Download Metallogeny of the Shilu Fe–Co–Cu deposit, Hainan Island, South

Ore Geology Reviews 57 (2014) 351–362
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Ore Geology Reviews
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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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J. Yu et al. / Ore Geology Reviews 57 (2014) 351–362
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
J. Yu et al. / Ore Geology Reviews 57 (2014) 351–362
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).
353
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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J. Yu et al. / Ore Geology Reviews 57 (2014) 351–362
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).
J. Yu et al. / Ore Geology Reviews 57 (2014) 351–362
355
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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J. Yu et al. / Ore Geology Reviews 57 (2014) 351–362
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).
J. Yu et al. / Ore Geology Reviews 57 (2014) 351–362
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.
357
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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J. Yu et al. / Ore Geology Reviews 57 (2014) 351–362
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
359
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.
360
J. Yu et al. / Ore Geology Reviews 57 (2014) 351–362
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.
J. Yu et al. / Ore Geology Reviews 57 (2014) 351–362
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
361
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.
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