Ore Geology Reviews 68 (2015) 59–78
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Geological characteristics and genesis of the Laoshankou Fe–Cu–Au deposit in Junggar, Xinjiang, Central Asian Orogenic Belt Qiang Li a, Shujun Lü b, Fuquan Yang a,⁎, Xinxia Geng a, Fengmei Chai c a b c
MLR Key Laboratory of Metallogeny and Mineral Assessment, Institute of Mineral Resources, Chinese Academy of Geological Sciences, 26 Baiwanzhuang Road, Beijing 100037, China Team 407 of Hunan Geological and Mineral Exploration and Developing Bureau, 308 Huaixi Road, Huaihua, Hunan 418000, China Xinjiang Key Laboratory for Geodynamic Processes and Metallogenic Prognosis of the Central Asian Orogenic Belt, Xinjiang University, 1230 Yanan Road, Urumqi, Xinjiang 830046, China
a r t i c l e
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Article history: Received 17 October 2014 Received in revised form 4 January 2015 Accepted 6 January 2015 Available online 13 January 2015 Keywords: Fluid inclusions Stable isotopes Geochronology Metallogenic process Fe–Cu–Au deposit Laoshankou China
a b s t r a c t The Laoshankou Fe–Cu–Au deposit is located at the northern margin of Junggar Terrane, Xinjiang, China. This deposit is hosted in Middle Devonian andesitic volcanic breccias, basalts, and conglomerate-bearing basaltic volcanic breccias of the Beitashan Formation. Veined and lenticular Fe–Cu–Au orebodies are spatially and temporally related to diorite porphyries in the ore district. Wall–rock alteration is dominated by skarn (epidote, chlorite, garnet, diopside, actinolite, and tremolite), with K–feldspar, carbonate, albite, sericite, and minor quartz. On the basis of ﬁeld evidence and petrographic observations, three stages of mineralization can be distinguished: (1) a prograde skarn stage; (2) a retrograde stage associated with the development of Fe mineralization; and (3) a quartz–sulﬁde–carbonate stage associated with Cu–Au mineralization. Electron microprobe analysis shows that garnets and pyroxenes are andradite and diopside-dominated, respectively. Fluid inclusions in garnet yield homogenization temperatures (Th) of 205–588 °C, and salinities of 8.95–17.96 wt.% NaCl equiv. In comparison, ﬂuid inclusions in epidote and calcite yield Th of 212–498 and 150–380 °C, and salinities of 7.02–27.04 and 13.4–18.47 wt.% NaCl equiv., respectively. Garnets yield values of 6.4‰ to 8.9‰ δ18Oﬂuid, whereas calcites yield values of −2.4‰ and 4.2‰ δ18Oﬂuid, and −0.9‰ to 2.4‰ δ13CPDB, indicating that the ore-forming ﬂuids were dominantly magmatic ﬂuids in the early stage and meteoric water in the late stage. The δ34S values of sulﬁdes range from −2.6‰ to 5.4‰, indicating that the sulfur in the deposit was probably derived from deep-seated magmas. The diorite porphyry yields LA–MC–ICP–MS zircon U–Pb age of 379.7 ± 3.0 Ma, whereas molybdenites give Re–Os weighted mean age of 383.2 ± 4.5 Ma (MSWD = 0.06). These ages suggest that the mineralizationrelated diorite porphyry was emplaced during the Late Devonian, coincident with the timing of mineralization within the Laoshankou Fe–Cu–Au deposit. The geological and geochemical evidence presented here suggest that the Laoshankou Fe–Cu–Au deposit is a skarn deposit. © 2015 Elsevier B.V. All rights reserved.
1. Introduction Skarns is a relatively simple rock type deﬁned by its mineralogy and is usually dominated by calc-silicate minerals (Ray and Webster, 1991; Meinert, 1992; Meinert et al., 2005). Skarns typically form during metamorphism and metasomatism, and both of these processes are generally spatially and temporally associated with felsic intrusions (Einaudi et al., 1981; Bowman et al., 1985; Brown et al., 1985; Brown and Essene, 1985; Kwak, 1987; Meinert, 1992). Skarns that contain a variety of metals (e.g., Fe, W, Cu, Pb, Zn, etc.) are called skarn deposits (Meinert et al., 2005). These deposits are one of the most important ore types in the world. Fe skarn deposits are the largest, locally containing signiﬁcant amounts of Cu and Au (Meinert et al., 2005), and constitute one of the most important source for high-grade Fe ores in China (Zhao et al., 2004). Fe deposits associated with extensive skarns in China can ⁎ Corresponding author. E-mail address: [email protected]
http://dx.doi.org/10.1016/j.oregeorev.2015.01.006 0169-1368/© 2015 Elsevier B.V. All rights reserved.
be classiﬁed into two types: submarine volcanogenic iron oxide (SVIO) deposits (Hou et al., 2014a) and intrusion-related skarn deposits (Mao et al., 2005; Yang et al., 2010a). The former type lacks any spatial link with intrusive rocks and is usually hosted by submarine volcanic rocks, such as Chagangnuoer in western Tianshan (Hong et al, 2012), Yamansu in Eastern Tianshan (Hou et al., 2014b) and Dahongshan in southern Yangtze carton (Zhao, 2010). The latter type of skarns has a clear spatial association with intrusions, such as Laoshankou and Qiaoxiahala Fe–Cu–Au deposits in the northern margin of the Junggar Terrane. The Laoshankou and Qiaoxiahala Fe–Cu–Au deposits are spatially related to diorite porphyry intrusions, contain signiﬁcantly different mineral associations from those of Cu- or Fe-dominated deposits, and comprise submarine volcanic wall rocks that differentiate them from typical skarn deposits (carbonate sedimentary wall rocks). Research has been focused on the geology, geochemistry, and isotopic composition of these two deposits (Yu et al., 2001; Yan et al., 2005; Zhou et al., 2005; Ying et al., 2006a, b). However, the origin of the Laoshankou
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deposit remains controversial, with previous works proposing various models of formation, including stratabound skarn, volcanogenic middle–low–temperature magmatic–hydrothermal, volcano–sedimentary and late-stage hydrothermal alteration, iron oxide–copper–gold (IOCG), and skarn types (Li, 2002; Li and Wang, 2009; Yan et al., 2005; Lu et al., 2009). As a result, it becomes the key point of contention whether the Fe–Cu–Au mineralization is genetically related to diorite porphyry or to coeval submarine volcanism. Here, we report zircon U–Pb ages for diorite porphyry that are spatially linked with Fe–Cu–Au mineralization and Re–Os ages for molybdenite that is associated with Cu–Au mineralization. The magnetite and skarn mineralogical characterization, detailed ﬂuid inclusion microthermometry, and S, C, O, and H stable isotope analysis undertaken during this study enable us to constrain the origin of the Laoshankou deposit.
Fault strikes 345°, dips approximately 80° to the east, crosscuts all other regional structures, and separates the Irtysh–Mayinebo Fault by 30 km. The Tuogeertuobie Au, Kalatongke Cu–Ni, and Laoshankou Fe–Cu–Au deposits are located at the intersections between the Kalaxiange'er–Ertai and other NW–SW-striking faults, indicating that the Kalaxiange'er–Ertai Fault has exerted an important control on metallogenesis in the study area. Early Devonian to Permian intrusions are widespread throughout the study area and are dominated by granite, diorite, diorite porphyry, and monzonitic granite. These intrusions occurred between 390–370 and 320–270 Ma (Han et al., 2004; R.H. Zhou et al., 2005; Zhang et al., 2006; Yang et al., 2008; G. Zhou et al., 2009). 3. Geology of the Laoshankou Fe–Cu–Au deposit 3.1. Stratigraphy
2. Regional geology The northern margin of the Junggar Terrane, located along the boundary between the Siberian and Kazakhstan–Junggar terranes, is an important part of the Central Asian Orogenic Belt (CAOB) (De Grave et al., 2013; Xiao et al., 2013), and hosts a series of Cu, Ni, Fe, and Au deposits, including maﬁc–ultramaﬁc Cu–Ni, porphyry Cu–(Mo), skarn Cu–Mo and Fe–Cu–Au, and orogenic Au deposits. These deposits are exempliﬁed by the large-scale Kalatongke Cu–Ni deposit; the medium-scale Qiaoxiahala Fe–Cu–Au, Halasu Cu, Yulekenhalasu Cu–(Mo), Suoerkuduke Cu–Mo, and Xilekuduke Mo– Cu deposits; and the small-scale Kekekuduke Au, Kalasayi Cu, Laoshankou Fe–Cu–Au, and Saerbulake Au deposits (Yan et al., 2003, 2005; C.M. Han et al., 2006; Han et al., 2007; Ying et al., 2006a; Yang et al., 2010b, 2012a; Z.H. Zhang et al., 2008; Wang et al., 2009), and indicate that the Junggar Terrane is highly prospective for mineral targeting. The Laoshankou Fe–Cu–Au deposit is located about 41 km southwest of Qinghe County, and occurs at the southeast of the Late Paleozoic Sawu'er island arc (Fig. 1), bordered to the north by the southern part of the Irtysh fault zone, which separates the Altay Orogenic Belt from the Junggar Terrane (He et al., 2004). The northern margin of Junggar is predominantly underlain by Late Paleozoic volcanic, sedimentary, and intrusive rocks. Early Paleozoic units in the area are rarely exposed, but where present, they are mainly represented by pyroclastic rocks metamorphosed to lower greenschist facies. The Upper Ordovician Jiabosaer Formation is dominated by shallow and paralic continental clastic and carbonate sediments, and the Devonian and Carboniferous units in the area are predominantly volcanic and volcano–sedimentary sequences. The Lower Devonian Tuoranggekuduke Formation consists of shallow marine sedimentary clastic rocks intercalated with carbonates and tuffs. The Middle Devonian units consist of intermediate–maﬁc volcanic lava, pyroclastic rocks, and cherts of the lower Beitashan Formation and intermediate–maﬁc volcanic rocks intercalated with minor sedimentary rocks of the upper Yundukala Formation. The upper Devonian Kaxiweng Formation consists of shallow and paralic continental clastic sediments. The overlying Lower Carboniferous Nanmingshui Formation includes paralic continental clastic rocks intercalated with slight alkaline volcanic rocks. The Upper Carboniferous Haerjiawu Formation contains continental clastic rocks, and is overlain by alternating continental and marine clastic rocks, and continental volcanic rocks of the Batamayineishan Formation (Li et al., 2003; Yan et al., 2003; C.M. Han et al., 2006; Zhang et al., 2009). The main faults in the study area strike predominantly NW–SE, and include the regional deep seated Irtysh–Mayinebo, Kalaxiange'er–Ertai, and Wulunguhe faults (Fig. 2A). These faults control the distribution of hydrothermal alteration and a series of intermediate–maﬁc volcanic rocks and felsic intrusions. The Irtysh–Mayinebo Fault is 45 km long, strikes NW–SW, dips at 60°–70° to the NE, and is a thrust fault that has been active over multiple periods. The Kalaxiange'er–Ertai
