Metallogeny of the Shilu Fe–Co–Cu deposit, Hainan Island, South China: Constraints from fluid inclusions and stable isotopes

Ore Geology Reviews 57 (2014) 351–362

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Metallogeny of the Shilu Fe–Co–Cu deposit, Hainan Island, South China: Constraints from fluid inclusions and stable isotopes Jinjie Yu a,⁎, Jingwen Mao a, Fuxiong Chen b, Yonghui Wang b, Linrui Che c, Tiezhu Wang a, Jiang Liang b a b c

Key Laboratory of Metallogeny and Mineral Assessment, Institute of Mineral Resources, Chinese Academy of Geological Sciences, Beijing 100037, People's Republic of China Hainan Mining United Co. Ltd., Hainan 572700, People's Republic of China Faculty of Geosciences, China University of Geosciences, Beijing 100083, People's Republic of China

a r t i c l e

i n f o

Article history: Received 8 April 2013 Received in revised form 14 August 2013 Accepted 20 August 2013 Available online 2 September 2013 Keywords: Mineralization stage Fluid inclusions Stable isotopes Metallogeny Shilu deposit South China

a b s t r a c t The Shilu deposit is a world-class Fe–Co–Cu orebody located in the Changjiang area of the western part of Hainan Island, South China. The distribution of Fe, Co, and Cu orebodies is controlled by strata of the No. 6 Formation in the Shilu Group and the Beiyi synclinorium. Based on a petrological study of the host rocks and their alteration assemblages, and textural and structural features of the ores, four mineralization stages have been identified: (1) the sedimentary ore-forming period; (2) the metamorphic ore-forming period; (3) the hydrothermal mineralization comprising the skarn and quartz–sulfide stage; and (4) the supergene period. The fluid inclusions in sedimentary quartz and/or chert indicate low temperatures (ca. 160 °C) and low salinities from 0.7 to 3.1 wt.% NaCleq, which corresponds to densities of 0.77 to 0.93 g/cm3. CO2-bearing or carbonic inclusions have been interpreted to result from regional metamorphism. Homogenization temperatures of fluid inclusions for the skarn stage have a wide range from 148 °C to 497 °C and the salinities of the fluid inclusions range from 1.2 to 22.3 wt.% NaCleq, which corresponds to densities from 0.56 to 0.94 g/cm3. Fluid inclusions of the quartz–sulfide stage yield homogenization temperatures of 151–356 °C and salinities from 0.9 to 8.1 wt.% NaCleq, which equates to fluid densities from 0.63 to 0.96 g/cm3. Sulfur isotopic compositions indicate that sulfur of the sedimentary anhydrite and Co-bearing pyrite, and the quartz–sulfide stage, was derived from seawater sulfate and thermochemical sulfate reduction of dissolved anhydrite at temperatures of 200 °C and 300 °C, respectively. H and O isotopic compositions of the skarn and quartz– sulfide stage demonstrate that the ore-forming fluids were largely derived from magmatic water, with minor inputs from metamorphic or meteoric water. The Shilu iron ore deposit has an exhalative sedimentary origin, but has been overprinted by regional deformation and metamorphism. The Shilu Co–Cu deposit has a hydrothermal origin and is temporally and genetically associated with Indosinian granitoid rocks. © 2013 Elsevier B.V. All rights reserved.

1. Introduction The Shilu Fe–Co–Cu orebodies are located at the intersection of the E–W-trending Changjiang–Qionghai and NE–SW-trending Gezhen fault zones (Fig. 1b). The ores are hosted in deformed Meso- to Neoproterozoic metamorphic rocks (Fig. 1b). The Shilu orebodies are the largest iron-rich ore deposit in Asia and comprise 38 iron orebodies, 17 cobalt orebodies, and 41 copper orebodies (Xiao et al., 2010; Xu et al., 2009). The orebodies have past production and indicated reserves of 460 Mt at 51.2 wt.% iron, 4.1 Mt at 0.29 wt.% cobalt, and 6.7 Mt at 1.2 wt.% copper (Xiao et al., 2010; Xu et al., 2009). Large economic orebodies include the Beiyi, South 6, and Fengshuxia iron orebodies, the No. 1 copper orebody, and the No. 1 cobalt orebody (MRESBGPHP, 2010).

⁎ Corresponding author. E-mail address: [email protected] (J. Yu). 0169-1368/$ – see front matter © 2013 Elsevier B.V. All rights reserved. http://dx.doi.org/10.1016/j.oregeorev.2013.08.018

Previous studies of the genesis of the Shilu Fe–Co–Cu orebodies have produced contrasting models for their origin(s) including hydrothermal replacement(No. 934 Geological Team, BMGGP 1954; Zhao et al., 2008) and a syngenetic seawater model with superimposed mineralization events due to regional metamorphism and skarns associated with granitoid intrusions (SCISTCAS, 1986; Shen et al., 2005). Recently, D.R. Xu et al. (2008) and Du et al. (2012) proposed that the Shilu Fe–Co–Cu deposit belongs to the IOCG clan (Hitzman et al., 1992). Xu et al. (2013) suggested that the Shilu Fe deposit is a structurally and hydrothermally reworked and re-enriched BIF-type ore deposit. Although a number of studies have been carried out on the deposit, the ore-forming material, fluid source, and mechanism of deposit formation are poorly constrained largely due to a lack of detailed fluid inclusion and isotope systematics data. Herein, we describe the geology and present fluid inclusion, isotopic geochemistry, and geochronology of the Shilu Fe–Co–Cu deposit, and integrate our results with previously published research (SCISTCAS, 1986; Xu et al., 2009), in order to better understand the genesis of the Shilu Fe–Co–Cu deposit. These new

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Fig. 1. (a) Structural outline of Southeast Asia showing the major tectonic blocks, boundaries and faults/shear zones (after Zhang et al., 2011). The region with pink color shows the Triassic Lancangjiang igneous zone. (b) Geological map of Hainan Island (after Guangdong BGMR, 1988; Zhang et al., 2011).

