore, D-quartzcontains abundant L2 and V1 inclusions.

In the vein quartz occurred in the stage 1B potassic zone, fluidinclusions have more vapor inclusions and lower salinity and temper-ature than that in D-quartz. Almost quartz-sulfide veins lack alterationenvelopes indicating the chemical equilibrium between fluid andearly-formed potassic altered rocks. It also indicates that thecompositions of the major vapors or liquids do not vary much duringthe hydrothermal evolution of the system during stage 1B and theformation of vein quartz during stage 1B did not modify the early-formed fluid chemistry significantly. The homogenization tempera-ture in the vein quartz is near to that in the D-quartz. Estimate pres-sures show a much wider scatter from 20 to 260 bars which indicatesthatmultiple cycles of sealing of the fracture system caused a repeatedincrease in pressure, probably due to mineral precipitation. The tran-sition between disseminated and vein mineralization in stage 1Binvolves essentially isochemical cooling at pressures under the two-phase surface (Fig. 12).

In the vein quartz occurred in the stage 1C phyllic zone, fluidinclusions contains most abundant vapor inclusions without brineinclusions in Baogutu. The fluid inclusions consist of most abundantlow-salinity (b23 wt.% NaCl equiv) fluids. Most vapor-rich inclusionsresults from the renewed hydraulic fracturing (pressure loss) whichproduced low density vapor. However, the vein quartz in stage 1Cand 1B have the similar ranges of the homogenization temperatures(Fig. 10c, d) and relatively clear pressure variations (Fig. 12). It shows

2002, 2007), showing all assemblages and single-phase assemblages. Red andblack squares


that the change physical process (mainly pressure loss) contribute tothe transition from potassic to sericitic alteration.

Quartz-sulfide veins formed in stage 1C cut the potassic zone andhave phyllic alteration envelopes. It indicates that there are a chemicalunequilibrium between fluid in stage 1C and early-formed potassicaltered rocks in stage 1B. The fluid inclusions lack high-salinitymagmatic fluids. In addition, the fluid in stage 1C contains CH4 andH2O and detected CO2 which could belong to a NaCl–H2O–CH4–CO2

system. Therefore, the transition between vein mineralization instages 1B and 1C involves different chemical process and lowerpressures at under the two-phase surface. This part of the fluidevolution is associated with feldspar-destructive alteration.

The changes in alteration style from potassic to phyllic can beexplained by transfer of CO2 to the vapor. It is supported by the factthat oxidation of CH4 to CO2 and H2O by chloritization and muscoviteof biotite in stage 1C. The early-formedpotassic alteration assemblageshave been overprinted by a pervasive phyllic alteration (sericite–quartz–pyrite–chlorite–carbonate) assemblage. Where biotite hasbeen sericitized and partly chloritized and plagioclase has been alteredto sericite (fine-grained muscovite), calcite, and chlorite. Theoxidation of methane into carbon dioxide and water liberates 8 molof e− per mole of methane. These electrons are captured during ironreduction related to chloritization of biotite (Tarantola et al., 2007,2009):

CH4 þ 2O2➝CO2 þ 2H2O ð1aÞ

12Fe2O3ðbiotiteÞ➝8Fe3O4ðchloriteÞ þ 2O2 ð1bÞ

CH4 þ 12Fe2O3ðbiotiteÞ➝CO2 þ 8Fe3O4ðchloriteÞ þ 2H2O ð1cÞ

The hydrothermal evolution at Baogutu is characterized by thechemical reaction of a potassic to a phyllic alteration due to a fluidcomposition evolution. Pressure variations exceeding the lithostaticto hydrostatic difference are documented in some porphyry-type oredeposits (e.g., Bajo de la Alumbrera: Ulrich et al., 2002; El Teniente:Klemm et al., 2007). In contrast to it, the small pressure fluctuationshave been indicated at Baogutu, suggesting that the fracture systemcaused a change in pressure weakly developed, which can caused theunobvious phase separation or boiling. It coincide with the geologicalfacts that dominant dissemination and minor vein mineralization atBaogutu.

