Geological and geochemical studies of the Shujiadian porphyry Cu deposit, Anhui Province, Eastern China: Implications for ore genesis

Journal of Asian Earth Sciences xxx (2014) xxx–xxx

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Geological and geochemical studies of the Shujiadian porphyry Cu deposit, Anhui Province, Eastern China: Implications for ore genesis Shiwei Wang a, Taofa Zhou a,⇑, Feng Yuan a, Yu Fan a, Noel C. White a,b, Fengjie Lin a a b

School of Resources and Environmental Engineering, Hefei University of Technology, Hefei 230009, China Centre of Excellence in Ore Deposits (CODES), University of Tasmania, Private Bag 79, Hobart, Tasmania 7001, Australia

a r t i c l e

i n f o

Article history: Received 17 March 2014 Received in revised form 23 July 2014 Accepted 2 August 2014 Available online xxxx Keywords: Geological features Fluid inclusions Geochemistry Porphyry Cu deposit Intracontinental settings

a b s t r a c t Most porphyry deposits in the world occur in magmatic arc settings and are related to subduction of oceanic plates. A small proportion of porphyry deposits occur in intracontinental settings, however they are still poorly understood. Shujiadian, a newly-discovered porphyry Cu deposit, is located in the Middle– Lower Yangtze River Valley metallogenic belt and belongs to the intracontinental class. The deposit has classic alteration zones defined by a core of potassic alteration and local Ca-silicate alteration, which is overprinted by a feldspar-destructive alteration zone and cut by veins containing epidote and chlorite. Wallrocks of the deposit are unreactive quartz-rich sedimentary rocks. Three main paragenetic stages have been recognized based on petrographic observations; silicate stage, quartz-sulfide stage, and sulfide-carbonate stage. Quartz + pyrite + chalcopyrite ± molybdenite veins, and quartz + chalcopyrite + pyrite veins of the quartz-sulfide stage contribute most of the copper, and chalcopyrite + chlorite ± pyrite ± pyrrhotite ± quartz ± illite veins of the sulfide-carbonate stage also contribute part of the copper; all the mineralized veins are associated with feldspar-destructive alteration. Investigations on the fluid inclusions in Shujiadian indicate that the ore-forming fluids had four evolutionary episodes: immiscibility and overpressure in the silicate stage, boiling in the quartz-sulfide stage and mixing with meteoric water in the sulfide-carbonate stage. Sulfur and strontium isotope studies suggest that ore metals were mainly derived from magmatic–hydrothermal fluids, and combined with our study of fluid inclusions, we infer that decompression, changes in oxygen fugacity and sulfur content were the main factors that caused Cu precipitation. Compared with porphyry deposits in magmatic arc settings, there are some differences in the ore-bearing rock, alteration, and the composition of ore-forming fluids. Ó 2014 Elsevier Ltd. All rights reserved.

1. Introduction More than 97% of the world’s giant porphyry deposits occur in island arcs (e.g., Indonesia and the Philippines) and continental margin arcs (e.g., central Chile and southwestern United States), which are closely related to subduction (Sillitoe, 1972; Mitchell, 1973; Cooke et al., 2005; Hou et al., 2011a; Sun et al., 2010; Mao et al., 2014). However, porphyry deposits can also occur in nonarc settings, such as post-collisional extensional settings (e.g., Yulong porphyry Cu belt, China) and intracontinental settings (e.g., porphyry deposits in the Middle–Lower Yangtze River Valley metallogenic belt, MLYR; Hou et al., 2011a; Chang et al., 1991; Zhai et al., 1992; Pan and Dong, 1999; Mao et al., 2006, 2011; Deng et al., 2011). Much research has focused on the magma source and evolution, emplacement, ore-forming fluid exsolution and ⇑ Corresponding author. Tel.: +86 551 62901525. E-mail address: [email protected] (T. Zhou).

evolution, and alteration mineralogy of porphyry deposits in subduction settings (Sillitoe, 1973, 2010; Gustafson and Hunt, 1975; Titley and Beane, 1981; Dilles and Einaudi, 1992; Hedenquist and Lowenstern, 1994; Ulrich et al., 1999; Ulrich and Heinrich, 2001; Heinrich, 2005; Richards, 2003, 2005; Cooke et al., 2005), whereas there was little work on porphyry deposits in non-arc settings. The MLYR is one of the most important metallogenic belts in China, located at the northern margin of the Yangtze Craton. The MLYR is confined by the Xiangfan-Guangji fault (XGF), Tan-Lu fault (TLF), Yangxin fault (YCF) and hosts a series of ore deposit clusters, such as Edong, Jiurui, Anqing-guichi, Luzong, Tongling, Ningwu and Ningzhen. These ore deposit clusters are located in faulted uplift zones and faulted basins (Fig. 1). There are in total about 200 Cu–Fe polymetallic deposits, including skarn Cu–Fe deposits, porphyry Cu–Au deposits and vein-type hydrothermal Cu–Au deposits (e.g., Chang et al., 1991; Tang et al., 1998; Wu et al., 1999; Xing and Xu, 1996; Ren et al., 1991; Wang and Zhao, 2001; Chu, 2003; Chen et al., 2001; Zhou et al., 2005, 2008, 2011) in the MLYR.

http://dx.doi.org/10.1016/j.jseaes.2014.08.004 1367-9120/Ó 2014 Elsevier Ltd. All rights reserved.

Please cite this article in press as: Wang, S., et al. Geological and geochemical studies of the Shujiadian porphyry Cu deposit, Anhui Province, Eastern China: Implications for ore genesis. Journal of Asian Earth Sciences (2014), http://dx.doi.org/10.1016/j.jseaes.2014.08.004

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S. Wang et al. / Journal of Asian Earth Sciences xxx (2014) xxx–xxx

Fig. 1. Geological map of magmatic rocks and deposits in the Middle–Lower Yangtze River Valley metallogenic belt (modified after Zhou et al., 2008; Mao et al., 2011) TLF: Tancheng-Lujiang fault; XGF: Xiangfan-Guangji fault; YCF: Yangxing-Changzhou fault.

Shujiadian is a recently discovered porphyry Cu deposit in the MLYR (Fig. 1). Its current resource is more than 0.7 Mt Cu, while it is still in the exploration stage. To date, only few studies (Wang, 2010; Wang et al., 2011, 2012; Lai et al., 2012) have been carried out on this deposit, mainly focused on petrology, geochemistry, isotopic geochemistry, and geochronology of the porphyritic rocks related to mineralization. Little previous work has been carried out on the geologic character of the deposit and the evolution of ore-forming fluids in Shujiadian deposit. In this paper, we not only provide a detailed description of the geological features of Shujiadian deposit, but also establish a model for evolution of the ore-forming fluid based on studies of fluid inclusions, trace elements, and sulfur and strontium isotopic data. In addition, we compare Shujiadian deposit with other porphyry deposits in magmatic arc settings to assist better understanding of porphyry deposits in intracontinental settings. 2. Regional geological setting Shujiadian deposit (latitude 30°550 5300 N, longitude 0 00 118°02 34 E) is located in the Tongling region, which hosts many Cu–Au deposits (e.g., Dongguashan Cu–Au deposit, Xishizishan Cu–Au deposit). The Tongling region contains marine clastic sedimentary rocks, carbonates, and evaporites of Silurian to Middle Triassic age, with a depositional gap in the Middle–Late Devonian. Mesozoic sedimentary-volcanic basins are widely distributed on these marine deposits (Fig. 2A; Chang et al., 1991; Pan and Dong, 1999). From the Jurassic to Cretaceous, the Tongling region experienced an intraplate deformational stage, developing NNE folds and multistage faults. More than 70 intrusions and associated Cu–Au deposits were emplaced, and controlled by the E–W-trending Tongling-Nanling deep fault (Fig. 2A; Chang et al., 1991). 3. Deposit geology 3.1. Stratigraphy and structures In the Shujiadian district, there are Silurian to Quaternary strata (Fig. 2B). The Shujiadian anticline is the main structure. The flanks of the anticline consist of Triassic, Permian, Carboniferous, Devonian and Silurian strata, and the core consists of Silurian Fentou

and Gaojiabian group strata (Fig. 2B). In the periphery of the Shujiadian deposit there are NW and N–S oriented faults, possibly post-dating the Shujiadian deposit, and most faults cut the Shujiadian anticline (Fig. 2B). Two intrusions (701 intrusion and Shujiadian intrusion) outcrop in the Shujiadian district, where they intruded into the Fentou formation. The 701 intrusion consists dominantly of quartz diorite porphyry and pyroxene diorite, and the Shujiadian intrusion is mainly composed of granodiorite, pyroxene diorite and minor quartz diorite porphyry. Granodiorite is the earliest magmatic rock with an outcrop area of 0.63 km2, and it is intruded by quartz diorite porphyry and contains no mineralization. Quartz diorite porphyry mainly occurs in the 701 intrusion as scattered outcrops (about 4500 m2), with minor outcrops in the Shujiadian intrusion. Pyroxene diorite, striking 50° has an outcrop area of 0.77 km2 (Wang, 2010); it crosscuts or parallels granodiorite and quartz diorite porphyry and is the ore-bearing rock. Later dykes, including quartz diorite, quartz syenite porphyry, and diabase porphyry (Table 1), crosscut quartz diorite porphyry and pyroxene diorite. 3.2. Alteration and mineralization 3.2.1. Wallrock alteration The wallrock alteration associated with Shujiadian deposit is dominated by potassic and propylitic alteration, with local Ca-silicate alteration and feldspar-destructive alteration (Figs. 3 and 4). 3.2.1.1. Ca-silicate alteration. Ca-silicate alteration is the earliest alteration, and is characterized by development of garnet, diopside, and minor quartz, anhydrite, epidote and carbonate. It occurs as pervasive replacement and veins (Figs. 4A, B and 5A), mainly located in the contact zone between pyroxene diorite and sandstone of the Fentou formation (Fig. 3). Selectively pervasive alteration is characterized by partial replacement of plagioclase or other magmatic minerals by garnet, with minor diopside. Vein garnet is the most obvious Ca-silicate alteration mineral, consisting of garnet ± minor fine-grained quartz. The garnets have zoned texture, and are usually replaced by chlorite and carbonate (Fig. 5A). 3.2.1.2. Potassic alteration. The potassic alteration is generally developed pervasively and/or occurring as veins, veinlets and

Please cite this article in press as: Wang, S., et al. Geological and geochemical studies of the Shujiadian porphyry Cu deposit, Anhui Province, Eastern China: Implications for ore genesis. Journal of Asian Earth Sciences (2014), http://dx.doi.org/10.1016/j.jseaes.2014.08.004

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Fig. 2. Geological sketch map of the Tongling metallogenic district (A) and Shujiadian deposit (B), Anhui Province, Eastern China.

alteration halos (Figs. 4C, D, E, G and I). Potassic alteration, characterized by K-feldspar and biotite, accompanied by abundant quartz and minor magnetite, can be subdivided into K-feldspar and biotite alteration. These two kinds of alteration display no consistent time–space relationship. K-feldspar alteration may have developed at the same time as the Ca-silicate alteration, but in different parts of the deposit. It was earlier than the biotite alteration (Fig. 4B). K-feldspar alteration occurs as pervasive or selectively pervasive replacement (Figs. 4C and 5J), veins, veinlets and alteration halos (Fig. 4D and E). Pervasive replacement is commonly developed in the quartz diorite porphyry and pyroxene diorite, and is characterized by partial replacement of plagioclase by K-feldspar (Fig. 4C). Selectively pervasive alteration is characterized by

plagioclase replaced by K-feldspar on their rims, twinning planes or fractures. Strongly pervasive alteration is characterized by the complete destruction of the original textures of plagioclase and the formation of an assemblage of K feldspar and quartz. The residual plagioclase and hydrothermal K-feldspar are mostly replaced by later sericite, epidote, carbonate and clay minerals (Fig. 4C), which leads to hand specimens appearing red and very soft (Fig. 4D). The alteration halo mode is dominant in Shujiadian deposit; it only occurs at depth in the pyroxene diorite and is characterized by K-feldspar replacing plagioclase along veins. In the third alteration mode, discontinuous hairline K-feldspar veinlets cut plagioclase phenocrysts, and continuous veins filled with

Please cite this article in press as: Wang, S., et al. Geological and geochemical studies of the Shujiadian porphyry Cu deposit, Anhui Province, Eastern China: Implications for ore genesis. Journal of Asian Earth Sciences (2014), http://dx.doi.org/10.1016/j.jseaes.2014.08.004

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Table 1 Petrographic characteristics and relationships of the different porphyry intrusions at Shujiadian. Type of rocks

Crosscutting relationships

Mineral content (%)

Quartz vein density (%)

