Fluid inclusions and C–H–O–S–Pb isotope systematics of the Caixiashan sediment-hosted Zn-Pb deposit, eastern Tianshan, northwest China: Implication for ore genesis

Journal Pre-proofs Fluid inclusions and C−H−O−S−Pb isotope systematics of the Caixiashan sediment-hosted Zn-Pb deposit, eastern Tianshan, Northwest China: Implication for ore genesis Kang Wang, Yin-Hong Wang, Chun-Ji Xue, Jia-Jun Liu, Fang-Fang Zhang PII: DOI: Reference:

S0169-1368(19)30436-6 https://doi.org/10.1016/j.oregeorev.2020.103404 OREGEO 103404

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Ore Geology Reviews

Received Date: Revised Date: Accepted Date:

9 May 2019 2 February 2020 9 February 2020

Please cite this article as: K. Wang, Y-H. Wang, C-J. Xue, J-J. Liu, F-F. Zhang, Fluid inclusions and C−H−O−S −Pb isotope systematics of the Caixiashan sediment-hosted Zn-Pb deposit, eastern Tianshan, Northwest China: Implication for ore genesis, Ore Geology Reviews (2020), doi: https://doi.org/10.1016/j.oregeorev.2020.103404

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1

Fluid

inclusions

and

C−H−O−S−Pb

isotope

systematics

of

the

Caixiashan

2

sediment-hosted Zn-Pb deposit, eastern Tianshan, Northwest China: Implication for ore

3

genesis

4 5

Kang Wanga, b, Yin-Hong Wanga, b*, Chun-Ji Xuea, b, Jia-Jun Liua, b, Fang-Fang Zhanga, b

6 7

aSchool

8

China

9

bState

10

of Earth Sciences and Resources, China University of Geosciences, Beijing 100083,

Key Laboratory of Geological Processes and Mineral Resources, China University of

Geosciences, Beijing 100083, China

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*Corresponding author at: State Key Laboratory of Geological Processes and Mineral

13

Resources, China University of Geosciences, 29 Xue-Yuan Road, Haidian District Beijing

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100083, China.

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Tel.: +86 10 82322346 (office)

16

E-mail address: [email protected]

17 18

Abstract: The Caixiashan sediment-hosted Zn-Pb deposit (131 Mt at 3.95% Zn + Pb) is

19

located in the western segment of eastern Tianshan, on the southern margin of the Central

20

Asian Orogenic Belt, Xinjiang, northwest China. Zinc and lead mineralization is mainly

21

hosted in the dolomite marble of the Mesoproterozoic Kawabulake Group. Four stages (I to

22

IV) of hydrothermal activity have been identified, i.e., calcite + dolomite + quartz + pyrite

23

stage I, calcite + dolomite + quartz + spahlerite + pyrrhotite ± arsenopyrite stage II, calcite +

24

dolomite + quartz + galena + pyrite ± chalcopyrite stage III, and late quartz + calcite stage IV.

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Sphalerite and galena mainly occur in the vein ores of stage II and III. Five types of fluid

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inclusions are distinguished in the calcite- and quartz-bearing veins, i.e., liquid-rich two-phase

27

(L-type), pure-liquid phase (PL-type), vapor-rich two-phase (V-type), pure-vapor phase

28

(PV-type), and halite-bearing (H-type) inclusions. Fluid inclusions of stages I to IV were

29

homogenized at temperatures of 336−494 °C, 240−357 °C, 140−297 °C, and 71−156 °C, with

30

salinities of 4.0−18.0 wt.% NaCl equiv., 4.2−17.3 wt.% NaCl equiv., 0.2−13.5 wt.% NaCl 1

31

equiv., and 1.2−7.6 wt.% NaCl equiv., respectively. The ore-forming fluids at the Caixiashan

32

deposit are characterized by high- to moderate temperatures, moderate salinities, and low

33

densities, belonging to the H2O−NaCl system. Hydrogen and oxygen isotope data indicate

34

that the ore-forming fluids at Caixiashan have a dominantly metamorphic signature and were

35

diluted by meteoric water. Carbon and oxygen isotope compositions of calcite, dolomite, and

36

marble demonstrate that the ore-forming fluids were primarily sourced from the dissolution

37

and low-temperature alteration of carbonates. Sphalerite, galena, pyrrhotite, and pyrite

38

samples from stage I to III record high δ34SV-CDT values between 11.2 and 16.1‰, indicating a

39

predominant sulfur source from the Precambrian marine sulfates by thermochemical sulfate

40

reduction. The 206Pb/204Pb, 207Pb/204Pb, and 208Pb/204Pb ratios of sulfide samples are similar to

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those of marble and carbonaceous slate of the Kawabulake Group, whereas they differ from

42

those of granitoid rocks, suggesting that ore-forming metals may have been primarily derived

43

from the Precambrian basement. All of these observations combined with the stable and

44

radiogenic isotope results reveal that the formation of the Caixiashan Zn-Pb deposit was

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dominated by a metamorphic system, and the ore-forming components were sourced from the

46

Precambrian basement.

47 48

Keywords: Fluid inclusions; Isotope systematics; Caixiashan Zn-Pb deposit; Eastern Tianshan;

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Northwest China

50 51

1. Introduction

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Sediment-hosted Pb-Zn deposits are a significant type of ores mainly hosted by siliceous

53

clastic rocks and carbonates that generally show no direct genetic association with intrusions,

54

and they account for a remarkable proportion of the lead and zinc production worldwide

55

(Leach et al., 2005, 2010). Most of them are further described as sedimentary exhalative

56

(SEDEX) deposits, which occur in passive margins, back-arcs, continental rifts, and sag

57

basins, and Mississippi Valley type (MVT) deposits, which occur in platform carbonate

58

sequences, typically in passive margins (Kucha et al., 2010; Leach et al., 2005, 2010), with a

59

small portion as the new Jinding-type deposits (Xue et al., 2007; Wang et al., 2014a). These 2

60

Pb-Zn deposits generally have no direct spatial or temporal relationship to the igneous

61

activity, which separates them from skarn or carbonate replacement ore deposits (Heta et al.,

62

2019).

63

The Central Asian Orogenic Belt (CAOB), known as one of the largest accretionary

64

orogens worldwide (Windley et al., 2007; Yang et al., 2009; Pirajno, 2010; Xiao et al., 2010),

65

extends more than 5000 km from Kazakhstan in the west to Siberia in the east (Fig. 1A; Jahn

66

et al., 2000; Seltmann et al., 2014; Deng and Wang, 2016), and is bounded by the

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Tarim-North China cratons to the south and by the Siberian craton to the north (Pirajno et al.,

68

2008; Wilhem et al., 2012; Wang et al., 2017b). The eastern Tianshan orogenic belt, located

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on the southern margin of the CAOB (Fig. 1B), is one of the most important polymetallic

70

belts in China and hosts more than 90 significant Cu, Mo, Au, Pb, Zn, Ag, Ni, and Fe deposits

71

(Mao et al., 2005; Shen et al., 2014a, 2014b; Wang et al., 2018c; Deng et al., 2017; Chen et

72

al., 2018; Zhao et al., 2016, 2019). Six sediment-hosted Pb-Zn(-Ag) deposits (i.e., Caixiashan,

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Jiyuan, Yuxi, Hongyuan, Shaquanzi, and Hongxingshan) occur in the Central Tianshan

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Terrane, whereas one Pb-Zn deposit (Aqishan) lies in the Aqishan-Yamansu arc belt (Liang et

75

al., 2005; Lu et al., 2012; Chen et al., 2012a; Li et al., 2016a, 2018).

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The Caixiashan Zn-Pb deposit, located in the western part of the eastern Tianshan

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orogenic belt, was discovered by the No. 1 Geological Party of Xinjiang Bureau of Geology

78

and Mineral Exploration in 2002. This deposit contains 131 million tons of ore reserves with

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an average Zn + Pb grade of 3.95% (Ag > 200 t), presenting the largest sediment-hosted

80

Zn-Pb deposits in eastern Tianshan. To characterize the mineralization of this deposit, several

81

geological studies have been conducted mainly focusing on the ore geology (Peng et al., 2006;

82

Liang et al., 2008), geochronology (Li et al., 2016b), stable isotope analyses (Gao et al.,

83

2007a), and fluid inclusion geochemistry (Li et al., 2016a). The source of ore-forming fluids

84

has been discussed with some controversy (Gao et al., 2006, 2007a; Cao et al., 2013; Li et al.,

85

2016a), and there also lacks a detailed geological work. Thus far, the source(s) of fluids and

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ore metals, and mineralization processes still remain poorly understood. In this contribution,

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we propose a detailed description of the geology and mineralization styles in the Caixiashan

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deposit based on field observations, detailed logging of drill holes, and petrographic analyses.

89

We further use fluid inclusions, and stable and radiogenic isotopes (hydrogen-oxygen isotopes 3

90

in quartz veins, carbon-oxygen isotopes in calcite and dolomite veins, and marble, and sulfur

91

and lead isotopes in sulfides) to better constrain the source(s) of ore-forming fluids and metals.

92

This new comprehensive dataset allows us to define the origin and evolution of the

93

hydrothermal system and to provide constraints on the possible sources of ore-forming metals

94

of the Caixiashan deposit. This information will also provide new clues for exploration of

95

sediment-hosted Zn-Pb deposits in eastern Tianshan.