The lithologies exposed in the Laoshankou ore district are mainly the Middle Devonian Beitashan Formation and Quaternary sediments (Fig. 2B). The Beitashan Formation outcrops in the central ore district, and can be divided into three sections. The lower part consists of basalts, pyroxene–phyric basalts, basaltic andesites, basaltic volcanic breccias, andesites, andesitic volcanic breccias, marbles, siltstones, and minor picrites. The middle part of the formation is dominated by siltstones, tuffs, siliceous rocks, and breccia-bearing sandstones that are locally intercalated with basalts. The upper part is mainly comprised of siliceous rocks. The Tuoranggekuduke Formation outcrops in the northeastern part of the ore district, and comprises continental clastic rocks of the lower part and volcanic rocks, volcanic clastic rocks of the upper part, dominantly by basalts, andesites, basaltic or andesitic volcanic breccias, tuffs, marbles, and sandstones. The orebodies of the Laoshankou Fe–Cu– Au deposit are generally hosted by the lower part of the Beitashan Formation. 3.2. Structures and intrusions There are two main faults in the ore district, namely the northernmost Kalaxiange'er–Ertai Fault (F1) and the southernmost Shanqian Fault (F2). The former strikes about 340° and dips at 60°–80° to the east, and the latter is parallel to the former and dips at approximately 70°. A series of E–W-striking secondary faults are present between the two main faults and separate the ore district into a number of rhombshaped sectors. The location of volcanic ediﬁces and magmatism in general was controlled by NW–SE-striking faults, whereas the location of hydrothermal alteration and mineralization was controlled by E–Wstriking faults. NW–SE-striking intermediate–maﬁc and intermediate–felsic intrusive rocks are exposed in the central ore district, and are dominated by diorite, monzonite, monzonitic granite, granite, and syenite dykes and stocks, and diorite porphyry, andesite porphyry, dacite porphyry, syenite porphyry, and diabase dykes, all of which were emplaced into units of the Beitashan Formation. The Fe–Cu–Au mineralization is related to intermediate–felsic intrusions. 3.3. Mineralization The Laoshankou Fe–Cu–Au deposit can be divided into four parts (I, II, III, IV). This paper focuses on part IV (or Tuosibasitao Fe–Cu–Au deposit), which contains an upper series of Cu–Au-bearing magnetite orebodies (I1), and a lower series of Au-bearing copper orebodies (I2; Fig. 3). Fe–Cu–Au orebodies are associated with abundant skarndominated alteration at the contact between diorite porphyry and the surrounding wall rocks, and are closely spatially related to the diorite porphyry (Fig. 2B). The I1 orebody is veined and 200 m long, has an average thickness of 8.34 m, trends 290°–300°, and dip at 55°–70° to the north, and is hosted by skarn formed between andesitic volcanic
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Fig. 1. (a) Relationship of study area with the Central Asian Orogenic Belt (Modiﬁed from Jahn et al., 2000); (b) relationship of study area with the Chinese Altay (Modiﬁed from He et al., 2004); (c) regional geologic map of the southeastern Altay orogenic belt and northeastern Junggar terrane, northern Xinjiang. Modiﬁed from Zhang et al. (2009).
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Fig. 2. Geological map of the Laoshankou Fe–Cu–Au deposit. Modiﬁed from the Geophysical Prospecting Team of Xinjiang Geoexploration Bureau for non-ferrous Metals (2003).
breccias, conglomerate-bearing basaltic volcanic breccias, and the diorite porphyry. The underlying I2 orebody is lenticular and 110 m long, has an average thickness of 4.60 m, trends 290°, and dip at 18°–50° to the north. The average Fe, Cu and Au grades of the I1 orebody are 36.42 wt.%, 0.28 wt.%, and 0.49 g/t respectively. The average Cu and Au grades of the I2 orebody are 0.41 wt.% and 1.31 g/t (Geophysical
Prospecting Team of Xinjiang Geoexploration Bureau for non-ferrous Metals, 2003). The Laoshankou Fe–Cu–Au deposit contains four different types of ores according to mineralogy, mineral assemblages, and grade, as follows. (1) Low-grade garnet–diopside–magnetite ores that are dominated by conglomerate, irregular stockwork, and disseminated
Fig. 3. Cross-section along the No. 4 prospecting line in the Laoshankou Fe–Cu–Au deposit. Modiﬁed from the Geophysical Prospecting Team of Xinjiang Geoexploration Bureau for non-ferrous Metals (2003).
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Fig. 4. Features of the mineralization within the Laoshankou Fe–Cu–Au deposit. (A) Low-grade garnet–diopside–magnetite ore; (B) banded epidote-magnetite ore; (C) magnetite–pyrite– chalcopyrite ore; (D) quartz–calcite–chalcopyrite–pyrite ore; (E) breccia magnetite ore; (F) pyrite–chalcopyrite vein cross-cutting magnetite ore; (G) chalcopyrite of stockwork in reﬂected light; (H) hematite and chalcopyrite cross-cutting and replacing magnetite in reﬂected light; (I) Pyrite–chalcopyrite vein in reﬂected light.
magnetite mineralization in an andradite, diopside, and minor idocrase gangue (Fig. 4A). (2) Epidote–actinolite–tremolite–magnetite ores; this type of ore dominates the Tuosibasitao deposit and forms Fe-rich mineralization. The main ore mineral within this assemblage is magnetite in an epidote, tremolite, and actinolite gangue, and this ore is usually banded (Fig. 4B). (3) Magnetite–pyrite–chalcopyrite ores dominated by magnetite-hosted pyrite and chalcopyrite conglomerates, ﬁnegrained veins, and disseminations (Fig. 4C). (4) Quartz–calcite–chalcopyrite–pyrite–native gold ores that are dominated by pyrite, chalcopyrite, pyrrhotite, bornite, native gold, sphalerite, and galena in a calcite, sericite, and quartz gangue (Fig. 4D). Fe–Cu–Au ores within the Laoshankou deposit have massive, mottled, banded, disseminated, veined, and brecciated structures (Fig. 4B– F). Massive ores contain ﬁne-grained dense magnetite, whereas mottled ores contain magnetite, chalcopyrite, and pyrite, and banded ores contain alternating bands of mineralization and epidote-, chlorite- and actinolite-dominated skarn alteration. Brecciated ores contain magnetite and calcite, disseminated magnetite is hosted by both skarns and wall rocks, and veined ores consist of calcite, pyrite, chalcopyrite, and native gold. These ores have euhedral or subhedral granular, selforganized, and metasomatic relict textures (Fig. 4G–I); for example, chalcopyrite has both subhedral and anhedral granular textures, is associated with magnetite and pyrite, and cross-cuts magnetite mineralization, indicating that chalcopyrite formation slightly postdated magnetite formation. Pyrite is the most common sulﬁde and has euhedral, subhedral, and anhedral granular textures. Gold is present as native gold, has subhedral and anhedral microgranular textures, is generally 0.01 mm in size or less, and is hosted predominantly by pyrite,
chalcopyrite, and magnetite. The ore minerals within the deposit are magnetite, chalcopyrite, pyrite, hematite, pyrrhotite, and native gold, with lesser bornite, sphalerite, galena, and molybdenite, in gangue assemblages of epidote, chlorite, garnet, actinolite, calcite, tremolite, amphibole, diopside, K–feldspar, albite, apatite, sericite, and quartz. 3.4. Wall–rock alteration Skarn (epidote, chlorite, garnet, diopside, actinolite, and tremolite) is the most important wall–rock alteration in the Laoshankou Fe–Cu– Au deposit, followed by K–feldspar, albite, carbonate, sericite, and quartz (Fig. 5). Skarn is present in both the intrusion and orebodies (Fig. 5B), or the contact between the diorite (porphyry) and intermediate–maﬁc volcanic rocks (Fig. 5C). The epidote and chlorite alteration is closely related to magnetite mineralization, whereas copper and gold mineralization is associated with siliciﬁcation and carbonate alteration (Fig. 5D). 3.5. Mineralization stages Based on crosscutting relationships of ore veins, mineral assemblages, paragenetic sequence, ore structures, and wall–rock alteration characteristics, the ore-forming process can be divided into three stages: Stage I is a prograde skarn stage that is characterized by garnet and diopside (Fig. 4A). Both coarse- and ﬁne-grained garnet formed during this stage; the former has euhedral or subhedral granular texture, is 0.15–0.3 mm in size, has a light brown color, and is zoned (Fig. 5E),
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Fig. 5. Wall–rock alteration associated with the Laoshankou Fe–Cu–Au deposit. (A) Garnet is present as residual phenocrysts within epidote and magnetite; (B) epidote skarnization of diorite porphyry; (C) skarn is present in both the intrusion and Fe orebody; (D) pyrite- and garnet-bearing calcite vein cross-cutting epidote–magnetite ore; (E) zoned coarse-grained garnet in plane-polarized light; (F) magnetite-containing ﬁne-grained garnet in plane-polarized light; (G)–(I): characteristics of diopside, epidote, and magnetite in crossed polars.
operating conditions were 15 kV accelerating voltage, 1 μm beam diameter and 18 × 10−8 A beam current.
whereas the latter has subhedral and anhedral granular texture, is 0.05–0.15 mm in size, black-red in color, and contains magnetite inclusions (Fig. 5 F). Diopside is subhedral and granular, is present as short prismatic crystals that are 0.1–1.25 mm in size (Fig. 5G, H). Stage II is a retrograde stage that is marked by the formation of a magnetite–epidote–chlorite assemblage and minor tremolite and actinolite (Fig. 5A, I). This stage represents the main stage of magnetite mineralization, and is associated with euhedral or subhedral magnetite (partly replaced by hematite) that is 0.05–0.25 mm in size (Fig. 4H). Epidote is present as tabular crystals of 0.05–0.15 mm (Fig. 5I). Chlorite is schistose and tabular. Amphibole is granular or with long prismatic crystals, and occasionally with visible cleavage. Stage III is a quartz–sulﬁde–carbonate stage that is characterized by the formation of native gold–bearing sulﬁde assemblages, and minor disseminated, conglomerate, and veined pyrite and chalcopyrite within magnetite ores and calcite veins (Fig. 5D). This stage is also associated with sphalerite, galena, pyrrhotite, bornite, and molybdenite mineralization, and represents the main stage of copper and gold mineralization.