7.3. Evolution model of methane-rich magmatic–hydrothermal system

The conclusions drawn in the following sections are based on alimited fluid inclusion database and could change if more data werecollected. A plausible model can be constructed from a consideration ofthe phase relationships in the H2O–NaCl system (Fig. 12). The inter-pretation offluid evolution in Fig. 13 is only onepossible scenario guidedby average trapping temperatures and salinities observed in the fluidinclusions as well as the characters of the alteration and mineralizationat Baogutu. Temporal evolution of the Baogutu porphyry Cu–Mo–Audeposit is illustrated in Fig. 13.

7.3.1. Baogutu stock emplacementIn the Late Paleozoic, oceanic crust of the Junggar terrain was

subducted beneath the Altai orogen and Kazakhstan plate (Chen andArakawa, 2005; Xiao et al., 2008, 2009). Volatiles, copper and othermetals were extracted from the hydrated tholeiitic basalt and pelagicsediments of the subducting oceanic crust. This metal- and volatile-rich melt then ascended into the overlying mantle wedge, inducingpartial melting that generated an intermediate magma. The dioriticmagma ascended from the mantle wedge and confining pressuresdecreased rapidly (e.g., Mungall, 2002; Sillitoe, 2010). This allowedvolatiles and metals to be released and transported, with minor

crustal interaction and assimilation (Shen et al., 2009), to the uppercrust. Dioritic magma accumulated in upper crustal magma chambersbeneath the site now occupied by the Baogutu intrusive complex. TheBaogutu complex (temperatures N400 °C) occupies a structuralintersection of N- and ENE-trending faults within the Xibeikulasisynclinorum at depths of 5 to 10 km (Figs. 2 and 13a).

7.3.2. Early K silicate alteration and formation of the dissemination andveins

Porphyry Cu (Mo–Au) deposits form in the upper crust, where fluidoverpressures which generated in the cupola of a crystallizing magmahydrofracture the overlying rock (Burnham, 1979), allowing magmaand fluids to intrude, forming porphyry dikes surrounded by stockworkfractures (Rusk et al., 2008; Hou and Cook, 2009). Fluids permeate thefractured rock forming veins which led to the dominant vein-stylemineralization developed in most porphyry Cu (Mo–Au) deposits.Therefore, sulfides in most porphyry Cu (Mo–Au) deposits occurprimarily in veinlets or breccia pipes (e.g. Kay and Mpodozis, 2001;Richards et al., 2001; Richards, 2005; Cooke et al., 2005; Davies et al.,2008). Baogutu differs in that Cu–Fe sulfide mineralization formedmostly as disseminations in the biotite-altered rocks, with lesseramounts in vein stockworks and hydrothermal breccias. It resultsfrom theweak tockwork fracturing and hydrothermal brecciationwhenhydrothermal fluid produced concentric potassic and propylitic alter-ation zones (Fig. 13b).

At the site of formation of the biotite–magnetite–quartz–sulfidesin dissemination, hydrothermal fluids had temperatures of b400 °Cand pressures of b320 bars, and consisted of hypersaline brine andlow-saline fluid. Fluid decompression can account for the presence ofbrine inclusions that homogenize by halite dissolution in quartzsamples from the D-quartz in stage 1B. These D-quartzes did notcontain brine inclusions that homogenize by disappearance of thevapor bubble, which account for without phase separation when therock fractured. Biotite alteration and propylite alteration in stage 1Bformed when the magmatic–hydrothermal brines interacted with thediorite and wall rocks of the Baogutu Group, respectively (Fig. 13b).

The highest concentrations of chalcopyrite occur in quartz veins instage 1B within the Baogutu diorites based on our logging cores. Fluidinclusion microthermometry suggests that chalcopyrite was deposit-ed with the quartz at temperatures between 180° and 420 °C andpressures b260 bars. These fluids were in similar temperatures andslightly lower pressure than those that caused the biotite–magnetite–quartz–sulfides alteration. Thus chalcopyrite-bearing veins mostlikely formed as the Baogutu diorite cooled and crystallized anddenuded (Fig. 13c).