data provide answers to the following questions: 1) How were the high-grade iron ores produced? 2) Did the Fe and Cu–Co ores form simultaneously? 3) What was the source of the ore-forming material and fluid? and 4) What are the differences between the Fe and Cu–Co mineralization? 2. Geological setting Hainan Island is a continental island separated from mainland South China by the Qiongzhou Strait (Fig. 1a). There are a number of different views on the tectonic setting of the island. Some studies have proposed that Hainan Island is part of the Cathaysian block within the South China Block (Guangdong BGMR, 1988; Shui, 1987; X.H. Li et al., 2002; Z.X. Li et al., 2002), whereas others studies have suggested that it is associated with the Indochina Block (Hsű et al., 1990; Chen et al., 1994). It has also been proposed that north and south Hainan Island, which are inferred to be separated by the Changjiang–Qionghai Fault, have affinities with the Cathaysian and Indochina blocks, respectively (Li et al., 2000; X.H. Li et al., 2002; Z.X. Li et al., 2002; Zhang et al., 2011; D. Xu et al., 2008). It has previously been proposed that the E–W-trending Wangwu– Wenjiao, Changjiang–Qionghai, Jianfeng–Diaoluo, and Jiusuo–Linshui brittle faults extend from north to south on Hainan Island (Guangdong BGMR, 1988; Fig. 1b). However, sparse geophysical and structural data are available to verify this proposition. The Western Island contains the early Mesozoic NE–SW-trending Gezhen–Lingao and Baisha faults (Fig. 1b) that might have experienced late-stage reactivation (Xia et al., 1990, 1991; Wang et al., 1991, 1992). The stratigraphic succession on Hainan Island includes the Mesoproterozoic Baoban Group, the Meso- to Neoproterozoic Shilu Group, Paleozoic shallow marine strata, and Mesozoic terrestrial

strata (Ma et al., 1998). The Baoban Group (also termed the Baoban Complex) comprises greenschist- to amphibolite-facies metamorphic rocks (Guangdong BGMR, 1988) that were intruded by 1.43 Ga granitoids (X.H. Li et al., 2002; Z.X. Li et al., 2002), which collectively are thought to be the oldest basement rocks in this region (Wang et al., 1991, 1992). The Shilu Group is unconformably overlain by Sinian (Neoproterozoic) neritic siliciclastic rocks that only crop out in the northwestern part of the island, and comprise iron-rich siliciclastic, carbonate, and minor volcanic rocks with flysch-like characteristics (Guangdong BGMR, 1988; Lu, 1988; Wang et al., 1991, 1992). The Cambrian and Ordovician low-grade metamorphic successions are predominantly found in the areas south of the Changjiang–Qionghai Fault (Fig. 1b). These rocks comprise shale, sandstone, siltstone, slate, and minor amounts of interbedded limestone. A shallow marine sandstone of Silurian age has also been identified on Hainan Island. The Upper Paleozoic successions are characterized by Carboniferous slate and metamorphosed volcanic rocks, Lower Permian limestone, and Middle Permian sandstone. These rocks predominantly crop out in the areas north of the Jiusuo–Linshui Fault. Middle Permian conglomerates unconformably overlie lower Permian limestone in the central part of Hainan Island. Upper Triassic sandstones are only found in the Anding and Qionghai areas in the eastern and northern parts of the island (Fig. 1b), and unconformably overlie the pre-Triassic sequence and, in turn, are overlain by Lower Cretaceous terrestrial siliciclastic rocks. A number of metabasite lenses are present within Late Paleozoic metasediments in the Chenxing (Tunchang) and Bangxi (Changjiang) areas along the Changjiang–Qionghai Fault (Fig. 1b), and these rocks have typically undergone amphibolite-facies metamorphism. The metabasites are considered to be fragments of Paleozoic ophiolitic rocks and mélange sheets that are remnants of the eastern part of the

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Paleotethys (Guangdong BGMR, 1988; Wang et al., 1991, 1992; X.H. Li et al., 2002; Z.X. Li et al., 2002). X.H. Li et al. (2002) obtained an Sm– Nd isochron age of 333 ± 12 Ma and Xu et al. (2007b) reported a SHRIMP zircon U–Pb age of 269 ± 4 Ma for the Bangxi metabasites. Geochemical and Sm–Nd isotopic data indicate that these rocks were generated in a restricted oceanic setting (Li et al., 2000; X.H. Li et al., 2002; Z.X. Li et al., 2002), which was either a back-arc basin or rifting environment (Xu et al., 2007b). Foliated and unfoliated granites crop out extensively on Hainan Island and cover ca. 37% of the land area (Wang et al., 1991; Fig. 1b). Strongly foliated granites are found over an area of ca. 800 km2 and are concentrated in the Ledong–Wuzhishan–Wanning areas in the central and southern parts of the island. The granites contain abundant mafic magmatic and paragneissic enclaves that are several tens of centimeters to 10 m in size. They were previously considered to be components of the Mesoproterozoic Baoban Group or misidentified as the “Shang'an migmatite” of unknown age (Wang et al., 1991). However, recently published geochemical and geochronological data show that most of the foliated granites are Middle Permian (zircon U–Pb ages of 262–272 Ma) syntectonic biotite granites and granodiorites (Xie et al., 2005; Li et al., 2006; Xie et al., 2006b, Fig. 1b), and Li et al. (2006) renamed these rocks as the Wuzhishan orthogneiss. The Wuzhishan orthogneiss was intruded by the Qiongzhong granitic batholith, which yields a SHRIMP zircon U–Pb age of 226–234 Ma (Xie et al., 2005; Li et al., 2005). Some of the foliated granites might be of Grenvillian origin, despite uncertainties about the spatial distribution of these granites (X.H. Li et al., 2002; Z.X. Li et al., 2002). Unfoliated granites include the Triassic Qiongzhong, Jianfengling and Changjiang granitic batholiths (Xie et al., 2005; Li et al., 2005; Xie et al., 2006a), and Jurassic–Cretaceous (zircon U–Pb ages of 150–60 Ma) medium- to coarse-grained monzogranite (Wang et al., 1991, 1992; Ge, 2003).

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comprises phyllite, slate, and dolomite. The Upper Carboniferous Qingtianxia Formation crops out in the eastern part of the Shilu ore district and comprises dolomite and dolomitic limestone intercalated with siltstones and phyllites. The Lower Permian units can be divided into sandstone and phyllite intercalated with limestone (Echa Formation), and limestone and phyllite (Eding Formation). The Lower and Upper Permian Nanlong Formations comprise siltstone, phyllite, and limestone. 3.2. Structural geology The main structures within the Shilu ore district are faults and the NW–SE-trending Beiyi synclinorium (Fig. 2). The Beiyi synclinorium includes the Baoxiu syncline, the Sanlengshan syncline, the Jixinling anticline, the Beiyi syncline, the Hongfangshan anticline, the Shihuiding syncline, the Fengshuxia anticline, and the Fengshuxia syncline. The core of the synclinorium is mainly composed of the Shihuiding Formation and the upper unit of the No. 6 Formation, and the limbs consist of the middle and lower units of the No. 6 Formation and No. 5 Formation (Fig. 3). Three groups of faults that trend NW–SE to NNW–SSE, NE– SW, and NNE–SSW to N–S are extensively developed in the Shilu ore district (Fig. 2). The dominant NW–SE-trending Beiyi synclinorium and NW–SE-trending faults developed during the period 830–840 Ma (Xu et al., 2013). NE–SW-trending and NNE–SSW to N–S-trending faults crosscut the NW–SE-trending folds and faults, and possibly formed during the Yanshanian (Zhang et al., 2011). The metamorphism in the Shilu ore district is mainly of the lower greenschist facies and locally of the lower amphibolite facies. Contact metamorphism has resulted in the formation of hornfels. Andalusite and tourmaline are present in the country rocks of the No. 1 to No. 5 formations of the Shilu Group (SCISTCAS, 1986; Xu et al., 2009). Veined skarn minerals such as garnet and tremolite are also found in the host rocks as a result of Indosinian granite intrusions.