Alteration

Mineralization

Granodiorite

Intruded by all porphyries

Plagioclase: 36–56 Quartz: 5–15 Hornblende: 2–8 Biotite: 2–5 Matrix: 35–55

1–5

Weak-moderate chlorite ± epidote, feldspar destructive alteration

Barren to weak

Quartz diorite porphyry

Cut by pyroxene diorite

Plagioclase: 50 Quartz: 15 Hornblende: 10 Matrix: 25

1–5

Weak potassic, moderate chlorite ± epidote and feldspar destructive alteration

Barren to weak

Pyroxene diorite

Intruded by quartz diorite, quartz syenite porphyry, diabase porphyry

Plagioclase: 65 Quartz: 10 Hornblende: 15

20–40

Moderate-intensely potassic, propylitic and feldspar destructive alteration, with local Ca-silicate alteration

Strong

Quartz diorite

No timing relationships with later porphyry

Plagioclase: 60 Quartz: 15 Hornblende: 20

1–5

Weak chlorite ± epidote alteration

Barren

Quartz syenite porphyry

Cut by diabase porphyry

Orthoclase: 30 Plagioclase: 10 Quartz: 5 Biotite: 2

No

Intensely feldspar destructive

Barren

Diabase porphyry

Later dyke

Plagioclase: 45 Pyroxene: 40 Calcite amygdale: 10

No

Intensely feldspar destructive

Barren

coarse-grained quartz and K-feldspar and microcrystalline biotite are commonly cut by later epidote and anhydrite (Fig. 4E). Biotite alteration occurs as pervasive or selectively pervasive replacement, veins and veinlets (Fig. 4B and E), mainly within the pyroxene diorite. Hydrothermal biotite in the veins is mediumto coarse-grained (0.2–1.5 mm), and has a yellow–brown color1 with obvious cleavage. Pervasive or selectively pervasive biotite alteration is characterized by the replacement of hornblende phenocrysts and earlier hydrothermal garnet and diopside, and can be easily distinguished from primary biotite by its shreddy texture. Some breccias are located at the contact between pyroxene diorite and quartz diorite porphyry in the deposit (Fig. 4I). Clasts consisting of pyroxene diorite, quartz diorite porphyry with K-feldspar alteration and Fentou formation pelitic siltstone with fine-grained biotite alteration are cemented by pyrite, anhydrite and other hydrothermal minerals (Fig. 4J and K). 3.2.1.3. Propylitic alteration. Propylitic alteration, very well developed, and later than Ca-silicate and potassic alteration (Fig. 4A, C and E), occurs in the pyroxene diorite, quartz diorite porphyry and quartz diorite. It is characterized by an assemblage of actinolite, chlorite, epidote, carbonate and albite, accompanied by a small amount of anhydrite, pyrite and chalcopyrite. Propylitic alteration occurs in two main modes: disseminated or clots replacing hornblende, plagioclase and earlier hydrothermal minerals (e.g., K-feldspar, biotite; Fig. 4F, H and M); continuous vein fillings (Figs. 4E, 7D and E) or discontinuous hairline veinlets (Fig. 5K). In the first mode of alteration, disseminated propylitic alteration occurs throughout pyroxene diorite and locally in quartz diorite porphyry, and clots of alteration occur in fractures in the pyroxene diorite. The second mode of alteration occurs only in the pyroxene diorite, and it consists of quartz + actinolite + pyrite ± chalcopyrite veins (Fig. 5D and E) at depth and epidote + chlorite ± pyrite ± chalcopyrite ± carbonate veins (Fig. 5K) at shallow levels. 3.2.1.4. Feldspar-destructive alteration. Feldspar-destructive (phyllic ± argillic) alteration occurs in all intrusive rocks in Shujiadian 1 For interpretation of color in Fig. 4, the reader is referred to the web version of this article.

deposit, where it typically overlaps earlier alteration (Fig. 3). It consists of quartz, illite, muscovite, chlorite, kaolinite, pyrite, and minor carbonate and gypsum. Feldspar-destructive alteration mainly occurs as pervasive or selective replacements of magmatic minerals and earlier hydrothermal minerals or as halos around earlier planar veins. Weak pervasive or selective replacement is characterized by illite, muscovite and calcite along hairline cracks in large plagioclase phenocrysts, with primary polysynthetic twinning and zoning still visible, and in K-feldspar it occurs as hairline fillings of clay minerals. Where feldspar-destructive alteration is intense, all silicate minerals except quartz are completely destroyed, and the rock becomes white (Fig. 4O). This mainly occurs in fracture zones in the pyroxene diorite. Feldspar minerals are pervasively replaced by fine-grained illite, muscovite and abundant ultra-fine granular quartz, with minor anhydrite, calcite and clay minerals (Fig. 4Q and R), or mainly by illite, muscovite and kaolinite (Fig. 4P). Crystal forms of plagioclase phenocrysts are still preserved (Fig. 4P). Hornblende and hydrothermal biotite are replaced by chlorite, sericite and quartz along rims and cleavage, and cleavages are still visible. Some quartz in the matrix of quartz diorite porphyry was probably resulted from the breakdown of feldspars. Pyrite as disseminated grains in the veins and veinlets is a ubiquitous part of the feldspar-destructive alteration. 3.2.2. Types and characteristics of veins Veins in Shujiadian deposit are diverse and multi-stage. Based on their composition and crosscutting relationships (Figs. 4 and 5), we divide the veins into three stages: Silicate stage: Silicate stage veins are characterized by Ca-silicate and potassic alteration. Garnet ± diopside veins accompanying Ca-silicate alteration, commonly overprinted by later carbonate and chlorite alteration, are irregular to planar and characterized by abundant garnet and minor diopside (Fig. 5A). The major types of veins associated with potassic alteration, from early to late, are K-feldspar ± quartz, magnetite, biotite and quartz ± K-feldspar. K-feldspar ± quartz veins (0.5–20 mm), occurring at depth in the pyroxene diorite, are irregular to planar and characterized by abundant K-feldspar (50–90%) accompanying clay mineral alteration and quartz (10–50%; Fig. 4F). They are commonly overprinted by epidote, chlorite, illite, muscovite, calcite, and cut by later veins,

Please cite this article in press as: Wang, S., et al. Geological and geochemical studies of the Shujiadian porphyry Cu deposit, Anhui Province, Eastern China: Implications for ore genesis. Journal of Asian Earth Sciences (2014), http://dx.doi.org/10.1016/j.jseaes.2014.08.004

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Fig. 3. Alteration and mineralization distribution on section A–B of Shujiadian deposit the location of the section A–B is as shown in Fig. 2B.

which make the K-feldspar pink or white (Fig. 4E). Magnetite veins (0.5–5 mm) are irregular to planar and consist of xenomorphic fine-grained (0.02–2 mm) magnetite (95%) and pyrite (5%), and they are mostly overprinted by later chlorite. Irregular biotite veinlets (Fig. 4I, 0.5–10 mm wide), consisting of brown leaf-like biotite, only occur in the pyroxene diorite and are overprinted by later epidote and/or chlorite. They typically cut plagioclase phenocrysts, but no clear timing relationships have been observed with other veins. Irregular to planar quartz ± K-feldspar veins (Fig. 4D, 0.5–12 mm wide) are the most important type of K-feldspar alteration and consist of quartz (80%) and K-feldspar (20%), and are commonly superposed by later feldspar-destructive alteration with minor pyrite. The quartz ± K-feldspar veins cut K-feldspar ± quartz veins and are truncated by later veins, in which Kfeldspar can be in the veins or in the alteration halos. Quartz-sulfide stage: The veins of this stage typically show faint bleached halos or lack alteration halos; they are characterized by abundant quartz, sulfide (pyrite, chalcopyrite, pyrrhotite) and minor epidote or actinolite, and consist of barren quartz veins, quartz + pyrite + chalcopyrite ± molybdenite, quartz + actinolite + pyrite + pyrrhotite ± chalcopyrite, epidote + chlorite ± q uartz ± anhydrite ± pyrite ± chalcopyrite ± pyrrhotite, qu artz + chalcopyrite + pyrite ± pyrrhotite and quartz + pyrrhotite ± pyrite veins, from early to late. Barren quartz veins (1–25 mm wide) mostly occur in the deep parts of the pyroxene diorite and in the contact zone with quartz diorite, and are irregular to planar and characterized by abundant fine-grained quartz (0.15 mm) or coarse-grained quartz (0.7 mm). Quartz + pyrite + chalcopyrite ± molybdenite veins (2–40 mm wide), occurring mainly in the pyroxene diorite and quartz diorite porphyry, are irregular to planar and characterized by abundant quartz and minor pyrite, chalcopyrite, molybdenite,

and pyrrhotite. Pyrite, chalcopyrite, molybdenite, and pyrrhotite are disseminated (Figs. 4G, 5G and J) or as discontinuous lines (Fig. 5C) in the veins. They cut earlier veins with Ca-silicate and potassic alteration, and are truncated by later veins (Figs. 4G and 5J). Quartz + actinolite + pyrite + pyrrhotite ± chalcopyrite veins (2–20 mm wide) are typically planar and characterized by abundant quartz, actinolite, pyrrhotite, pyrite and minor chalcopyrite (Fig. 5D–F). Under the microscope, we observed that actinolite replaces quartz (Fig. 5E) and pyrite and chalcopyrite replace pyrrhotite (Fig. 5F). No clear timing relationships have been observed with earlier veins, but they are truncated by later veins (Fig. 5D). Epidote + chlorite ± quartz ± anhydrite ± pyrite ± chalcopyrite veins (0.5–2 mm wide), are throughout the pyroxene diorite and less abundant in the quartz diorite porphyry and the contact zone between quartz diorite and pyroxene diorite. They are irregular to planar and characterized by abundant epidote, chlorite and pyrite, occasionally with quartz, anhydrite, pyrite, pyrrhotite and chalcopyrite. They cut earlier veins and are commonly truncated by veins of the sulfide-carbonate stage (Figs. 4E and 5K). Quartz + chalcopyrite + pyrite veins (5–25 mm wide), are planar and characterized by abundant quartz, chalcopyrite, minor pyrite and pyrrhotite. Quartz + pyrrhotite ± pyrite veins (2–25 mm wide), are planar and characterized by abundant quartz, pyrrhotite and lesser pyrite, and cut quartz + chalcopyrite + pyrite veins, but show no clear timing relationships with other veins. Sulfide-carbonate stage: Veins of this stage are associated with feldspar-destructive alteration, occur at shallow depths and consist of chalcopyrite + chlorite ± pyrite ± pyrrhotite ± quartz ± illite, pyrite ± specularite ± quartz ± gypsum ± calcite ± chlorite ± illite and gypsum + calcite ± quartz ± pyrite ± illite veins. Chalcopyrite + pyrite ± pyrrhotite ± quartz ± chlorite ± illite veins (<15 mm

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Fig. 4. Typical alteration and veins of Shujiadian deposit A: Clots of skarn alteration, superposed by star-like epidote; B: Biotite replacing garnet (cross-polarized light); C: Plagioclase phenocrysts and matrix of quartz diorite porphyry with pervasive K-feldspar alteration; D: Pyroxene diorite with irregular K-feldspar vein and barren quartz vein with K-feldspar alteration; E: Pyroxene diorite with irregular K-feldspar vein, superposed by pyrite, anhydrite and epidote; F: Microphotograph of the vein in Fig. E (crosspolarized light), in which all the K-feldspars with argillic alteration and epidote replacing K-feldspar are cut by later anhydrite; G: Pyroxene diorite with intensive biotite alteration, pyrite vein and quartz + pyrite + chalcopyrite vein, and pyrite vein, accompanying epidote alteration replacing hydrothermal biotite, cut by quartz + pyrite + chalcopyrite vein; H: Chlorite replacing hydrothermal biotite along its rim (plane-polarized light); I: Pyroxene diorite, with biotite–anhydrite vein, in which the biotites are automorphic with invisible cleavage; J: Pyroxene diorite and siltstone fragments cut by irregular anhydrite and pyrite veinlet, cemented by anhydrite, pyrite and other hydrothermal minerals; K, L: The microphotos of Fig. J, in which silty mudstone breccia has become mica schist after thermal metamorphism; M: Pyroxene diorite with clots propylitic alteration consisting of epidote, chlorite and minor anhydrite, quartz, carbonate and pyrite; N: Pyroxene diorite with pyrite + chlorite + illite vein, with carbonate alteration halo; O: Intensive feldspar-destructive alteration occurs mainly in pyroxene diorite with pyrite + chlorite + illite vein and quartz + chlorite + illite + pyrite, and is characterized by complete destruction of all silicate minerals and bleaching of the rock; P: Microphotos of Fig. O, in which pyroxene diorite is pervasively replaced mainly by sericite and kaolinite. Crystal forms of plagioclase are preserved; Q: In shallow depths pyroxene diorite is pervasively replaced by abundant ultra-fine granular quartz and clay minerals; R: In fracture zones, the rock develops carbonate alteration. Anh: Anhydrite, Bt: Biotite, Cal: Calcite, Chl: Chlorite, Clay: Clay minerals, Cp: Chalcopyrite, Ep: Epidote, Gt: Garnet, Ill: Illite, Kf: K-feldspar, Pl: Plagioclase, Py: pyrite, Qtz: Quartz, Ser: Sericite.