7.3.3. Phyllic alteration and formation of the veinsIntense phyllic alteration overprinted the potassic alteration zone

and produced additional Cu–Mo ore at Baogutu. The highest concen-trations of molybdenite occur in quartz veins within the Baogutudiorite. Fluid inclusion microthermometry suggests that molybdenitewas deposited with the quartz at temperatures between 160° and400 °C. These fluids have lower pressure (b230 bars) than that thequartz-sulfide veins and alteration (b260 bars) in stage 1B. Thusmolybdenite-bearing veins most likely formed as the Baogutu dioritedenuded continually (Fig. 13d).

8. Conclusions

Detailed fluid inclusion petrography and microthermometry fol-lowed by Raman Spectroscopy analyses and Quadrupole MassSpectrometer analyses allowed construction of the evolution of themethane-rich magmatic–hydrothermal fluids at Baogutu.

The fluids in latemagmatic stage and stage 1B rich in reductive CH4

plus H2O, while the fluids in stage 1C rich in reductive CH4 plus H2O

Fig. 13. Schematic diagram depicting the evolution of the magmatic–hydrothermal system at Baogutu. (a) In this model, a large pluton at depth develops a volatile-charged dioritecupola at its top. At depth, where the country rock is hottest, little cooling of fluid occurs and lithostatic pressure dominates. (b) The concentric potassic and propylitic alterationzone, indicating the initial hydrothermal system. At slightly depth, where the country rock lacks significant fracturing and disseminated mineralization well developed. All fluidinclusions are trapped in disseminated quartz under these conditions. (c) B1-, B2-veinlets and B3-veins prograde upward forming from fluids cooling and depressurizing andoverprint in early disseminatedmineralization. Rates of fluid cooling and depressurization were greater than in earlymineralization. (d) C1 and C2 veins form from fluids cooling anddepressurizing and cutting all earlier vein types. Rates of fluid depressurization relative to cooling were greater than in stage 1B mineralization.


and CO2. Oxidization of the CH4 to CO2 characterizes the transitionfrom stage 1B (potassic alteration) to stage 1C (phyllic alteration).

The Baogutu dioritic magma emplaced at unusually depth (5–10 km)and the estimated magmatic emplacping conditions are at TN400 °C andP=1500 to 3100 bar. The dominant disseminated mineralization andminor vein mineralization indicates small temperature and pressurefluctuations from stage 1B to stage 1C. The estimated fluid trapping

conditions are at T=200 to 400 °C and P=50 to 320 bars in D-quartz instage 1B, T=180 to 400 °C and P=20 to 263 bars in vein quartz in stage1B, andT=170to400 °CandP=10to230 bars inveinquartz in stage1C.

The overall fluid path at Baogutu is toward lower pressure andsmall changed temperature, with some composition transit of fluidfrom a NaCl–H2O–CH4 system to NaCl–H2O–CH4–CO2 system and ledto the formation of Cu–Mo–Au from stage 1B to stage 1C.


Acknowledgments

Thanks are given to Prof. Hongrui Fan and Zhenhao Duan forthoughtful discussions. We express our gratitude to Editor Prof. DonaldB. Dingwell and two anonymous reviewers for their critical reviews andexcellent suggestions on improvement of the manuscript. We would liketo thank themembers of the Institute ofGeology, XinjiangGeoexplorationBureau for Non-ferrous Metals, especially Zhang Rui, Zhang Yunxiao, andWang Jiang, for providing access to the mine and core. This paper wasfinancially supported by theNational Science Fund (GrantNo. 40972064),Innovative Project of the Chinese Academy of Sciences (Grant No. KZCX2-YW-107), and National 305 Project (Grant No. 2006BAB07B01-01).