96

2. Regional geology

97

The eastern Tianshan orogenic belt lies between the Junggar Basin to the north and the

98

Tarim Basin to the south (Fig. 1B), and is a typical Palaeozoic island arc system (Xiao et al.,

99

2004; Charvet et al., 2007; Wang et al., 2015b). From the north to the south, the eastern

100

Tianshan orogenic belt consists of the Bogeda-Haerlike Belt, the Jueluotage Belt, and the

101

Central Tianshan Terrane (Fig. 1C; Wang et al., 2014b, 2016c; Zhang et al., 2016a). The

102

Bogeda-Haerlike Belt is composed of well-developed Ordovician-Carboniferous volcanic

103

rocks, granites, and minor mafic-ultramafic intrusions, and hosts few porphyry Cu and Au

104

occurrences (Fig. 1C; Goldfarb et al., 2014; Wang et al., 2016c; Zhu et al., 2018). The

105

Jueluotage Belt, further subdivided into the Dananhu-Tousuquan arc in the north, the

106

Kanggur-Huangshan ductile shear zone in the centre, and the Aqishan-Yamansu arc in the

107

south (Zhang et al., 2016b; Ding et al., 2018; Wang et al., 2018a), is mainly composed of the

108

Middle Paleozoic volcanic and sedimentary rocks which were intruded by voluminous

109

Carboniferous-Jurassic felsic and mafic-ultramafic complexes (Zhou et al., 2010; Qin et al.,

110

2011; Gao et al., 2015; Wang et al., 2016d), and hosts many significant Cu, Mo, Au, and Ni

111

deposits (Zhang et al., 2003, 2006; Zhou et al., 2004; Qin et al., 2009; Wang et al., 2016a,

112

2018b, 2018c; Xiao et al., 2018). The Central Tianshan Terrane is mainly comprised of the

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Precambrian basement, which consists of the metamorphic rocks of the Xingxingxia Group,

114

including gneiss, quartz schist, migmatite, and marble, as well as low-grade metamorphic

115

clastic rocks and magnesium-rich carbonates of the Kawabulake Group (Qin et al., 2002; Cao

116

et al., 2013). This belt hosts some hydrothermal magnetite deposits (Jiang et al., 2002; Zhang

117

et al., 2005; Zhang et al., 2016b) and several sediment-hosted Zn-Pb deposits as mentioned

118

above (Wang et al., 2006; Xiao et al., 2009).

119

The eastern Tianshan orogenic belt has experienced a complex tectonic evolution from 4

120

the Late Paleozoic to the Mesozoic, involving the subduction of the paleo-Tianshan Ocean,

121

collision-accretionary, strike-slip motion, post-collisional, and intracontinental extension

122

between the Tarim Basin and the Junggar Basin (Zhang et al., 2008; Pirajno et al., 2011; Xiao

123

et al., 2013; Wang et al., 2016b). The main structures within eastern Tianshan exhibit a series

124

of approximately EW-trending faults, including the regional-scale Dacaotan, Kanggur,

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Yamansu, and Aqikuduke faults, and some sub-faults (Fig. 1C; Chen et al., 2012a; Zhang et

126

al., 2015; Wang et al., 2015a), and they generally control the various types of mineralization.

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The Central Tianshan Terrane is situated in the southern segment of the eastern Tianshan belt,

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clearly distinct from other composite arcs in eastern Tianshan for its residual Precambrian

129

basement, which is generally absent in other regional arc belts. The basement of central

130

Tianshan was dominantly formed in the early Mesoproterozoic, with significant crustal

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growth occurring in the Neoproterozoic (He et al., 2015; Huang et al., 2015), and underwent

132

extensive tectono-thermal-magmatic events in the late Neoproterozoic and Paleozoic, with the

133

metamorphism grade of basement rocks up to amphibolite facies and accompanied by

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multistage magmatic events (Hu et al., 1986, 2006, 2010). The tectonic evolution of the

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Central Tianshan Terrane in the Mesozoic is generally characterized by intraplate extension

136

and deformational activity. Coincident with complex tectonic evolution and extensive

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magmatic activity, metallogenic events occurred in the central Tianshan basement, resulting

138

in large scale Zn-Pb ore mineralization.

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3. Ore deposit geology

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The Caixiashan Zn-Pb deposit is in close proximity to the EW-trending Aqikuduke Fault

141

and is located at the western segment of the Central Tianshan Terrane (Fig. 1C). Orebodies

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are mainly hosted in the siliceous siltstone and dolomite marble of the first lithologic section

143

of the Kawabulake Group (Fig. 2A, B; Cao et al., 2013). The first lithologic section of the

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Kawabulake Group, exposed in the Caixiashan area, generally strikes to the northeast and

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nearly east-west and dips 45°~75° to the south. This section is composed stratigraphically

146

from bottom to top of lower dolomite marble, carbonaceous and siliceous siltstone, and

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mudstone with minor siliceous rocks interlayers (unit 1), and upper quartz sandstone locally

148

with marble lenses (unit 2). The Caixiashan deposit is mainly hosted in unit 1 (Fig. 2A). Li et

149

al. (2016b) reported the pyrite Re-Os ages of 1019 ± 70 Ma, 859 ± 79 Ma, and 837 ± 39 Ma, 5

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which are interpreted to be the depositional age of the Kawabulake group, the Zn-, and

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Pb-mineralization age, respectively, suggesting that there was a syn-sedimentary

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mineralization event in the early Neoproterozoic era.

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To date, the Caixiashan mining district possesses a reserve of ~131 Mt at 3.95% Zn + Pb,

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with orebodies outlined by a cut-off of 0.5% Zn + Pb (Li et al., 2016a). Four Zn-Pb

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mineralization zones have been identified, Zone I, II, III, and IV (Fig. 2A). The orebody II3

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strikes northeast and dips 65°~80° to the southeast, and is about 200 m long with thickness of

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~6 m. Its average grade of Zn + Pb is 3.7%, with accompanying silver as high as 19.2 g/t

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(Liang et al., 2005).

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There are three episodes of faults crosscutting the ore and alteration zones and wall rocks

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(Fig. 2A), among which the earliest one (F1) strikes northeast and dips 75° to the southeast,

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and might be synsedimentary faults contemporaneous with deposition of the Kawabulake

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Group (Li et al., 2016a). The second episode fault (F2) includes a series of faults that strike

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northeast and dip 75° to the south, partly crosscutting Zones I and IV. The third episode fault

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(F3) strikes north-northeast, significantly cutting ore and alteration Zone II and the wall rocks.

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Magmatism is well developed throughout the Caixiashan area. The intrusions mainly

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include granite, granodiorite, quartz diorite, microdiorite, and allgovite dikes (Gao et al., 2006;

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Wang et al., 2008; Cao et al., 2013). Liang et al. (2005) obtained a Rb-Sr age of 323 ± 6 Ma

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of the microdiorite at Caixiashan, which locally intruded the carbonate rocks of the

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Kawabulake Group, but showed no direct spatial contact relationship with Zn-Pb orebodies

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(Gao et al., 2007a, 2007b). Besides, Li et al. (2016c) reported that high Mg dioritic dikes

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within this area were emplaced in the Early Carboniferous at 353−348 Ma.

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The orebodies at Caixiashan include four morphological types: stratiform, stratiform-like,

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veins, and lenticular. Sulfide mineralization is characterized by massive veins and laminated

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ores. The Zn-Pb mineralization is hosted in the dolomite marble and the contact zone between

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carbonate and clastic rocks. The host rocks have experienced dolomitization, silicification,

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tremolization, chloritization alterations, and as shown in the cross section (Fig. 2B), the line

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44 illustrates a remarkable spatial relationship between the Zn-Pb mineralization and

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dolomitization. Sulfide minerals are dominated by sphalerite, galena, pyrite, and pyrrhotite,

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with minor arsenopyrite and chalcopyrite (Fig. 3A, E; Fig. 4B, C, G), and gangue minerals 6

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mainly include calcite, dolomite, quartz, tremolite, and chlorite (Fig. 3K, L). From the west to

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the east of the deposit area, the main sulfides change from pyrite, sphalerite and pyrrhotite to

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galena, exhibiting a mineralogical zonation.

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Based on field investigations and microscopic observations of sulfide and gangue

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mineralogy, textural relationships, and paragenetic sequence, Zn-Pb mineralization can be

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divided into four stages (Fig. 5). Stage I is characterized by abundant pyrite with variable

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calcite, quartz, and dolomite (Fig. 3A, B). Anhedral pyrite is generally cemented by calcite

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and quartz in the host rocks (Fig. 4A), and is locally replaced by sphalerite and pyrrhotite.

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Stage II is defined as the Zn-mineralization stage, represented by a widespread sulfide mineral

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assemblage of sphalerite + pyrrhotite ± arsenopyrite, with minor galena, and gangue minerals

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in this stage include calcite, dolomite, and quartz. This type of mineralization commonly

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occurs as veins (Fig. 3C, F), laminated veins (Fig. 3D), and massive ores (Fig. 3E). Large

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amounts of pyrrhotite replace pyrite (Fig. 4B) and some of them coexisting with euhedral

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arsenopyrite are replaced by sphalerite (Fig. 4C). Massive sphalerite is commonly replaced by

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galena (Fig. 4D). Stage III is the Pb-mineralization stage, containing an assemblage of

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abundant galena and pyrite, with minor chalcopyrite and sphalerite, and is accompanied by

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gangue minerals including calcite, dolomite, and quartz. In this stage, mineralization is

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represented as massive galena and pyrite ores (Fig. 3G), and quartz−galena−pyrite veins (Fig.

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3H). Galena veins commonly crosscut and replace early sulfides such as pyrrhotite and

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sphalerite, which show residual textures by replacement (Fig. 4E, G). It was also observed

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that minor chalcopyrite in this stage replaces earlier pyrrhotite (Fig. 4F). Notable among these

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minerals is the pyrite occurring in stage III, which is significantly different from earlier pyrite

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and exhibits subhedral to euhedral crystals; some of them replace pyrrhotite (Fig. 3I; Fig. 4H),

203

or occur as euhedral crystals in silicified marble (Fig. 3J). Stage IV quartz−calcite veins (Fig.

204

3K, L; Fig. 4I) with tremolite, dolomite, and locally with chlorite, are barren and represent the

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last hydrothermal stage. The supergene minerals near the surface in the Caixiashan deposit,

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such as jarosite and anglesite, generally record a post-ore secondary oxidation.