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wide) are irregular to planar and characterized by abundant chalcopyrite, chlorite and minor pyrite, pyrrhotite, quartz, and illite. They usually cut earlier veins (Fig. 5K) and are truncated by later pyrite ± quartz ± gypsum ± calcite ± chlorite ± illite veins. Pyrite ± specularite ± quartz ± gypsum ± calcite ± chlorite ± illite veins (0.5–30 mm wide) occur throughout the pyroxene diorite and quartz diorite porphyry. They cut earlier veins (Figs. 4D, 5D and J) and are truncated by gypsum + calcite ± quartz ± pyrite ± illite veins. Gypsum + calcite ± quartz ± pyrite ± illite veins, occurring in all magmatic rocks, are the latest veins in Shujiadian deposit and cut all other veins (Fig. 5O). In Shujiadian deposit, quartz + pyrite + chalcopyrite ± molybdenite and quartz + chalcopyrite + pyrite veins of the quartz + sulfide stage contribute most of the copper and mainly occur in the feldspar-destructive alteration zone. Chalcopyrite + chlorite ± pyrite ± pyrrhotite ± quartz ± illite veins of the sulfide-carbonate stage also contribute part of the copper, and they are concentrated in feldspar-destructive alteration zone.

4. Isotopes and trace elements studies 4.1. Sulfur isotopes Twenty samples from drill holes ZK601, ZK1814, ZK1815 and ZK2002 (Fig. 2B) were collected for sulfur isotope analyses using elemental analyzers (EuroVector EA3000) and Isotope-MASS (GV Instruments IsoPrime) in the Institute of Geochemistry, Chinese Academy of Sciences. Two molybdenite samples have d34S values of 6.5‰ and 6.7‰, with a mean of 6.6‰, whereas three pyrrhotite samples have d34S values of 5.3‰, 5.1‰, and 5.7‰, and four chalcopyrite samples have d34S values of 4.3‰, 4.8‰, 5.1‰ and 4.7‰, and eight pyrite samples have d34S values between 3.6‰ and 6.0‰, with a mean of 4.9‰. All seventeen samples fall in a narrow range from 3.6‰ to 6.7‰. Three anhydrite samples have d34S values of 13.9‰, 13.1‰, and 20.6‰. All the sulfur isotopes of sulfides and sulfates in the same stage conform to the rule of sulfate > molybdenite > pyrite > pyrrhotite > chalcopyrite (Zheng and Chen, 2000; Table 2, Fig. 6).

4.2. Trace elements and strontium isotopes of anhydrite We collected four anhydrite samples: (a) from the quartz-sulfide stage coexisting with clots of epidote (Fig. 4M, 1814-204 (1), from 204 m depth in drill hole ZK1814); (b) from an anhydrite + chlorite vein from the sulfide-carbonate stage (601–773, from 773 m depth in drill hole ZK601); (c) with feldspar-destructive alteration superimposed on a K-feldspar vein from the sulfide-carbonate stage (Fig. 5E, 601-569, from 569 m depth in drill hole ZK601); and (d) from an anhydrite-carbonate vein from the sulfide-carbonate stage (Fig. 5O, 1814-51.5, from 51.5 m depth in drill hole ZK1814). Trace element and strontium isotope analyses were done on the four samples in the Guizhou Tuopu Resource and Environment Analysis Center. Specific test methods are described in Qi and Zhou (2008). (1) Trace elements Rare earth elements of hydrothermal anhydrites and of the pyroxene diorite are enriched in LREEs and depleted in HREEs (Table 3, Fig. 7). RREE of anhydrites from early to late are 15.2  106, 3.6  106, 63.4  106 and 243.6  106; LREE/HREE are 108.1, 11.7, 43.5 and 5.5, and dEu are 2.0, 1.5, 0.9 and 1.0. They show the earlier anhydrites have higher RREE, and lower LREE/ HREE and dEu.

7

(1) Strontium isotopes As the ionic radius of Ca2+ (1.38  1010m) is similar to Sr2+ (1.27  1010m), strontium can replace Ca in minerals such as anhydrite and fluorite. More importantly, anhydrite does not contain radioactive 87Rb, implying that anhydrite preserves its initial strontium, which can act as an indicator of source. The 87Sr/86Sr values of three anhydrite samples fall in a narrow range from 0.708138 to 0.708939 (Fig. 8), whereas one anhydrite from an anhydrite-calcite vein has a 87Sr/86Sr value of 0.710487 (Table 4). 5. Fluid inclusions To study the evolution of the ore-forming fluid in Shujiadian deposit, we collected different stage veins with quartz and anhydrite (except veins with Ca-silicate alteration) to study fluid inclusions. Microthermometric Analyses were undertaken using a Linkam THMS 600 programmable heating-freezing stage at the Fluid Inclusion Laboratory at Hefei University of Technology and the State Key Laboratory for Mineral Deposits Research, Department of Earth Sciences, Nanjing University. The stage has a maximum temperature of 600 °C and was calibrated using synthetic fluid inclusions, and heating rates of 0.1–1.0 °C/min below 10 °C, and 3–5 °C/min between 10 and 31 °C, with a reproducibility of ±0.1 °C. A heating rate of 5–10 °C/min was used at higher temperatures (>100 °C), with reproducibility of ±2 °C. Single fluid inclusion compositions were determined using a Reinishaw RM2000 Laser Raman Spectrometer at the State Key Laboratory for Mineral Deposits Research, Department of Earth Sciences, Nanjing University. The compositions of single fluid inclusions were identified using a Renishaw RM2000 Raman microprobe with an Ar ion laser with a surface power of 5 mW to excite the radiation (514.5 nm), and the scanning range of spectra was set between 1000 and 4000 cm1 with an accumulation time of 30 s for each scan. Salinities of halite-undersaturated fluid inclusions, reported as weight percent NaCl equivalent (wt% NaCl equiv), were calculated from freezing point depression temperatures (final ice melting), using the method of Potter et al. (1978). Weight percent NaCl equivalent was calculated for halite-saturated and homogenization by vapor disappearance fluid inclusions by the algorithm of Sterner et al. (1988), whereas halite-saturated and homogenization by salt dissolution fluid inclusions were calculated by the method of Lecumberri-Sanchez et al. (2012). 5.1. Populations The fluid inclusion types observed in the Shujiadian deposit are summarized in Fig. 9 and Table 5. Fluid inclusion sample numbers are listed in Appendix A and sample locations are shown in Fig. 2B. Three major types of fluid inclusions and 4 subtypes have been identified. They have been classified by the number of phases present at room temperature, using the criteria of Nash (1976). Type I fluid inclusions are liquid rich, halite undersaturated, and homogenize by disappearance of the vapor phase (Fig. 9A and B). Type II fluid inclusions are vapor rich and homogenize by expansion of the vapor phase, and in the silicate stage contain an opaque metallic mineral (Fig. 9D–F). Type II fluid inclusions are generally filled entirely by vapor, whereas larger Type II fluid inclusions typically contain a small amount of liquid. Type III fluid inclusions are halite saturated and homogenize either by disappearance of the vapor phase (Type IIIa) or dissolution of the halite daughter crystal (Type IIIb). Type III fluid inclusions may contain only a single soluble daughter mineral (e.g., halite; Fig. 9G) or may contain halite plus one or more insoluble daughter minerals. Some Type III fluid

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Fig. 5. Photographs showing different vein characteristics and stages of Shujiadian deposit A: Pyrite veinlet and garnet vein in the pyroxene diorite superposed by carbonate and chlorite alteration; B: Barren quartz stockwork vein with faint bleached halos; C: Quartz + pyrite vein with faint bleached halos, discontinuous linear pyrite in the vein; D: Quartz + actinolite + pyrite + pyrrhotite + chalcopyrite vein with faint bleached halos, truncated by pyrite veinlet; E, F: Microphotograph of quartz + actinolite + pyrite + pyrrhotite + chalcopyrite vein in Fig. D, in which actinolite replaces quartz and pyrite, chalcopyrite replaces pyrrhotite; G: Quartz + molybdenite + pyrite vein with faint bleached halos, in which molybdenite and pyrite occur disseminated in the vein; H: Pyrrhotite + quartz vein cutting quartz + chalcopyrite vein, with feldspar-destructive alteration halos; I: Epidote + chlorite + chalcopyrite + pyrite + calcite vein with faint bleached halo; J: Pyroxene diorite with pervasive K-feldspar alteration and quartz + pyrite + chalcopyrite vein (pyrite and chalcopyrite occur disseminated in the vein) and pyrite stockwork vein. Pyrite stockwork vein with epidote alteration cutting through quartz + pyrite + chalcopyrite vein; K: Epidote + pyrite vein with faint bleached halos truncated by pyrite + chalcopyrite veinlet; L: Cross-polarized light microphotograph of pyrite + chalcopyrite veinlet in K, in which the plagioclase breaks into fine-grained quartz along the vein; M: Pyrite + calcite + quartz vein characterized by abundant pyrite, calcite and minor quartz in the fracture zone; N: Specularite ± pyrite vein occurs in shallow levels of the drill hole, commonly have holes from leaching; O: Gypsum + calcite vein in the outcrop of pyroxene diorite with pervasive feldspar-destructive and pyrite alteration. Act: Actinolite; Anh: Anhydrite, Bt: Biotite, Cal: Calcite, Chl: Chlorite, Clay: Clay minerals, Cp: Chalcopyrite, Ep: Epidote, Gt: Garnet, Gy: Gypsum, Ill: Illite, Kf: K-feldspar, Mo: Molybdenite, Pl: Plagioclase, Py: Pyrite, Pyr: Pyrrhotite; Qtz: Quartz, Ser: Sericite, Shem: Specularite.

inclusions contain halite and sylvite, as well as multiple insoluble daughter minerals (Fig. 9F and H). Most early veins in the Shujiadian deposit are cut by later veins and overprinted by later hydrothermal minerals (Figs. 4 and 5). This has resulted in complex overprinting of various fluid inclusion populations, and it is almost impossible to establish timing relationships among the different fluid inclusion types within any single quartz vein. It is not possible to distinguish conclusively

between overprinting and simultaneous trapping of different fluids. For this reason, fluid inclusions have not been classified as primary, secondary, or pseudosecondary. Instead, we have adopted a method of relating groups of fluid inclusions to the various stages of vein development similar to that outlined by Bodnar and Beane (1980) and Bloom (1981). Fluid inclusion populations containing multiple inclusion types in any one vein must satisfy the following criteria (Masterman et al., 2005), if they are considered to have

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S. Wang et al. / Journal of Asian Earth Sciences xxx (2014) xxx–xxx Table 2 Sulfur isotope compositions of sulfides and sulfates from the Shujiadian deposit (‰). Sample no.

Samples description

Stage

d34SMo

601-13 1815-386 601-344 1815-360 601-210

Quartz + pyrite + molybdenite vein Quartz + molybdenite vein Quartz + pyrite vein Quartz + pyrite + pyrrhotite vein Pyrite vein with epidote and chlorite alteration halo Clots epidote, chlorite, pyrite and anhydrite Quartz + pyrite vein with K-feldspar alteration halo Pyrite + quartz vein Pyrite veinlet (quartz diorite porphyry) Calcite + pyrite vein Quartz + pyrrhotite vein Pyrrhotite + quartz vein Chalcopyrite + pyrrhotite + pyrite vein Disseminated Chalcopyrite, pyrrhotite accompanying feldspar-destructive alteration Quartz + pyrite + chalcopyrite vein Chalcopyrite + quartz vein Chalcopyrite + pyrite vein Clots epidote, chlorite, anhydrite and pyrite alteration Anhydrite with feldspar-destructive alteration superposed K-feldspar vein Anhydrite + calcite vein

Qtz-sulfide stage

6.49 ± 0.05 6.65 ± 0.15

1814-204(2) 1814-33 2002-25 2002-286 2002-51 1814-998 1814-119 1815-814 1507-23 1815-459 1815-814 1815-498 1814-204(1) 601-569 601-773 Average CH-1

d34SPy

d34SPyr

d34SCp

d34SAnh

3.58 ± 0.01 4.69 ± 0.10 3.68 ± 0.03 5.48 ± 0.02 5.81 ± 0.08

Silicate stage Sulfide-carbonate stage

5.18 ± 0.06 5.95 ± 0.04 4.67 ± 0.06

Qtz-sulfide stage Sulfide-carbonate stage

5.29 ± 0.04 5.05 ± 0.02 5.73 ± 0.02 4.34 ± 0.02

Qtz-sulfide stage Sulfide-carbonate stage

4.79 ± 0.11 5.11 ± 0.02 4.73 ± 0.02

Qtz-sulfide stage

13.86 ± 0.06

Sulfide-carbonate stage

13.08 ± 0.20

6.57

4.88

Triassic gypsum-salt layer

5.35

4.74

20.62 ± 0.19 15.85 25.41 ± 0.01

Fig. 6. Sulfur isotopes of sulfide and sulfate in Shujiadian deposit.

Table 3 Trace elements of anhydrites of different stage in Shujiadian deposit (106). Sample no.

