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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+CO2composition.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.
+86 10 62010846.
ll rights reserved.
© 2010 Elsevier B.V. All rights reserved.
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.,
mailto:[email protected]://dx.doi.org/10.1016/j.chemgeo.2010.04.016http://www.sciencedirect.com/science/journal/00092541
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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
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80 P. Shen et al. / Chemical Geology 275 (2010) 78–98
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Fig. 2. Geological map of the Baogutu porphyry copper belt in
the West Junggar (modified after Shen and Jin, 1993).
81P. Shen et al. / Chemical Geology 275 (2010) 78–98
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.
82 P. Shen et al. / Chemical Geology 275 (2010) 78–98
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.
83P. Shen et al. / Chemical Geology 275 (2010) 78–98
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
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84 P. Shen et al. / Chemical Geology 275 (2010) 78–98
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.
85P. Shen et al. / Chemical Geology 275 (2010) 78–98
-
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.
86 P. Shen et al. / Chemical Geology 275 (2010) 78–98
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.
-
87P. Shen et al. / Chemical Geology 275 (2010) 78–98
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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
ZK211-330 ZK211 330 D-quartz(stage 1B)
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)
ZK211-245 ZK211 245 D-quartz(stage 1B)
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
ZK211-408 ZK211 408 D-quartz(stage 1B)
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)
Potassic L1 215–295 (6) −7.2∼−8.3 (3) 10.73–12.05 19–20 LV
curve
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
ZK211-330 ZK211 330 B1-veinlets(stage 1B)
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
ZK203-208 ZK203 208 B1-veinlets(stage 1B)
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)
ZK203-260 ZK203 260 B1-veinlets(stage 1B)
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)
ZK203-271 ZK203 271 B1-veinlets(stage 1B)
Potassic L1 210–300 (10) −0.8∼−4.1 (8) 1.4–6.6 18–41 LV
curve
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
ZK211-1 ZK211 373 B2-veinlets(stage 1B)
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
ZK102-530 ZK102 530 B2-veinlets(stage 1B)
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
ZK203-199 ZK203 199 B2-veinlets(stage 1B)
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
88 P. Shen et al. / Chemical Geology 275 (2010) 78–98
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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
ZK211-509 ZK211 509 C1-vein(stage 1C)
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)
ZK211-511 ZK211 511 C1-vein(stage 1C)
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.
89P. Shen et al. / Chemical Geology 275 (2010) 78–98
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.
-
90 P. Shen et al. / Chemical Geology 275 (2010) 78–98
-
Fig. 10. Histograms of homogenization temperatures and
salinities (wt.% NaCl equivalent) from microthermometric data for
all fluid inclusions from different quartzes.
91P. Shen et al. / Chemical Geology 275 (2010) 78–98
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.
92 P. Shen et al. / Chemical Geology 275 (2010) 78–98
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 CH4and 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.
93P. Shen et al. / Chemical Geology 275 (2010) 78–98
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
andRicard, 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
-
94 P. Shen et al. / Chemical Geology 275 (2010) 78–98
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
-
95P. Shen et al. / Chemical Geology 275 (2010) 78–98
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–CO2system. 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
CH4plus 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.
96 P. Shen et al. / Chemical Geology 275 (2010) 78–98
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.
-
97P. Shen et al. / Chemical Geology 275 (2010) 78–98
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).
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