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4. Sampling and analytical methods

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Fluid inclusions in quartz and calcite from various kinds of veins or veinlets selected 7

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from the drill cores of different hydrothermal stages (I, II, III, and IV) were chosen for

210

microthermometric measurements and Laser Raman spectroscopic analyses. Polished thin

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sections were examined under microscope to characterize the phase, shape, size, and

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distribution of fluid inclusions. Fluid inclusion microthermometric measurements were

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performed at the Resources Exploration Laboratory of China University of Geosciences at

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Beijing, using a Linkam MDSG 600 heating-freezing stage, equipped with a Zeiss microscope.

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The stage enables measurements within the range of −196 to +600 °C. The measurements

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comprise ice-melting temperature (Tm,ice), halite dissolution temperature (Ts,halite), and

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homogenization temperatures of fluid phases in fluid inclusions (Th,total). Fluid inclusions were

218

initially cooled to about −190 °C at a rate of −5 °C/min and held for 5 min in order to make

219

sure the components in the inclusion were frozen. Heating rate was generally 1−5 °C/min

220

during the initial stages of each heating run and reduced to 0.3−1 °C/min close to the phase

221

change points. The Laser Raman spectroscopic analyses of selected inclusions were carried

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out using a Renishaw System-2000 microscopic confocal Laser Raman spectrometer, at the

223

Ore-Forming Laboratory of Institute of Mineral Resources, Chinese Academy of Geological

224

Sciences at Beijing, operating with an excitation wave length of 514.5 nm, a laser beam with

225

a power of 20 mW, and a spot size of 1 μm.

226

Salinities of aqueous (NaCl−H2O) inclusions, expressed as wt.% NaCl equiv., were

227

acquired by calculating via the final melting temperatures of ice (Bodnar, 1993), while the

228

salinities of halite daughter mineral-bearing inclusions were calculated using the method

229

outlined by Hall et al. (1988). Densities of fluid inclusions were calculated using the Flincor

230

program (Brown, 1989); details of this calculation method were given by Brown and Lamb

231

(1989). Homogenization conditions and isochores of individual fluid inclusion were

232

calculated with the HokieFlincs_H2O−NaCl program (Steele-MacInnis et al., 2012).

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Five quartz samples from the quartz−sulfide veins as well as three quartz samples from

234

the quartz−calcite veins were selected for H-O isotopic analyses. Hydrogen isotope

235

compositions of fluid inclusions in the quartz were analyzed using a MAT-253EM mass

236

spectrometer, while oxygen isotope compositions of quartz were determined by a Delta v

237

advantage mass spectrometer, at the Analytical Laboratory of the Beijing Research Institute

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of Uranium Geology. Oxygen was extracted from quartz by reaction with BrF5, and converted 8

239

to CO2 on a platinum-coated carbon rod for oxygen isotope analyses (Clayton and Mayeda,

240

1963). Hydrogen were measured on water in fluid inclusions hosted in quartz and the water

241

was released by heating the quartz to above 500 °C through an induction furnace, and then

242

reacted with chromium powder at 800 °C to obtain hydrogen for isotopic analyses (Wan et al.,

243

2005), which largely eliminated the potential effect of secondary inclusions on the H isotope

244

values. The isotope data are expressed in the delta (δ) notation as per mil (‰) deviation

245

relative to Vienna Standard Mean Ocean Water (V-SMOW); the analytic precisions are ±2‰

246

for δD (1σ) and ±0.2‰ for δ18Ο (2σ).

247

Seven calcite samples from quartz−calcite veins and two dolomite samples from

248

dolomite−calcite veins were selected for C-O isotopic analyses. Carbon and oxygen isotope

249

compositions were obtained using a Finnigan MAT-253 mass spectrometer at the Analytical

250

Laboratory of the Beijing Research Institute of Uranium Geology. Calcite and dolomite were

251

reacted with pure phosphoric acid to produce CO2. Carbon and oxygen isotope data are

252

reported in per mil (‰) relative to the Pee Dee Belemnite limestone (PDB) standard. The

253

analytical reproducibilities are ±0.1‰ for δ13C (2σ) and ±0.2‰ for δ18O (2σ). The value of

254

δ18OSMOW was calculated by δ18OSMOW = 1.03086 × δ18OPDB + 30.86 (Friedman and O'Neil,

255

1977).

256

Eighteen sulfide samples from sulfide-bearing veins and ores of different mineralization

257

stages were chosen for sulfur isotopic analyses. These sulfide samples were crushed, cleaned,

258

and sieved to 40 to 60 mesh, and then sulfide grains were carefully handpicked under

259

abinocular microscope to guarantee the purity of single sulfide separates (> 99%). Sulfur

260

isotope compositions were determined using a Delta v plus mass spectrometer, at the

261

Analytical Laboratory of the Beijing Research Institute of Uranium Geology. The sulfur

262

isotope compositions of sulfides were measured on SO2 obtained by placing the sulfide−CuO

263

composite (at weight ratio of 1/7) into a vacuum system heated to 980 °C (Robinson and

264

Kusakabe, 1975). Sulfur isotope data are reported as δ34S relative to the Vienna Canyon

265

Diablo Troilite (V-CDT); the analytical reproducibilities are ±0.2‰.

266

Eighteen samples including thirteen sulfides and five granitoid rocks were chosen for

267

lead isotopic analyses. Lead isotope compositions of most samples were measured by an

268

ISOPROBE-T Thermal Ionization Mass Spectrometer instrument, while those of five sulfide 9

269

samples (Table 3; 18CXS-1, 18CXS-33, and 18CXS-37) were determined by a Phoenix

270

Thermal Ionization Mass Spectrometer instrument, at the Analytical Laboratory of the Beijing

271

Research Institute of Uranium Geology. The external reproducibilities (2σ) of the NBS

272

SRM981 standard are 0.12% for 206Pb/204Pb, 0.10% for 207Pb/204Pb, and 0.21% for 208Pb/204Pb,

273

respectively. The procedural blanks were between 1 and 2 ng for Pb.

274

5. Fluid inclusion results

275

5.1. Classification strategy

276

Fluid inclusion studies were performed in the quartz- and calcite-bearing veins from the

277

hydrothermal stages. According to Chi and Lu (2008), we selected the inclusions occurring

278

as isolated, random, clusters, or along growth zones, including fluid inclusion assemblages

279

(FIAs), which are interpreted to be primary inclusions, for microthermometric study, in order

280

to make sure that the data are valid and truly represent the ore-forming fluids trapped in fluid

281

inclusions. Secondary inclusions, occurring in annealed trails and penetrating crystal

282

boundaries, were avoided. Based on phase proportion at room temperature and phase

283

transformations during cooling and heating processes (Chi et al., 2017; Goldstein and

284

Reynolds, 1994), five types of primary inclusions in the quartz and calcite veins have been

285

identified. They are classified as: liquid-rich two-phase inclusions (L-type), pure-liquid

286

phase inclusions (PL-type), vapor-rich two-phase inclusions (V-type), pure-vapor phase

287

inclusions (PV-type), and halite-bearing two- or three-phase hypersaline inclusions (H-type).

288

L-type inclusions are the most common type in the quartz and calcite in all stages (Fig.

289

6G, K), generally occurring as FIAs (Fig. 6A, H, J, L), isolated (Fig. 6I), some of which occur

290

along growth zones. These inclusions contain a liquid phase and a vapor phase, with the vapor

291

bubbles occupying 5−45% of the inclusion volume, and a few of them have an opaque

292

daughter mineral (Fig. 6E). They are usually oval, elongated, rhombic, and irregular in shape,

293

and range from 2−15 μm in size, most of which between 5 and 10 μm. PL-type inclusions are

294

widespread in the quartz and calcite crystals. They are distributed randomly, occurring as oval,

295

square or irregular, and usually coexist with L-type inclusions but are much smaller than these

296

two-phase inclusions (Fig. 6K, L), ranging from 1−6 μm in diameter. There are also some

297

secondary PL-type inclusions, which always occur in a certain array and penetrate crystal 10

298

boundaries. V-type inclusions are generally observed in stages I and II quartz- and

299

calcite-bearing veins, with the vapor phase ratios in the range of 50−80% volumetrically.

300

They usually range from 6−12 μm in diameter and are elliptical, rounded, and irregular in

301

shape, occurring as isolated entities (Fig. 6B) or coexisting with H-type inclusions in stage I

302

(Fig. 6C, D), while as an isolated one (Fig. 6F) or coexisting with L-type inclusions in stage II

303

calcite crystals (Fig. 6G). PV-type inclusions were occasionally observed in the quartz and

304

calcite veins, and occur as oval in shape with dark colour, generally coexisting with V-type

305

inclusions (Fig. 6C). H-type inclusions are rare and occur only in the stage I quartz veins,

306

with a range of 5 to 12 μm in size, usually coexisting with V-type inclusions (Fig. 6C, D).

307

These inclusions consist of liquid phase, vapor phase, and also a halite crystal (Fig. 6C), some

308

of which are comprised of liquid + halite mineral (Fig. 6D). The halite-bearing three-phase

309

inclusions are homogenized by halite dissolution after bubble disappearance.

310

Primary L-type inclusions are prevalent from stage I to stage IV, whereas the V-type

311

inclusions are generally present in stages I and II, and the H-type inclusions are only observed

312

in stage I quartz veins. In the late quartz−calcite stage, only the L-type inclusions are

313

observed and measured.

314

5.2. Microthermometry and Raman spectroscopy

315

Microthermometric measurements were conducted on the L-, V-, and H-type fluid

316

inclusions hosted in the quartz and calcite minerals. All the L-type fluid inclusions

317

homogenized to liquid phase, whereas V-type inclusions homogenized to the vapor phase in

318

the process of heating. The microthermometric data and calculated parameters for single

319

inclusion of different paragenetic stages are shown in Table 1 and graphically illustrated in

320

Figures 7 and 8.