Garnet, epidote, and calcite samples were selected from the stages I, II and III respectively. The petrographic characteristics of ﬂuid inclusions were studied, and their types and assemblages were distinguished by optical microscopy of doubly polished sections (~ 200 to 300 μm in thickness). The microthermometric measurements of ﬂuid inclusions follow the procedure of Shepherd et al. (1985). Microthermometric analysis was undertaken using a LINKAM THMGS–600 programmable heating and freezing stage with a temperature range of − 196 °C to 600 °C at the China University of Geosciences, Beijing, China. The reproducibility of these measurements is ±0.1 °C below 30 °C, and heating and freezing temperatures were reproducible within ±1 °C and ±0.1 °C, respectively.
4. Samples and analytical techniques
4.3. Oxygen, hydrogen, carbon, and sulfur isotopes
4.1. Electron microprobe analysis
A total of 12 garnet and calcite samples were used for oxygen, hydrogen, and carbon isotope analysis, and 24 pyrite, 7 chalcopyrite, and 3 pyrrhotine samples in the ores and carbonate veins from the stage III were used for sulfur isotope analysis. Mineral separation was carried out at the Langfang Geochemical Laboratory in Hebei Province, China. All mineral separates were examined using a binocular microscope prior to isotope analysis to ensure 99% purity.
A total of 25 garnet, diopside and epidote samples were selected from the prograde skarn stage and the retrograde stage, and 10 magnetite samples from the retrograde stage. Electron microprobe analysis was undertaken using electron microprobe EPMA–1600 at the Geological Lab Center of the China University of Geosciences, Beijing, China. The
4.2. Microthermometric measurements
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Isotope analyses were performed using a MAT–253 EM spectrometer at the Isotopic Laboratory of the Institute of Mineral Resources, Chinese Academy of Geological Sciences (CAGS), Beijing, China. The analytical uncertainty was ± 0.2‰ for oxygen, carbon, and sulfur isotope analyses, and ±2‰ for hydrogen isotope analysis. Oxygen isotopic analysis was undertaken using the BrF5 method (Clayton and Mayeda, 1963), where pure garnet is reacted with BrF5 for 15 h to produce oxygen, before this oxygen is transferred to a CO2 transformation system at a temperature of 700 °C for CO2 collection after 12 min, as outlined in Mao et al. (2002). Hydrogen isotope analysis involved the release of water from ﬂuid inclusions by crushing; this water was then reacted with zinc for 30 min at a temperature of 400 °C to produce hydrogen (Coleman et al., 1982), which was transferred to a sample bottle ﬁlled with activated carbon after freezing in liquid nitrogen. Oxygen and carbon isotope analyses used an approach where calcite was reacted with phosphoric acid to release CO2 at 25 °C (McCrea, 1950). Chinese high-quality carbonate reference materials for carbon and oxygen (GBW04416 and GBW04417) were used as a working standard, and the δ13CPDB and δ18OPDB values of GBW04416 and GBW04417 were 1.61‰ and −11.59‰, and −6.06‰ and −24.12‰, respectively. The δ18OPDB value of calcite was directly measured from the CO2 obtained during the reaction between calcite and phosphoric acid, and the δ18OSMOW = 1.03086*δ18OPDB + 30.86 (Friedman and O'Neil, 1977) equation was used to determine δ18OSMOW values from δ18OPDB. Sulﬁdes were prepared using Cu2O as an oxidant (Robinson and Kusakabe, 1975), and sulfate minerals were puriﬁed to pure BaSO4 using a carbonate–zinc oxide semi–melt method, before SO2 was extracted using V2O5 as an oxidant; this SO2 was used for sulfur isotope analysis.
4.4. Zircon U–Pb dating One sample (diorite porphyry) for LA–MC–ICP–MS zircon U–Pb dating was collected from exposures in the Laoshankou ore district. Sample LSK–37 (90°05′50.2″E, 46°28′06.5″N) was collected from a diorite porphyry that is closely spatially related to Fe–Cu–Au mineralization. The diorite porphyry is massive and porphyritic, is tens to thousands of meters long, and is several to several hundreds of meters wide. The unit contains euhedral–subhedral, tabular phenocrysts of plagioclase (35%–45%) that are 0.5–8.5 mm in size and are partially altered to zoisite and subhedral, columnar phenocrysts of hornblende (5%–10%) that are 0.3–1.5 mm in size and are altered to chlorite, in a matrix (45%– 60%) that consists of approximately 0.15 mm plagioclase and minor amphibole that has undergone partial alteration to chlorite (Fig. 6A, B). Zircons were separated from an approximately 10 kg sample using conventional crushing and sieving techniques, and standard magnetic and heavy liquid separation, before puriﬁcation using cold HF and HNO3. Single zircon was hand–picked under a binocular microscope before mounting in epoxy resin. The sample mount was polished to expose the centers of individual zircons before examination under transmitted and reﬂected light and using cathodoluminescence (CL) imaging. CL imaging was undertaken at the Beijing Ion Microprobe Centre, Beijing, China, using a CAMECA SX–50 microprobe. U–Pb dating was undertaken by laser ablation multicollector inductively coupled plasma mass spectrometry (LA–MC–ICP–MS) at the Institute of Mineral Resources, CAGS, Beijing, China. Detailed operating conditions for the laser ablation system, the MC–ICP–MS instrument, and the data reduction approach used are given in Hou et al. (2009). Laser sampling was performed using a New Wave UP 213 laser ablation system, and a
Fig. 6. Samples of diorite porphyry and molybdenite from the Laoshankou Fe–Cu–Au deposit. (A) Massive diorite; (B) plagioclase and amphibole phenocrysts; (C)–(F) molybdenite of ﬁlmlike surface coatings, dissemination, ﬁne vein, and lumpy in epidote–chlorite–magnetite, pyrite–chalcopyrite-bearing magnetite ores, and in diorite porphyry.
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Thermo Finnigan Neptune MC–ICP–MS instrument was used to acquire ion-signal intensities. An array of four multi-ion-counters and three Faraday cups was allowed for simultaneous detection of 202Hg (on IC5), 204 Hg, 204Pb (on IC4), 206Pb (on IC3), 207Pb (on IC2), 208Pb (on L4), 232 Th (on H2), and 238U (on H4) ion signals. Helium was used as a carrier gas, and argon was used as a make-up gas and mixed with the carrier gas via a T-connector before entering the ICP. Each analysis incorporated a background acquisition of approximately 20–30 s (gas blank) followed by 30 s data acquisition from the sample. Off-line raw data selection, integration of background and analyte signals, timedrift corrections, and quantitative calibration for U–Pb dating were undertaken using the ICPMSDataCal (Liu et al., 2010). The Zircon GJ1 standard was used during external calibration of U– Pb dating, and was analyzed twice every 5–10 analyses. Time–dependent U–Th–Pb isotopic ratio drifts were corrected using a linear interpolation (with time) for every 5–10 analyses according to the variations of GJ1 (i.e., 2 zircon GJ1 analyses + 5–10 unknown analyses + 2 zircon GJ1 analyses; Liu et al., 2010). The U–Th–Pb isotopic ratios used for the GJ1 standard are those given by Jackson et al. (2004). The uncertainties on the preferred values used for the external GJ1 standard were propagated through during calculation of the ﬁnal uncertainties on these analyses. Common Pb corrections were not necessary as all analyzed zircons had low common 204Pb signals and high 206Pb/204Pb ratios. U, Th, and Pb concentrations were calibrated using the zircon M127 standard with U = 923 ppm, Th = 439 ppm, and Th/U = 0.475 (Sláma et al., 2008). Concordia diagrams were constructed and weighted mean age calculations were undertaken using the program Isoplot/Ex_ver3 (Ludwig, 1999). A plesovice standard zircon was also dated as an unknown sample and yielded a weighted mean 206Pb/238U age of 337 ± 2 Ma (2σ, n = 12), which is in good agreement with the recommended 206 Pb/238U age of 337.13 ± 0.37 Ma (2σ; Sláma et al., 2008). 4.5. Re–Os isotope analysis The Re–Os dating undertaken during this study used 6 molybdenite samples from the Laoshankou deposit. These samples not only occur as ﬁlm-like surface coatings and as disseminations in epidote–chlorite– magnetite and pyrite–chalcopyrite–bearing magnetite ores, but also occur as ﬁne veins within fractures and lumpy shapes in diorite porphyry (Fig. 6C–F). Mineral separates were prepared using standard heavy liquid and Frantz magnetic separation techniques before puriﬁcation under a binocular microscope. Re–Os analysis was undertaken at the National Research Center of Geoanalysis, Beijing, China. Samples were digested in a sealed Carius tube using the sample digestion and separation of Re and Os techniques outlined in Shirey and Walker (1995), Du et al. (2001), and Mao et al. (2003). Re isotopic ratios were measured using a TJA PQExCell ICP– MS, and Os isotopic ratios were determined using a Finnigan HR ICP– MS. During Re isotope analysis, mass numbers 185 and 187 were used for Re, and 190 was used to detect Os, whereas during Os isotope analysis, mass numbers 186, 187, 188, 189, 190, and 192 were used for Os, and 185 was used to detect Re. Total procedural blanks were ~ 2.8 pg for Re and ~0.1 pg for common Os, far lower than the Re and Os concentrations in the analyzed samples, indicating that the inﬂuence of contamination on Re and Os isotope measurements is negligible. All Re– Os ages are reported with 2σ uncertainties. 5. Results 5.1. Minerals chemistry 5.1.1. Garnet Nine garnet compositions of Laoshankou deposit are listed in Table 1 and plotted in Fig. 7A. These samples are dominated by Si, Al, Ca and Fe, with minor amounts of Ti, Mn, Cr and Mg. End-members mainly comprise andradite (31.79–93. 81 mol%, averaging 68.24 mol%), followed