1814-204(1)

601-773

Stage La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Y RREE LREE/HREE LaN/YbN dEu

Qtz-sulfide stage 6.108 7.059 0.511 1.268 0.099 0.059 0.072 0.011 0.025 0.004 0.019 0.001 0.007 0.001 0.15 15.24 108.15 587.35 2.04

Sulfide-carbonate stage 0.813 25.5 1.477 25.9 0.161 2.26 0.641 7.07 0.131 0.979 0.055 0.257 0.082 0.635 0.017 0.08 0.084 0.32 0.014 0.06 0.04 0.17 0.005 0.02 0.031 0.12 0.005 0.02 0.62 1.72 3.56 63.39 11.77 43.48 17.61 143.60 1.52 0.94

601-569

1814-51.5 40.700 91.400 11.500 46.200 12.512 3.956 12.660 2.080 11.100 2.080 5.170 0.640 3.180 0.400 72.800 243.58 5.53 8.65 0.95

Fig. 7. Chondrite-normalized REE patterns for anhydrites from Shujiadian deposit (Chondrite data after Taylor and McLennan, 1985).

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S. Wang et al. / Journal of Asian Earth Sciences xxx (2014) xxx–xxx

Fig. 8. Strontium isotopes of anhydrites, garnet and pyroxene diorite in Shujiadian deposit.

coexisted at the time of trapping (e.g., Type II and Type III inclusions trapped during phase separation): (1) they occur in proximity not separated by more than 5–50 lm, (2) they have overlapping temperatures of homogenization, and (3) they do not occur in younger vein stages. 5.2. Fluid inclusions of different stages Silicate stage: Because we did not find fluid inclusions in any transparent mineral associated with Ca-silicate alteration, there are only data for fluid inclusions associated with potassic alteration in this stage. Quartz ± pyrite veins with a K-feldspar alteration halo containing all types of fluid inclusions represent the fluid composition of the silicate stage. Quartz ± pyrite veins contain abundant Type I fluid inclusions that homogenize between 375 and 556 °C, with a mean value of 446 °C (Appendix A; Fig. 10A). Salinities of the Type I fluid inclusions vary from 8.0 to 23.1 wt% NaCl equiv, with a mean value of 18.3 wt% NaCl equiv (Appendix A; Fig. 10B). Type II fluid inclusions homogenize between 450 and 596 °C (three fluid inclusions >600 °C), with a mean of 532 °C (Appendix A; Fig. 10A). Salinities of the Type II fluid inclusions vary from 0.2 to 2.2 wt% NaCl equiv, with mean of 1.2 wt% NaCl equiv (Appendix A; Fig. 10B). Most Type III hypersaline fluid inclusions homogenized by halite dissolution, and the remaining Type III hypersaline fluid inclusions homogenized by vapor disappearance. Final homogenization temperatures by vapor disappearance (Type IIIa) ranged from 450 to 562 °C (one fluid inclusion >600 °C) and averaged 502 °C (Appendix A; Fig. 10A), corresponding to calculated salinities of 42.4–53.7 wt% NaCl equiv (av. 49.3 wt% NaCl equiv; Appendix A; Fig. 10B). Type IIIb fluid inclusions with final homogenization temperatures by halite dissolution homogenize between 400 and 594 °C, with a mean of 481 °C (six fluid inclusions >600 °C; Appendix A; Fig. 10A), corresponding to calculated salinities of 48.0–70.2 wt% NaCl equiv (av. 56.8 wt% NaCl equiv; Appendix A; Fig. 10B). Quartz-sulfide stage: Veins of this stage contain abundant Type I and Type II fluid inclusions. Type I fluid inclusions from barren

quartz veins homogenize between 304 and 442 °C, with a mean of 377 °C (Appendix A), and salinities of the Type I fluid inclusions vary from 21.2 to 23.0 wt% NaCl equiv, with a mean of 22.2 wt% NaCl equiv (Appendix A). Quartz + pyrite + molybdenite veins (pyrite and molybdenite occur disseminated in the veins) contain Type I and Type II fluid inclusions. Fluid inclusions of Type I homogenize from 319 to 449 °C (av. 407 °C) and Type II homogenize from 318 to 411 °C (av. 365 °C), corresponding to calculated salinities of 14.6–22.4 wt% NaCl equiv (av. 19.6 wt% NaCl equiv; Appendix A) and 0.2–6.0 wt% NaCl equiv (av. 2.2 wt% NaCl equiv; Appendix A), respectively. Type I fluid inclusions from quartz + pyrite + chalcopyrite veins homogenize between 314 and 436 °C, with a mean of 383 °C (Appendix A), and salinities vary from 15.8 to 22.9 wt% NaCl equiv, with a mean of 21.0 wt% NaCl equiv (Appendix A). Homogenization temperatures of Type I fluid inclusions from quartz + pyrite veins (Fig. 5C) range from 356 to 496 °C, with a mean of 429 °C (Appendix A), and salinities vary from 14.2 to 18.7 wt% NaCl equiv, with a mean of 16.5 wt% NaCl equiv (Appendix A). Quartz + molybdenite veins (Fig. 5G) contain abundant Type II vapor-rich fluid inclusions that homogenize between 294 and 480 °C (av. 363 °C; Appendix A). Salinities of Type II fluid inclusions vary from 0.5 to 7.2 wt% NaCl equiv (av. 3.4 wt% NaCl equiv; Appendix A). We also found two closely-spaced fluid inclusions that homogenize at 399 and 480 °C with corresponding calculated salinities of 14.8–17.3 wt% NaCl equiv. Type I fluid inclusions from quartz + chalcopyrite + pyrite veins (Fig. 5H) homogenize between 364 and 396 °C, with a mean of 377 °C (Appendix A), and salinities vary from 14.9 to 21.5 wt% NaCl equiv, with a mean of 19.1 wt% NaCl equiv (Appendix A). Homogenization temperatures of Type I fluid inclusions from quartz + actinolite + pyrite + pyrrhotite + chalcopyrite veins (Fig. 5D) range from 291 to 496 °C, with a mean of 386 °C (Appendix A), and salinities vary from 19.9 to 23.2 wt% NaCl equiv, with a mean of 22.2 wt% NaCl equiv (Appendix A). In general, Type I fluid inclusions in the quartz-sulfide stage homogenize between 291 and 496 °C, with a mean of 395 °C (Appendix A), corresponding to calculated salinities of 14.2– 23.2 wt% NaCl equiv, with a mean of 20.2 wt% NaCl equiv (Appendix A; Fig. 9C and D). Type II fluid inclusions homogenize between 295 and 480 °C, with a mean of 364 °C (Appendix A), corresponding to calculated salinities of 0.2–7.2 wt% NaCl equiv, with a mean of 3.1 wt% NaCl equiv (Appendix A; Fig. 9C and D). Sulfide-carbonate stage: Veins of this stage contain only Type I fluid inclusions, which homogenize from 157 to 379 °C, corresponding with calculated salinities of 0.9–19.0 wt% NaCl equiv. Pyrite + chalcopyrite + quartz veins, consisting of coarse-grained pyrite (80%) and chalcopyrite (10%) with minor quartz (8%), contain abundant Type I inclusions that homogenize between 333 and 379 °C, with a mean of 356 °C (Appendix A), and salinities that vary from 13.2 to 19.0 wt% NaCl equiv, with mean value of 16.3 wt% NaCl equiv (Appendix A). Type I fluid inclusions from pyrite + chalcopyrite + quartz veins that contain fine-grained pyrite (90%) and quartz (8%) and minor chalcopyrite (10%) homogenize from 237 to 291 °C, with a mean of 254 °C (Appendix A), and salinities that range from 8.6 to 16.9 wt% NaCl equiv, with a mean of 14.8 wt% NaCl equiv (Appendix A). Homogenization temperatures of fluid inclusions from quartz + carbonate + pyrite veins

Table 4 Strontium isotopes of anhydrite and garnet of Shujiadian deposit. Sample no.

BCR-1

BHVO-2

NISTSRM-987

Recommendation 87

Sr/86Sr 2SE

0.70502 0.00004

Notes: Z-94-70 is the

0.703456 0.000017

0.710250 0.00001

987

Z-94-70

601-569

Measured

Garnet

Anhydrite

0.710231 0.000013

0.709486 0.000028

0.708138 0.000009

1814-51.5

1814-204(1)

601-773

0.710487 0.000017

0.708745 0.000015

0.708939 0.000008

87

Sr/86Sr value of garnet, data from Zhao and Zhao (1997).

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S. Wang et al. / Journal of Asian Earth Sciences xxx (2014) xxx–xxx

Fig. 9. Photomicrographs of fluid inclusions in vein quartz from Shujiadian deposit (plane-polarized light) A: Type I fluid inclusion from pyrite + calcite + quartz vein of the sulfide-carbonate stage; B: Type I fluid inclusion with opaque from quartz + pyrite vein (pyrite occurs disseminated in the vein); C: Type I fluid inclusion from anhydrite with feldspar-destructive alteration superposed K-feldspar vein; D: Boiling fluid inclusion assemblage in quartz + molybdenite vein of the quartz-sulfide stage; E: Type II fluid inclusion from quartz + pyrite vein with K-feldspar alteration halo (pyrite occur disseminated in the vein) of the silicate stage; F: Immiscible fluid inclusion assemblage (Type I and IIIb fluid inclusions) from quartz + pyrite + K-feldspar vein (pyrite occurs disseminated in the vein) of the silicate stage; G: Type IIIa fluid inclusion from quartz + pyrite + K-feldspar vein (pyrite occur disseminated in the vein) of the silicate stage; H: Type IIIb fluid inclusion consists of sylvite, halite and opaque from quartz + pyrite + K-feldspar vein (pyrite occur disseminated in the vein) of the silicate stage; V: Vapor; L: Liquid; H: Halite; S: Sylvite; O: Opaque.

Table 5 Summary of fluid Inclusion types, morphology, phases present at 25 °C, and homogenization behavior. Type

Inclusion shape

Vapor–liquid ratio

Dominant

Types

Homogenization behavior

I

Elongate, negative; crystals or rounded, less; commonly irregular Stubby negative crystals, less commonly rounded

10–40%

Liquid

Liquid + vapor ± opaque

Vapor disappearance

50–90%

Vapor

Vapor + liquid ± opaque

Liquid disappearance

Negative crystals, rounded or irregular Negative crystals, rounded or irregular

5–25% 5–25%

Liquid Liquid

Liquid + vapor + halite ± opaque Liquid + vapor + halite ± sylvite ± opaque

Vapor disappearance Salt dissolution

II III a b

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S. Wang et al. / Journal of Asian Earth Sciences xxx (2014) xxx–xxx

Fig. 10. Homogenization temperature and salinity histograms for all inclusion types separated into different stages and host minerals.

are between 248 and 294 °C, and salinities vary from 6.3 to 15.1 wt% NaCl equiv. Type I fluid inclusions from anhydrite superimposed on K-feldspar veins (Fig. 4E and F) homogenize between 157 and 286 °C, with a mean of 219 °C (Appendix A), and salinities vary from 0.9 to 9.7 wt% NaCl equiv, with a mean of 4.7 wt% NaCl equiv (Appendix A).

5.3. Laser Raman spectroscopic analysis Representative fluid inclusions were selected for Laser Raman microspectroscopy to constrain their compositions. The results show that the liquid phases in all kinds of fluid inclusions are dominated by H2O with minor HCO 3 in Type I fluid inclusions in anhydrite in the sulfide-carbonate stage (Fig. 11); however the composition of vapor phases vary in different mineralization stages. The vapor phases of Type II and III fluid inclusions of the Ca-silicate and quartz-sulfide stages consist of H2O and CO2, and Type I fluid inclusions from the two stages show similar characters of minor H2O and CO2, whereas the vapor phase of Type I fluid inclusions of the sulfide-carbonate stage contains only H2O. The CO2 contents of fluid inclusions of different stages show a decreasing trend from early to late.