321

For the early ore stages, L-, V-, and a small number of H-type fluid inclusions were

322

observed in quartz and calcite crystals. The L-type inclusions yielded homogenization

323

temperatures of 336−488 °C, variable salinities of 4.0−16.1 wt.% NaCl equiv., and estimated

324

densities of 0.55−0.81 g/cm3. The V-type inclusions yielded relatively high homogenization

325

temperatures of 382−494 °C, salinities of 12.9−18.0 wt.% NaCl equiv., and densities of

326

0.57−0.78 g/cm3. Three H-type fluid inclusions in this stage yielded halite dissolution 11

327

temperatures of 327−344 °C, corresponding to salinities values of 40.4−41.9 wt.% NaCl

328

equiv., with densities of 1.31−1.32 g/cm3.

329

Homogenization temperatures of L-type inclusions in the Zn-mineralization stage are in

330

the range of 240 to 357 °C, with corresponding salinities from 4.2 to 17.1 wt.% NaCl equiv.

331

and the densities from 0.61 to 0.93 g/cm3. The V-type inclusions in this stage yielded

332

homogenization temperatures of 255 to 350 °C, with salinities and densities of 8.0−17.3 wt.%

333

NaCl equiv., 0.77−0.89 g/cm3, respectively.

334

In the Pb-mineralization stage, the microthermometric measurements were only

335

conducted on L-type fluid inclusions, which yielded homogenization temperatures of 140 to

336

297 °C, with peak Th values of 160 to 240 °C, with corresponding salinities of 0.2 to 13.5 wt.%

337

NaCl equiv. The estimated fluid densities are between 0.80 and 1.01 g/cm3.

338

Only L-type inclusions were observed and analysed in the stage IV quartz−calcite veins.

339

These inclusions yielded Th and salinities ranging from 71 to 156 °C, and from 1.2 to 7.6 wt.%

340

NaCl equiv., respectively. The estimated densities of these inclusions are in the range of

341

0.94−1.02 g/cm3.

342

Some of the representative fluid inclusions were chosen for the Laser Raman analyses to

343

constrain their gaseous and liquid compositions, and the results are shown in Fig. 9. The

344

vapor phases of the L-type inclusions in quartz from the ore stages are dominated by H2O,

345

with two low-intensity peaks at 3080 and 1120 unidentified (Fig. 9A, B). Another L-type

346

inclusion shows that the compositions of vapor phases are mainly H2O with trace amount of

347

SO2 (Fig. 9D). The liquid phases of the L-type inclusions consist mainly of H2O but also

348

contain minor CO32- and SO2 (Fig. 9C). The Raman results herein suggest that the

349

ore-forming fluids at Caixiashan are dominated by the H2O−NaCl system.

350

5.3. Trapping pressure and depth estimation

351

Trapping pressure can be estimated only when the exact trapping temperature is known,

352

or if fluid boiling or immiscibility occurred in the system at the time of fluid entrapment

353

(Roedder and Bodnar, 1980; Brown and Hagemann, 1995). Pressures determined for

354

non-boiling assemblages are derived from the homogenization temperatures and represent

355

minimum

estimation

values

(Rusk

et 12

al.,

2008).

The

excel

spreadsheet

356

HokieFlincs_H2O−NaCl was used to constrain the trapping temperature−pressure conditions

357

and estimate the ore formation depth (Steele-MacInnis et al., 2012). Minor H-type inclusions

358

are identified in the stage I, and they sometimes coexist with V-type inclusions (Fig. 6C, D),

359

indicating that these inclusions record a fluid immiscibility event. The coexistence of V-type

360

with L-type inclusions in stage II (Fig. 6G), which yield quite similar homogenization

361

temperatures, presents evidence that there also has been an existence of fluid immiscibility in

362

the main stage. Thus, homogenization temperatures of these ore stages are interpreted to

363

closely approximate actual trapping temperatures (Roedder and Bodnar, 1980). For the late

364

stage IV, only the liquid-rich fluid inclusions have been recognized and the lack of evidence

365

for fluid boiling suggests that the estimated pressures could only represent the minimum

366

values (Rusk et al., 2008; Wang et al., 2018c).

367

Four fluid inclusions of stage I give high trapping pressures of 497 to 537 bars,

368

corresponding to 1.8 to 2.0 km when assuming a lithostatic condition (a rock density of 2.75

369

g/cm3). Meanwhile, trapping pressures estimated for the remaining inclusions of stage I are

370

between 99 and 320 bars (Fig. 10; average of ~201 bars), corresponding to 0.4−1.2 km under

371

lithostatic conditions, or 1.0−3.3 km under hydrostatic pressure conditions. The trapping

372

pressures of inclusions in ore-forming stage II are estimated to range from 31 to 162 bars,

373

with an average of 94 bars, corresponding to depths of 0.3−1.7 km under the hydrostatic

374

conditions, which is consistent with a trapping depth of the temporally earlier pyrite

375

mineralization stage (∼0.4−2.0 km, under lithostatic pressure), indicating there might be an

376

occurrence of pressure transition from lithostatic to hydrostatic system during the deposition

377

of sphalerite minerals in main ore stage II. The inclusions in stage III then yield much lower

378

pressures ranging from 3 to 78 bars, corresponding to 0.1−0.8 km under hydrostatic pressure.

379

When the fluids cooled to 71−157 °C (stage IV), inclusions in the quartz−calcite veins yielded

380

trapping pressures below 10 bars (Fig. 10), with corresponding depths close to the surface. In

381

general, the data presented together with the discussions above suggest that the ore

382

mineralization at Caixiashan occurred at the depths of ~0.8 to 2.0 km or even shallower, as an

383

estimated minimum depth of ore formation.

384

6. Isotope results

13

385

6.1. Sulfur isotope compositions

386

Sulfur isotope data are shown in Table 2 and plotted in Fig. 11. Eighteen sulfide samples

387

from Caixiashan give the δ34SCDT values of 11.2 to 16.1‰. The δ34SCDT values of sphalerite

388

and galena separates range from 13.5 to 15.5‰ and from 11.2 to 13.8‰, averaging of 14.3‰

389

(n=6) and 12.3‰ (n=3), respectively. Pyrrhotite separates have δ34SCDT values of 12.0 to 16.0‰

390

with an average of 13.9‰ (n=4). Pyrite grains yield δ34SCDT values of 11.2 to 16.1‰, with an

391

average of 15.0‰ (n=5). In short, sulfur isotope compositions at the Caixiashan deposit are

392

characterized by considerably positive values, and the δ34S values of sulfides from different

393

ore stages display a gradual decreasing trend as showed in Fig. 11A.

394

6.2. Lead isotope compositions

395

Lead isotope data of sulfides, granitoids and rocks from the strata are shown in Table 3 206Pb/204Pb, 207Pb/204Pb,

396

and Fig. 12. Thirteen sulfide samples at Caixiashan have

397

208Pb/204Pb

398

Five granitoid samples yield

399

ratios from 15.574 to 15.655, and

400

dolomite marble sample from the Kawabulake Group yields

401

208Pb/204Pb

402

interlayers sample has corresponding ratios of 17.226, 15.531, and 37.083, respectively. As

403

depicted in Fig. 12, except the granitoid samples, all of these sulfides show similar lead

404

isotope compositions.

405

6.3. Hydrogen and oxygen isotope compositions

406

and

ratios of 17.173 to 17.807, 15.507 to 15.613, and 36.959 to 38.016, respectively. 206Pb/204Pb

ratios ranging from 18.195 to 20.013,

208Pb/204Pb

207Pb/204Pb

ratios from 38.186 to 39.572, respectively. A 206Pb/204Pb, 207Pb/204Pb,

and

ratios of 17.184, 15.527, and 37.013, while one carbonaceous slate with siltstone

Hydrogen and oxygen isotope compositions obtained from quartz veins of different ore

407

stages are listed in Table 4 and plotted in Fig. 13. The measured δ18Oquartz

408

different mineralization stage display a relatively homogeneous range from +13.0 to +15.5‰

409

whereas the δDV-SMOW values show a variable range from −104.5 to −74.7‰. The δ18O values

410

of hydrothermal fluids were calculated using the equation of Clayton et al. (1972),

411

1000lnaqtz−water = 3.38 × 106 × T−2 − 3.40, together with the measured δ18Oquartz values, and the

412

correspondingly average homogenization temperatures of the fluid inclusions from quartz 14

(SMOW)

from

413

minerals in each stage were used to represent that T mentioned above (in degrees Kelvin).

414

Consequently, the calculated δ18OH2O values of fluids from stage I to IV are +10.0‰, from

415

+5.3 to +8.1‰, from +3.4 to +3.8‰, and from −4.4 to −2.3‰, respectively (Table 4).

416

6.4. Carbon and oxygen isotope compositions

417

Carbon and oxygen isotope data are reported in Table 5 and plotted in Fig. 14, some of

418

which were taken from Gao et al. (2007a) and Cao et al. (2013). Seven calcite samples in this

419

study give the δ13CPDB values of −3.8 to −2.2‰ and δ18OPDB values of −19.8 to −17.4‰. The

420

two dolomite samples yield the δ13CPDB and δ18OPDB values of −3.0 to −2.6‰ and −14.6 to

421

−14.2‰, respectively. According to the equation (δ18OSMOW = 1.03086 × δ18OPDB + 30.86)

422

provided by Friedman and O'Neil (1977), the calculated δ18OSMOW values of calcite, dolomite,

423

and marble are from +10.5 to +17.4‰, from +15.9 to +16.2‰, and from +14.7 to +19.1‰,

424

respectively.

425

7. Discussion

426

7.1. Sources of the metals and ore-forming fluids

427

All sulfur isotope compositions of sulfides from Caixiashan display remarkably positive

428

δ34S values of 11.2 to 16.1‰ (Table 2), significantly differing from those of magmatic

429

hydrothermal deposits (−3 to +1‰; Hoefs, 2009). The δ34S values of hydrothermal fluids can

430

be estimated from those of sulfides and sulfates based on the oxygen fugacity (ƒO2) and

431

temperature during mineral precipitation (Ohmoto and Goldhaber, 1997; Ohmoto and Rye,

432

1979). The existence of the ƒO2 indicator mineral pyrrhotite (pH > 6, T < 500°) suggests that

433

there is a negligible δ34S difference between the sulfides and fluids (Zheng and Chen, 2000).