3. Geology of the Shilu deposit 3.3. Igneous rocks 3.1. Stratigraphy The stratigraphic sequence in the Shilu ore district comprises the Meso- to Neoproterozoic Shilu Group, the Sinian (Neoproterozoic) Shihuiding Formation, the Lower Carboniferous Nanhao and Upper Carboniferous Qingtianxia Formations, the Lower Permian Echa and Eding Formations, and the Lower to Middle Permian Nanlong Formation (Fig. 2). The Meso- to Neoproterozoic Shilu Group crops out in the central and south regions of the Shilu ore district and comprises lower greenschist-facies and locally developed amphibolite-facies metamorphic rocks (e.g. banded diopsidic and tremolitic rocks) with siltstone, mudstone, carbonate, and rhyolite protoliths. These rocks are suggestive of formation in a flysch-like setting that may have formed in a shallow marine and/or lagoon environment (SCISTCAS, 1986; Zhang et al., 1992; Xu et al., 2009). The Shilu Group can be further divided into six Formations (Nos. 1 to 6, from base to top). The No. 1, No. 3, and No. 5 Formations are quartz–sericite schist. The No. 2 Formation comprises dolomite, diopsidic and tremolitic rocks, and the No. 4 Formation consists of quartz schist and quartzite. The No. 6 Formation can be subdivided into three units: the lower unit hosts the Cu–Co ores and comprises dolomite, diopsidic and tremolitic dolomite, and diopside- and tremolite-rich rocks; the middle unit hosts the iron ores and comprises banded diopsidic and tremolitic rocks that may be garnet-bearing, diopsidic and tremolitic dolomites, and ferruginous phyllites intercalated with jasperite and rhyolitic pyroclastic rocks; and the upper unit comprises carbonaceous and/or pelitic dolomite intercalated with carbonaceous slate or phyllite. The Sinian Shihuiding Formation overlies the Neoproterozoic Shilu Group along fault contacts (Fig. 2), and comprises quartz sandstone that is locally interbedded with phyllites. These Precambrian rocks are unconformably overlain by Late Paleozoic shallow marine carbonate and clastic rocks. The Lower Carboniferous Nanhao Formation is found in the southeastern and central parts of the Shilu ore district and

Multiple intrusive events have been recognized in the Shilu ore district. These intrusions include Mesoproterozoic biotite granite with a gneissic foliation defined by oriented K-feldspar phenocrysts and biotite crystals, Indosinian biotite monzonitic granite, and Yanshanian granitic porphyry (Fig. 2). The Indosinian biotite monzonitic granite is part of the Changjiang granitic batholith, has a high-K calc-alkaline and shoshonitic composition, and intruded the western, northern, and southern parts of the Shilu ore district. This granite has yielded a zircon U–Pb age of 230–248 Ma (Yu et al., 2012) by laser ablation–inductively coupled plasma–mass spectrometry (LA–ICP–MS), and has been suggested to have formed in a post-orogenic environment. The Yanshanian biotite granite and granitic porphyry are exposed in the southeastern parts of the Shilu ore district. Wang et al. (2011) reported a LA–ICP– MS zircon U–Pb age of 93 ± 2 Ma for the granite porphyry. 4. Orebodies 4.1. Orebodies A total of 38 iron orebodies, 17 cobalt orebodies, and 41 copper orebodies have been identified in the Shilu ore district (Xu et al., 2009; Xiao et al., 2010). The large iron orebodies include Beiyi, Beiyi East, South 6, and Fengshuxia, which account for 90% of the total iron reserves (MRESBGPHP, 2010). The Beiyi and Beiyi East orebodies are subdivided into different orebodies due to different stages of exploration. The Beiyi East orebody is the eastern extension of the Beiyi orebody and, in fact, they are the same orebody and as such are referred to as the Beiyi orebody in this paper. The Beiyi orebody is the largest iron orebody in the Shilu ore district. The No. 1 copper and cobalt orebodies are the largest in the district and contain 92% and 98% of the total copper and cobalt reserves, respectively.

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Fig. 2. Simplified geological map of the Shilu deposit (modified from MRESBGPHP, 2010).

The main iron-producing part of the Shilu ore district is the Beiyi sector, located in the northwestern part of the Shilu ore district. The Beiyi iron orebody is located in the core of the Beiyi syncline (Figs. 2 and 3). The No. l copper orebody occurs in both the limb and core of the Beiyi syncline, and the No. 1 cobalt orebody is found in the southern limb and core of the Beiyi syncline. These NW–SE-striking Fe–Co–Cu orebodies are partly to fully stratiform (Fig. 4). The Beiyi iron orebody is 1280 m long with a thickness of 430 m in the west and 60–80 m in the east, is 704 m in vertical extent and has an average iron grade of 58 wt.% (MRESBGPHP, 2010). The No. l Cu orebody is 440 m long, 3.0 to 35.7 m thick (average = 6.1 m), 290 m in vertical extent, and has an average copper grade of 1.69 wt.%. The No. 1 cobalt orebody ranges in thickness from 2 to 5 m (average = 4.4 m) and can generally be followed along strike for ca. 1200 m. The cobalt orebody is 562 m in vertical extent and has average cobalt and copper grades of 0.31 and 0.75

wt.%, respectively. The South 6 iron orebody is present in the southern limb of the Shihuiding syncline and in the core of the Fengshuxia anticline (Fig. 2), and is 930 m long, 15 m thick, 274 m in vertical extent, and has an average iron grade of 51.8 wt.%. The Fengshuxia iron orebody is located in the southern limb and in the core of the Fengshuxia syncline (Figs. 2 and 3), and is 1800 m long, 14 to 222 m thick, 208 m in vertical extent, and has an average iron grade of 51.6 wt.%. In general, the Cu–Co and Fe orebodies are discrete (Fig. 4), although there are some locally mixed orebodies within the Beiyi iron orebody. The host rocks of the iron ores are banded diopsidic and tremolitic rocks of the middle unit of the No. 6 Formation of the Shilu Group. The Cu–Co ores are hosted by dolomite, and diopsidic and tremolitic dolomite of the lower unit of the No. 6 Formation of the Shilu Group. The Cu–Co orebodies are stratigraphically below the Fe orebodies (Fig. 4).