5.4. Pressure estimation Laser Raman spectroscopic analyses of fluid inclusions show that the fluid inclusions contain CO2, but we cannot see CO2 phases at room temperature and use the NaCl–H2O system to discuss the evolution of the ore-forming fluid at Shujiadian. The trapping pressure can be estimated only when the exact trapping temperature is known, or fluid inclusions are trapped during phase separation (Roedder and Bodnar, 1980; Brown and Hagemann, 1995). Experimental (Haas, 1976; Bodnar and Sterner, 1985; Sterner et al., 1988; Bischoff and Pitzer, 1989; Bodnar, 1992) and theoretical (Bischoff and Pitzer, 1989; Anderko and Pitzer, 1993) PVTX data are available for the vapor-saturated halite solubility curve and the liquid–vapor surfaces for the system H2O–NaCl. Type IIIb fluid inclusions homogenize along the halite liquidus and traverse the liquid + halite field after the vapor bubble disappears (Becker et al., 2008). PVTX data along the halite liquidus and in the liquid + halite field are only available for a composition of 40 wt% NaCl equiv (Bodnar, 1994). Therefore, we can roughly estimate the trapping pressure for Type I, Type II and Type III fluid inclusions through the H2O–NaCl system (Driesner and Heinrich, 2007). Fluid inclusions that homogenize by halite dissolution (Type IIIb) also can provide a tool to constrain the pressure of fluid

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S. Wang et al. / Journal of Asian Earth Sciences xxx (2014) xxx–xxx

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Fig. 11. Laser Raman spectra of fluid inclusions from Shujiadian deposit A: Vapor of Type III fluid inclusion (final homogenization temperatures by halite dissolution) from quartz + pyrite vein (accompanying K-feldspar alteration halo; Fig. 4D) of the silicate stage consist of H2O and CO2; B: Vapor of Type II fluid inclusion from the same vein with the fluid inclusion in the Fig. A contains H2O and CO2; C: Vapor of Type I fluid inclusion from the same vein with the fluid inclusion in the A consist of H2O and minor CO2; D: Vapor of Type II fluid inclusion from quartz + pyrite + molybdenite vein (pyrite and molybdenite occur disseminated) of the quartz-sulfide stage contain H2O and CO2; E: Type I fluid inclusion from the same vein with the fluid inclusion in D have less obvious peak of H2O and CO2; F: Type II fluid inclusion from quartz + molybdenite vein (Fig. 5G) show less obvious peak of CO2 and apparent H2O; G: Fluid inclusion from quartz + carbonate + pyrite only show the peak of H2O; H: Liquid of fluid inclusion from anhydrite superposed K-feldspar vein (Fig. 4F) show the peak of HCO 3.

entrapment when microthermometric data are available. As an example, Hurai et al. (2002) and Milovsky´ et al. (2012) determined the pressure at the temperature of halite dissolution by determining the intersection of the metastable P–T path followed by fluid inclusions in the liquid field with the measured halite dissolution temperature. The method of Hurai et al. (2002) can be applied to inclusions that can be cooled after homogenization to re-nucleate the bubble without nucleating halite, and which can then be reheated to the temperature of metastable vapor dissolution in the absence of halite. In that case, the halite-absent homogenization temperature yields a metastable dP/dT slope that can be used to determine the salinity and pressure at the temperature of halite dissolution. Becker et al. (2008) developed an equation that expresses pressure at halite dissolution as a function of the halite dissolution temperature (Tm) and temperature of bubble disappearance (Th). The model of Becker et al. (2008) can be used to

interpret microthermometric data from many fluid inclusions that homogenize by halite disappearance, but there are also many fluid inclusions that have Tm and Th values outside the range of validity of the model. For example, fluid inclusions that have Th < 300 °C and that homogenize by halite disappearance are common in porphyry copper deposits (Roedder and Bodnar, 1997). Building on the work of Becker et al. (2008), Lecumberri-Sanchez et al. (2012) developed a model that is applicable over the range 100–600 °C for calculation of salinity and estimation of pressure in fluid inclusions that homogenize by halite dissolution. We estimate the pressure of Type IIIb fluid inclusions based on this model. Fig. 12A shows the distribution of Type I, Type II and Type IIIa fluid inclusions as a homogenization temperature versus salinity plot, in which isobars are calculated after Driesner and Heinrich (2007). Fig. 12B shows the distribution of Type IIIb fluid inclusions as a halite dissolution temperature versus vapor disappearance

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temperature plot, in which isobars are calculated after LecumberriSanchez et al. (2012). Pressure of Type I fluid inclusions of the silicate stage are mostly between 200 and 800 bars, with an average value of 500 bars, corresponding to 2 km under lithostatic conditions. The pressure of Type II and Type IIIa fluid inclusions of the silicate stage are mostly between 300 and 700 bars, with an average value of 500 bars, corresponding to 2 km under lithostatic conditions, whereas the pressure of Type IIIb fluid inclusions of the silicate stage are higher and mostly range from 800 to 2000 bars. Type II fluid inclusions of the quartz-sulfide stage were trapped at 100–300 bars (av. 200 bars), corresponding to 2 km under hydrostatic conditions. Type I fluid inclusions from barren veins of the quartz-sulfide stage were trapped at 150–400 bars and from ore-bearing veins were trapped at 100–300 bars. The pressure of fluid inclusions of the sulfide-carbonate stage vary from 20 to 200 bars, with an average value of 110 bars, corresponding to 1.1 km under hydrostatic conditions. 6. Discussion 6.1. Origin of ore-forming fluids Many studies have been carried out on the sulfides and sulfates of sedimentary strata in the Tongling ore district, in which the Permian black shale and Carboniferous limestone and dolomite of have d34S values between 13.0 and 38.6‰ (Gu, 1984; Liu et al., 1984; Zhou, 1984), with an average of approximately 28.4‰ (Fig. 6), implying that they are of biogenic origin (Xu et al., 2010). But there are also d34S values of sedimentary sulfides showing positive values. For example, pyrites from sandstone, siliceous rocks and disseminated, lamellar sulfides have d34S values between +3.9 and +8.7‰, with an average of approximately +5.6‰ (Gu, 1984; Anhui Provincial Bureau of Geology 321 Geological Team, 1995; Pan and Dong, 1999). Pyrites from dolomite of the Huanglong and Chuanshan formations (C2+3) of Carboniferous age have d34S values of +4.1‰ (Fig. 6; Anhui Provincial Bureau of Geology 321 Geological Team, 1995). Anhydrites from sedimentary strata of Carboniferous and Permian age have d34S values between +14.8‰ and 25.4‰ (Anhui Provincial Bureau of Geology 321 Geological Team, 1995; Pan and Dong, 1999; Hou et al., 2011b), similar to the value of Carboniferous seawater (Claypool et al., 1980), implying that the anhydrites were derived from Carboniferous seawater (Xu et al., 2010).

d34S values of sulfides in Shujiadian deposit are between +3.6‰ and +6.7‰, and d34S values of anhydrites range from 13.1‰ to 20.6‰. We cannot directly determine whether the sulfides and the anhydrites are sedimentary or hydrothermal in origin by simple analogy. d34S values of sulfides and anhydrites are characterized by equilibrium fractionation. Therefore, we can discuss the evolution of sulfur isotopes in Shujiadian deposit based on the model of Ohmoto and Rye (1979) and the diagrams of sulfur isotope evolution in equilibrium with hydrothermal systems (Chu et al., 1984). Whole-rock sulfur isotopic compositions of contemporaneous magmatic rocks in the Tongling region (quartz diorite porphyry, pyroxene diorite and granodiorite porphyry) range from d34S +0.3‰ to +6.9‰ (Anhui Provincial Bureau of Geology 321 Geological Team, 1995; Li et al., 1997), which are consistent with the sulfur isotopic composition of magmatic melts (3‰ to +7‰; Ohmoto, 1986; Ohmoto and Goldhaber, 1997). Therefore, we infer that the pyroxene diorite of Shujiadian deposit had d34S between +0.3 and +6.9‰. Accessory minerals of the pyroxene diorite mainly consist of apatite and zircon, with no quartz, magnetite, hematite, etc. During the process of hydrous silicate melts producing a magmatic hydrothermal fluid, we infer that the system cooled from 1100 °C to 700 °C and the resulting aqueous fluids mainly contained H2S and SO2 (Ohmoto and Rye, 1979). Sulfur isotopic values are distributed in the gray shaded area of the log f o2–T diagram (34Sfluid-melt = +0.6 to +3‰; Fig. 13A), and the magmatic hydrothermal fluid after the melt-hydrothermal fluid fractionation had d34S values of +0.9 to +9.9‰. The silicate stage at Shujiadian deposit is characterized by Casilicate alteration with hydrothermal garnet, diopside, K-feldspar, biotite, quartz, minor magnetite and pyrite, which is mainly distributed in the light gray area of the log f o2–T diagram (Fig. 13B). The quartz-sulfide stage began with abundant pyrite and ended with pyrrhotite, chalcopyrite and minor anhydrite, which extends to the pyrite–pyrrhotite equilibrium area of the log f o2–T diagram (Fig. 13B). We have determined that the fluid inclusions of the quartz-sulfide stage homogenized between 300 and 500 °C, mainly from 350 to 500 °C (Fig. 10C), and their sulfur species were dominated by H2S (Ohmoto and Rye, 1979). The hydrothermal fluid of the quartz-sulfide stage is mainly distributed in the dark area in Fig. 13B, assuming that SO2–H2S concentration is at the maximum limit for hydrothermal fluids of the quartz-sulfide stage, and sulfur isotope fractionation values (D34Spyrite-fluid) between sulfides

Fig. 12. Pressure estimation for fluid inclusions of Shujiadian deposit A: Type IIIa and Type II inclusions of the silicate stage were trapped under immiscible conditions, Type I and Type II inclusions of the quartz-sulfide stage were trapped under boiling conditions, thus the estimated pressures indicate the actual trapping pressures for both. Isobars were calculated from the equations of Driesner and Heinrich (2007). B: Type IIIb fluid inclusions of the silicate stage homogenized via halite dissolution and therefore their estimated pressures indicate their minimum trapping pressures. Phase diagram used was from Lecumberri-Sanchez et al. (2012).

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S. Wang et al. / Journal of Asian Earth Sciences xxx (2014) xxx–xxx

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Fig. 13. Equilibrium system diagrams of Sulfur isotopes A: Sulfur isotopic fractionation between magmatic hydrothermal and hydrous silicate melt (base map Ohmoto and Rye, 1979); B: Sulfur isotopic fractionation log f o2–T diagram of pyrite–fluid in equilibrium with high temperature hydrothermal system (Chu et al., 1984); C: Sulfur isotopic fractionation log f o2–T diagram of sulfates–fluid in equilibrium with high temperature hydrothermal system (Chu et al., 1984) Q: Quartz, M: Magnetite, H: Hematite, F: Fayalite, Py: Pyrite, Pyr: Pyrrhotite.

(pyrite) and hydrothermal fluid vary from 1.0‰ to 0‰. Combined with the sulfur isotopic composition of magmatic hydrothermal fluid (d34S = +0.9‰ to 9.9‰), we calculate d34S values of 0.1‰ to +9.9‰. The quartz-sulfide stage has anhydrite coexisting with epidote (Fig. 4M), corresponding to the dark area in the log f o2–T diagram (Fig. 13C), and sulfur isotope fractionation value (D34Ssulfate-fluid) between sulfate (anhydrite) and the hydrothermal fluid varies from +7.0 to +19.0‰. Taking the average value (+5.1‰) of the magmatic hydrothermal fluid to represent the initial value of the hydrothermal fluid, the calculated d34S value for anhydrite ranges from +12.1 to +24.1‰. Thus, the sulfur isotopic compositions of sulfides and anhydrites from the silicate and quartz-sulfide stages are consistent with calculated values, indicating that they mainly come from a magmatic hydrothermal fluid. Sulfides of the sulfide-carbonate stage have similar sulfur isotopic values to the early stage (Table 2), which may imply they also come from a magmatic hydrothermal fluid. Anhydrite superimposed on K-feldspar veins (Fig. 4E and F) of the sulfide-carbonate stage has d34S values of 13.1 ± 0.2‰, similar to anhydrite of the quartz-sulfide stage (13.9 ± 0.1‰), whereas anhydrite from anhydrite-calcite veins of the sulfide-carbonate stage have d34S values of 20.6 ± 0.2‰, higher than anhydrite superimposed on K-feldspar veins, indicating that they were the result of hydrothermal fractionation or addition of sulfur from another source, such as the regional gypsum-salt layer (Zhou et al., 2007; Zhang, 1981; Chu et al., 1986). In porphyry Cu (-Mo) deposits, along with an increase over time in the proportion of meteoric fluid, the ore-forming fluid system will decrease in temperature and pH and increase in fluid/rock ratio, which can lead to leaching of REE (especially of HREE; Zhao, 1997). The earlier anhydrites at Shujiadian have higher RREE, and lower LREE/HREE and dEu, which may be caused by REE leaching. Overall, RREE of hydrothermal anhydrites shows similar character to the pyroxene diorite (Fig. 7), possibly inheriting the REE characteristics of the magmatic rocks. Strontium isotope ratios of anhydrites are mainly concentrated in the range 0.708138–0.708939 (Fig. 8), similar to the ratios of pyroxene diorite (0.707023–0.707392; our unpublished data), which suggests that they originated from magmatic fluids. Anhydrite from anhydrite + calcite veins of the sulfide-carbonate stage has an 87Sr/86Sr value of 0.710487, higher than anhydrites from

the quartz-sulfide and sulfide-carbonate stages (Fig. 8), which implies that strontium isotopes from other sources were added. There are three possible sources: wall rocks, contemporaneous seawater, or meteoric water. Anhydrites of different stages have similar compositions of strontium isotopes, slightly higher than the pyroxene diorite, which possibly results from mixing between magmatic fluid and another fluid. Strontium isotopes of contemporaneous seawater range from 0.706500 to 0.708000 (Huang et al., 2005) and the 87Sr/86Sr values of anhydrites have an increasing trend (Fig. 8), which suggests that anhydrite + calcite veins in the shallow levels of Shujiadian deposit have interacted with meteoric water. 6.2. Evolution of the ore-forming fluid An interpretation of fluid evolution of the Shujiadian hydrothermal system is illustrated in Fig. 14, which shows the pressure–temperature paths of two different fluid-flow stages characterized by different fluid inclusion assemblages. Episodic fluid activity resulted in development of alteration minerals and deposition of metals. 6.2.1. Silicate stage At shallow levels (6–8 km) in the earth’s crust, the initial supercritical fluid with intermediate salinity (10 wt% NaCl equiv, >600 °C) exsolved from the magma chamber (Bodnar, 1995; Masterman et al., 2005; Richards, 2005). Fast decompression and slow cooling resulted in this fluid intersecting the two-phase surface (Shinohara and Hedenquist, 1997), and hence fluid immiscibility between brine and low-salinity vapor occurred, as observed in many porphyry deposits and widely documented (Baker and Lang, 2003; Masterman et al., 2005). The silicate stage at Shujiadian shows immiscible features, evident in the coexistence of Type II and Type III inclusions with similar homogenization temperatures. At 510 °C, phase separation of the initial supercritical fluid produced a 50–57 wt% NaCl equiv brine and a low-density vapor with salinity of less than 2 wt% NaCl equiv as inferred from Fig. 14, and most Ca-silicate and potassic alteration minerals formed from this high-salinity liquid (Meinert et al., 1997; Shinohara and Hedenquist, 1997; Shu et al., 2013).