by grossularite (2.73–59.23 mol%, averaging 23.74 mol%) and almandite (2.27–8.41 mol%, averaging 5.1 mol%), with less than 3 mol% spessartine, pyrope and uvarovite. 5.1.2. Pyroxene Six pyroxene compositions are given in Table 2 and graphically presented in Fig. 7B. The pyroxene from the Laoshankou deposit is diopside (Di)-dominated as demonstrated by the end-member calculations that mostly falls in the diopside ﬁeld (compositional variation ranging from Hd18Di81 to Hd46Di52), with minor johannsenite. 5.1.3. Epidote Representative EPMA data of 10 epidote samples are given in Table 3. The epidote is rich in FeOT (13.27–19.73 wt.%), Al2O3 (19.34–22.81 wt.%) and CaO (20.23–22.28 wt.%), lack in SiO2 (35.99–37.65 wt.%), TiO2 (0.07–0.38 wt.%), MnO (0.08–0.6 wt.%) and MgO (0.04–0.48 wt.%). 5.1.4. Magnetite EPMA data of 10 magnetite samples are given in Table 4 and plotted in Fig. 8. These samples are dominated by FeOT (89.47–93.51 wt.%), with minor amounts of TiO2 (0.02–0.26 wt.%), Al2O3 (0.12–0.93 wt.%), MnO (0.01–0.43 wt.%) and MgO (0.02–0.22 wt.%). 5.2. Fluid inclusion analysis 5.2.1. Inclusion types and characteristics This study focused on ﬂuid inclusions within 7 garnet, 4 epidote and 5 calcite samples of the stages I, II and III respectively. Fluid inclusions are abundant in garnet, epidote and calcite that are undeformed and have not been recrystallized. These ﬂuid inclusions can be classiﬁed into H2O–NaCl and H2O–CO2–(±CH4/N2)–NaCl types of ﬂuid based on phase relationships and chemical compositions at room temperature. H2O–NaCl ﬂuid inclusions can be further divided into liquid-rich, vapor-rich and daughter crystal-bearing polyphase inclusions. The H2O–CO2–(±CH4/N2)–NaCl ﬂuid inclusions occur as three-phase CO2 type inclusions. These four types of ﬂuid inclusions are here termed types I, II, III and IV, respectively. These types of ﬂuid inclusion and their characteristics are listed in Table 5 and shown in Fig. 9. Fluid inclusions in garnet of the stage I are dominantly type I inclusions, with minor type II and III inclusions. These inclusions are primary or pseudosecondary, and are interpreted to contain the ﬂuid from which garnet were precipitated during stage I. Type I inclusions contain both vapor and liquid at room temperature with a liquid volume of 50%–95%. Type II inclusions have a liquid volume of b 50%, and homogenize into the vapor phase when heated. Type III inclusions (Fig. 9C) comprise vapor, liquid and daughter crystals at room temperature with a liquid volume of 60%–95%. The daughter crystals occur generally as cube-shaped, with minor round ones. Fluid inclusions in epidote of the stage II belong to type I. These inclusions are isolated, thought to be primary, and interpreted as containing the ﬂuid from which epidote formed. Type I inclusions have ellipsoidal and irregular shapes, contain both vapor and liquid, with a liquid volume of 50%–97% and are 2–10 μm in size. Fluid inclusions in calcite of stage III are predominantly type I inclusions, with rare types III and IV. These inclusions are primary or pseudosecondary, and are interpreted to contain the ﬂuid from which calcite were precipitated. Type I inclusions contain both vapor and liquid with a liquid volume of 65%–97%, and coexist with other types of inclusions. Minor type III inclusions contain one or more cubic daughter crystals. These inclusions are isolated and rare. Type IV inclusions (Fig. 9H) are three-phase (liquid water, liquid CO2 and CO2-rich vapor) at room temperature, with some two-phase ﬂuid inclusions also nucleating a third phase during slight cooling below room temperature, contain 5%–10% CO2 by volume.
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Table 1 Electron microprobe analyses, ion proportions and end members of the representative garnet (element site allocation based on computer program of Lu, 2004). Sample no.
35.66 0.2 0.72 bd.l. 30.04 0.27 0.41 32.22 bd.l. bd.l. bd.l. 99.74
35.08 bd.l. 0.44 0.18 31.06 0.29 0.56 32.56 bd.l. bd.l. bd.l. 99.74
37.3 bd.l. 5.87 0.11 24.09 0.88 bd.l. 31.47 bd.l. bd.l. bd.l. 99.74
37.13 bd.l. 5.65 bd.l. 25.31 0.89 bd.l. 30.98 0.07 bd.l. bd.l. 99.74
38.23 0.05 14.58 0.27 12.26 1.13 0.01 32.8 bd.l. bd.l. bd.l. 99.74
36.72 2.23 8.83 bd.l. 18.84 0.62 0.36 31.38 0.12 bd.l. bd.l. 99.74
38.31 1.67 10.33 0.27 16.16 0.68 0.24 32.33 bd.l. 0.44 bd.l. 99.74
36.12 3.17 8.55 0.02 17.86 0.52 0.48 32.04 bd.l. bd.l. 0.02 99.74
Number of ions on the basis of 12 atoms of oxygen Si 3.07 2.95 Ti 0.00 0.01 Al 0.41 0.07 Cr 0.00 0.00 1.54 1.96 Fe3+ Fe2+ 0.18 0.12 Mn 0.08 0.02 Mg 0.01 0.05 Ca 2.66 2.85
2.90 0.00 0.04 0.01 2.01 0.13 0.02 0.07 2.88
3.01 0.00 0.56 0.01 1.43 0.20 0.06 0.00 2.72
3.00 0.00 0.54 0.00 1.46 0.25 0.06 0.00 2.68
3.00 0.00 1.35 0.02 0.64 0.17 0.08 0.00 2.76
2.94 0.13 0.83 0.00 1.12 0.15 0.04 0.04 2.69
3.01 0.10 0.96 0.02 0.96 0.10 0.05 0.03 2.72
2.90 0.19 0.81 0.00 1.13 0.07 0.04 0.06 2.75
Garnet components (mole fraction, %) Uvt 0.13 Adr 78.68 Prp 0.21 Sps 2.85 Grs 11.98 Alm 6.15
0.57 92.31 2.22 0.65 5.05 4.25
0.35 71.58 0.00 2.02 19.28 6.78
0.00 73.05 0.00 2.03 16.51 8.41
0.84 31.79 0.04 2.50 59.23 5.60
0.00 57.24 1.47 1.44 34.85 5.00
0.87 49.62 0.97 1.56 43.42 3.56
0.07 58.31 1.97 1.21 36.18 2.27
Oxide composition (wt.%) 37.84 SiO2 bd.l. TiO2 4.34 Al2O3 Cr2O3 0.04 ⁎ FeO 25.38 MnO 1.22 MgO 0.05 CaO 30.69 0.17 Na2O NiO bd.l. bd.l. K2O Total 99.74
0.00 93.81 1.66 0.62 2.73 3.91
Abbreviations: Uvt = uvarovite, Adr = andradite, Prp = pyrope, Sps = spessartine, Grs = grossular, Alm = almandine. bd.l. means lower than detecting limit. FeO⁎ as total iron. Analysis: Yin Jingwu.
5.2.2. Microthermometric results 126.96.36.199. H2O–NaCl ﬂuid inclusions. A total of 100 NaCl–H2O inclusions in 16 samples were analyzed. These inclusions contain 84 liquid-rich, 13 daughter-bearing and 3 vapor-rich inclusions, and the results are summarized in Fig. 10. The homogenization temperatures (Th) of the 32 liquid-rich inclusions in garnet are highly variable, ranging from 205 °C to 588 °C, with the majority clustering around 210 °C–310 °C. Ice-melting temperatures for these inclusions range from −14.2 °C to −5.8 °C, and salinities range from 8.95 to 17.96 wt.% NaCl equiv. (Bodnar, 1992), with a peak at approximately 16.5 wt.% (Fig. 10). Based on the NaCl–H2O reference table of Liu and Shen (1999) and the homogenization temperatures and salinities outlined above, we obtain aqueous phase densities of 0.60–1.00 g/cm3. 4 type III inclusions give homogenization temperatures of 241–588 °C, with 3 inclusions N550 °C, and all bubbles disappear before daughter crystals, indicating that they are supersaturated
inclusions. The only one type II inclusion analyzed in garnet yields homogenization temperatures of 284 °C and homogenizes into a vapor. The Th values of the 27 type I inclusions in epidote range from 212 °C to 498 °C, with the majority clustering around 210 °C–330 °C. Icemelting temperatures range from − 18.1 °C to − 4.4 °C, determining the ice-melting temperature of 7.02 to 27.04 wt.% NaCl equiv. and the aqueous phase densities of 0.60–0.95 g/cm3. 1 type III inclusions give homogenization temperature of 230 °C, with 2 inclusions N550 °C, and all bubbles disappear before daughter crystals. 2 type II inclusions analyzed in epidote yield homogenization temperatures of 370 °C and 562 °C. The Th values of the 25 type I inclusions in calcite range from 150 °C to 380 °C, with most deﬁning peaks at 150 °C and 230 °C. Salinity range from 13.4 to 18.47 wt.% NaCl equiv., and ﬂuid densities of 0.75–1.10 g/cm3. 3 supersaturated type III inclusions give homogenization temperatures of 150–250 °C with daughter crystals having higher melting temperatures of 370–376 °C.
Fig. 7. (A) Ternary diagrams showing compositional variations of garnet from the Laoshankou deposit; (B) ternary diagrams showing compositional variations of pyroxene from the Laoshankou deposit.
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Table 2 Electron microprobe analyses, ion proportions and end members of the representative pyroxene (element site allocation based on computer program of Lu, 2004). Sample no.