References

Bodnar, R.J., 1994. Synthetic fluid inclusions: XII. The system H2O–NaCl. Experimentaldetermination of the halite liquidus and isochores for a 40 wt.% NaCl solution.Geochimica et Cosmochimica Acta 58, 1053–1063.

Bouzari, F., Clark, A.H., 2006. Prograde evolution and geothermal affinities of a majorporphyry copper deposit: the Cerro Colorado hypogene protore, I Región, northernChile. Economic Geology 101, 95−1.

Burke, E.A., 2001. Raman microspectrometry of fluid inclusions. Lithos 55, 139–158.Burnham, C.W., 1979. Magmas and hydrothermal fluids, In: Barnes, H.L. (Ed.),

Geochemistry of hydrothermal ore deposits, 2nd edition. John Wiley and Sons, NewYork, pp. 71–136.

Catling, D.C., Zahnle, K.J., McKay, C.P., 2001. Biogenic methane, hydrogen escape, andthe irreversible oxidation of early Earth. Science 293, 839–843.

Charlou, J.L., Donval, J.P., Fouquet, Y., Jean-Baptiste, P., Holm, N., 2002. Geochemistry ofhigh H2 and CH4 vent fluids issuing from ultramafic rocks at the Rainbowhydrothermal field (36°14′N, MAR). Chemical Geology 191, 345–359.

Chen, B., Arakawa, Y., 2005. Elemental and Nd–Sr isotopic geochemistry of granitoidsfrom the West Junggar foldbelt (NW China), with implications for Phanerozoiccontinental growth. Geochimica et Cosmochimica Acta 69, 1307–1320.

Chen, B., Jahn, B-m, 2004. Genesis of post-collisional granitoids and basement nature ofthe Junggar Terrane, NW China: Nd–Sr isotope and trace element evidence. Journalof Asian Earth Sciences 23, 691–703.

Cheng, Y., Zhang, R., 2006. Mineralization regularity of Cu–Au deposits in the Baogutuarea, western Junggar, Xinjiang. Geology and Prospecting 42, 11–15 (in Chinesewith English abstract).

Cloke, P.L., Kesler, S.E., 1979. The halite trend in hydrothermal solutions. EconomicGeology 74, 1823–1831.

Cooke, D.R., Wilson, A.J., Davies, A.G.S., 2004. Characteristics and genesis of porphyrycopper-gold deposits: University of Tasmania. Centre for Ore Deposit ResearchSpecial Publication 5, 17–34.

Cooke, D.R., Hollings, P., Walshe, J.L., 2005. Giant porphyry deposits: characteristics,distribution, and tectonic controls. Economic Geology 100, 801–818.

Davies, A.G.S., Cooke, D.R., Gemmell, J.B., van Leeuwen, T., Cesare, P., Hartshorn, G., 2008.Hydrothermal breccias and veins at the Kelian gold mine, Kalimantan, Indonesia:genesis of a large epithermal gold deposit. Economic Geology 103, 717–757.

Driesner, T., Heinrich, C.A., 2002. Revised critical curve for the system H2O–NaCl 12thAnnual Goldschmidt Conference, Davos. Geochimica et Cosmochimica Acta A196.

Driesner, T., Heinrich, C.A., 2007. The system H2O–NaCl. Part I. Correlation formulae forphase relations in temperature-pressure-composition space from 0 to 1000 °C, 0 to5000 bar, and 0 to 1 XNaCl. Geochimica et Cosmochimica Acta 4880–4901.

Dubessy, J., Buschaert, S., Lamb, W., Pironon, J., Thie´ry, R., 2001. Methane-bearingaqueous fluid inclusions: raman analysis, thermodynamic modeling and applica-tion to petroleum basins. Chemical Geology 173, 193–205.

Fan, H.R., Xie, Y.H., Wang, K.Y., Tao, K.J., Wilde, S.A., 2004. Methane-rich fluid inclusions inskarn near the giant REE–Nb–Fe deposit at Bayan Obo, Northern China. Ore GeologyReviews 25, 301–309.