Fig. 3. E11 cross-section (located in Fig. 2) in the Shilu mine showing the relationship of the iron orebodies to the Beiyi synclinorium (modified from MRESBGPHP, 2010).

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Fig. 4. Section of profile A–A′ (located in Fig. 2) in the Shilu mine showing the Fe–Co–Cu orebodies with Co–Cu in the depth and Fe upward (modified from MRESBGPHP, 2010).

4.2. Mineralization Detailed petrological studies of the host rocks, textural and structural features of the ores, and alteration mineral assemblages have led to the recognition of sedimentary, metamorphic, hydrothermal, and supergene mineralization stages. The hydrothermal ore-forming stage related to Cu–Co mineralization can be further subdivided into skarn and quartz–sulfide stages (Table 1). The sedimentary ore-forming stage took place in the No. 6 Formation of the Meso- to Neoproterozoic Shilu Group and is defined by the deposition of carbonates and evaporates with minor clastic rocks and jasperite (Fig. 5a). The iron ores that resulted from this stage include laminated or banded oolitic and blastopsammitic hematite ores (SCISTCAS, 1986; Xu et al., 2013). Colloidal and cryptocrystalline to microcrystalline Cobearing pyrites were also formed in this stage. The main minerals that developed during this stage were dolomite, chert/quartz, hematite, anhydrite, barite, and colloidal and cryptocrystalline to microcrystalline pyrite. The banded diopsidic and tremolitic rocks (Fig. 5c) and garnetbearing diopsidic and tremolitic rocks were formed due to metamorphism. At scales ranging from hand samples to orebodies Fe-ores often exhibit banded and/or massive structures and some, especially those strongly deformed, have a schistose (Fig. 5b) and brecciated appearance (Xu et al., 2013). On a microscopic scale, original sedimentary textures are preserved in less metamorphosed ores. However, where more intensively metamorphosed, fine-grained, lepidoblastic textures are increasingly abundant (Xu et al., 2013). The main minerals that formed during the metamorphic ore-forming stage were garnet, diopside, tremolite, and hematite. The Fe ores are cut by veined skarn minerals and sulfides (Fig. 5d–i), indicating that the Cu–Co ores were formed after the Fe ores. The Cu–Co ores formed during the quartz–sulfide stage. Garnet and tremolite veins (Fig. 5d–f) formed in the skarn stage. The skarn minerals are different to those produced by regional metamorphism in that the former occur as veins, whereas the latter are banded (Fig. 5c). The late quartz–sulfide stage associated with Cu–Co mineralization involved the formation of Co-bearing pyrite and pyrrhotite (Fig. 5h and i), chalcopyrite, cobaltite, quartz, calcite, dolostone, and chlorite. Azurite, malachite and limonite were formed during the supergene period. 5. Fluid inclusions 5.1. Analytical methodology Fluid inclusions were characterized by optical microscopy observations of doubly polished sections of 200–300 μm in thickness.

Microthermometric measurements were carried out on a Linkam THMSG 600 programmable heating–freezing stage (−196 °C to +600 °C) in the Fluid Inclusion Laboratory of China University of Geosciences, Beijing, China. The stage was calibrated using synthetic

Table 1 Generalized paragenetic sequence for the Shilu deposit.

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Fig. 5. Photographs showing the Fe–Co–Cu ores and country rocks in the Shilu mine. (a) Jasperite crosscut by calcite–pyrite stockwork. (b) schistose hematite-rich ore; (c) Banded diopsidic and tremolitic rocks; (d) Hematite cut by veined garnet; (e) Hematite cut by quartz–garnet stringer; (f) Hematite cut by quartz–tremolite vein; (g) Hematite cut by sulfide stockwork; (h) Co-bearing pyrrhotite in dolomite; (i) Co-bearing pyrite and pyrrhotite in dolomite. Mineral/rock abbreviations: Jas, jasperite; Cal, calcite; Gar, garnet; Dip, diopside; Tre, tremolite; Q, quartz; Do, dolomite; Po, pyrrhotite.

fluid inclusions. The heating rate was 0.1 to 1 °C/min below 10 °C and ca. 3 to 5 °C/min from 10 °C to 30 °C, with a temperature reproducibility of ±0.1 °C. The heating rate was 5 to 10 °C/min at high temperatures (N100 °C) with a temperature reproducibility of ±2 °C.

5.2. Fluid inclusion characteristics Fluid inclusions were examined in 26 samples from the sedimentary, skarn, and quartz–sulfide stages. Fluid inclusions are abundant in quartz

Table 2 Characteristics of microthermometric data of fluid inclusions from the Shilu deposit. Period/stage

Sedimentary

Mineral Inclusion type a Number of inclusions Vapor/liquid ratio(%) Tmice (°C) TmCO2 (°C) TmC (°C) ThCO2(L) (°C) Th(LV−L) (°C) Salinity (wt.% NaCl equiv.) b Density of CO2 phase (g/cm3) c Bulk density (g/cm3) c

Quartz I 19

Metamorphism

Skarn

III 14

Quartz I 72

Quartz II 46 10 to 95

−8.8 to −0.7

−8.8 to −2.1 −60.0 to −57.7 1.1 to 8.7 16.2 to 31.0 248 to 497 234 to 389 2.6 to 14.3 3.6 to 12.6 0.56 to 0.82 0.56 to 0.94 0.62 to 0.92

−1.8 to −0.5 −60.0 to −58.0 9.2 to 30.9 144 to 268 0.71 to 3.1

148 to 348 1.2 to 12.6 0.57 to 0.87

0.77 to 0.93

0.71 to 0.94

Quartz–sulfide Garnet I 32

Epidote I 10

Quartz I 183

Dolomite I 33

Calcite I 16

−19.9 to −3.2 −5.2 to −0.5 −4.7 to −1.4 −2.4 to −0.5

288 to 452 5.3 to 22.3

151 to 356 0.9 to 8.1

171 to 297 2.4 to 7. 5

191 to 282 0.9 to 4.0

0.70 to 0.89

0.63 to 0.96

0.78 to 0.92

0.74 to 0.89

Tmice — melting temperature of ice; TmCO2 — CO2 melting temperature; TmC — CO2 clathrate melting temperature; ThCO2(L) — CO2 temperature of homogenization to liquid; Th(LV−L) — temperature of homogenization to liquid. a I-two-phase liquid-vapor inclusion; II-CO2-bearing three-phase inclusion; III-CO2-bearing or carbonic inclusion. b Salinity calculated from Tmice (Bodnar, 1993). c Density of CO2 phase and Bulk density estimated following the method of Liu and Shen (1999).