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Fig. 14. Fluid evolution paths of the Shujiadian deposit (phase diagram after Fournier, 1987). The initial fluids for silicate and quartz stage are both from magma chamber with salinity of 10 wt% NaCl equiv, but fluids for these two stages follow different cooling paths. In silicate stage, the trajectory hits the solvus above 510 °C and at a pressure of 500 bars (equal to lithostatic depth of about 2 km). In quartzsulfide stage, the trajectory hits solvus at 370 °C and pressure of 200 bars (equal to hydrostatic depth of about 2 km). For details see the text.

Type I fluid inclusions are observed in every vein stage at Shujiadian, this may reflect complex overprinting of different fluid types. Type I inclusions of the silicate stage homogenize between 375 and 556 °C, corresponding to calculated salinities of 8.0– 23.1 wt% NaCl equiv, and Type I inclusions of the quartz-sulfide stage homogenize between 291 and 496 °C, corresponding to calculated salinities of 13.2–23.18 wt% NaCl equiv (Fig. 10). Veins of the silicate stage usually are overprinted by later alteration (Fig. 4E and F). Thus, we conclude that Type I inclusions of the silicate stage with high homogenization temperatures are the product of fluid immiscibility and inclusions with low homogenization temperatures may be caused by later fluids. Halite saturated inclusions with final homogenization by halite dissolution (Type IIIb) can either be produced by direct exsolution from the parent magma (e.g., Cline and Bodnar, 1994; Zhang et al., 2007), or through a process of postentrapment modification (e.g., Klemm et al., 2008), or else, trapping after undergoing overpressuring (e.g., Baker and Lang, 2003; Masterman et al., 2005). When studying the fluid inclusions of the Questa porphyry Mo deposit, Cline and Bodnar (1994) found that there was a lack of evidence for fluid boiling, therefore, they attributed the homogenization of halite-saturated inclusions through halite dissolution to entrapment of a high-salinity single phase fluid directly exsolved from a silicate melt. However, Klemm et al. (2008) re-evaluated this deposit and discovered unambiguous boiling assemblages, within which coexisting lowsalinity inclusions and halite saturated inclusions homogenizing through vapor disappearance are both abundant and they proposed that the vapor and brine were separated from a singlephase fluid of intermediate salinity. They argued that brine inclusions that homogenized by halite dissolution in Questa were generated from post-entrapment modification of brine inclusions homogenized through vapor disappearance because of loss of H2O (Audétat and Günther, 1999). Shu et al. (2013) found halite saturated inclusions that final homogenized by halite dissolution (Type IIIb) to be later than halite saturated inclusions with final homogenization by vapor disappearance (Type IIIa) and inclu-

sions that homogenized by expansion of the vapor phase (Type II), indicating that they were trapped at anomalously high pressures. In the Shujiadian deposit, we have found evidence for fluid immiscibility, and we can therefore rule out the likelihood of direct exsolution from a silicate melt for Type IIIb fluid inclusions. We have no evidence suggesting that the Type IIIb inclusions were not modified from Type IIIa inclusions. Besides, most of the Type IIIb inclusions have rounded or negative crystal shapes, in contrast with the irregular, flat shape that indicates post-entrapment modification (Bodnar, 2003; Klemm et al., 2008). This suggests that the Type IIIb inclusions in the Shujiadian deposit were unlikely to have been generated by a significant modification process, but can be better explained as a result of entrapment under overpressuring conditions. Burnham (1985) pointed out that release and expansion of volatiles during the emplacement and crystallization of volatile-rich magmas could produce high fluid pressure gradients, which have been documented in many deposits (e.g., Baker and Lang, 2003; Oliver et al., 2006; Bertelli and Baker, 2010). In the Shujiadian deposit, local overpressuring, which could be caused by continuous fluid immiscibility in the silicate stage, also occurred and resulted in the entrapment of the Type IIIb inclusions under anomalously high pressures (Fig. 12B). 6.2.2. Quartz-sulfide stage Type I fluid inclusions in the quartz-sulfide stage homogenize between 291 and 496 °C, corresponding with calculated salinities of 14.2–23.2 wt% NaCl equiv, and Type II fluid inclusions homogenize between 295 and 480 °C, with a mean value of 364 °C, corresponding with calculated salinities of 0.2–7.2 wt% NaCl equiv. (Appendix A; Fig. 10C and D). They have lower homogenization temperatures than Type II inclusions in the veins of the silicate stage. We therefore infer that the vapor-rich fluids trapped in the veins of the quartz-sulfide stage were different from those trapped in the veins of the silicate stage, and may have been produced by a second exsolution from the silicate melt. 6.2.3. Sulfide-carbonate stage Veins of this stage contain only Type I fluid inclusions, which homogenize from 157 to 379 °C, corresponding with calculated salinities of 8.6–19.0 wt% NaCl equiv and the homogenization temperatures of barren veins range from 157 to 294 °C, with salinities of 0.9–15.1 wt% NaCl equiv. The salinities show an obvious positive linear relationship with homogenization temperatures (Fig. 12A), which suggests addition of external meteoric water, consistent with the results of sulfur and strontium isotopes of this study. Homogenization temperatures and salinities of ore-bearing veins are higher than those of the barren veins, possibly implying that the metal mainly came from the magma. 6.3. Migration and precipitation of ore-forming fluid Krauskopf (1957, 1964) concluded that the solubility of most metals in aqueous vapor is negligible (Hg, As, and Sb are exceptions), even at temperatures as high as 800 °C. However, his estimates were based on data for the vapor pressure of metallic species over the corresponding solids (i.e., their dry sublimation or volatility) and ignored the possibility that interactions with the solvent might enhance solubility in aqueous vapor (Williams-Jones and Heinrich, 2005). In recent years, a number of experimental studies have investigated the stability of metallic species in aqueous vapors, and in each case the solubility is orders of magnitude higher than that predicted from volatility data (cf. Williams-Jones et al., 2002). Even at the low temperatures (6360 °C) of these experiments, the measured metal concentrations in the vapor phase would be sufficient to permit

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formation of economic ore deposits, and experiments at magmatic conditions (e.g., Simon et al., 2005a,b) have yielded metal concentrations similar to those reported in vapor inclusions from natural systems (cf. Heinrich et al., 1999). In Shujiadian deposit, fluid inclusions from ore-bearing veins (quartz + pyrite + chalcopyrite veins, quartz + actinolite + pyrrhotite + chalcopyrite + pyrite veins) of the quartz-sulfide stage and pyrite + chalcopyrite + quartz veins have higher salinities than fluid inclusions from barren veins of the same stage; we therefore infer that the brine is probably the main carrier of metals. Experimental studies have indicated that many factors (e.g., temperature, pressure, salinity, oxygen fugacity, etc.) can cause Cu precipitation, of which decrease of temperature seems to be the most important (Hezarkhani et al., 1999), and has been documented in many deposits (e.g., Ulrich et al., 2001; Redmond et al., 2004; Landtwing et al., 2005). But this experimental study (Hezarkhani et al., 1999) is based on the Cu being mainly transported in the form of a chloride complex (CuCl 2 ). Williams-Jones and Heinrich (2005) conducted an experimental study that suggests that Cu, Fe, Zn of fluids are enriched in the brine when the porphyry system is rich in Cl and poor in S, whereas Cu preferentially partitions into vapor and can be carried as a Cl complex or HS, or SO2 3 complex, or as CuClmnH2O. As a result, changes in pressure and sulfur content are the important controlling factors for precipitation of Cu in low density vapor (e.g., Qulong porphyry deposit in Tibet, Yang and Hou, 2009). In the quartz-sulfide stage (the main stage of mineralization), as the underlying magma continuously cooled and crystallized, it would approach the end stages of stagnant crystallization (Shinohara and Hedenquist, 1997). As a result, the late exsolved fluid would cool faster upon ascent due to the lack of concurrent dike intrusion (Shinohara and Hedenquist, 1997), and therefore it would not intersect the two-phase surface again because of the more gentle cooling path (path 2, Fig. 14). However, when the temperature decreased below 400 °C, which is thought to be the transition temperature from lithostatic to hydrostatic conditions caused by the change from ductile to brittle behavior of the system (Fournier, 1992, 1999), the fluid pressure would fall rapidly, consequently resulting in the supercritical fluid intersecting the two-phase surface again and boiling, which would result in the precipitation of sulfides (Shu et al., 2013). Homogenization temperatures of ore-bearing veins from the quartz-sulfide stage are less than 400 °C and pressures are 200 bars (same depth as the silicate stage, Fig. 12A), which indicates sulfides (e.g., chalcopyrite) mainly precipitated below 400 °C with fluid boiling. This also suggests that pressure change is one of the important factors for mineralization. Eu is an important Rare Earth Element with multiple valence states, Eu3+ and Eu2+. The redox potential of Eu3+/Eu2+ has a positive correlation with increasing temperature and no obvious change with increasing pH and pressure. Under most hydrothermal conditions and metamorphism, Eu is divalent in the fluid, and is preferentially absorbed by some minerals, therefore it can show differentiation relative to other REE (Sverjensky, 1984). The decreasing trend of dEu values of anhydrites indicates that the ore-forming fluid of quartz-sulfide stage were relatively reduced compared to other stages. Veins of the quartz sulfide stage commonly contain chalcopyrite, pyrite and other sulfides, which indirectly suggests that changes of oxygen fugacity may have caused the precipitation of sulfides. In addition, the precipitation of sulfides in Shujiadian deposit may be related to the deposition of anhydrite. No anhydrite has been found in the silicate stage, where sulfides are absent. The deposition of sulfides began with precipitation of anhydrite in the late quartz-sulfide stage. The precipitation of anhydrite can not only reduce the sulfur content and the solubility of Cu

17

in the fluid, and it also promotes the disproportionation of SO2 (Hemley and Hunt, 1992), both of which result in an increase of H2S content and the precipitation of sulfides. After decompression and decrease of sulfur content in the quartz-sulfide stage, Cu was mainly transported as a Cl complex in the sulfide-carbonate stage, which is consistent with higher salinity (Fig. 12A) and the absence of Type II fluid inclusions in the ore-bearing veins. Decrease of temperature and salinity may be the main factor for the precipitation of sulfides in the sulfide-carbonate stage (Fig. 12A). 6.4. Deposit erosion and possible genesis of skarn 6.4.1. Deposit erosion If we can identify the trapping pressure of fluid inclusions, we can estimate the depth of deposition (Hedenquist et al., 1998; Proffett, 2009). Type IIIa and Type II inclusions of the silicate stage were trapped under immiscible conditions, thus the estimated pressures approximate the actual trapping pressures. The silicate stage occurred at least 2 km deep, but most of the samples were collected about 1 km below the present-day surface; we therefore conclude that the shallow parts of Shujiadian deposit have been denuded at least 1 km, and that is supported by the distribution of biotite alteration in the outcrop. 6.4.2. Possible genesis of skarn Shujiadian deposit was mainly emplaced in siltstones and sandy shales of Silurian age, with no carbonate, but we found skarn alteration (garnet, diopside), which requires an external source to provide calcium. There are other porphyry–skarn deposits (e.g., Dongguashan, Xishizishan), which were emplaced in Devonian Wutong group sandstones and Carboniferous, Permian, and Triassic carbonates. The 87Sr/86Sr ratio of garnets in Shujiadian deposit is 0.709486 (Zhao and Zhao, 1997), higher than the value of the pyroxene diorite, which is approximately equal to the strontium isotope value of early anhydrite and is lower than anhydrite + carbonate veins, indicating that the garnets are the product of magmatic and other fluids. The strontium isotope value of garnet in Shujiadian is similar to that in garnet of other deposits (e.g., Dongguashan: 0.709597; Zhao and Zhao, 1997), which suggests garnets in Shujiadian are the product of magmatic fluid interacting with a carbonate formation. As we have not found any carbonate formation at Shujiadian, we propose two possible explanations. Firstly, the Ca could be derived from xenoliths of deep carbonate strata (e.g., carbonate of deep Sinian, Cambrian, or Ordovician units). Residual time of strontium (1 Ma) in seawater is much longer than mixing time (1000a; McArthur et al., 1992), therefore strontium isotopes of seawater are homogeneous in any one epoch (McArthur et al., 1992). The Yangtze Craton was stable from the Sinian to Cambrian, with strontium isotope ratios of 0.708110– 0.708780 (Zhang, 1995), and strontium isotope ratios of global seawater vary from 0.707760 to 0.707920 (McArthur et al., 2001); we therefore conclude that deep carbonate xenoliths could not form garnets with the observed strontium isotopic composition. Secondly, calcium may have been derived from carbonate strata of Carboniferous, Permian, or Triassic that may have been intruded by the pyroxene diorite. However we have not observed these carbonate strata in the shallow parts of the porphyry or skarn orebodies (e.g., Dongguashan porphyry Cu–Au deposit). There is also a possibility that the previous shallow skarn orebodies or skarn alteration in Shujiadian deposit have been eroded off, which is supported by the presence of carbonate strata in the periphery of Shujiadian deposit (Fig. 2B) and the calculated palaeodepth of Shujiadian deposit based on fluid inclusion data.