434

Meanwhile, a trend of δ34SGn < δ34SPo < δ34SSp < δ34SPy given by sulfides from Caixiashan also

435

suggests an overall equilibrium between the sulfides and H2S components in the fluid system,

436

indicating that the δ34S values of the sulfides could actually represent the δ34S values of

437

hydrothermal fluids. Thus, the δ34S values of hydrothermal fluids in the Caixiashan deposit,

438

with a limited range from 11.2‰ to 16.1‰, are close to those of the Precambrian seawater

439

sulfates (δ34S = 15−23‰; Holser, 1977; Lu et al., 2018), and the country rocks at Caixiashan

440

(δ34S = 6.2−17.0‰; Li et al., 2018). Several possible mechanisms resulting in sulfur isotope 15

441

fractionation include organic sulfate reduction (OSR), bacterial sulfate reduction (BSR) and

442

thermochemical sulfate reduction (TSR) (Zheng and Chen, 2000; Hoefs, 2009). In contrast to

443

OSR, particularly from ~100 to 150 ºC (Basuki et al., 2008), BSR would only be possible

444

below 120 °C (Jorgenson et al., 1992; Dixon et al., 1996), resulting in large and

445

heterogeneous

446

occurs at higher temperatures and may lead to smaller and more homogeneous sulfur

447

fractionations (Hoefs, 2009). Considering the temperatures for ore formation at Caixiashan

448

are mainly in the range of ~160−340 ºC, as well as the very positive δ34S values for sulfides,

449

TSR is the most likely mechanism of sulfide formation. This interpretation for the δ34S values

450

is also supported by the sulfur analyses of other sediment-hosted Zn-Pb(-Ag) deposits in the

451

Central Tianshan Terrane, such as Hongyuan (15.8 to 16.8‰; Lu et al., 2018), Hongxingshan

452

(8.8 to 12.7‰; Xiao et al., 2009), and Yuxi (−2.6 to +15.6‰; Zhou et al., 1999). Therefore,

453

the δ34S values of hydrothermal fluids at Caixiashan were most likely inherited from the

454

reduction of isotopically heavy Precambrian marine sulfates by TSR reactions.

34S-depletions

generally dominated by negative δ34S values. TSR generally

455

The lead isotope compositions of various sulfides, granitoids, and host rocks are listed in

456

Table 3. As depicted in the Pb isotope diagrams (Fig. 12), these lead compositions fall near

457

the field of the orogen evolution curve in the

458

while appearing between the mantle and the upper crust and also near the orogen line in the

459

207Pb/204Pb

460

nature, and may have been sourced from a mixture of upper crust and possible mantle

461

components (Zartman and Doe, 1981; Zartman and Haines, 1988). Moreover, the sulfides,

462

marble, and carbonaceous slate show similar Pb isotopic distributions with a restricted

463

variation but distinct from those of granitoids. The carbonates at Caixiashan and the

464

Precambrian basement were characterized by remarkably high Pb and Zn contents (Peng et al.,

465

2007), which could contribute enough lead, zinc, and other metals to the ore-forming fluids.

466

Therefore, it is strongly suggested that the ore-forming metals at Caixiashan were mainly

467

derived from the Precambrian basement.

vs.

206Pb/204Pb

208Pb/204Pb

vs.

206Pb/204Pb

diagram (Fig. 12A),

diagram (Fig. 12B). These suggest that the lead was of mixture

468

The δDH2O and calculated δ18OH2O values vary from −104.5 to −74.7‰, and from −4.4 to

469

+10.0‰, respectively (Table 4). The δ18OH2O value for stage I (+10.0‰) is slightly higher

470

than those of magmatic water (+7 to +9‰; Taylor, 1974; Taylor and Sheppard, 1986), more 16

471

likely in line with metamorphic water, rather than a typical magmatic origin. The δ18OH2O

472

values of fluids for stages II and IV veins range from +5.3 to +8.1‰, from +3.4 to +3.8‰,

473

and from −4.4 to −2.3‰, respectively, with a gradual decreasing trend. In the δD-δ18O

474

diagram (Fig. 13), the isotope compositions of fluids for stage I plot near the lower limit of

475

metamorphic water box, whereas those of stage II to IV show a decreasing trend towards the

476

meteoric water line. All isotope compositions of quartz gradually decrease from early to late,

477

indicative of a possible involvement of meteoric water (Taylor, 1974). Moreover, the

478

Precambrian

479

Neoproterozoic (Cao et al., 2013). As illustrated in the δD vs. δ18O diagram (Fig. 13), both

480

hydrogen and oxygen isotope values gradually decrease from stage I to IV, indicating that the

481

ore-forming fluids in the Caixiashan Zn-Pb deposit have a dominantly metamorphic signature

482

and were diluted by abundant meteoric water later.

basement

has

undergone

extensive

regional

metamorphism

in

the

483

On the basis of measured carbon and oxygen isotope data (Table 5), we have further

484

evaluated the source of carbon in the ore-forming fluids. Nine calcite and dolomite samples

485

from the Caixiashan deposit yield δ13CPDB values of −3.7 to −2.2‰, together with the δ13CPDB

486

values of −2.4 to +0.1‰ from Gao et al. (2007a) and −6.7 to −0.5‰ from Cao et al. (2013),

487

which show a narrow range of carbon isotopes. In the δ13CPDB vs. δ18OSMOW diagram (Fig. 14),

488

C-O isotopic compositions of calcite, dolomite, and marble at Caixiashan are distinctly

489

different from those of sedimentary organic matter, magma-mantle carbonate rocks, but are

490

slightly lower than and broadly consistent with those of marine carbonates (~0‰; Hoefs,

491

2009). The stage I samples with higher δ13CPDB and δ18OSMOW values are similar to those

492

expected from dissolution of carbonate, whereas the stages II and III calcite samples yield

493

lighter isotope values, which may reflect the effect of low-temperature alteration. In addition,

494

the lowest δ18OSMOW value given by one stage IV calcite may reflect significant addition of

495

meteoric water. C-O isotope analyses demonstrate that the marine carbonates are principal

496

carbon source by the dissolution and low-temperature alteration, which coincides with the

497

well-developed carbonate platform of the Kawabulake Group in the central Tianshan in the

498

Mesoproterozoic (Cai et al., 2013). Thus, we propose the carbon in hydrothermal fluids at

499

Caixiashan was mainly sourced from dissolution and low-temperature alteration of carbonates

500

with little sedimentary contamination, such as from clastic wall rocks. 17

501

7.2. Evolution of ore-forming fluids

502

Petrographic studies of fluid inclusions indicate that the major types of inclusions (L-, V-,

503

and H-) were present in hydrothermal calcite- and quartz-bearing veins during the different

504

mineralization stages at the Caixiashan Zn-Pb deposit. The nature and evolution of

505

hydrothermal fluids in terms of their composition and temperature, salinity, and associated

506

mineralization, is reconstructed to establish a model of fluid evolution and help reveal the ore

507

formation process.

508

In the earlier ore stage I, the hydrothermal fluids that have extracted the ore-forming

509

elements (e.g., Zn, Pb, and Fe) from the basement ascended, filtered through, and dissolved

510

the calcites and dolomites, which subsequently caused the dolomitization alteration. Abundant

511

pyrites were then precipitated with gangue minerals such as calcite, dolomite, and quartz.

512

Fluid inclusions in the quartz and calcite crystals of this stage are mainly of the V- and

513

L-type, yielding relatively high homogenization temperatures (336−494 °C), and intermediate

514

salinities (4.0−18.0 wt.% NaCl equiv.). Three H-type inclusions give halite dissolution

515

temperatures of 327−344 °C, with salinities of 40.4−41.9 wt.% NaCl equiv. The coexistence

516

of vapor-rich inclusions and halite-bearing hypersaline inclusions suggests a fluid

517

immiscibility event in this stage, also evidenced by the coexistence of L-type with V-type

518

inclusions yielding relatively similar homogenization temperatures. However, the vapor-rich

519

inclusions generally have similar or even higher salinities than liquid-rich inclusions, which

520

cannot be well explained by fluid boiling (Chi et al., 2017), and they might not represent two

521

immiscible phases produced by a boiling event.

522

The L-type fluid inclusions in the Zn-mineralization stage are homogenized at

523

240−357 °C, with moderate salinities of 4.2−17.1 wt.% NaCl equiv., whereas V-type

524

inclusions yielded homogenization temperatures of 255 to 350 °C, and salinities of 8.0 to 17.3

525

wt.% NaCl equiv. Abundant sphalerite and pyrrhotite with limited galena sulfides were

526

precipitated from the fluids in this stage. As for the Pb-mineralization stage, characterized by

527

the precipitation of galena and galena-rich ores, measurements on the L-type fluid inclusions

528

show lower homogenization temperatures of 140−297 °C, and salinities of 0.2−13.5 wt.%

529

NaCl equiv. From stage I to III, there is a gradually decreasing trend in homogenization

530

temperatures and salinities (Fig. 8), suggesting a dominant cooling process during the 18

531

evolution of the ore-forming fluids (Shepherd et al., 1985). When the fluids subsequently

532

ascended to shallower depths in the late stage, L-type fluid inclusions in stage IV quartz and

533

calcite have much lower homogenization temperatures of 71−156 °C, with salinities of

534

1.2−7.6 wt.% NaCl equiv. In response to the decrease in temperature and pressure of the fluid

535

system, quartz, calcite, and tromelite gradually precipitated from the fluids.

536

Therefore, the hydrothermal fluids of the Caixiashan deposit are characterized by

537

relatively high- to moderate temperatures and moderate salinities in early ore stage, with a

538

gradual decline in temperatures and salinities from main stages to the late quartz−calcite stage.