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of the sedimentary ore-forming stage; in garnet, epidote, and quartz of the skarn stage; and in quartz and calcite of the quartz–sulfide stage. Fluid inclusion data are listed in Table 2 and shown in Fig. 6.

5.3. Microthermometry measurements Fluid inclusions include two-phase (L + V L = liquid phase, V = vapor phase) (70% of all inclusions; Fig. 7a–d), CO2-bearing threephase (L + V + LCO2, LCO2 = liquid CO2 phase) (20%; Fig. 7f), and two-phase (LCO2 + VCO2, VCO2 = vapor CO2 phase) CO2-bearing or carbonic inclusions (10%; Fig. 7e). A total of 425 two-phase NaCl–H2O, CO2-bearing three-phase, and two-phase (LCO2 + VCO2) CO2 inclusions were measured. All of the two-phase NaCl–H2O inclusions homogenized into a liquid phase.

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Homogenization temperatures (Th(LV−L)) of the sedimentary oreforming quartz range from 144 °C to 268 °C, with most being between 150 °C and 210 °C and a peak at 160 °C (Fig. 6a). The ice-melting temperatures range from −1.8 °C to −0.5 °C. Salinities determined from fluid inclusion ice-melting temperatures vary from 0.7 to 3.1 wt.% NaCleq (Bodnar, 1993; Fig. 6b and Table 2). Based on the final homogenization temperatures and the salinities, the corresponding densities of the aqueous phase can be calculated to range between 0.77 and 0.93 g/cm3 using the NaCl–H2O reference table of Liu and Shen (1999). The quartz-hosted CO2-bearing or carbonic inclusions yield CO2 melting temperatures from −58 °C to −60 °C (Table 2), which is lower than the melting temperature of pure CO2 (−56.6 °C). This temperature range is consistent with the presence of small amounts of CH4 in the inclusions in addition to CO2. CO2 homogenization occurs at temperatures from 9.2 °C to 30.9 °C, which corresponds to densities of 0.57

Fig. 6. Histogram of homogenization temperature and salinities for the fluid inclusions in the Shilu deposit. Th — homogenization temperature.

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Fig. 7. Photographs of fluid inclusion in quartz, garnet and epidote from the Shilu deposit. (a) Two-phase fluid inclusions in jasperite/chert; (b) Two-phase fluid inclusions in quartz from skarn stage; (c) Two-phase fluid inclusions in garnet from skarn stage; (d) Two-phase fluid inclusions in epidote from skarn stage; (e) two-phase CO2-bearing or carbonic inclusions in quartz; and (f) CO2-bearing three-phase inclusions in quartz from skarn stage.

to 0.87 g/cm3. The CO2-bearing or carbonic inclusions have been interpreted to have been produced by regional metamorphism. Th(LV−L) for the skarn stage displays a wide range of values from 148 °C to 497 °C with peaks at 250 °C to 310 °C and 370 °C to 410 °C (Fig. 6c). Salinities determined from the melting temperatures of the fluid inclusions range from 1.2 to 22.3 wt.% NaCleq with two peaks between 4 and 7 wt.% and 21 and 22 wt.% NaCleq (Fig. 6d), which correspond to fluid densities between 0.56 and 0.94 g/cm3. Microthermometry measurements of fluid inclusions in quartz, dolomite, and calcite from the quartz–sulfide stage yielded Th(LV−L) from 151 °C to 356 °C, with most temperatures between 170 °C and 270 °C (Fig. 6e) and a peak at ca. 200 °C. These inclusions have salinities of 0.9 to 8.1 wt.% NaCleq (Fig. 6f) and densities of 0.63 to 0.96 g/cm3. 6. Stable isotopes 6.1. Samples and analytical methods S, H, and O isotopic analyses were performed in the Stable Isotope Laboratory of the Institute of Mineral Resources, Chinese Academy of Geological Sciences, Beijing, China. Samples were crushed and passed through 60–80 mesh sieves. Single minerals were handpicked under a binocular microscope to a purity N99%. After handpicking the minerals were powdered and oxidized to SO2 at a temperature of 1020 °C, approximately 20 to 100 μg of S-bearing minerals was weighed out for S isotopic analysis. All analytical results are reported relative to the VCDT international standard and the analytical precision on δ34S values is ≤0.2‰. A Finnigan MAT-253 mass spectrometer was used for the stable isotope measurements of S. CO2 was prepared using the BrF5 method for oxygen isotope analysis of silica. Water in the primary fluid inclusions was released by the thermal decrepitation method for hydrogen isotope analysis. The water was then reacted with Zn at 400 °C to produce H2, which was collected in sample tubes with activated charcoal at the temperature of liquid N2. H and O isotope analyses of H2 and CO2 were conducted with a Finnigan MAT 253EM mass spectrometer. Analytical reproducibilities for O and H isotope measurements were ±0.2‰ and ±1‰, respectively. 6.2. Results The S isotope results and comparative data from other studies of the Shilu deposit are listed in Table 3. δ34S values of anhydrite in the No. 6

Formation of the Shilu Group range from +21.4‰ to +21.8‰, with a mean value of +21.6‰. Sulfides in the Cu–Co ores have δ34S values from +8.1‰ to +21.2‰, with a mean value of +15.4‰. The H and O isotope results are given in Table 4. δ18OSMOW values of quartz for the skarn stage range from +12.5‰ to +17.1‰ (mean = +15.7‰), and those for the quartz–sulfide stage vary from +11.9‰

Table 3 Sulfur isotopic compositions (‰) of sulfides in the Shilu deposit. No.

Sample no.