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6.5. Comparison with the magmatic arc porphyry deposits Much work has been done on porphyry Cu deposits in magmatic arc settings (Sillitoe, 1972, 2010; Mitchell, 1973; Cooke et al., 2005; Richards, 2005), and the classic models have long been established (Lowell and Guilbert, 1970). The Shujiadian porphyry Cu deposit, which is located in an intracontinental setting, has some similar and some different characteristics when compared with porphyry Cu deposits in magmatic arc settings. In Shujiadian deposit, the ore-bearing pyroxene diorite does not have an aplitic groundmass texture, which is different from porphyry deposits in magmatic arc settings. This possibly suggests the pyroxene diorite magma ascended slowly or had low volatile content (Burnham, 1967). In magmatic arc settings, the sequence of intrusions generally changes from dioritic to monzonitic rocks, commonly with later latitic to rhyolitic or ‘‘quartz porphyry’’ intrusions (Lowell and Guilbert, 1970). In contrast, the late intrusive rock at Shujiadian is mafic pyroxene diorite, possibly suggesting mixing of magmas in a deep magmatic chamber, as documented for Dongguashan deposit (Wang et al., 2015). This is one of the key factors in generating the fertile parental magma in intracontinental settings (Wang et al., 2015), which may be the main difference from porphyry Cu deposits in magmatic arc settings. Mostly, the Shujiadian deposit alteration zone is defined by a core of potassic alteration with local Ca-silicate alteration, which is overprinted by a propylitic alteration zone and a feldspardestructive alteration zone, similar to most porphyry deposits in the magmatic arc environment (Lowell and Guilbert, 1970; Sillitoe, 1973; Beane and Titley, 1981). But potassic alteration at Shujiadian is characterized by a lot of biotite and minor K-feldspar, which may be due to the more mafic pyroxene diorite hostrock (Sillitoe, 2010). Argillic alteration at Shujiadian is mainly controlled by faults, and does not occur around the phyllic alteration as illustrated by Lowell and Guilbert (1970) in their model. It is difficult to distinguish argillic from phyllic alteration in the deposit, and we combine them as feldspar-destructive alteration (e.g., Yang et al., 2009). Metal assemblages in the Shujiadian deposit are characterized by Cu and Mo, again fitting with porphyry deposits in magmatic arc settings (Lowell and Guilbert, 1970). Ore-forming fluids exsolved from the magma in Shujiadian deposit contain CO2, which is an important characteristic of porphyry deposits in intracontinental settings, and different from porphyry deposits in the magmatic arc environment (Chen and Li, 2009). Porphyry deposits in magmatic arc settings may have shallow advanced argillic lithocaps and shallow intermediatehigh sulfidation epithermal systems, whereas they are absent at Shujiadian. In summary, Shujiadian deposit has fundamentally the same geological characteristics as porphyry deposits in magmatic arc

Sample no.

Host mineral

Silicate stage: Quartz + pyrite 1612-575 Quartz 1612-575 Quartz 1612-575 Quartz 1612-575 Quartz 1612-575 Quartz 1612-575 Quartz 1612-575 Quartz 1612-575 Quartz 1612-575 Quartz

Type

Proportion of bubble (%)

Homogenization temperature of H2O phases

settings. But there are some differences in style of alteration, composition of ore-forming fluid and ore-bearing rock between Shujiadian and porphyry deposits in magmatic arc settings. These could result from their different tectonic setting, magmatic source (Wang et al., 2015) and other factors (cf. Gustafson, 1978). These remain unclear at present and further work is required. 7. Conclusions (1) The Shujiadian deposit alteration zone has a core of potassic alteration and local Ca-silicate alteration, which is overprinted by a propylitic alteration zone and a feldspardestructive alteration zone. Biotite is the main potassic alteration mineral, which it typical of porphyry deposits in mafic protoliths. (2) The hydrothermal evolution of Shujiadian deposit can be divided into four episodes: immiscibility and overpressuring in the silicate stage, boiling in the quartz-sulfide stage and meteoric water mixing in the sulfide-carbonate stage. (3) Metallic elements in the deposit were mainly derived from magmatic hydrothermal fluids, and brine is probably the main carrier of metals. Decompression, and changes of oxygen fugacity and sulfur content are the main factors for Cu precipitation. (4) The Shujiadian deposit has the same basic geological characteristics as the porphyry deposits in magmatic arc settings, but some difference in style of associated intrusions, alteration and the composition of the ore-forming fluid. Acknowledgements Thanks to Liang Qi for trace element and strontium isotope analyses of anhydrite. We are indebted to two anonymous reviewers and the Editor-in-Chief Bor-ming Jahn for their detailed comments and constructive suggestions, which greatly improved the original draft of the manuscript. We are grateful for the editorial handling by Guest-Editor Yanbo Cheng. This research was financially supported by the National Natural Science Foundation of China (Grant Nos. 41320104003; 40830426; 41172084; 41172086), China Geological Survey (Grant Nos. 1212011121115; 1212011220369; SinoProbe-03-02-05), and Public Welfare Project of Anhui Province (Grant No. 2009-g-22), CODES Funding (Project No. P2.B1B.), Centre of Excellence, University of Tasmania. Appendix A. Analytical data of fluid inclusions from Shujiadian deposit

Hydrohalite melting temperature

vein, with K-feldspar alteration halo (from 575 m depth in drill hole 1612) I 10 404 I 30 428 I 25 389 I 18 380 I 25 404 I 15 482 I 25 524 I 25 397 I 15 438

Final melting temperature 13.6 10.2 5.1 17.6 16.7 14.3 19.4 13.8 18.6

Dissolution temperature of halite

Salinity

17.4 14.2 8.0 20.7 20.0 18.0 22.0 17.6 21.4

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S. Wang et al. / Journal of Asian Earth Sciences xxx (2014) xxx–xxx (continued) Sample no.

Host mineral

Type

Proportion of bubble (%)

Homogenization temperature of H2O phases

1612-575 1612-575 1612-575 1612-575 1612-575 1612-575 1612-575 1612-575 1612-575 1612-575 1612-575

Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz

III-b III-b III-b III-b III-b III-b III-b III-b III-b III-b III-b

9 18 15 10 15 10 10 5 10 10 15

359 378 370 388 415 345 325 305 370 329 572

Silicate stage: Quartz + pyrite 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz 1814-33 Quartz

Hydrohalite melting temperature

vein, with K-feldspar alteration halo (from 33 m depth in drill hole 1814) I 20 442 I 18 415 I 15 488 I 10 462 32.9 I 18 452 35.8 I 10 479 I 25 464 I 15 407 33.7 I 15 385 I 16 375 I 15 473 I 20 445 I 20 416 I 25 515 I 20 434 I 30 424 I 25 441 III-a 10 562 III-a 5 495 III-a 15 467 III-a 15 481 III-a 15 538 III-a 10 450 III-b 5 355 III-b 15 370 III-b 12 478 III-b 15 424 III-b 9 402 III-b 15 424 III-b 9 402 III-b 5 464 III-b 10 345 III-b 15 376 III-b 13 369 III-b 18 379 III-b 20 514 III-b 15 382 III-b 50 356 III-b 5 501 III-b 22 >600 II 85 496 II 70 539 II 75 528 II 75 533 II 75 520 II 75 495 II 75 450 II 80 >600 II 80 >600

Silicate stage: Quartz + K-feldspar vein (from 104 m depth in drill hole 601) 601-78 Quartz I 10 397 601-78 Quartz I 10 403 601-78 Quartz I 35 421 601-78 Quartz I 30 412 601-78 Quartz I 15 494 601-78 Quartz I 15 411 601-78 Quartz I 15 410 601-78 Quartz I 25 450 601-78 Quartz I 20 515 601-78 Quartz I 28 408 601-78 Quartz I 35 419

34 23 32.7 34

35.1 33.6

Final melting temperature

Dissolution temperature of halite

Salinity

526 501 521 512 515 421 400 >602 >600 >600 >600

61.3 58.8 60.9 60.1 60.7 50.1 48.0

13.1 12.2 9.2 14.1 13.4 14.2 12.8 19.7 20.8 12.2 13.6 15.4 7.9 17.9 17.8 15.3 16.9 350 399 425 450 454 420 429 529 594 432 439 490 481 533 464 464 476 401 552 462 430 >600 >601

17.0 16.2 13.1 17.9 17.3 18.0 16.7 22.2 22.9 16.2 17.4 19.0 11.6 20.9 20.8 18.9 20.2 42.4 47.4 50.3 53.3 53.7 49.7 51.0 61.8 70.2 51.6 52.2 58.1 56.8 63.2 54.7 54.9 56.1 48.1 66.1 54.6 51.1

0.5 1 0.5 0.7

0.9 1.7 0.9 1.2

0.6

1.1

1.3 1.2

2.2 2.1

13.7 12.9 12.6 14.3 14.9 13.7 11.3 12.1 15.5 16.2 13.6

17.5 16.8 16.5 18.0 18.6 17.5 15.3 16.1 19.1 19.6 17.4

(continued on next page) Please cite this article in press as: Wang, S., et al. Geological and geochemical studies of the Shujiadian porphyry Cu deposit, Anhui Province, Eastern China: Implications for ore genesis. Journal of Asian Earth Sciences (2014), http://dx.doi.org/10.1016/j.jseaes.2014.08.004

20

S. Wang et al. / Journal of Asian Earth Sciences xxx (2014) xxx–xxx

(continued) Sample no.

Host mineral

Type

Proportion of bubble (%)

Homogenization temperature of H2O phases

Hydrohalite melting temperature

Final melting temperature

601-78 601-78 601-78 601-78 601-78 601-78 601-78 601-78 601-78 601-78 601-78 601-78 601-78 601-78 601-78 601-78 601-78 601-78 601-78 601-78 601-78 601-78 601-78 601-78 601-78 601-78 601-78 601-78 601-78 601-78 601-78 601-78 601-78 601-78

Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz

I I I I I I I I I I I I I I I I I I I III-a III-a III-a III-a III-a III-a II II II II II II II II II

35 25 25 25 25 20 20 15 20 15 35 25 30 15 25 25 8 15 15 25 10 20 20 20 25 70 65 60 80 75 70 70 90 70

440 463 477 556 518 465 495 476 524 478 480 423 469 454 432 442 490 406 399 496 478 561 499 494 >600 546 528 586 520 576 512 596 562 >600

34.1 30.5 35.2 33.1

15.2 16.3 17.1 15.1 16.5 13.9 15.3 15.4 19.7 11.5 17.9 16.1 10.8 17.9 9.2 15.1 20.0 20.0 21.0

Quartz-sulfide stage: quartz + molybdenite vein (from 389 m depth in drill hole 1814) 1814-389 Quartz I 40 480 1814-389 Quartz I 50 399 1814-389 Quartz II 80 302 1814-389 Quartz I 15 295 1814-389 Quartz I 5 300 1814-389 Quartz II 70 457 1814-389 Quartz II 75 480 1814-389 Quartz II 70 296 1814-389 Quartz II 80 399 1814-389 Quartz II 50 325 1814-389 Quartz II 70 378 1814-389 Quartz II 50 415 1814-389 Quartz II 65 338 1814-389 Quartz II 65 354 1814-389 Quartz II 60 405 1814-389 Quartz II 75 346

30 30.4 33.4 39.6 37.2 31.8 36.5 30.9 35.2

Dissolution temperature of halite

402 408 400 398 436 440

Salinity 18.8 19.7 20.3 18.7 19.8 17.7 18.9 19.0 22.2 15.5 20.9 19.5 14.8 20.9 13.1 18.7 22.4 22.4 23.1 47.7 48.3 47.4 47.3 51.6