Oxide composition (wt.%) 53.59 53.12 SiO2 TiO2 0.08 0.06 0.45 0.55 Al2O3 bd.l. bd.l. Cr2O3 FeO⁎ 7.40 7.48 MnO 0.30 0.51 MgO 14.37 14.37 CaO 23.47 23.65 0.27 0.38 Na2O NiO bd.l. bd.l. K2O bd.l. bd.l. Mn/Fe 0.04 0.07 Total 99.94 100.13
52.4 bd.l. 0.73 0.09 8.59 0.48 13.98 22.37 0.29 0.05 0.02 0.06 99.00
53.54 0.16 0.41 0.21 7.86 0.45 14.80 22.50 0.39 bd.l. bd.l. 0.06 100.31
48.87 1.23 4.60 0.07 17.32 0.99 13.19 10.33 1.46 0.25 0.63 0.06 98.92
53.65 0.15 0.37 0.13 6.19 0.08 15.66 23.46 0.13 bd.l. bd.l. 0.01 100.25
Number of ions on the basis of 6 atoms of oxygen Si 1.99 1.98 1.98 1.99 Al(iv) 0.01 0.02 0.02 0.01 Al(vi) 0.01 0.00 0.01 0.00 Ti 0.00 0.00 0.00 0.00 Cr 0.00 0.00 0.00 0.01 3+ 0.02 0.07 0.04 0.04 Fe 2+ 0.21 0.17 0.23 0.21 Fe Mn 0.01 0.02 0.02 0.01 Mg 0.80 0.80 0.79 0.82 Ca 0.94 0.94 0.91 0.89 Na 0.02 0.03 0.02 0.03 K 0.00 0.00 0.00 0.00 Pyroxene components (mole fraction, %) Di 75.97 76.70 72.94 Hd 23.12 21.75 25.64 Jo 0.90 1.55 1.42
75.68 23.01 1.31
1.88 0.12 0.09 0.04 0.00 0.13 0.42 0.03 0.76 0.43 0.11 0.03
1.98 0.02 0.00 0.00 0.00 0.02 0.17 0.00 0.86 0.93 0.01 0.00
52.06 45.71 2.22
81.33 18.44 0.24
5.3. Stable isotope geochemistry 5.3.1. Sulfur isotope compositions A total of 33 sulﬁdes yielded δ34S values from −2.6‰ to 5.4‰, with a mean of 1.4‰, and with only three samples having δ34S values of N 5‰. The δ34S values of 24 pyrites range from −2‰ to 5.4‰ with a mean of 2.1‰, whereas the δ34S values of 7 chalcopyrite are slightly lower, varying from −2.6‰ to 2.5‰, with a mean of −0.9‰. The δ34S values of 2 pyrrhotine are 1.9‰ and 2.4‰ (Table 6; Fig. 11). 5.3.2. Carbon, oxygen, and hydrogen stable isotope analysis The carbon, oxygen, and hydrogen isotope compositions of samples from the stages I and III in the Laoshankou Fe–Cu–Au deposit are listed in Table 7 and plotted in Fig. 12. The δDSMOW values of 7 garnet samples vary from −110‰ to −84‰, whereas the δ18OSMOW values fall in a consistently narrow range from 5.2‰ to 6.8‰. Using the garnet–water fractionation equation of 1000lnα = 1.22 × 106 T−2 − 4.88 (Taylor, 1974) and the average homogenization temperature of ﬂuids inclusions of the same sample, the δ18Oﬂuid values of the ore-forming ﬂuids are calculated to be 6.4‰ to 8.9‰. The δ13CPDB, δDSMOW, and δ18OSMOW values of 5 calcites are −0.9‰ to 2.4‰, − 144‰ to − 92‰, and 7.1‰ to 13.3‰, respectively. These values, combined with the calcite–water fractionation equation (1000lnα = 2.78 × 106 T−2 − 3.39; O'Neil et al., 1969) and the average homogenization temperature of ﬂuid inclusions in the same calcite sample, yield δ18Oﬂuid values of between −2.4‰ and 4.2‰. 5.4. Zircon U–Pb and molybdenite Re–Os dating
Abbreviations: Di = diopside, Hd = hedenbergite, Jo = johannsenite. bd.l. means lower than detecting limit. FeO⁎ as total iron. Analysis: Yin Jingwu.
188.8.131.52. H2O–CO2–(± CH4/N2)–NaCl ﬂuid inclusions. A total of 6 type IV inclusions in calcite give solid CO2 melting temperatures of around − 57.7 °C, which is lower than the standard triple point (56.6 °C) of the pure CO2, suggesting minor contamination of CO2 with other volatiles, such as CH4 or N2 (Burruss, 1981). Clathrate melting temperatures range from − 3.9 °C to 3.2 °C. Using the approach of Bozzo et al. (1973) and assuming no effects associated with the presence of minor methane or nitrogen, we determined the salinities of 11.61–19.05 wt.% NaCl equiv. (Fig. 10). The partial CO2 and ﬁnal homogenization temperatures range from 23.9 °C to 30.7 °C and from 190 to 266 °C, respectively.
Zircons from the diorite porphyry with well-developed magmatic oscillatory zoning, low numbers of inclusions, and lacking in bright rims or ﬁssures were chosen for U–Pb dating. These zircons are transparent or translucent, have an adamantine luster. The majority of zircons are euhedral–subhedral, short prismatic, and bipyramidal in shape, and have a prolate axis length of 100–400 μm, with length/ width ratios of between 1:1.5 and 3:1. Cathodoluminescence images reveal some zircons have distinct contacts between rounded cores and rims that have well-developed oscillatory zoning (Fig. 13), suggesting that the rounded cores may be remnant zircons that were inherited or captured during transport before overprinting by the growth of a rim under magmatic conditions. A total of 19 zircons from diorite porphyry LSK–37 were used for U– Pb dating by LA–ICP–MS during this study; the results of these analyses are given in Table 8. These zircons have Th and U concentrations of 14 to 1967, and 36 to 1856 ppm, respectively, yielding Th/U ratios of N0.1 (0.18–2.39). This, together with the fact that these zircons are euhedral and oscillatory zoned, suggests that they are magmatic (Claesson et al., 2000; Rayner et al., 2005). These analyses can be divided into three groups, as follows. The ﬁrst group (spots 4, 6, 7, 8, 12, 14, 15, 16, and 19) has consistent 206Pb/238U ages between 387 and 376 Ma that
Table 3 Electron microprobe analyses of the representative epidote. Sample no.
Oxide composition (wt.%) SiO2 37.22 0.21 TiO2 21.76 Al2O3 bd.l. Cr2O3 FeO⁎ 14.84 MnO 0.09 MgO 0.08 CaO 22.28 Total 96.48
37.12 0.07 21.81 bd.l. 14.93 0.17 0.04 21.62 95.76
37.14 0.28 21.19 bd.l. 15.63 0.18 bd.l. 21.61 96.02
36.85 0.17 21.78 bd.l. 15.04 0.46 0.16 21.76 96.23
37.55 bd.l. 22.31 bd.l. 14.39 0.16 0.17 21.79 96.21
35.99 0.14 21.76 bd.l. 15.98 bd.l. 0.10 21.06 95.02
37.39 0.37 20.60 bd.l. 15.99 0.21 0.19 22.19 96.93
36.64 0.38 22.61 bd.l. 13.97 0.08 0.34 21.64 95.64
37.65 0.09 22.81 bd.l. 13.27 0.45 0.48 21.37 96.13
36.24 0.24 19.34 bd.l. 19.73 0.60 0.30 20.23 96.69
bd.l. means lower than detecting limit. FeO⁎ as total iron. Analysis: Yin Jingwu.
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Table 4 Electron microprobe analyses of the representative magnetite. Sample no.
Oxide composition (wt.%) 0.61 SiO2 0.02 TiO2 0.26 Al2O3 ⁎ FeO 90.89 MnO 0.43 MgO 0.04 CaO 0.05 bd.l. Na2O bd.l. K2O Total 92.29
1.98 0.08 0.93 89.47 0.28 0.22 0.21 bd.l. bd.l. 93.16
0.02 bd.l. 0.24 92.92 bd.l. 0.07 0.16 bd.l. bd.l. 93.41
1.19 0.26 0.41 91.94 bd.l. 0.10 bd.l. 0.28 0.14 94.32
1.25 bd.l. 0.37 91.36 0.34 bd.l. 0.09 bd.l. bd.l. 93.41
0.06 bd.l. 0.12 92.57 0.01 bd.l. 0.10 bd.l. bd.l. 92.87
0.44 0.23 0.46 91.66 0.39 0.08
0.65 bd.l. bd.l. 93.51 0.29 0.17 0.34 bd.l. bd.l. 94.96
1.35 bd.l. 0.24 90.36 0.15 0.02 0.27 bd.l. bd.l. 92.39
0.95 0.04 0.28 92.04 0.12 bd.l. 0.14 bd.l. bd.l. 93.56
bd.l. bd.l. 93.26
bd.l. means lower than detecting limit. FeO⁎ as total iron. Analysis: Yin Jingwu.
yielded a weighted mean age of 379.7 ± 3.0 Ma (MSWD = 0.48; Fig. 14). As such, we consider 379.7 ± 3 Ma to represent the timing of crystallization of the diorite porphyry. The second group (spots 2, 5, 10, 11) gave 207Pb/206Pb ages of 1946 Ma, 1780 Ma, 528 Ma and 1890 Ma, respectively, indicative of remnant zircons that were inherited or captured from older units. The third group (spots 1, 3, 9, 13, 17, 18) gave 206Pb/238U ages between 107 and 283 Ma; these zircons were probably affected by postmagmatic hydrothermal alteration, causing Pb loss that resulted in these younger ages. The results of Re–Os analysis of six molybdenite samples from the Laoshankou Fe–Cu–Au deposit are given in Table 9. These samples have fairly uniform Re and 187Re concentrations (442.0–629.4 and 277.8–395.6 ppm, respectively), and have Os and 187Os concentrations of 0.143–4.478 and 1782–2523 ppb, respectively. A regression analysis was performed using the equation t = 1/λ[ln (1 + 187Os/187Re)], with λ (decay constant of 187Re) = 1.666 × 10−11a−1 (Smoliar et al., 1996) and using the program ISOPLOT (Model 3; Ludwig, 1999) to calculate an isochron age of 383 ± 26 Ma (MSWD = 0.11; Fig. 15), consistent with the narrow range of model ages (between 381.0 ± 6.2 and 384.5 ± 5.2 Ma), and with a weighted mean age of 383.2 ± 4.5 Ma (MSWD = 0.063). Similar concentrations of Re and Os may cause the large error of isochron age.