Mineral

δ34SCDT (‰)

1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22 23 24 25 26 27 28 29 30 31 32 33

SL-77-139-2B SL-77-114 SL-77-115 F9-04 F9-09 F9-11 N21-10 SL-77-37-4 SL-77-140-2 SL-77-224-1-6 SL-77-224-7A SL-77-224-7B SL-77-92 SL10-146 SL10-151 SL10-152 SL10-153 SL10-156 SL-77-37-1 SL-77-74-1 SL-77-140-1 SL10-143 SL10-145-1 SL10-147 SL10-149-1 SL10-155-1 SL-77-239 N14-09 SL10-145-2 SL10-149-2 SL10-150 SL10-154 SL10-155-2

Sedimentary Co-bearing pyrite in host rock Sedimentary anhydrite in host rock Sedimentary anhydrite in host rock Co-bearing pyrite in Co–Cu orebody Co-bearing pyrite in Co–Cu orebody Co-bearing pyrite in Co–Cu orebody Co-bearing pyrite in Co–Cu orebody Co-bearing pyrite in Co–Cu orebody Co-bearing pyrite in Co–Cu orebody Co-bearing pyrite in Co–Cu orebody Co-bearing pyrite in Co–Cu orebody Co-bearing pyrite in Co–Cu orebody Co-bearing pyrite in Co–Cu orebody Co-bearing pyrite in Co–Cu orebody Co-bearing pyrite in Co–Cu orebody Co-bearing pyrite in Co–Cu orebody Co-bearing pyrite in Co–Cu orebody Co-bearing pyrite in Co–Cu orebody Chalcopyrite in Co–Cu orebody Chalcopyrite in Co–Cu orebody Chalcopyrite in Co–Cu orebody Chalcopyrite in Co–Cu orebody Chalcopyrite in Co–Cu orebody Chalcopyrite in Co–Cu orebody Chalcopyrite in Co–Cu orebody Chalcopyrite in Co–Cu orebody Co-bearing pyrrhotite in Co–Cu orebody Co-bearing pyrrhotite in Co–Cu orebody Co-bearing pyrrhotite in Co–Cu orebody Co-bearing pyrrhotite in Co–Cu orebody Co-bearing pyrrhotite in Co–Cu orebody Co-bearing pyrrhotite in Co–Cu orebody Co-bearing pyrrhotite in Co–Cu orebody

21.2 21.8 21.4 15.7 15.9 16.9 15.3 17.4 15.6 11.1 17.5 16.6 17.4 15 17.3 12.7 12.4 18.4 13.8 13.5 15.7 8.1 13.3 12.6 17 15.4 14.7 17.2 13.4 17.6 15.2 17.7 15.4

Data source: Numbers 4 to 7, and No. 28 are from Xu et al. (2009), Numbers 1 to 3, 8 to 13, 19 to 21, 27 to 28 are from SCISTCAS (1986) and others from this study.

J. Yu et al. / Ore Geology Reviews 57 (2014) 351–362 Table 4 Hydrogen and oxygen isotopic compositions (‰) of quartz in the Shilu deposit. Sample no.

Stage

δ18Oquartz(‰)

Th(°C)

δD(‰)

δ18Ofluid(‰)

SL10-27 SL10-43 SL10-44 SL10-45 SL10-46 SL10-49 SL10-75 SL10-77 SL10-47 SL10-87 SL10-88 SL10-91 SL10-117 SL10-118 SL10-119 SL10-123

Skarn stage Skarn stage Skarn stage Skarn stage Skarn stage Skarn stage Skarn stage Skarn stage Skarn stage Quartz–sulfide stage Quartz–sulfide stage Quartz–sulfide stage Quartz–sulfide stage Quartz–sulfide stage Quartz–sulfide stage Quartz–sulfide stage

12.5 16.2 15.7 15.6 14.9 16.2 17.1 16.7 16 17.9 18.1 17 11.9 16.8 17 16.4

235 238 238 215 266 281 297 205 220 209 202 172 218 188 204 215

−65 −80 −61 −72 −80 −74 −54 −70 −73 −66 −63 −66 −68 −63 −83 −79

2.8 6.7 6.2 4.8 6.7 8.6 10.1 5.3 5.5 6.8 6.5 3.3 1.3 4.3 5.5 5.6

Th average value of homogenization temperature of fluid inclusions in quartz.

to +18.1‰ (mean = +16.4‰). Using the quartz–water fractionation equation, 1000lnα = 3.38 × 106/T2 − 3.40 (Clayton, 1972), and the average homogenization temperature of fluid inclusions in quartz in the same samples, δ18Ofluid values of the mineralizing fluids are calculated to be +2.8‰ to +10.1‰ (mean = +6.3‰) for the skarn stage. δ18Ofluid values for the quartz–sulfide stage vary from +1.3‰ to +6.8‰ (mean = +4.8‰). δD values of quartz for the skarn stage range from −80‰ to −54‰ (mean = − 69.9‰), and those for the quartz–sulfide stage vary from − 63‰ to − 83‰ (mean = − 69.7‰). 7. Discussion 7.1. Source of sulfur A mean δ34S value of +17.5 ± 3‰ has been proposed for Neoproterozoic seawater at the time of deposition of the Shilu Group (Claypool et al., 1980; El Desouky et al., 2010; Holser and Kaplan, 1966; McGowan et al., 2006). This value is supported by δ34S analyses of anhydrite from the Shilu Group, which have values between 21.4‰ and 21.8‰ (Fig. 8). The δ34S value of sample SL-77-139-2B (Co-bearing

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pyrite) from the sedimentary ore-forming stage is 21.2‰ (Table 3), which is similar to that of the sedimentary anhydrite and implies that both share a common origin and sulfur source. Open system thermochemical reduction of sulfate (TRS) in the Shilu region at temperatures of 100–200 °C resulted in ΔSO4−sulfides values of between 10‰ and 15‰ (Machel et al., 1995; McGowan et al., 2006). These fractionations are not consistent with ΔSO4−sulfides values for the Cu–Co orebodies in the Shilu deposit, which are typically only 6.2‰ lighter and. In fact, δ34S values of sulfides overlap with the projected seawater sulfate value (Fig. 8). This may indicate that the Cu–Co orebodies originated from an anhydrite source that was consumed in a closed system and/or that thermochemical reduction took place at temperatures N200 °C. The latter possibility is supported by the fact that fluid inclusion homogenization temperatures of N 200 °C were obtained for the quartz–sulfide stage. There are two broad possible sulfur sources for the sulfides in the Cu–Co ores: (1) Neoproterozoic seawater sulfate for the sedimentary anhydrite and Co-bearing pyrite; and (2) TRS of the dissolved anhydrite at 200 °C to 300 °C for the quartz–sulfide stage.