19 15.4 33.4 79 31.5 77.4 8.6 20.8 76.8

0.15 0.4 0.3 0.6 1.3 0.1 0.7 0.9 0.9

0.3 0.7 0.5 1.1 2.2 0.2 1.2 1.6 1.6

27.3 38.2 27.8 25.8 32.2 24.4 27.3 27.3 28.2 22.5 23 22

13.5 10.8 0.3

17.3 14.8 0.5

3.7 3.5 1.3

6.0 5.7 2.2

1.1 1.4 1.2 1.5 1.9

1.9 2.4 2.1 2.6 3.2

4.5

7.2

14.5 15.0 18.7 18.1 10.9 17.2

18.2 18.6 21.5 21.0 14.9 20.4

24.8 25.7

Quartz-sulfide stage: quartz + chalcopyrite + pyrite vein (from 739 m depth in drill hole 1814) 1814-739 Quartz I 18 364 1814-739 Quartz I 20 380 38.2 1814-739 Quartz I 10 375 25 1814-739 Quartz I 15 396 29.7 1814-739 Quartz I 10 365 20.3 1814-739 Quartz I 15 381

Quartz-sulfide stage: quartz + actinolite + pyrite + pyrrhotite + chalcopyrite vein (from 1104 m depth in drill hole 1816) 1816-1104 Quartz I 15 454 16.6 1816-1104 Quartz I 10 342 17.7 1816-1104 Quartz I 15 352 18.9 1816-1104 Quartz I 25 388 19.6 1816-1104 Quartz I 18 392 19.7 1816-1104 Quartz I 20 407 20.1 20.3 1816-1104 Quartz I 5 383 20.8 1816-1104 Quartz I 15 340 21.0 1816-1104 Quartz I 8 369 21.2 1816-1104 Quartz I 25 345 21.0 1816-1104 Quartz I 5 430 18.5 1816-1104 Quartz I 20 340 20.7 1816-1104 Quartz I 30 461 1816-1104 Quartz I 20 496 1816-1104 Quartz I 15 291 21.0

19.9 20.8 21.6 22.1 22.2 22.6 22.9 23.1 23.2 23.1 21.3 22.9

23.1

Please cite this article in press as: Wang, S., et al. Geological and geochemical studies of the Shujiadian porphyry Cu deposit, Anhui Province, Eastern China: Implications for ore genesis. Journal of Asian Earth Sciences (2014), http://dx.doi.org/10.1016/j.jseaes.2014.08.004

21

S. Wang et al. / Journal of Asian Earth Sciences xxx (2014) xxx–xxx (continued) Sample no.

Host mineral

Type

Proportion of bubble (%)

Homogenization temperature of H2O phases

Hydrohalite melting temperature

Quartz-sulfide stage: quartz + pyrite + chalcopyrite vein (from 313 m depth in drill hole 1816) 1816-313 Quartz I 25 431 20.3 1816-313 Quartz I 20 432 35.6 1816-313 Quartz I 35 389 33.8 1816-313 Quartz I 5 314 24.2 1816-313 Quartz I 10 316 1816-313 Quartz I 15 436 1816-313 Quartz I 10 363 40.1 1816-313 Quartz I 15 389 1816-313 Quartz I 15 380

Final melting temperature

Dissolution temperature of halite

Salinity

22.9 18.6 15.8 22.0 21.6 22.0 21.9 22.4 22.2

20.7 14.9 11.8 19.5 18.9 19.5 19.3 20 19.8

Quartz-sulfide stage: quartz + pyrite + molybdenite vein (pyrite and molybdenite occur disseminated in the veins) (from 453 m depth in drill hole 1816) 1816-453 Quartz I 13 413 20 1816-453 Quartz I 15 438 19.1 1816-453 Quartz I 15 419 41.2 12.7 1816-453 Quartz I 15 319 11.6 1816-453 Quartz I 20 449 17.7 1816-453 Quartz I 15 413 19.4 1816-453 Quartz I 18 351 18.4 1816-453 Quartz I 10 443 22.1 13.5 1816-453 Quartz I 13 445 23.4 15.4 1816-453 Quartz I 25 440 22.5 10.6 1816-453 Quartz I 28 389 19.4 1816-453 Quartz I 10 367 19.2 1816-453 Quartz II 75 356 25 3.7 1816-453 Quartz II 65 411 26.8 0.3 1816-453 Quartz II 70 374 26.8 0.1 1816-453 Quartz II 85 318

22.4 21.8 16.6 15.6 20.8 22.0 21.3 17.3 19.0 14.6 22.0 21.8 6.0 0.5 0.2

Quartz-sulfide stage: quartz + pyrite vein (from 52 m depth in drill hole 1816) 1816-52 Quartz I 15 356 1816-52 Quartz I 15 496 1816-52 Quartz I 20 438 1816-52 Quartz I 25 428 1816-52 Quartz I 25 445 1816-52 Quartz I 20 362 1816-52 Quartz I 25 480

13.4 15.1 13.1 11.8 10.2 12 12.8

17.3 18.7 17.0 15.8 14.2 16.0 16.7

Quartz-sulfide stage: barren quartz vein (from 823 m depth in drill hole 1816) 1816-823 Quartz I 10 416 1816-823 Quartz I 11 373 1816-823 Quartz I 5 356 1816-823 Quartz I 15 442 1816-823 Quartz I 5 304 1816-823 Quartz I 10 371 1816-823 Quartz I 15 378

20.8 18.4 19.1 20.8 20.1 20.9 18.3

22.9 21.33 21.8 22.9 22.4 23.0 21.2

5.1

8.0

11.1 6.3 4.1 3.9

15.1 9.6 6.6 6.3

11.7 5.5 11.7

15.7 8.6 15.7

11 11.5 11.5 13 11.8 10.4

15.0 15.5 15.5 16.9 15.8 14.4

6.4 4.6 1.2 2.9 0.5

9.7 7.3 2.1 4.8 0.9

Sulfide-carbonate stage: quartz + carbonate + pyrite vein (from 765 m depth in drill hole 2001) 2001-36 Quartz I 20 294 18.1 2001-36 Quartz I 10 248 15.8 2001-36 Quartz I 20 294 2001-36 Quartz I 15 290 22.6 2001-36 Quartz I 15 269 25.8 2001-36 Quartz I 30 288 24.7 2001-36 Quartz I 30 278 27.9 2001-36 Quartz I 30 267 2001-36 Quartz I 30 250 Sulfide-carbonate stage: pyrite + chalcopyrite + quartz vein (from 556 m depth in drill hole 1814) 1814-76 Quartz I 25 240 19.5 1814-76 Quartz I 25 244 19.7 1814-76 Quartz I 20 237 24.1 1814-76 Quartz I 20 257 17.9 1814-76 Quartz I 25 291 27.4 1814-76 Quartz I 20 240 32.4 1814-76 Quartz I 20 264 32.4 1814-76 Quartz I 25 240 34.6 1814-76 Quartz I 20 272 34.6 1814-76 Quartz I 20 252 27 Sulfide-carbonate stage: anhydrite superimposed on K-feldspar (from 504 m depth in drill hole 601) 601-504 Anhydrite I 25 241 30.5 601-504 Anhydrite I 20 193 39.7 601-504 Anhydrite I 20 182 25.9 601-504 Anhydrite I 20 157 28.7 601-504 Anhydrite I 30 207 35.8

(continued on next page)

Please cite this article in press as: Wang, S., et al. Geological and geochemical studies of the Shujiadian porphyry Cu deposit, Anhui Province, Eastern China: Implications for ore genesis. Journal of Asian Earth Sciences (2014), http://dx.doi.org/10.1016/j.jseaes.2014.08.004

22

S. Wang et al. / Journal of Asian Earth Sciences xxx (2014) xxx–xxx

(continued) Sample no.

Host mineral

Type

Proportion of bubble (%)

Homogenization temperature of H2O phases

Hydrohalite melting temperature

Final melting temperature

601-504 601-504 601-504 601-504 601-504 601-504

Anhydrite Anhydrite Anhydrite Anhydrite Anhydrite Anhydrite

I I I I I I

25 25 30 30 15 35

193 220 257 279 195 286

19.2 26.6

2.9 2.1

4.8 3.6

13.2 12.3 11.6 11.2 10.9 14.4 15.4 13.5 9.3 12.8

17.1 16.2 15.6 15.2 14.9 18.1 19.0 17.3 13.2 16.7

Anderko, A., Pitzer, K.S., 1993. Equation-of-state representation of phase equilibria and volumetric properties of the system NaCl–H2O above 573 K. Geochim. Cosmochim. Acta 57, 1657–1680. Anhui Provincial Bureau of Geology 321 Geological Team, 1995. Exploration Research on the Copper and Associated Ore Minerals of the Important Metallogenic Area along Yangtze River, Anhui Province (in Chinese). Audétat, A., Günther, D., 1999. Mobility and H2O loss from fluid inclusions in natural quartz crystals. Contrib. Miner. Petrol. 137, 1–14. Baker, T., Lang, J.R., 2003. Reconciling fluid inclusions, fluid processes and fluid source in skarns: an example from the Bismark skarn deposit, Mexico. Miner. Deposita 38, 474–495. Beane, R.E., Titley, S.R., 1981. Porphyry Copper Deposits. Part II. Hydrothermal Alteration and Mineralization. Economic Geology 75th Anniversary Volume, pp. 235–269. Becker, S.P., Fall, A., Bodnar, R.J., 2008. Synthetic fluid inclusions. XVII. PVTX properties of high salinity H2O–NaCl solutions (>30 wt% NaCl): application to fluid inclusions that homogenize by halite disappearance from porphyry copper and other hydrothermal ore deposits. Econ. Geol. 103, 539–554. Bertelli, M., Baker, T., 2010. A fluid inclusion study of the Suicide Ridge breccia pipe, Cloncurry district, Australia: implication for breccia genesis and IOCG mineralization. Precambr. Res. 179, 69–87. Bischoff, J.L., Pitzer, K.S., 1989. Liquid–vapor relations for the system NaCl–H2O: summary of the P–T–X surface from 300 °C to 500 °C. Am. J. Sci. 289, 217–248. Bloom, M.S., 1981. Chemistry of inclusion fluids: stockwork molybdenum deposits from Questa, New Mexico, and Hudson Bay Mountain and Endako, British Columbia. Econ. Geol. 76, 1906–1920. Bodnar, R.J., 1992. Can we recognize magmatic fluid inclusions in fossil hydrothermal systems based on room temperature phase relations and microthermometric behavior? Geol. Surv. Jpn. 279, 26–30. Bodnar, R.J., 1994. Synthetic fluid inclusions. XII. Experimental determination of the liquidus and isochores for a 40 wt.% H2O–NaCl solution. Geochim. Cosmochim. Acta 58, 141–148. Bodnar, R.J., 1995. Fluid-inclusion evidence for a magmatic source for metals in porphyry copper deposits. Mineral. Assoc. Canada Short Course Series 23, 139– 152. Bodnar, R.J., 2003. Re-equilibration of fluid inclusions. Mineral. Assoc. Canada Short Course Series 32, 213–230. Bodnar, R.J., Beane, R.E., 1980. Temporal and spatial variations in hydrothermal fluid characteristics during vein filling in pre-ore cover overlying deeply buried porphyry copper-type mineralization at Red Mountain, Arizona. Econ. Geol. 75, 876–893. Bodnar, R.J., Sterner, S.M., 1985. Synthetic fluid inclusions in natural quartz. II. Application to PVT studies. Geochim. Cosmochim. Acta 49, 1855–1859. Brown, P.E., Hagemann, S.G., 1995. MacFlincor and its application to fluid in Archean lode-gold deposits. Geochim. Cosmochim. Acta 59, 3943–3952. Burnham, C.W., 1967. Hydrothermal fluids at the magmatic stage. In: Barnes, H.L. (Ed.), Geochemistry of Hydrothermal Ore Deposits. Holt, Rinehart and Winston, New York, pp. 34–76. Burnham, C.W., 1985. Energy release in subvolcanic environments: implications for breccia formation. Econ. Geol. 80, 1515–1522. Chang, Y.F., Liu, X.P., Wu, Y.C., 1991. The Copper–Iron Belt of the Lower and Middle Reaches of the Changjiang River. Geological Publishing House, Beijing, pp. 1– 379 (in Chinese with English abstract).

Salinity

27.6

Sulfide-carbonate stage: pyrite + chalcopyrite + quartz vein (from 734 m depth in drill hole 2001) 2001-39 Quartz I 25 345 2001-39 Quartz I 25 353 2001-39 Quartz I 20 372 2001-39 Quartz I 15 333 2001-39 Quartz I 10 360 2001-39 Quartz I 15 358 2001-39 Quartz I 10 359 2001-39 Quartz I 15 363 2001-39 Quartz I 25 379 2001-39 Quartz I 25 338

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Dissolution temperature of halite

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Please cite this article in press as: Wang, S., et al. Geological and geochemical studies of the Shujiadian porphyry Cu deposit, Anhui Province, Eastern China: Implications for ore genesis. Journal of Asian Earth Sciences (2014), http://dx.doi.org/10.1016/j.jseaes.2014.08.004