539

The fluids responsible for ore formation at Caixiashan belong to a H2O−NaCl system as

540

evidenced by the Laser Raman analyses, and the evolution of the hydrothermal fluids is

541

dominated by a significant fluid cooling process.

542

7.3. Ore-forming mechanisms and genetic model

543

Generally, the transport of zinc and lead in hydrothermal fluids mainly occurs through

544

chloride and hydrosulfide complexes, which is demonstrated by many studies on the

545

migration and precipitation of zinc and lead in ore-forming fluid systems (Ruaya and Seward,

546

1986; Sverjensky et al., 1997; Tagirov et al., 2007; Tagirov and Seward, 2010; Seward et al.,

547

2014; Mei et al., 2015; Zhong et al., 2015). For moderate- to high-salinity hydrothermal fluids,

548

the existence of chloride complexes is of great importance to the transport and deposition of

549

lead and zinc (Seward, 1984; Chen et al., 2014). These chloride complexes are much more

550

stable than those of hydrosulfide or sulfide species under high-temperature and

551

moderate-salinity conditions (Zhang et al., 2016). Given that the fluids for Zn- and

552

Pb-mineralization at Caixiashan are marked by medium- to high temperatures mainly of

553

160−340 ºC, the lead and zinc were likely transported primarily by chloride complexes in the

554

ore-forming fluids. This would naturally be accompanied by minor presence of hydrosulfide

555

complexes during the migration processes.

556

The precipitation of sulfides from the hydrothermal fluids is generally related to the

557

processes that can influence the instability of metal complexes transported by ore-forming

558

fluids. The main mechanisms that could cause the destabilization of metal complexes and ore

559

precipitation include fluid boiling (immiscibility), fluid mixing, decreases in temperature, 19

560

pressure, and salinity, and fluid-rock interactions (Hemley et al., 1992; Zhai et al., 2011; Seo

561

et al., 2012; Peng et al., 2016). Fluid immiscibility events were unlikely applied to most fluid

562

inclusions analyzed at Caixiashan, though petrographic studies illustrate that immiscibility

563

events were locally identified in the ore stages I and II, not pervasively occurring in all stages.

564

These observations suggest that fluid immiscibility was not a key parameter controlling ore

565

precipitation. As shown in Table 1 and Fig. 7, the hydrothermal fluid temperatures of the

566

Caixiashan deposit decreased strikingly from stage I (336−494 °C) to III (140−297 °C).

567

However, the salinities decreased less significantly, from stage I (4.0−18.0 wt.% NaCl equiv.)

568

to III (mainly 0.8−13.5 wt.% NaCl equiv.). These results indicate that the significant

569

temperature decrease is more likely to account for the ore precipitation. Ruaya and Seward

570

(1986) have demonstrated that the stability of chloride complexes of zinc and lead generally

571

varies with temperature change. Moreover, the temperature decrease could reduce the

572

solubility of the ore-forming substance in the fluid system (Brimhall and Crerar, 1987). With

573

the significant decline in temperature, the chloride complexes of zinc and lead destabilize

574

(Reed and Palandri, 2006); meanwhile, thermochemical sulfate reduction of marine sulfates

575

occur, which could produce large amounts of S2− into the fluids, accompanied by the sulfur

576

contributions from country rocks, consequently resulting in the large-scale deposition of

577

abundant ore sulfides such as sphalerite (ZnS, usually Fe-rich) and galena (PbS), along with

578

other sulfides.

579

As depictured in Fig. 8, the evolution of hydrothermal fluids is mainly characterized by a

580

process of fluid cooling, but the plots also show an overall trend of fluid mixing. Hydrogen

581

and oxygen isotope values at Caixiashan gradually decrease from stage I to IV (Fig. 13),

582

indicative of fluid mixing event. When the hydrothermal fluids with a dominantly

583

metamorphic signature migrated upward, there would be a subsequent involvement of

584

meteoric water. Mixtures of different fluid signatures and metal contents contributed to the

585

deposition of sulfide ores from the hydrothermal system. As mentioned above, dolomitization

586

is one of the common hydrothermal alteration types in the Caixiashan deposit, which could

587

significantly increase permeability of the wall rocks, and is further favorable to the migration

588

of ore-forming fluids and sulfide precipitation. Furthermore, the pervasive dolomitization

589

could, to some extent, increase the pH of the fluid system and thus, promote the precipitation 20

590

of sphalerite, galena, and other sulfides from the fluids. It is also observed that several

591

orebodies are hosted in dolomitization zones exhibiting evidence that the ore formation at

592

Caixiashan is closely associated with this type alteration. The gradually decreasing O isotope

593

compositions with relatively stable C isotopes of dolomite, calcite, and marble are also in

594

favor of the water-rock reactions between carbonate rocks and the fluids (Liu et al., 2004),

595

consistent with the recrystallization of dolomite and calcite. Therefore, significant

596

temperature decrease and fluid mixing were considered to be effective causes of the formation

597

of the giant Caixiashan Zn-Pb deposit, while dolomitization provided additional favorable ore

598

precipitation conditions.

599

The Caixiashan Zn-Pb deposit shares some similar geological characteristics with the

600

typical SEDEX, MVT, and Irish-type deposits that could be viewed as an important

601

transitional ore type between SEDEX and MVT (Wilkinson, 2014). However, the Caixiashan

602

deposit differs from those types in some respects. SEDEX deposits are generally considered

603

to be syngenetic, but geological and geochemical studies illustrate that the formation of the

604

Caixiashan deposit is related to the dissolution and alteration of carbonates and is later than

605

the diagenetic processes, which suggests an epigenetic nature. The fluid temperatures of the

606

Caixiashan ore formation are mainly of 160 to 340 ºC, generally higher than those of typical

607

MVT (50−250 °C; Leach et al., 2005) and Irish-type deposits (100−240 °C; Wilkinson, 2003).

608

The salinities of most MVT fluids are typically from 10 to 30‰ (Leach et al., 2010), which

609

are higher than those of Caixiashan. Furthermore, the sulfur source for the Irish-type deposit

610

is dominantly of bacteriogenic origin indicated by negative δ34S values (δ34S = −15‰ ± 10‰;

611

Wilkinson, 2003, 2005), differing from the sulfur isotope data at Caixiashan with a range of

612

11.2 to 16.1‰. Therefore, these features suggest that the Caixiashan deposit is likely not a

613

typical SEDEX, MVT, or Irish-type deposit.

614

It was suggested that the regional Zn-Pb mineralization in central Tianshan was

615

genetically associated with the Carboniferous granitic intrusions and the ore-forming fluids

616

were magmatic in origin (Xiao et al., 2009; Cao et al., 2013). However, several researchers

617

have revealed some details of tectonic evolution of the Central Tianshan Terrane, which is

618

characterized by several episodes of magmatism in the Mesoproterozoic and Neoproterozoic

619

(Huang et al., 2015; Wang et al., 2017a). Li et al. (2016b) also reported a Neoproterozoic age 21

620

for the Caixiashan ore deposition by pyrite Re-Os dating, which shed new light on the

621

understanding of the regional Zn-Pb mineralization. The ore mineralization at Caixiashan

622

displays remarkable stratigraphic control and is dominantly hosted by dolomite marble of the

623

Kawabulake Group, occurring as massive, veins, and less stratiform-like orebodies, with an

624

ore mineral assemblage of pyrite + sphalerite + galena + pyrrhotite (± chalcopyrite ±

625

arsenopyrite) closely associated with dolomitization. Combined with these above-mentioned

626

characteristics, isotope studies and fluid inclusion results, a simplified genetic model for

627

Caixiashan deposit is proposed as follows (Fig. 15).

628

The heat produced by magmatic activity during the Mesoproterozoic to Neoproterozoic

629

triggered the evolution of the metamorphosed water, forming the initial hydrothermal fluids.

630

Then the fluids gradually migrated within the permeable basement and ascended, extracting

631

the metals such as Zn, Pb, and Fe from the basement rocks. Subsequently, the evolved

632

ore-forming fluids, primarily characterized by metamorphic water which experienced an

633

involvement of meteoric water later, filtered into the dolomite marble through

634

syn-sedimentary faults and dissolved calcites and dolomites, which resulted in pervasive

635

dolomitization and provide favorable conditions for mineralization. This, coupled with

636

efficient S sources from thermochemical sulfate reduction, caused large-scale deposition of

637

the sulfide ores. During Early Carboniferous, these Zn-Pb orebodies might have been

638

reworked by the granitic or dioritic dikes (Li et al., 2016a). We suggest that the

639

Mesoproterozoic-Neoproterozoic magmatic activity acted as a trigger for the regional Zn-Pb

640

mineralization by producing heat but did not provide main ore-forming materials into the

641

fluids, which is in accordance with the notion that the Proterozoic era is one important epoch

642

of large-scale Zn-Pb mineralization in China (Wang et al., 2014a), and the formation of the

643

Caixiashan deposit might be coeval or slightly anterior to the Neoproterozoic.

644

8. Conclusions

645

(1) The Caixiashan Zn-Pb deposit, located in the western part of the Central Tianshan

646

Terrane, is mainly hosted by the dolomite marble of the Mesoproterozoic Kawabulake Group,

647

and its hydrothermal mineralization processes could be divided into four stages, i.e., calcite +

648

dolomite + quartz + pyrite stage (I), calcite + dolomite + quartz + spahlerite + pyrrhotite ± 22

649

arsenopyrite stage (II), calcite + dolomite + quartz + galena + pyrite ± chalcopyrite stage (III),

650

and late quartz + calcite stage (IV).

651

(2) Hydrogen and oxygen isotope data at Caixiashan indicate that the ore-forming fluids

652

had a dominantly metamorphic signature and were then diluted by meteoric water. Carbon

653

and oxygen isotope compositions of calcite, dolomite, and marble suggest that hydrothermal

654

fluids were primarily sourced from the dissolution and low-temperature alteration of

655

carbonates. Sulfur and lead isotope results reveal that the ore-forming components were

656

sourced from the Precambrian basement.