Gemmell, J.B., Large, R.R., 1992. Stringer system and alteration zones underlying theHellyer volcanogenic massive sulfide deposit, Tasmania, Australia. EconomicGeology 87, 620–649.

Guo, L.S., Liu, Y.L., Wang, Z.H., Song, D., Xu, J.F., Su, L., 2010. The zircon U-Pb LA-ICP-MSgeochronology of volcanic rocks in DBaogutu areas, western Junggar. ActaPetrologica Sinica 26, 471–477 (in Chinese with English abstract).

Han, B.F., Ji, J.Q., Song, B., Chen, L.H., Zhang, L., 2006. Late Paleozoic vertical growth ofcontinental crust around the Junggar Basin, Xinjiang, China (Part II): timing of post-collisional plutonism. Acta Petrologica Sinica 22, 1077–1086 (in Chinese withEnglish abstract).

Hedenquist, J.W., Arriba, A., Reynolds, T.J., 1998. Evolution of intrusion-centeredhydrothermal system: far Southeast–LepantoporphyryandepithermalCu–Audeposits,Philippines. Economic Geology 93, 373–404.

Heinhorst, J., Lehmann, B., Ermolov, P., 2000. Paleozoic crustal growth and metallogenyof Central Asia: evidence from magmatichydrothermal ore systems of CentralKazakhstan:Tectonophysics 328, 6–87.

Heinrich, C.A., 2005. The physical and chemical evolution of low-salinity magmaticfluids at the porphyry to epithermal transition: a thermodynamic study.Mineralium Deposita 39, 864–889.

Hou, Z.Q., Cook, N.J., 2009. Metallogenesis of the Tibetan collisional orogen: a reviewand introduction to the special issue. Ore Geology Reviews 36, 2–24.

Hurai, H., Kihle, J., Kotulova, J., Marko, F.S., Wierczewska, A., 2002. Origin of methane inquartz crystals from the Tertiary accretionary wedge and fore-arc basin of theWestern Carpathians. Applied Geochemistry 17, 1259–1271.

Kay, S.M., Mpodozis, C., 2001. Central Andean ore deposits linked to evolving shallowsubduction systems and thickening crust. GSA Today 11, 4–9.

Klemm, L.M., Pettke, T., Heinrich, C.A., Campos, E., 2007. Hydrothermal evolution of theEl Teniente deposit, Chile: porphyry Cu–Mo ore deposition from low-salinitymagmatic fluids. Economic Geology 102, 1021–1045.

Kudryavtsev, Y.K., 1996. The Cu–Mo deposits of Central Kazakhstan. In: Shatov, V.,Seltmann, R., Kremenetsky, A., Lehmann, B., Popov, V., Ermolov, P. (Eds.), Granite-related ore deposits of Central Kazakhstan and adiacent Areas. Glagol PublishingHouse, St. Petersburg. 119–144 pp.

Landtwing, M.R., Pettke, T., Halter, W.E., Heinrich, C.A., Redmond, P.B., Einaudi, M.T.,Kunze, K., 2005. Copper deposition during quartz dissolution by cooling magmatic–hydrothermal fluids: the Bingham porphyry. Earth and Planetary Science Letters235, 229–243.

Landtwing, M.R., Furrer, C., Redmond, P.B., Pettke, T., Guillong, M., Heinrich, C.A., 2010.The Bingham Canyon porphyry Cu–Mo–Au deposit. III. Zoned copper–gold oredeposition by magmatic vapor expansion. Economic Geology 105, 91–118.

Lécuyer, C., Ricard, Y., 1999. Long-term fluxes and budget of ferric iron: implication forthe redox states of the Earth's mantle and atmosphere. Earth and Planetary ScienceLetters 165, 197–211.

Li, H.Q., Chen, F.W., 2004. Isotopic geochronology of regional mineralization in Xinjiang,China.Geologicalpublishinghouse,Beijing. 19–42pp. (inChinesewithEnglishabstract).