6. Discussion 6.1. Skarn formation The magnetite within the Laoshankou deposit has similar TiO2 and Al2O3 but lower MnO and MgO concentrations compared with magnetite within contact metasomatic calcic skarn deposits (0.07–0.4 wt.% TiO2, 0.04–0.8 wt.% Al2O3, 0.1–2.15 wt.% MnO, and 0.35–11.51 wt.% MgO; Xu and Shao, 1979). Fig. 8 indicates that the majority of magnetite analyzed plots in the hydrothermal and calcic skarn type ﬁelds, with two samples plotting in a transitional area, indicating that the magnetite in the study area is genetically associated with the formation of the calcic skarns. Calcic skarns have a characteristic calc-silicate mineralogy dominated by Fe-rich prograde garnet and pyroxene and retrograde amphibole, epidote, actinolite, and chlorite (Purtov et al., 1989; Oyman, 2010), whereas magnesian skarns comprise Fe-poor forsterite, diopside, periclase, talc, and serpentine (Hall et al., 1988). Zhao (1987) suggested that the composition of pyroxenes in calcic skarn iron deposits was characterized by enrichments in diopside components (Di 50–90 mol%). Electron microprobe analysis shows that the garnets and pyroxenes are dominated by andradite and diopside compositions (Fig. 7), respectively, similar to the Mengku skarn iron deposit in the Altay Orogenic Belt (Xu et al., 2010). Nakano et al. (1994) used the relationship between pyroxene Mn/Fe ratios and the metals contained within 46 skarn deposits to show that skarn Cu–Fe deposits have low Mn/Fe ratios (b0.1). Pyroxenes within the Laoshankou deposit have Mn/Fe ratios of 0.013–0.068, with an average of 0.049, consistent with the classiﬁcation of Nakano et al. (1994). These data suggest that the Laoshankou Fe– Cu–Au deposit skarn assemblages belong to calcic skarn. 6.2. Age constraints on the timing of intrusion and mineralization
Fig. 8. TiO2–Al2O3–MgO diagram of magnetite from the Laoshankou deposit. 1 Granite area; 2 basalt area; 3 gabbro area; 4 peridotite area; 5 carbonatite area; 6 kimberlite area; 7-1 amphibolite area; 7-2 diorite area; 8 transition area; 9 hydrothermal and calcic skarn area; 10 hydrothermal and magnesian skarn area; 11 sedimentary metamorphism and hydrothermal superimposition area. Modiﬁed from Chen et al. (1987).
6.2.1. Intrusion age constraints The spatial relation between mineralization and diorite porphyry (Fig. 2) indicates that these intrusions might play an important role in the formation of the Laoshankou deposit. However, the relative timing of magma emplacement and hydrothermal mineralization is currently unknown. Here, we present a new zircon LA–ICP–MS U–Pb age of 379.7 ± 3.0 Ma for the diorite porphyry (Fig. 14). Previous research suggests that the Junggar Terrane underwent tectonism associated with ocean expansion, plate subduction, collision, and a postorogenic setting (X.C. Xiao et al., 1992; He et al., 1994; Jahn et al., 2001, 2004; Windley et al., 2002; Li et al., 2003; W.J. Xiao et al., 2004; Wang et al., 2006). These processes were associated with a significant amount of granitic magmatism that peaked at 390–370 and 320–270 Ma, with the most voluminous magmatism occurring between the late Carboniferous and the early Permian around 305 Ma. This is consistent with research by B.F. Han et al., 2006, who suggested that postcollision stage magmatism within the eastern Junggar Terrane
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Table 5 Inclusion types and characteristics of the Laoshankou Fe–Cu–Au deposit. Type
Aqueous liquid and vapor bubble
1–25, mainly 2–7
1–10, mainly 3–8
daughter crystal-bearing type
Vapor bubble and aqueous liquid Aqueous liquid, daughter crystal and vapor bubble
Elliptical, polygonal, columnar, irregular and negative crystal shapes Elliptical shapes
Three-phase CO2 type
Aqueous liquid, carbonic liquid and vapor
peaked at 325–310 and 305–280 Ma. Xiao et al. (1992) obtained an age of 394 Ma for an M-type granite related to ocean expansion stage of tectonism, and Chen and Jahn (2004) proposed an age of 294 Ma for collision and postcollision related granites. In addition, 381–374 Ma plate subduction-related granites are exposed in the Kalaxiange'er and Qiaoxiahala–Laoshankou areas within the northern margin of the Junggar, including a mineralized diorite porphyry (379 Ma), and a porphyritic quartz monzonite (375–374 Ma) in the Yulekenhalasu copper deposit (Yang et al., 2012b), 381–375 Ma granodiorite porphyries in the Halasu ore district (Zhang et al., 2006; Wu et al., 2008), a 376 Ma monzonite diorite porphyry in the Kalasayi area (Zhang et al., 2006), and a 377 Ma diorite porphyry in the Qiaoxiahala ore district (Li et al., 2014). Han et al. (1997) suggested that these intrusions were associated with the subduction of ancient oceanic crust of the Junggar Terrane. These intrusions have similar ages with the diorite porphyry
Degree of ﬁll
Number of inclusions observed
Elliptical, polygonal, columnar, and irregular shapes
Elliptical, and irregular shapes
Garnet, epidote, and calcite grains formed in stages I, II and III Garnet, and epidote grains formed in stages II and III Garnet, epidote, and calcite grains formed in stages I, II and III Calcite grains formed in stage III
(379.7 ± 3.0 Ma) in Laoshankou ore district, indicating that the diorite porphyry formed during the subduction stage of the Junggar Terrane. 6.2.2. Mineralization age constraints Molybdenite contains signiﬁcant amounts of Re, and the Re–Os geochronometer in molybdenite is remarkably robust and is not affected by prolonged (i.e., 2–8 Ma) and high-temperature (400 °C–500 °C) hydrothermal activity (Selby and Creaser, 2001; Selby et al., 2002), post-ore metamorphism and/or tectonism (Stein et al., 1998). Molybdenite Re–Os ages can directly record the timing of primary sulﬁde mineralization because the Re–Os system has closure temperatures that reach up to 500 °C (Suzuki et al., 1996; Selby and Creaser, 2001; Stein et al., 2001). Six molybdenite samples from the Laoshankou ore district yielded a weighted mean age of 383.2 ± 4.5 Ma (MSWD = 0.06), within error of the zircon LA–ICP–MS U–Pb age (379.7 ± 3.0 Ma) of
Fig. 9. Photomicrographs of ﬂuid inclusions from the Laoshankou Fe–Cu–Au deposit. (A) Vapor-rich inclusion in garnet; (B) liquid-rich inclusion in garnet; (C) daughter-bearing inclusion in garnet; (D) vapor-rich inclusion in epidote; (E) daughter-bearing inclusion in epidote; (F) liquid-rich inclusion in epidote; (G) daughter-bearing inclusion in calcite; (H) three-phase CO2-type inclusion in calcite; (I) liquid-rich inclusion in calcite.
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Fig. 10. Histogram of homogenization temperature and salinity for the ﬂuid inclusions in the Laoshankou Fe–Cu–Au deposit.
mineralization-related diorite porphyry. The molybdenite coexists with chalcopyrite, pyrite, and native gold-bearing sulﬁde assemblages (Fig. 6), suggesting that these minerals formed nearly contemporaneously. In addition, the retrograde stage derived Fe mineralization should form a little earlier than Cu–Au mineralization. This indicates that the Late Devonian 383.2 ± 4.5 Ma molybdenite age most likely represents the timing of formation of the Laoshankou deposit. The age of
molybdenite in the Laoshankou deposit is also consistent with Re–Os ages for the Halasu porphyry Cu deposit (376.3 ± 2.3 Ma; Wu et al., 2008; 378.3 ± 5.6 Ma; Yang et al., 2012b), the Yulekenhalasu porphyry Cu deposit (373.9 ± 2.2 Ma; Yang et al., 2012b) in the Kalaxiange'er area, and the Qiaoxiahala Fe–Cu–Au deposit (375.2 ± 2.6 Ma, Li et al., 2014). Porphyry ore systems are always related to subduction or postcollision (Sillitoe, 2010). Laoshankou skarn Fe–Cu–Au is close to Kalaxiange'er porphyry Cu belt in space and time, and they form in
Table 6 Compositions of the sulfur isotope in the Laoshankou Fe–Cu–Au deposit. Ser. no.
1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17
LSK11-17 LSK11-18 LSK11-23 LSK11-24 LSK11-25 LSK11-26 LSK11-27 LSK11-28 LSK11-34-1 LSK74 LSK75 LSK76 LSK77 LSK78 LSK79 LSK80 LSK81
Pyrite Pyrite Pyrite Pyrite Pyrite Pyrite Pyrite Pyrite Pyrite Pyrite Pyrite Pyrite Pyrite Pyrite Pyrite Pyrite Pyrite
1.8 −1.8 4.8 5.4 5.4 2.3 4.0 4.4 1.5 2.7 2.4 1.8 3.1 1.5 1.1 0.8 0.2
18 19 20 21 22 23 24 25 26 27 28 29 30 31 32 33
LSK83 LSK85 LSK83 LSK87 LSK88 LSK89 LSK96 LSK313 LSK314 LSK316 LSK319 LSK320 LSK11-15 LSK11-34 LSK223 LSK224
Pyrite Pyrite Pyrite Pyrite Pyrite Pyrite Pyrite Chalcopyrite Chalcopyrite Chalcopyrite Chalcopyrite Chalcopyrite Chalcopyrite Chalcopyrite Pyrrhotite Pyrrhotite
−2.0 2.8 0.1 0.6 0.4 1.3 5.1 −0.2 −0.2 −2.1 −2.4 −2.6 −1.3 2.5 2.4 1.9 Fig. 11. Histogram of sulfur isotope compositions of the Laoshankou Fe–Cu–Au deposit.
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Table 7 Carbon, hydrogen and oxygen isotopic data of the Laoshankou Fe–Cu–Au deposit. Ser. no.