657

(3) The major types of fluid inclusions are the liquid-rich two-phase, vapor-rich

658

two-phase, and halite-bearing inclusions. The ore-forming fluids are characterized by

659

relatively high- to moderate temperatures, moderate salinities, and low densities, and are

660

dominated by the NaCl−H2O system. Combined with fluid inclusion results and isotopic

661

studies, we suggest that the temperature decrease, fluid mixing, and pervasive dolomitization

662

are the key factors resulting in large-scale ore precipitation.

663

Acknowledgments

664

This

research

was

supported

by

the

DREAM

project

of

MOST

China

665

(2017YFC0601202), the National Natural Science Foundation of China (41772073, 41572066,

666

and 41702079), the 111 Project of the Ministry of Science and Technology (BP0719021), and

667

the MOST Special Fund from the State Key Laboratory of Geological Processes and Mineral

668

Resources, China University of Geosciences (MSFGPMR201804). We thank Jing Feng,

669

Jun-Tao Yang, Jin-Liang Wang, and Jiang-Tao Tian for assistance during the field work. We

670

also appreciate the kind help of Hai-Xia Chu from Resources Exploration Laboratory of

671

China University of Geosciences at Beijing on the fluid inclusion microthermometric

672

measurements, Xin Xiong from Ore-Forming Laboratory of Institute of Mineral Resources of

673

Chinese Academy of Geological Sciences at Beijing on Laser Raman spectroscopic analyses,

674

and Mu Liu from Beijing Research Institute of Uranium Geology on carbon, hydrogen,

675

oxygen, sulfur, and lead isotopic analyses. Thorough and constructive reviews by two

676

reviewers, and editorial comments and suggestions by Editor-in-Chief Hua-Yong Chen and

677

Associate Editor Guo-Xiang Chi have been very helpful in our revision of the paper, which

678

are gratefully acknowledged. 23

679

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996

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997

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998

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999 1000

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1004

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1007

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1010

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1015

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1019

Huangshannan magmatic Ni-Cu sulfide deposit, Xinjiang. Acta Petrol. Sin. 32, 2086−2098 (in

1020

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1022

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1023 1024 1025

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1026

Zhou, M.F., Lesher, C.M., Yang, Z.X., Li, J.W., Sun, M., 2004. Geochemistry and petrogenesis of 270 Ma

1027

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1028

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1029

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35

1030

Zhou, T.F., Yuan, F., Zhang, D.Y., Fan, Y., Liu, S., Peng, M.X., Zhang, J.D., 2010. Geochronology,

1031

tectonic setting and mineralization of granitoids in Jueluotage area, eastern Tianshan, Xinjiang. Acta

1032

Petrol. Sin. 26, 478−502 (in Chinese with English abstract).

1033

Zhu, Z.X., Li, P., Zhao, T.Y., Wang, K.Z., Jin, L.Y., Zhu, Y.F., 2018. Tectonic magmatic evolution and

1034

mineralization of Bogeda-Harlik tectonic belt, Xinjiang, China. Xinjiang Geol. 36, 1−7 (in Chinese

1035

with English abstract).

36

1036

Figure captions

1037

Fig. 1. A. Schematic map showing the position of the Central Asian orogenic belt (modified

1038

from Jahn et al., 2000). B. Sketch map showing the tectonic units of the Tianshan Belt (from

1039

Chen et al., 2012b). C. Regional geological map of the Caixiashan Zn-Pb deposit, showing

1040

the distribution of some important Zn-Pb(-Ag) deposits (modifed after Wang et al., 2018b).

1041 1042

Fig. 2. A. Geological sketch map of the Caixiashan Zn-Pb deposit (modified from Cao et al.,

1043

2013). B. Cross-section along section line 44 in the Caixiashan Zn-Pb deposit (modified after

1044

Li et al., 2016a).

1045 1046

Fig. 3. Photographs of hand specimens showing ore mineralization and mineral assemblages

1047

in the Caixiashan Zn-Pb deposit. A. Stage I irregluar calcite−dolomite−pyrite vein. B. Stage I

1048

quartz−calcite−pyrite vein. C. Stage II laminated pyrrhotite ore with minor calcite and

1049

dolomite. D. Stage II laminated sphalerite ore with calcite and dolomite. E. Stage II massive

1050

pyrrhotite−sphalerite ore with minor galena. F. Stage II quartz−sphalerite vein in the wall

1051

rocks. G. Stage III massive galena−pyrite ore. H. Stage III quartz−pyrite−galena vein. I. Stage

1052

III subhedral to euhedral pyrite replace the sphalerite and pyrrhotite. J. Stage III euhedral

1053

pyrite crystals in the silicified marble. K. The carbonate rock showing a stockwork of stage

1054

IV quartz and calcite veins. L. Stage IV barren quartz−calcite vein. Abbreviations: Cc =

1055

calcite, Gn = galena, Po = pyrrhotite, Py = pyrite, Qz = quartz, Sp = sphalerite.

1056 1057

Fig. 4. Photomicrographs showing mineral assemblages and paragenetic relationships in the

1058

Caixiashan Zn-Pb deposit. A. Stage I pyrite in the wall rocks. B. Stage II pyrrhotite replaced

1059

Stage I pyrite, showing a reticular texture of replacement. C. Stage II pyrrhotite coexisting

1060

with euhedral arsenopyrite, repalced by sphalerite. D. A common stellate replacement of stage

1061

II massive sphalerite by galena. E. Stage III galena vein cutting and replacing the former

1062

sphalerite. F. Stage III chalcopyrite replacing the early pyrrhotite. G. Stage III massive galena

1063

with replacement resdual of stage II sphalerite and pyrrhotite. H. Stage III subhedral to

1064

euhedral pyrite replacing the former pyrrhotite. I. Stage IV quartz and calcite coexisting with 37

1065

tremolite. Abbreviations: Apy = arsenopyrite, Cp = chalcopyrite, Gn = galena, Po =

1066

pyrrhotite, Py = pyrite, Qz = quartz, Sp = sphalerite, Tr = tremolite.

1067 1068

Fig. 5. Paragenetic sequence for the Caixiashan Zn-Pb deposit.

1069 1070

Fig. 6. Photomicrographs of representative fluid inclusions. A. Liquid-rich two-phase fluid

1071

inclusion assemblage (L-type FIA) from stage I calcite. B. Individual vapor-rich two-phase

1072

(V-type) inclusion from stage I quartz. C. Halite-bearing three-phase (H-type) inclusion in

1073

stage I quartz, coexisting with vapor-rich two-phase aqueous (V-type) and pure-vapor phase

1074

(PV-type) inclusion. D. Coexisting vapor-rich two-phase inclusion and halite-bearing

1075

two-phase inclusion. E. Liquid-vapor two-phase (L-type) inclusion with opaque daughter

1076

mineral in stage II quartz, coexisting with L-type inclusion. F. Individual vapor-rich

1077

two-phase (V-type) inclusion from stage II calcite. G. Coexisting vapor-rich two-phase

1078

(V-type) fluid inclusions and L-type inclusion in stage II calcite. H. Liquid-rich two-phase

1079

fluid inclusion assemblage (L-type FIA) from stage II quartz. I. Isolated liquid-rich two-phase

1080

inclusions in stage III calcite. J. Liquid-rich two-phase fluid inclusion assemblage (L-type

1081

FIA) from stage III quartz. K. Square-like liquid-rich two-phase fluid inclusions in IV calcite,

1082

coexisting with pure-liquid phase (PL-type) inclusion. L. Coexisting liquid-rich two-phase

1083

inclusion assemblages (L-type FIA) and PL-type inclusion in IV calcite.

1084 1085

Fig. 7. Histograms of homogenization temperatures and salinities.

1086 1087

Fig. 8. Summary plot of homogenization temperatures and salinities of fluid inclusions in

1088

different stages of the Caixiashan Zn-Pb deposit.

1089 1090

Fig. 9. (A-D) Laser Raman spectra for fluid inclusions of the Caixiashan Zn-Pb deposit.

1091 1092

Fig. 10. Pressure estimation for ore stage I to IV fluid inclusions of the Caixiashan Zn-Pb

1093

deposit. Isobars were calculated from the equations of Driesner and Heinrich (2007).

1094 38

1095

Fig. 11. A. Histogram of δ34SV-CDT values of sulfides from Caixiashan deposit; B. Sulfur

1096

isotope compositions of sulfides and country rocks at Caixiashan in comparison with sulfides

1097

from the Hongyuan, Hongxingshan, and Yuxi deposits, seawater sulfates, and the typical

1098

SEDEX, MVT, and Irish-type Zn-Pb deposits worldwide. Data sources: 1, Hoefs, 2009; 2,

1099

Holser, 1977; 3, Lu et al., 2018; 4, Li et al., 2018; 5, Xiao et al., 2009; 6, Zhou et al., 1999; 7,

1100

Leach et al., 2005; 8, Wilkinson, 2003.

1101 208Pb/204Pb

vs.

206Pb/204Pb

(A) and

207Pb/204Pb

vs.

206Pb/204Pb

1102

Fig. 12. Plot of

(B) for the

1103

Caixiashan sulfides, local granitoids, and host rocks. The lead isotope data for host rocks were

1104

taken from Liang et al. (2005); the lead isotope curves for the mantle, orogen, and crust are

1105

from Zartman and Doe (1981).

1106 1107

Fig. 13. Plot of δDH2O versus δ18OH2O for the Caixiashan ore-forming fluids. Field of

1108

metamorphic water, magmatic water, and meteoric water line are from Taylor (1974); the

1109

meteoric water in Tianshan was adopted from Wang et al. (2016c); SMOW = Standard Mean

1110

Ocean Water.