Liu, W., Pan, X.F., 2006. Methane-rich fluid inclusions from ophiolitic dunite and post-collisionalmafic–ultramafic intrusion: themantle dynamics underneath the Palaeo-Asian Ocean through to the post-collisional period. Earth and Planetary ScienceLetters 242, 286–301.

Lowenstern, J.B., 2001. Carbon dioxide in magmas and implications for hydrothermalsystems. Mineralium Deposita 36, 490–502.

Lu, H.Z., Fan, H.R., Ni, P., Ou, G.X., Shen, K., Zhang, W.H., 2004. Fluid Inclusion. SciencePublishing House, Beijing, pp. 345–364 (in Chinese).

Mungall, J.E., 2002. Roasting the mantle: slab melting and the genesis of major Au andAu-rich Cu deposits. Geology 30, 915–918.

Nash, T.J., 1976. Fluid-inclusion petrology—data from porphyry copper deposits andapplications to exploration.U.S. Geological Survey Professional Paper 907-D,D1–D16.

Nie, F.J., Jiang, S.H., Zhang, Y., Liu, Y., Hu, P., 2004. Geological features and origin ofporphyry copper deposits in China–Mongolia border region and tis neighboringareas. Mineral Deposits 23, 176–189 (in Chinese with English abstract).

Qin, K.Z., Sun, S., Fang, T.H., Wang, S.L., Liu, W., 2002. Paleozoic epithermal Au andporphyry Cu deposits in North Xinjiang, China: epochs, features, tectonic linkageand exploration significance. Resource Geology 52, 291–300.

Redmond, P.B., Einaudi, M.T., Inan, E.E., Landtwing, M.R., Heinrich, C.A., 2004. Copperdeposition by fluid cooling in intrusion-centered systems: new insights from theBingham porphyry ore deposit, Utah. Geology 32, 217–220.

Richards, J.P., 2005. Cumulative factors in the generation of giant calc-alkaline porphyryCu deposits. In: Porter, T.M. (Ed.), Super porphyry copper & gold deposits, 1. PGCPublishing. 7–25 pp.

Richards, J.P., Boyce, A.J., Pringle, M.S., 2001. Geological evolution of the Escondidaarea, northern Chile: a model for spatial and temporal localization of porphyry Cumineralization. Economic Geology 96, 271–305.

Roedder, E., 1984. Fluid inclusions. Reviews in Mineralogy 12, 646.Rui, Z.Y., Huang, C.K., Qi, G.M., Xu, Y., Zhang, H.T., 1984. Porphyry copper (molybdenum)

deposits of China. Geological Publishing House, Beijing. 1–350 pp. (in Chinese).Rui, Z.Y., Wang, L.S., Wang, Y.T., Liu, Y.L., 2002. Discussion on metallogenic epoch of

Tuwu and Yandong porphyry copper deposits in East Tianshan Mountains,Xinjiang. Mineral Deposits 21, 16–22 (in Chinese with English abstract).

Rusk, B., Reed, M., 2002. Scanning electron microscope-cathodoluminescence analysisof quartz reveals complex growth histories in veins from the Butte porphyry copperdeposit, Montana. Geology 30, 727–730.

Rusk, B., Reed, M., Dilles, J., 2008. Fluid inclusion evidence for magmatic–hydrothermalfluid evolution in the porphyry copper–molybdenum deposit at Butte, Montana.Economic Geology 103 (2), 307–334.

Seedorff, E., Dilles, J.H., Proffett Jr., J.M., Einaudi, M.T., Zurcher, L., Stavast, W.J.A.,Johnson, D.A., Barton, M.D., 2005. Porphyry deposits: characteristics and origin ofhypogene features. Economic Geology 100, 251–298.

Shen, Y.C., Jin, C.W. (Eds.), 1993. Magmatism and Gold Mineralization in WesternJunggar. Beijing Science Press. 113–172 pp. (in Chinese with English abstract).