δDV − SMOW/‰
δ18OV − SMOW/‰
1 2 3 4 5 6 7 8 9 10 11 12
LSK58 LSK60 LSK61 LSK62 LSK63 LSK65 ZK8004-2 LSK49 LSK51 LSK96 LSK97 LSK107
Garnet Garnet Garnet Garnet Garnet Garnet Garnet Calcite Calcite Calcite Calcite Calcite
−101 −110 −105 −94 −104 −98 −84 −144 −110 −92 −95 −117
6.4 5.8 5.9 5.2 6.2 6.8 6.3 11.8 7.1 12.7 12.9 13.3
263 304 255 332 443 394 250 156 192 207 207 166
the same tectonic setting, indicating that Laoshankou may be part of a porphyry system. 6.3. Ore ﬂuid evolution Kwak (1986) summarized the majority of the skarn ﬂuid inclusion literature prior to the mid-1980s, and concluded that a signiﬁcant proportion of skarns were associated with high-temperature (N700 °C) and high-salinity (N50 wt.% NaCl equivalent, with multiple daughter minerals) ﬂuids. Takeno et al. (1999) also determined that most skarn types have ﬂuid inclusion homogenization temperatures up to and exceeding 700 °C, whereas the majority of Cu and Zn skarns are associated with lower-temperature ﬂuids that have homogenization temperatures of 300 °C to 550 °C. Meinert et al. (2005) suggested that this difference derived from relatively shallow and distal geologic settings. Singoyi and Zaw (2001) and Zürcher et al. (2001) proposed that prograde garnet and pyroxene formed at 500 °C to 700 °C from ﬂuids with 50 wt.% NaCl equivalent salinities, whereas retrograde epidote and crosscutting quartz veins formed at several hundred degree lower temperatures and from ﬂuids with salinities of b 25 wt.% NaCl equivalent. However, Yao et al. (2014) argued that the peak of large amounts of ore-forming element precipitation usually occurred much later in spite of the magmatic ﬂuid with high temperature and salinity being the beginning of hydrothermal mineralization. For example, the formation of the Mengku skarn-type Fe–(Cu) deposit in the Altay area involved the precipitation of a large quantity of magnetite between prograde (241 °C–500 °C ﬂuids with salinities of 9.6–12.9 wt.% NaCl equivalent) and sulﬁde (145 °C– 382 °C ﬂuids with salinities of 1.2–13.0 wt.% NaCl equivalent) stages
Fig. 12. δD–δ18Oﬂuid diagram of the Laoshankou Fe–Cu–Au deposit. Data of primary magmatic water are from Sheppard (1986).
δ13CV − PDB/‰
−0.4 2.4 0.6 0.8 -0.9
7.04 7.02 6.41 6.75 8.70 8.94 6.72 0.10 -2.36 4.03 4.23 2.27
(Xu et al., 2010), Fe mineralization formed during the retrograde stage of mineralization of the Yinan skarn-type Au–Cu–Fe deposit in Shandong Province at temperatures of 290 °C–438 °C, and involved ﬂuids with salinities of 9.7–23.1 wt.% NaCl equivalent, and was followed by a quartz–sulﬁde stage associated with Cu–Au mineralization, which occurred at temperatures of 150 °C–300 °C and involved ﬂuids with salinities of 6.5–17.3 wt.% NaCl equivalent (Zhang et al., 2011). These ﬂuid inclusion studies documented Fe and Cu–Au mineralization associated with similar homogenization temperatures and ﬂuid salinities to those recorded during the formation of the Laoshankou Fe–Cu–Au deposit. In the ore–forming process of the Laoshankou deposit, stage I shows high to moderate temperatures (205 °C–588 °C) and moderate to low salinity (8.95–17.96 wt.% NaCl equivalent) ﬂuids, indicative of magmatic ﬂuids. This was followed by Stage II, the main stage of Fe mineralization, which shows slightly lower temperatures (212 °C–498 °C) and salinity (7.02–27.04 wt.% NaCl equivalent) ﬂuids, indicative of mixing of minor amount of low-salinity meteoric waters. Stage III, the main stage of Cu–Au mineralization, is characterized by type III and type IV ﬂuid inclusions that are assumed to have been trapped during the crystallization of calcite from lower-temperature (clustering around 150 °C–250 °C) and moderate-salinity (13.4–18.47 wt.% NaCl equivalent) NaCl–H2O–CO2–(± CH4/N2) ﬂuids, suggesting an increased amount of meteoric water participation. 6.4. Source of ore ﬂuids The δ18Oﬂuid values of the seven garnet samples range narrowly from 6.4‰ to 8.9‰, falling exactly in the range of δ18Oﬂuid values of magmatic ﬂuids (5.5‰–9.5‰; Sheppard, 1986), whereas the δDSMOW values vary from − 110‰ to − 84‰, slightly lighter than that of magmatic ﬂuids (− 80‰ – − 40‰; Sheppard, 1986). The δ18Oﬂuid and δDSMOW values of the ﬁve calcite samples range from −2.4‰ to 4.2‰ and −144‰ to −92‰, respectively, both are lower than that of magmatic ﬂuids. The H–O isotope compositions of Laoshankou deposit are similar to the Yinan Au–Cu–Fe deposit in Shandong Province (anhydrous skarn stage associated with garnet with δ18Oﬂuid value of 6.8‰ and δDSMOW value of − 73‰, and hydrous skarn–magnetite stage associated with magnetite and specularite with δ18Oﬂuid values of 7.9‰–11.6‰ and δDSMOW values of − 112‰ to − 82‰; Zhang et al., 2011), the initial mixing magmatic water in granite of the Au–Cu and Fe–Co series (δ18O = 6.0‰–9.0‰, δD = − 110‰ to − 65‰) as deﬁned by Zhang (1985), and the Qiaoxiahala Fe–Cu–Au deposit (calcite with − 0.72‰ to 7.67‰ δ18Oﬂuid and − 136‰ to − 111‰ δDSMOW in quartz–sulﬁde– carbonate stage; Li et al., 2014). Taylor et al. (1983), Taylor (1986) and Hedenquist (1994) suggested that H isotopes are strongly fractionated during early magma degassing, with late-formed OH–bearing igneous minerals representing the isotopic compositions of the degassed melt rather than that of the initial magmatic water. Shen et al. (2007) thought the magma degassing and mixture of meteoric water are the two factors causing the anomalously low δD values in Kuoerzhenkuola and Buerkesidai deposits at the
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Fig. 13. Cathodoluminescence images of zircons from diorite porphyry in the Laoshankou Fe–Cu–Au deposit. Analyzed spots are circled.
northwestern margin of the Junggar Terrane. As discussed above, temperature and salinity in garnet samples in stage I show characteristics of magmatic ﬂuids, whereas ﬂuids in calcite samples in stage III show
characteristics of mixing of meteoric water. Thus, the lighter δDSMOW values in garnet can be interpreted as result of magma degassing, whereas the lighter δDSMOW and δ18Oﬂuid values can be interpreted as
Table 8 Zircon LA–ICP–MS U–Pb isotopic data for the diorite porphyry from the Laoshankou Fe–Cu–Au deposit. Spot
1.1 2.1 3.1 4.1 5.1 6.1 7.1 8.1 9.1 10.1 11.1 12.1 13.1 14.1 15.1 16.1 17.1 18.1 19.1
Concentration Isotope ratios (10−6) Th
589 14 206 47 106 540 1884 168 213 79 71 596 815 76 1450 155 1967 1370 356
689 36 552 92 151 702 1856 336 614 451 171 654 552 204 1378 675 822 771 625
0.0500 0.1193 0.0531 0.0542 0.1088 0.0655 0.0697 0.0554 0.0553 0.058 0.1157 0.0555 0.1005 0.0565 0.0544 0.0554 0.0471 0.1009 0.0574
0.0005 0.0009 0.0004 0.0008 0.0004 0.0013 0.0012 0.0004 0.0004 0.0004 0.0005 0.0004 0.0039 0.0006 0.0002 0.0004 0.0004 0.0031 0.0004
Apparent ages (Ma) 207
0.1757 5.6835 0.2844 0.4474 4.4173 0.5658 0.5949 0.4721 0.3423 0.6555 5.0927 0.3639 0.3276 0.4745 0.4553 0.4619 0.1288 0.2406 0.4824
0.0027 0.121 0.0031 0.0084 0.0478 0.0194 0.0167 0.006 0.0053 0.0087 0.0562 0.0065 0.0163 0.0088 0.0064 0.0052 0.0016 0.0094 0.0091
0.0256 0.3441 0.039 0.0601 0.2945 0.0608 0.0605 0.0618 0.0449 0.0818 0.3194 0.0605 0.0219 0.0611 0.0606 0.0604 0.0198 0.0167 0.0607
0.0004 0.0062 0.0004 0.0008 0.0032 0.0009 0.0007 0.0007 0.0006 0.0009 0.0035 0.0006 0.0003 0.001 0.0008 0.0005 0.0002 0.0002 0.001
0.0006 0.0181 0.0008 0.0028 0.0041 0.0012 0.001 0.0011 0.0008 0.0029 0.0058 0.001 0.0005 0.0023 0.001 0.002 0.0004 0.0004 0.0012
0.0001 0.0053 0.0002 0.0006 0.0008 0.0002 0.0002 0.0003 0.0002 0.0008 0.0014 0.0002 9.2668 0.0005 0.0002 0.0005 8.1372 0.0001 0.0003
195 1946 332 389 1780 791 920 428 433 528 1890 432 1635 472 391 428 54 1640 506
20 13 14 33 7 44 42 15 15 12 8 19 72 22 5 21 21 58 10
164 1929 254 375 1716 455 474 393 299 512 1835 387 288 394 381 386 123 219 340
2 18 2 6 9 13 11 4 4 5 9 4 12 6 4 4 2 8 6
163 1906 246 376 1664 380 379 387 283 507 1787 379 140 382 379 378 127 107 380
2 30 3 9 16 5 4 4 4 5 17 4 2 6 5 3 1 1 6
12 362 17 57 82 23 20 23 16 58 117 20 10 46 21 40 7 9 25
3 104 4 12 16 5 4 5 4 15 28 4 2 10 4 10 2 2 6
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interpreted the different mean values and clusters for the different minerals most likely relate to degassing during magma ascent. Therefore, the average δ34S value (1.4‰) of all sulﬁdes, similar with that of Qiaoxiahala deposit (–0.1‰, Li et al., 2014), indicates that the sulfur in the Laoshankou and Qiaoxiahala Fe–Cu–Au deposit is from a mantle-derived source (0 ± 3‰; Hoefs, 2004). Previous studies (Foster et al., 1996; Ruiz and Mathur, 1999; Mao et al., 2002) indicated that the Re–Os isotope system had been recognized as a geochemical tool, not only for directly dating mineralization, but also as a tracer for metal source area. For example, Stein et al. (2001) obtained the molybdenites from deposits with a mantle component that have signiﬁcantly higher Re contents than those derived from a crustal source. Mao et al. (1999) proposed that Re concentrations decrease from mantle to mixed mantle–crustal and to crustal sources, from several hundred ppm, to tens of ppm, and to several ppm, respectively, on the study of the some major molybdenite-bearing deposits in China; Z.H. Zhang et al. (2008) and Han et al. (2007) suggested that crustal contamination was an important factor for the genesis of magmatic sulﬁdes in the Kalatongke and Huangshan East Cu–Ni sulﬁde deposits in northern Xinjiang through the study of initial 187Os/188Os values. The ω(Re) of molybdenite from the Laoshankou Fe–Cu–Au deposit ranges between 442.0 ppm and 629.4 ppm, similar to Cu deposits in other parts of the world (Berzina et al., 2005), and also prove that some mantle source was probably involved in the ore-forming process of Laoshankou Fe–Cu–Au deposit as sulfur isotope did.