7.2. Source of ore fluids δ18Ofluid values of the mineralizing fluids for the skarn stage range from +2.8‰ to +10.1‰. Five samples fall in the range of magmatic water (5.5‰–9.5‰) as defined by Ohmoto (1986) and Sheppard (1986). δ18Ofluid values of two samples (5.3‰ and 4.8‰) are slightly lower than that of magmatic water. One sample has a δ18Ofluid value of 2.8‰ (Table 4), which is clearly lower than that of magmatic water, whereas one sample has a δ18Ofluid value of 10.1‰ (Table 4) that is slightly higher than that of magmatic water. δD values of water in quartz from the skarn stage vary from −54‰ to −80‰ and are similar to those of magmatic water (−40‰ to −80‰) as defined by Taylor (1986). On a plot of δD versus δ18Ofluid (Fig. 9), five analyses fall within the magmatic water field, indicating that the ore fluids of the skarn stage in the Shilu deposit were mainly derived from magmatic fluids. An analysis of one sample plots within the metamorphic water field and analyses of three samples plot near the magmatic water field, which implies that the ore-forming fluids relating to these samples may have experienced oxygen isotopic exchange with water from

Fig. 8. Histogram of S isotope for the sulfides in the Shilu deposit.

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possibly associated with the Indosinian granitoids found in the Shilu mine and surrounding region (Figs. 1 and 2), representing a synorogenic or post-orogenic event related to the final closure of the Paleotethys Ocean closure in southeastern Asia. 7.4. Genesis and tectonic setting of the Shilu Fe–Co–Cu deposit

Fig. 9. δD vs. δ18Ofluid diagram of sulfides of the Shilu deposit (after Taylor, 1986; Zheng and Chen, 2000).

metamorphic rocks or meteoric water during the late stage of skarn mineralization. δ18Ofluid values of the mineralizing fluids for the quartz–sulfide stage vary between 1.3‰ and 6.8‰. Four of the analyzed samples fall in the range of magmatic water and three samples have δ18Ofluid values that are slightly lower than those of magmatic water. δD values of water in quartz from the quartz–sulfide stage range from −63‰ to −79‰, which are similar to those of magmatic water, apart from one sample with a δD value of −83‰. Four analyses plot within or near the magmatic water field (Fig. 9), suggesting that the ore-forming fluid of the quartz–sulfide stage was mainly magmatic water. Two analyses plot within the metamorphic water field and one analysis fall outside the metamorphic water field, implying that the ore-forming fluids related to these samples may have experienced O isotopic exchange with water from metamorphic rocks or meteoric water during the late stages of sulfide mineralization. 7.3. Ages of the sedimentary iron ore and regional metamorphism Zhang et al. (1989) determined the age of the Shilu Group to be Neoproterozoic based on the presence of algae fossils such as the Chuaria–Tawuia life-form association. It is very unlikely that these algae fossils would survive amphibolite-facies metamorphism. Generally, algae and other organic matters that were cogenetic with sedimentation or diagenesis are metamorphosed and decomposed to form hydrocarbons or other volatiles. Xu et al. (2007a) conducted a U–Pb SHRIMP study of detrital zircons of the Shilu Group, yielding ages ranging from ca. 960 to 1300 Ma. This represents the time of deposition of the Shilu Group. The Sm–Nd isochron age of hematite in iron-rich ores from the Beiyi and Fengshuxia orebodies is 840.6 ± 20.4 Ma (Zhang et al., 1992), and Six hematite and diopsidic and tremolitic dolomite samples yielded a Sm–Nd isochron age of 830 ± 16 Ma (Xu et al., 2009, 2013). The Sm-Nd isotopic system has been regarded as having a high closing temperature, so unless it has been subjected to more intense thermal disturbances, the system tends to stay stable during low-pressure metamorphism (Hodges, 2003). Although the possibility of a mixed age cannot be excluded, Xu et al. (2013) interpreted the Sm–Nd isochron age of ca.830–840 Ma as a resetting age due to ca. 1.0–0.7 Ga Grenvillian orogenic event(s) which resulted in regional deformation and metamorphism of the Shilu Group. Four samples of garnet, iron-poor ore, tremolite, and diopsidic and tremolitic dolomite yielded a Sm–Nd isochron age of 212.9 ± 6.6 Ma (Xu et al., 2009, 2013). Hydrothermal tremolite in the diopsidic and tremolitic rocks yields a 40Ar/39Ar plateau age of 211.7 ± 4.2 Ma (Xu et al., 2009). These alterations of skarn minerals at about 212 Ma were

The genesis of the Shilu Fe–Co–Cu deposit has been a subject of debate for more than 30 years, and a number of hypotheses have been proposed to account for the Fe–Co–Cu mineralization, as mentioned in the Introduction. The geological characteristics of the Shilu iron deposit suggest that the geometry of the iron orebodies is controlled by the sedimentary beds of the lower and middle units of the No. 6 Formation of the Meso- to Neoproterozoic Shilu Group, and by the NW–SE-trending Beiyi synclinorium (Figs. 2–4). Laminar hematite is intercalated with ferruginous fine-grained sandstone and siltstone in the hematite-poor ores (Fe grade = 30–50 wt.%), and the hematite is also intercalated with anhydrite-bearing dolomite, sandstone, and mudstone with a thickness of N10 m for a single anhydrite layer (SCISTCAS, 1986). Colloidal and oolitic quartzes occur in the Fe–Co–Cu ores or are found as remnant enclaves within the massive hematite ore (SCISTCAS, 1986). Jasperite is also present in the Beiyi and South 6 orebodies (Fig. 5a). The fluid of the sedimentary ore-forming stage is indicative of low temperatures (ca.160 °C) and low salinities (0.7 to 3.1 wt.% NaCleq). All of these observations suggest that the iron orebodies have an exhalative sedimentary origin. Regional metamorphism resulted in overprinting of the hematite laminae and the growth of garnet, diopside, and tremolite/actinolite. The metamorphosed ore includes hematite, garnet, diopside, and tremolite/actinolite within the banded and schistose hematite ores (Fig. 5b and c). The schistosity of the hematite-rich ores is parallel to the axial plane of folds that trend NW–SE, indicating that the desilicification of the sedimentary ores took place during regional metamorphism and deformation at ca. 830–840 Ma. At scales ranging from hand samples to ore bodies, some of the iron ores, especially those strongly deformed, have a schistose appearance (Fig. 5b). On a microscopic scale, where more intensively metamorphosed, fine-grained lepidoblastic textures are increasingly abundant (Xu et al., 2013). The quartzhosted CO2-bearing or carbonic inclusions are interpreted to have been produced by regional metamorphism. All of these data suggest that the iron orebodies have been overprinted by regional deformation and metamorphism. Xu et al. (2013) described the ore types, texture and structure of the Shilu Fe–Co–Cu deposit. The iron ores can be grouped into five classes: Class 1, also called Fe-rich ore, is marked by N 58 wt.% total FeO and less than 12 wt.% SiO2, 0.15 wt.% S, and 0.05 wt.% P2O5. Classes 2 and 3 each contain more than 45 wt.% total FeO but have sulfur contents of b0.3 wt.% and N0.3 wt.%, respectively. Class 4, also called Fe-poor ore, has total FeO contents of 30–40 wt.%. Class 5 is the most Fe-poor ore with total FeO contents between 20 and 30 wt.%. Class 1 ores are the most abundant commodity, mainly occurring in central parts of larger ore bodies; minor Fe-poor ores are distributed at the margins. Ore minerals in Fe-rich ores are dominated by hematite (~85%), locally with minor magnetite (≤1%). Gangue minerals include quartz (~14%), anhydrite and barite (~1%). The Fe-rich ores have fine-grained lepidoblastic and schistose textures, indicating a sedimentary origin overprinted by regional metamorphism and deformation. However, the mineralogy of Fe-poor ores is more complex, comprising hematite (20%–40%), magnetite (20%–45%), garnet (20%–25%), quartz (~19%), feldspar (~5%), diopside and actinolite (~4%) and barite (~1%), as well as minor titanite and biotite. The Fe-poor ores show a banded structure and oolithic texture (Xu et al., 2013), indicating a sedimentary origin. The most Fe-poor ore (Class 5) occurs in the quartz sandstone and phyllite of the Sinian Shihuiding Formation, and in phyllite of the middle unit in the No. 6 Formation. Original blastopsammitic textures are preserved in Class 5 ore, indicating a sedimentary origin.