Shen,Y.C., Jin,C.W., Qi, J.Y., 1993.Ore-formingmodel andmechanismof goldmineralizationarea inwest Junggar. In: Guangchi, Tu (Ed.), Newdevelopments of Solid Earth Sciencesof Northern Xinjiang. Science Press, Beijing. 137–150 pp. (in Chinese).

Shen, Y.C., Xiao, Z., Liu, H., 1996. The characteristics of the Altay–Junggar goldmineralization belt. In: Jin, C. (Ed.), Geology of Main Gold Metallogenic Belts inNorthern Part of China. Seismological Press, Beijing. 7–31 pp. (in Chinese withEnglish abstract).

Shen, P., Shen, Y.C., Liu, T.B., Li, G.M., Zeng, Q.D., 2007. Genesis of volcanic-hosted golddeposits in the Sawur gold belt, northern Xinjiang, China: evidence from REE, stableisotopes, and noble gas isotopes. Ore Geology Review 32, 207–226.

Shen, P., Shen, Y.C., Liu, T.B., Lu, J.J., Wei, J.P., Song, G.X., Meng, L., 2008. Late Paleozoicgold and copper mineralization and tectonic evolution in Northwestern Xinjiang,China. Acta Petrologica Sinica 24, 1087–1100 (in Chinese with English abstract).

Shen, P., Shen, Y.C., Liu, T.B., Meng, L., Dai, H.W., Yang, Y.H., 2009. Geochemical signatureof porphyries in the Baogutu porphyry copper belt, western Junggar, NW China.Gondwana Research 16, 227–242.

Shepherd, T.J., Rankin, A.H., Alderton, D.H.M., 1985. A practical guide to fluid inclusionstudies. Blackie, London. 239 pp.


Sillitoe, R.H., 2010. Porphyry copper systems. Economic Geology 105, 3–41.Simakov, S.K., 1998. Redox state of Earth's upper mantle peridotites under the ancient

cratons and its connection with diamond genesis. Geochimica et CosmochimicaActa 62 (10), 1811–1820.

Song, H.X., Liu, Y.L., Qu, W.J., Song, B., Zhang, R., Cheng, Y., 2007. Geological characters ofBaogutu porphyry copper deposit in Xinjiang, NWChina. Acta Petrologica Sinica 23,1891–1988 (in Chinese with English abstract).

Sterner, S.M., Hall, D.L., Bodnar, R.J., 1988. Synthetic fluid inclusions. V: Solubilityrelations in the system NaCl-KCl-H2O under vapor-saturated conditions. Geochi-mica et Cosmochimica Acta 52, 989–1005.

Su, Y.P., Tang, H.F., Hou, G.S., 2006. Geochemistry of aluminous A-type granites alongDarabut tectonic belt in West Junggar, Xinjiang. Geochimica 35, 1–5 (in Chinesewith English abstract).

Tang, G.J., Wang, Q., Zhao, Z.H., Derek, A.W., Chen, H.H., Jia, X.H., Jiang, Z.Q., 2009.Geochronology and Geochemistry of the Ore-Bearing Porphyries in Baogutu Area(Western Junggar): Petrogenesis and their implications for tectonics and Cu-Aumineralization. Earth Science Journal of China University of Geosciences 34, 56–74(in Chinese with English abstract).

Tarantola, A., Mullis, J., Vennemann, T., Dubessy, J., de Capitani, C., 2007. Oxidation ofmethane at the CH4 / H2O–(CO2) transition zone in the external part of the CentralAlps, Switzerland: evidences from stable isotope investigations. Chemical Geology237, 329–357.

Tarantola, A., Mullis, J., Guillaume, D., Dubessy, J., de Capitani, C., Abdelmoula, M., 2009.Oxidation of CH4 to CO2 and H2O by chloritization of detrital biotite at 270±5 °Cin the external part of the Central Alps, Switzerland. Lithos 112, 497–510.