Fig. 14. Zircon LA–ICP–MS U–Pb concordia diagrams of diorite porphyry from the Laoshankou Fe–Cu–Au deposit.
result of mixture of magmatic water and temporally increasing amounts of meteoric water. Brown et al. (1985) and Baker and Lang (2003) indicated that δ13C values in calcite ranged from typical sedimentary δ13C values in limestone away from skarn to typical magmatic values in calcite interstitial to prograde garnet and pyroxene. The δ13CPDB values of the 5 calcite samples from calcite veins in the Laoshankou deposit range from −0.9‰ to 2.4‰, with an average of 0.5‰, higher than the isotopic compositions of organic carbon within sediments (average of approximately − 25‰; Clark et al., 2004), igneous rocks (− 30‰ to − 3‰; Hoefs, 2004), and mantle carbon (−5‰ ± 2‰; Hoefs, 2004). The values fall in the range of marine carbonate (−1‰ to 2‰; Rollinson, 1993), suggesting that the carbon was predominantly derived from marine carbonate during magmatic–hydrothermal activity. Han and Ma (2003) proposed that the carbonate carbon is derived from convective circulation of meteoric water in carbonate rocks. The δ13C values again support the presence of meteoric water in the late stage of ore–forming ﬂuid.
6.6. Genesis of deposit Early prograde stage (stage I) is associated with anhydrous minerals, such as garnet and pyroxene, which formed from high-temperature, hypersaline magmatic ﬂuids, whereas the later retrograde stage (stage II) is associated with hydrous minerals such as epidote, amphibole, and chlorite that partially or entirely overprint and replace early skarn minerals, and formed from lower-temperature and low-pH-magmatic or variable-salinity (high to low) ﬂuids (Kwak, 1986; Singoyi and Zaw, 2001; Meinert et al., 2003). In the Laoshankou deposit, abundant skarns (in which epidote predominates) are spatially associated with diorite porphyry, and stage I assemblages are characterized by garnet and pyroxene that were replaced by stage II epidote–chlorite assemblages, forming residual garnet and pyroxene phenocrysts under a relatively lower temperature. These skarns are derived from the metamorphism and metasomatism of intermediate–maﬁc volcanic wall rocks and the diorite porphyry, and are somewhat atypical compared with those which are related to limestone wall rocks. However, Meinert et al. (2005) summarized that limestone is not necessarily required to form skarn, because skarns can form in almost any rock type, including volcanic rocks. Transfer of heat, ﬂuid, and metals from a cooling magma to the surrounding rocks can lead to zoned systems in both space and time in most large skarn deposits, and typical skarn zonation usually occurs at the skarn-marble contact (Meinert et al., 2005). Laoshankou deposit has no clear skarn zonation, and contains obviously much more epidote than other skarn minerals. This is probably caused by volcanic wall rocks.
6.5. Source of ore–forming material The sulfur isotopic compositions of hydrothermal minerals depend on the δ34S values of the source and the physicochemical conditions under which S-bearing materials were deposited from hydrothermal ﬂuids (Ohmoto, 1972; Ohmoto and Rye, 1979). The average δ34S of minerals would represent the total sulfur in the hydrothermal ﬂuid in a relatively simple assemblage (Hoefs, 2004). In Laoshankou deposit, sulﬁde minerals comprise pyrite, chalcopyrite, and pyrrhotite, suggesting a relatively simple paragenesis. 33 samples clearly show that the δ34S ratios of sulﬁdes from the ores are clustered from 0‰ to 3‰ (Fig. 11), with average values of 1.4‰, 2.1‰, − 0.9‰ and 2.15‰ for the whole, pyrite, chalcopyrite, and pyrrhotite, respectively. Chen and Wang (2004)
Table 9 Re-Os isotope data for molybdenites from the Laoshankou Fe–Cu–Au deposit. Sample no.
LSK12-22 LSK12-19 LSK12-21 LSK12-22-1 LSK12-23 LSK12-23-1
597.484 487.406 629.428 585.377 449.582 442.004
4.363 4.171 5.660 6.105 3.905 3.310
0.277 4.478 0.143 0.342 1.272 0.890
0.293 6.146 0.320 0.361 1.511 0.229
375.531 306.344 395.608 367.921 282.571 277.809
2.742 2.622 3.557 3.837 2.454 2.081
2413 1951 2523 2361 1812 1782
19 22 21 18 17 14
384.5 381.0 381.6 383.9 383.8 383.8
5.2 6.2 5.6 5.9 5.8 5.2
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Fig. 15. Re–Os isochron diagram for molybdenites and weighted average model ages of the Laoshankou Fe–Cu–Au deposit.
In a word, the geological features above indicate that the Laoshankou belongs to skarn-type deposit despite the lack of zonation and limestone wall rocks. 6.7. Processes of mineralization Zhang et al. (2009) used the geochemistry of volcanic rocks to suggest that the Middle Devonian Beitashan Formation formed in an immature volcanic arc generated by the northward subduction of the ancient Junggar Ocean, consistent with the presence of boninites, picrites, Nbenriched basalts, and intermediate–maﬁc volcanic rocks within this formation (Yu et al., 1995; Chen et al., 2004). Diorite (porphyry) and granites intrusions were emplaced into volcanic rocks of the Beitashan Formation during the Late Devonian, and Zhang et al. (2006) and Chai et al. (2012) considered that the magmas that formed these intrusions were derived from the melting of subducted slab material derived from the ancient Junggar Ocean. The Fe–Cu–Au mineralization within the Laoshankou ore district formed as a result of the subduction. Layne and Spooner (1991) and Hemley et al. (1992) suggested that transport of metallic components is directly related to the physicochemical conditions of ore-forming ﬂuids, Rose and Burt (1979) and Simon et al. (2004) demonstrated that Fe is a more mobile element at high temperatures. Simon et al. (2004, 2006) proposed that signiﬁcant quantities of Fe could be transported as FeCl2 together with CuCl in brines, and other research (Holland, 1972; Candela and Holland, 1984; Urabe, 1985; Hedenquist and Lowenstern, 1994) suggested that Cu and Au are signiﬁcantly enriched in ﬂuid phases separated from melts and the partitioning process is substantially affected by the H2O content and the chlorine concentration in the magma. Eadington (1985) and Heinrich (1990, 1995) indicated that a number of processes can cause the precipitation of metals from ore-forming ﬂuids, including reaction of these ﬂuids with host rocks, boiling, redox-coupled precipitation, mixing of ﬂuids, or a combination of some of the above. Here, we used the results obtained during this study to outline the processes of mineralization involved in the formation of the Laoshankou deposit. Initial high-temperature contact metamorphism was caused by the intrusion of the diorite porphyry, altering the surrounding volcanic rocks into garnet- and pyroxene-bearing prograde anhydrous skarns. High-temperature and high-salinity magmatic ﬂuids were segregated from the deep-seated dioritic magma; these ﬂuids contained signiﬁcant quantities of Fe that was probably present as FeCl2 complexes. The oreforming ﬂuids ascended and mixed with small amounts of meteoric water, causing a decrease in both temperature and pressure, and resulting in the replacement of anhydrous silicates with an epidote– chlorite assemblage, the precipitation of large amounts of magnetite
under relatively oxidizing conditions, and favoring the Cu and Au during retrograde skarn formation. The ﬁnal late quartz–sulﬁde–carbonate stage of mineralization involved the mixing of ore-forming magmatic ﬂuids with more meteoric water (in signiﬁcant excess of the mixing during stage II), leading to rapid cooling and dilution and the deposition of sulﬁdes and Au under relatively reducing conditions. 7. Conclusions The Laoshankou Fe–Cu–Au deposit is hosted in the Beitashan Formation, and contains vein and lenticular orebodies. Wall–rock alteration is dominated by skarns, with K–feldspar, albite, carbonate, sericite, and minor quartz alteration. The ore-forming process comprises (1) a prograde skarn stage; (2) a retrograde stage; and (3) a quartz–sulﬁde–carbonate stage. Fe and Cu–Au mineralization formed during stages (2) and (3), respectively. Electron microprobe analysis shows that garnets and pyroxenes are andradite-dominated and diopside-dominated, respectively, indicating that the skarns are calcic. The prograde skarn stage indicates high to moderate temperature (205–588 °C) and moderate to low salinity (8.95–17.96 wt.% NaCl equiv.) ﬂuids. In comparison, the retrograde stage reveals lower temperature (212–498 °C) and salinity (7.02–27.04 wt.% NaCl equiv.) ﬂuids. Quartz–sulﬁde–carbonate stage is characterized by CO2-type inclusions and shows ﬂuids of lower temperature (cluster in 150–250 °C) and moderate salinity (13.4–18.47 wt.% NaCl equiv.). C–H–O isotope compositions suggest that the prograde skarn stage comprises dominantly magmatic ﬂuids, whereas quartz–sulﬁde–carbonate stage consists mainly of meteoric water. S and Re isotope analysis shows that the metallogenic materials were probably derived from mantle-related magmas. The diorite porphyry yielded zircon U–Pb ages of 379.7 ± 3.0 Ma, whereas molybdenite gave Re–Os age of 383.2 ± 4.5 Ma (MSWD = 0.06), suggesting that the mineralization formed during the Late Devonian. The geological and geochemical evidence presented in this paper suggest that the Laoshankou Fe–Cu–Au deposit is a skarn deposit. Acknowledgments This research was jointly supported by the Ministry of Land and Resources Public Welfare Industry Special Funds for Scientiﬁc Research Project (Grant No. 201211073), and the National Key Technologies R&D Program (Grant No. 2011BAB06B03–02). Jingwu Yin and Huiyan Zhu are thanked for assistance during electron microprobe analysis
Q. Li et al. / Ore Geology Reviews 68 (2015) 59–78
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