1111 1112

Fig. 14. Plot of δ13CV-PDB versus δ18OV-SMOW for the Caixiashan calcite, dolomite, and marble

1113

samples (Base map modified after Mao et al. 2002 and Liu et al. 2004). Three main carbon

1114

sources are marine carbonate (Baker and Fallick, 1989; Hoefs, 1997), sedimentary organic

1115

matter carbon (Hodson, 1977; Hoefs, 1997), and magma-mantle carbonate (Taylor et al.,

1116

1967; Ray et al., 1999).

1117 1118

Fig. 15. A brief genetic model for Caixiashan Zn-Pb mineralization.

1119

39

1120

Table captions

1121

Table 1 Summary of microthermometric data and calculated parameters for fluid inclusions in

1122

quartz and calcite of different ore stages at the Caixiashan Zn-Pb deposit

1123

Table 2 Sulfur isotope data of sulfides from the Caixiashan Zn-Pb deposit

1124

Table 3 Lead isotope data of sulfides, granitoids, and host rocks from the Caixiashan Zn-Pb

1125

deposit

1126

Table 4 Hydrogen and oxygen isotope data of quartz from the Caixiashan Zn-Pb deposit

1127

Table 5 Carbon and oxygen isotope data of calcite, dolomite, and marble from the Caixiashan

1128

Zn-Pb deposit

1129

Conflict of interest:

1130

We wish to confirm that there are no known conflicts of interest associated with this

1131

publication and there has been no significant financial support for this work that could have

1132

influenced its outcome.

1133 1134 1135

Highlights:

1136

1) The ore-forming fluids had a dominantly metamorphic signature and were diluted by

1137

meteoric water.

1138 1139

2) The ore-forming components were primarily sourced from the Precambrian basement.

1140 1141

3) The temperature decrease, fluid mixing, and pervasive dolomitization are the key factors

1142

resulting in large-scale ore precipitation.

1143

40

1144

41

1145

42

1146

43

1147

1148

44

1149

45

1150

1151

46

1152

1153

1154 47

1155

1156

48

1157 1158

Graphical Abstract:

1159 1160 1161

Table 1 Summary of microthermometric data and calculated parameters for fluid inclusions in quartz and calcite of different ore Stage

Samples

FI type1

N

Size (μm)

Stage I

Quartz–calcite–pyrite vein

H

3

5−12

L

40

4−15

49

Tm,ice (°C)

−13.4 to −2.4

Th,total (°C)

Ts,halite (°C ）

Salini

279−291

327−344

4

336–488

V

8

5−12

L

43

4−12

V

10

6−12

Quartz–calcite–galena vein

L

72

3−13

Barren quartz–calcite vein

L

25

2−14

Stage II

Quartz–calcite−sphalerite vein

Stage III Stage IV 1 Fluid

−14.3 to −9.0 −13.2 to −2.5 −13.4 to −5.1 −9.6 to −0.1 −4.8 to −0.7

382−494

1

240−357 255−350 140−297 71−156

inclusion type: L = liquid-rich; V = vapor-rich; H = halite-bearing

2 Methods

used for calculating the salinity and density of the fluid inclusions are described in the text

1162 1163

Table 2 Sulfur isotope data of sulfides from the Caixiashan Zn-Pb deposit. Sample no.

Mineral

Stage

CXS5701-1 CXS5701-3 CXSTV447-5 CXSTV447-6 CXS3102-1 CXS3303-2

Pyrite Pyrite Pyrite Pyrite Pyrite Sphalerite

CXS3303-3

Sphalerite

II

CXSШ4404-1 CXSIII4404-4 CXSIII4404-5 18CXS-37 CXSTV214-4 CXSTV447-1 18CXS-1 18CXS-33 18CXS-1 18CXS-33 18CXS-37

Sphalerite Sphalerite Sphalerite Sphalerite Pyrrhotite Pyrrhotite Pyrrhotite Pyrrhotite Galena Galena Galena

II

I I I I III II

II II II II II II II III III III

Sample description

δ34SV-CDT(‰)

Quartz–pyrite vein Quartz–pyrite vein Irregular pyrite vein Irregular pyrite vein Quartz vein with euhedral pyrite Laminated galena with sphalerite Laminated galena with sphalerite and minor pyrite Sphalerite vein Sphalerite vein with minor chalcopyrite Sphalerite vein Ores Pyrrhotite vein Pyrrhotite vein Ores Ores Ores Ores Ores

16.0 15.7 15.8 16.1 11.2 14.3 14.5 13.5 15.5 13.6 14.1 12.0 16.0 14.9 12.6 13.8 11.2 11.8

1164

Table 3 Lead isotope data of sulfides, granitiods, and host rocks from the Caixiashan Zn-Pb deposit. Sample no.

Mineral/Rock

208Pb/204Pb

50

2σ

207Pb/204Pb

2σ

206Pb/204P

CXS801-9

Granite porphyry

38.714

0.025

15.574

0.010

18.817

CXS801-20

Granite porphyry

39.572

0.018

15.655

0.007

20.013

CXSCL001-2

Monzonite porphyry

38.605

0.009

15.596

0.004

18.642

CXSCL001-12

Granodiorite

38.186

0.008

15.603

0.002

18.195

CXSCL001-15

Granodiorite

38.242

0.007

15.599

0.003

18.254

CXS3102-1

Pyrite

38.016

0.011

15.613

0.004

17.807

CXS5701-1

Pyrite

37.215

0.010

15.567

0.004

17.314

CXS5701-3

Pyrite

37.293

0.010

15.574

0.004

17.371

CXSTV447-6

Pyrite

37.169

0.008

15.536

0.003

17.333

CXS3303-2

Sphalerite

37.172

0.008

15.576

0.003

17.229

CXSIII4404-4

Sphalerite

37.060

0.007

15.538

0.003

17.201

CXSIII4404-5

Sphalerite

37.091

0.005

15.547

0.002

17.212

CXSШ4404-1

Sphalerite

37.100

0.006

15.549

0.002

17.204

18CXS-1

Galena

36.988

0.014

15.519

0.005

17.178

18CXS-33

Galena

36.964

0.012

15.512

0.005

17.173

18CXS-37

Galena

36.959

0.005

15.507

0.002

17.176

18CXS-1

Pyrrhotite

36.970

0.013

15.510

0.005

17.183

18CXS-33

Pyrrhotite

37.087

0.021

15.549

0.007

17.202

FcxIIzk3801-b10 Dolomite marble FcxIIzk3801-b18

Carbonaceous slate with siltstone interlayers

37.013

15.527

17.184

37.083

15.531

17.226

Note: Data of CXS801-9 to 18CXS-33 were from this studey; data of FcxIIzk3801-b10 and FcxIIzk3801-b18 were for al. (2005). 1165

Table 4 Hydrogen and oxygen isotope data of quartz from the Caixiashan Zn-Pb deposit. Th (°C)

δ18Oquartz(‰)

δ18OH2O(‰)

δDV-SMOW (‰)

I

354

15.2

10.0

-74.7

CXSCL6401-3-1 Quartz

II

304

14.8

8.1

-81.8

CXSCL6401-3-2 Quartz

II

304

13.8

7.1

-84.6

CXSCL6401-1

Quartz

II

264

13.6

5.3

-81.0

CXSII250-2

Quartz

III

200

15.1

3.4

-93.8

Sample no.

Mineral

Stage

CXS1504-3

Quartz

51

CXSII250-3

Quartz

III

200

15.5

3.8

-91.4

CXSTV214-1

Quartz

IV

130

15.1

-2.3

-104.5

CXSTV214-3

Quartz

IV

130

13.0

-4.4

-104.4

Note: Th is the average homogenization temperature values of the FIs in the quartz samples of the same stage. δ18OH2O values are calculated according to the quartz-water equilibrium temperature formula provided by Clayton et al. (1972) . 1166 1167 1168

Table 5 Carbon and oxygen isotope data of calcite, dolomite, and marble from the Caixiashan Zn-Pb deposit. Sample no.

Mineral

Stage

δ13CPDB(‰)

δ18OPDB(‰)

δ18OSMOW(‰)

CXSII250-8-1

Dolomite

I

-3.0

-14.6

15.9

CXSII250-8-2

Dolomite

I

-2.6

-14.2

16.2

CXS504-2

Calcite

II

-3.7

-18.2

12.2

CXSCL6401-1

Calcite

II

-3.7

-17.8

12.5

CXSCL6401-2

Calcite

II

-3.8

-18.0

12.3

CXSCL6401-3

Calcite

II

-3.3

-17.4

13.0

CXSII250-1

Calcite

III

-2.3

-19.8

10.5

CXS1504-1

Calcite

III

-2.2

-19.8

10.5

CXS1504-2

Calcite

III

-2.2

-19.6

10.6

CXS-KSD-18

Calcite

I

-1.3

-15.8

14.6

CXS-KSD-13

Calcite

II

-1.7

-17.1

13.3

CXS-ZK1502-1 Calcite

II

-6.7

-18.8

11.5

CXS-CT-1

Calcite

III

-0.5

-19.2

11.2

CXS-KSD-2

Calcite

IV

-4.9

-24.3

5.9

ZK3802-b9

Calcite

I

-0.8

-13.1

17.4

I

0.1

-11.5

19.1

I

-2.4

-12.7

17.8

I

-1.3

-15.7

14.7

ZK3801-b10 ZK3801-b20 ZK3801-b22

Alterated marble Alterated marble Alterated marble

52

Note: The data were reported in permil relative to the Pee Dee Belemnite limestone (PDB) standard with total uncertainties were estimated to be better than 0.2‰ for δ18O and 0.1‰ for δ13C. Data of CXSII250-8-1 to CXS1504-2 were from this study; data of CXS-KSD-18 to CXS-KSD-2 were from Cao et al. (2013); and data of ZK3802-b9 to ZK3801-b22 were from Gao et al. (2007b). δ18OSMOW values are calculated according to the equation (δ18OSMOW = 1.03086 × δ18OPDB + 30.86) provided by Friedman and O'Neil (1977). 1169

53