Ulrich, T., Günther, D., Heinrich, C.A., 2001. The evolution of a porphyry Cu–Au depositbased on LA-ICP-MS analyses of fluid inclusions. Bajo de la Alumbrera, Argentina.Economic Geology 96, 1743–1774.

Ulrich, T., Günther, D., Heinrich, C.A., 2002. Evolution of a porphyry Cu–Au deposit,based on LA-ICP-MS analysis of fluid inclusions: Bajo de la Alumbrera, Argentina.Economic Geology 97 (8), 1863–1920.

Wang, J.B., Xu, X., 2006. Postcollisional tectonic evolution and metallogenesis inNorthern Xinjiang, China. Acta Geologica Sinica 80, 23–31 (in Chinese with Englishabstract).

Wang, R., Zhu, Y.F., 2007. Geology of the Baobei gold deposit in Western Juggar andzircon SHRIMP age of its wall-rocks, Western Junggar (Xinjiang, NW China).Geological Journal of China Universities 13, 590–602.

Wang, J.B., Wang, Y.W.,Wang, L.J., 2004. The Junggar immature continental crust provinceand its mineralization. Acta Geologica Sinica (English edition) 78, 337–344.

Wilkinson, J.J., 2001. Fluid inclusions in hydrothermal ore deposits. Lithos 55, 229–272.Wilson Jr., J.W., Kesler, S.E., Cloke, P.L., Kelly, W.C., 1980. Fluid inclusion geochemistry of

the Granisle and Bell porphyry copper deposits, British Columbia. EconomicGeology 75, 45–61.

Wilson, A.J., Cooke, D.R., Harper, B.L., 2003. The Ridgeway gold-copper deposit: a high-grade alkalic porphyry deposit in the Lachlan fold belt, New SouthWales, Australia.Economic Geology 98, 1637–1666.

Xiao, W.J., Han, C.M., Yuan, C., Sun, M., Shoufa, L., Chen, H.L., Li, Z.L., Li, J.L., Sun, S., 2008.Middle Cambrian to Permian subduction-related accretionary orogenesis ofNorthern Xinjiang, NW China: implications for the tectonic evolution of centralAsia. Journal of Asian Earth Sciences 32, 102–117.

Xiao, W.J., Kröner, A., Windley, B.F., 2009. Geodynamic Evolution of Central Asia inthe Paleozoic and Mesozoic. International Journal of Earth Sciences 98, 1185–1188.

Xiao, W.J., Huang, B.C., Han, C.M., Sun, S., Li, J.L., 2010. A review of the western partof the Altaids: a key to understanding the architecture of accretionary orogens.Gondwana Research. doi:10.1016/j.gr.2010.01.007.

Zhang, R., Zhang, Y.X., Dong, G.S., 2005. Geological characteristics and origin of Baogutoucopperdeposit inXinjiang. 2005meetingofChineseScienceand technologyAssociation,Urümgi, pp. 1–6 (in Chinese).

Zhang, L.C., Wan, B., Jiao, X.I., Zhang, R., 2006a. Characteristics and geologicalsignificance of adakitic rocks in copper-bearing porphyry in Baogutu, westernJunggar. Geology in China 33, 626–631 (in Chinese with English abstract).

Zhang, R., Zhang, Y.X., Dong, G.S., Wang, J., Li, L.Q., 2006b. Major breakthrough in copperexploration in the Baogutu porphyry copper deposit, western Junggar, Xinjiang, andits significance. Geology in China 33, 1354–1360 (in Chinese with English abstract).

Zhu, Heping, Wang, Lijuan, 2000. Determining gaseous composition of fluid inclusionswithQuadrupoleMass Spectrometer. Science inChina,D31(7), 586–590 (inChinese).

Zhukov, N.M., Kolesnikov, V.V., Miroshnichenko, L.M., Egembayev, K.M., Pavlova, Z.N.,Bakarasov, E.V., 1997. Copper deposits of Kazakhstan. Reference Book Alma Ata 149(in Russian).

http://dx.doi.org/10.1016/j.gr.2010.01.007

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