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In general, the Fe-rich ores are of sedimentary origin and have been overprinted by regional metamorphism and deformation. An alternative explanation of their origin is desilicification of the sedimentary ores by regional metamorphism and deformation. The Cu–Co ores mainly have a veinlet and stockwork form (Fig. 5h and i) despite their semi-stratiform nature and restricted occurrence in the lower unit of the No. 6 Formation of the Shilu Group (Fig. 4). The Cu–Co orebodies cut the Fe ores (Fig. 5g) and therefore formed later than the Fe ores. The Cu–Co ore-forming fluids were mainly derived from magmatic water that had experienced isotopic exchange with water from metamorphic rocks or meteoric systems during the late stages of sulfide mineralization. The sulfur in the Cu–Co ores originated mainly from TRS of the dissolved anhydrite at 200 °C–300 °C during the quartz–sulfide stage. Mixing of the metal- and sulfur-bearing fluids resulted in the formation of the Cu-Co ores. Cu–Co mineralization in the Shilu ore district is temporally and genetically associated with Indosinian granitoids, although further research is required regarding the age of the Cu–Co ores. Paleozoic metabasites are widely exposed along the Changjiang– Qionghai Fault on Hainan Island (Fig. 1b). Geochemical and isotopic data indicate that these rocks are relics of fragmented oceanic crust of the eastern part of the Paleo-Tethys Ocean. The Song Ma ophiolite in northern Vietnam and the fragmented oceanic crust in Hainan Island are correlatives and represent the plate boundary between the Indochina and South China blocks (Lan et al., 2000; X. H. Li et al., 2002; Xu et al., 2007b; D. Xu et al., 2008; Zhang et al., 2011; Fig. 1a). Southwarddirected subduction of Paleo-Tethyan ocean crust beneath Indochina produced a Permian magmatic arc in southern Hainan Island (Li et al., 2006), and south of the Song Ma suture (Lan et al., 2000). The timing of collision between the Indochina and South China blocks remains poorly constrained. However, Lepvrier et al. (1997) showed that 40 Ar−39Ar dates of the Song Ma Ophiolite have the same metamorphic age of ca. 245 Ma, which is similar to the age of 242–250 Ma for the NW–SE to WNW–ESE-trending high-strain shear zones on Hainan Island (Zhang et al., 2011). These ages imply that suturing between Indochina and South China took place in the earliest Triassic (Li et al., 2006). On southern Hainan Island, the post-orogenic Jianfengling granite yields a SHRIMP zircon U–Pb age of 249 ± 5 Ma (Xie et al., 2006a), whereas the post-orogenic Qiongzhong and Changjiang granites yield zircon U–Pb ages of 226–234 and 230–248 Ma, respectively (Li et al., 2005; Yu et al., 2012; Fig. 1b). These granites were produced by collision between the Indochina and South China blocks. The Indosinian granite in the Shilu ore district is a part of the Changjiang granite (Yu et al., 2012). 8. Conclusions (1) The Shilu Fe ore is hosted by the middle unit of the No. 6 Formation of the Meso- to Neoproterozoic Shilu Group. The Shilu Fe deposit is of syngenetic sedimentary origin and has been overprinted by regional metamorphism and deformation at ca. 830–840 Ma that resulted in desilicification of the sedimentary ores and the formation of hematiterich ore with an Fe grade of N45% total FeO. (2) The lower unit of the No. 6 Formation hosts the Cu–Co orebodies. The Shilu Cu–Co deposit is a hydrothermal stratabound and stratiform mineral system. The Cu–Co ores are related to a quartz–sulfide mineralization stage associated with granitoids that intruded the Shilu Group. Acknowledgments This research was jointly financially supported by the National Key Basic Science Research Project of China (973 Program; 2012CB416803) and the National Research Program for Non-Profit Trades (sponsored by the Ministry of Land and Resources, China; Project No. 20091100724). We thank the geologists of Shilu Mine for their help during fieldwork. Chu Huiyan is thanked for her assistance with the fluid inclusion

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microthermometry measurements. We gratefully acknowledge manuscript review, constructive comments and helpful suggestions by professors Zhang Zhaochong, Li Jianwei and Li Houmin that have improved the earlier version of this paper. Managing editor Aaron Stallard is thanked for Science and English language improvements of the original and revised manuscripts. References Bodnar, R.J., 1993. Revised equation and table for determining the freezing point depression of H2O–NaCl solutions. Geochim. Cosmochim. Acta 57, 683–684. Chen, H.H., Sun, X., Li, J.L., Haag, M., Dobson, J., Hsu, J.H., Heller, F., 1994. Paleomagnetic constraints on early Triassic tectonics of South China. Sci. Geol. Sin. 29, 1–9. Claypool, G.E., Holser, W.T., Kaplan, I.R., Sakai, H., Zak, I., 1980. The age curves of sulphur and oxygen isotopes in marine sulphate and their mutual interpretation. Chem. Geol. 28, 199–260. Clayton, R.N., 1972. Oxygen isotope exchange between quartz and water. J. Geophys. Res. 77, 3057–3607. 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