Co-rich sulfides in mantle peridotites from Penghu Islands, Taiwan: Footprints of Proterozoic mantle plumes under the Cathaysia Block

Journal of Asian Earth Sciences 37 (2010) 229–245

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Co-rich sulfides in mantle peridotites from Penghu Islands, Taiwan: Footprints of Proterozoic mantle plumes under the Cathaysia Block Kuo-Lung Wang a,b,*, Suzanne Y. O’Reilly a, Masahiko Honda c, Takuya Matsumoto d,e, William L. Griffin a, Norman J. Pearson a, Ming Zhang a a

Australian Research Council National Key Centre for the Geochemical Evolution and Metallogeny of Continents (GEMOC), Department of Earth and Planetary Sciences, Macquarie University, Sydney, NSW 2109, Australia Institute of Earth Sciences, Academia Sinica, Taipei 11529, Taiwan c Research School of Earth Sciences, Australian National University, Canberra, ACT 0200, Australia d Department of Earth and Space Science, Graduate School of Science, Osaka University, Osaka 560-0043, Japan e Institute for Study of the Earth’s Interior, Okayama University, Tottori 682-0193, Japan b

a r t i c l e

i n f o

Article history: Received 11 June 2009 Received in revised form 7 August 2009 Accepted 24 August 2009

Keywords: Mantle sulfides In situ Os model ages Subchondritic Ni/Co Mantle plume Proterozoic mantle Cathaysia Block

a b s t r a c t Abundant primary sulfides occur as inclusions in silicates and as discrete grains in mantle-derived spinel lherzolite xenoliths from Miocene intraplate basalts on the Penghu Islands, Taiwan, which is located at the southeastern margin of Cathaysia Block. These sulfides are dominantly mixtures of Fe-rich and Nirich monosulfide solid solutions (MSS), with minor pentlandite, millerite and chalcopyrite, and are considered to represent sulfide melts crystallized at high temperatures (>900 °C). Some sulfides from the Tungchiyu (TCY) islet (37 out of 118 grains) have remarkably high Co contents resulting in subchondritic Ni/Co ratios (<21; 5–20, median = 12), distinct from the superchondritic values (Ni/Co = 48–157, median = 83) typical of mantle sulfides worldwide. The Co-rich nature of the TCY sulfides is considered to be a primary characteristic as no secondary processes can be identified to account for the feature. They are similar to Ni–Co-rich sulfides from Lac de Gras, Slave Craton (Aulbach et al. (2004) Chemical Geology 208, 61–88) interpreted as being derived from the lower mantle. Experimental studies suggest that the sulfide melt/silicate melt partition coefficient of Ni becomes lower than that of Co at pressures greater than 28 GPa, similar to recent estimates of the magma ocean conditions. Os model ages of the TCY Co-rich sulfides reveal four episodes of generation: 2.0, 1.7, 1.4 and 0.8 Ga; this is consistent with the age pattern of all Penghu sulfides, indicating significant lithosperic mantle formation, melt extraction or metasomatic events at these time periods. These events closely correspond to the global 1.9-Ga superplume event related to the assembly of the Nena/Columbia supercontinent, a minor 1.7-Ga superplume event in SW Laurentia prior to breakup of Nena/Columbia, the 1468 Ma Moyie event in the Belt Basin region in western Laurentia and the 0.8 Ga breakup of Rodinia, with which the Cathaysia Block was associated at various stages during its Proterozoic evolution (Li et al. (2008) Precambrian Research 160, 179–210 and references therein). Olivine in a peridotite sample from the TCY locality has distinctly high 3He/4He (11 RA), whereas other peridotites from the KP and TCY localities have 3He/4He 6.7 RA, lower than MORB. The high 3He/4He further suggests that materials from the deep mantle have interacted with the host peridotite of Co-rich sulfides. We thus propose that the Co-rich sulfide melts may have been trapped in the lower mantle during core–mantle differentiation and then transported to shallow depths by mantle plumes that entrained lower mantle materials at several different time periods. This study provides the first substantial evidence from the lithosperic mantle beneath the Cathaysia Block to support the activity of mantle plumes related to the breakup of the supercontinents Nena/Columbia and Rodinia in Proterozoic time. Ó 2009 Elsevier Ltd. All rights reserved.

1. Introduction

* Corresponding author. Address: Institute of Earth Sciences, Academia Sinica, P.O. Box 1-55, Nankang, Taipei 11529, Taiwan. Fax: +886 2 27839871. E-mail address: [email protected] (K.-L. Wang). 1367-9120/$ - see front matter Ó 2009 Elsevier Ltd. All rights reserved. doi:10.1016/j.jseaes.2009.08.008

The lithospheric mantle is one of Earth’s most significant geochemical reservoirs due to its non-convecting character. It is geochemically complex with a depleted composition overprinted by multiple melt/fluid processes, and carries the history of its stabilization and subsequent events that affected the whole lithosphere.

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This relatively rigid uppermost part of the mantle, the subcontinental lithospheric mantle (SCLM), acts as roots of the continents and contributes to the stability and longevity of continents (e.g. O’Neill et al., 2007; Poudjom Djomani et al., 2001; Griffin et al., 2008 and references therein). The SCLM may have formed as partial melting residues from primordial mantle, by cooling of upwelling asthenosphere, or by plume accretion to existing lithosphere (e.g. O’Reilly and Griffin, 2006 and references therein). Such processes could be related to the formation of the overlying crust and correspond to growth and evolution of the crust. Thus, the SCLM records the cumulative geochemical effects of large-scale tectonic events, and affects the response of the crust to tectonic stresses. The evolution of the SCLM is of essential importance to understanding the growth and stability of continents and longterm mantle evolution. Whole-rock Re–Os isotope data from mantle-derived peridotites have contributed much information on the age of the SCLM. However, Os in these rocks is concentrated in sulfide phases and that these can be mobile within the SCLM, so that whole-rock Re–Os model ages may reflect mixing processes, rather than single melting events (Alard et al., 2002; Griffin et al., 2002). The development of in situ analysis of elemental and Re–Os isotopic compositions in single sulfide grains makes it possible to explore the heterogeneity of different generations of sulfides and to unravel some of the complexity contained in mantle-derived samples (Alard et al., 2000, 2002; Aulbach et al., 2004; Griffin et al., 2002, 2004; Lorand and Alard, 2001; Luguet et al., 2001; Pearson et al., 2002). The base metal sulfides (BMS), in the form of monosulfide solid solution (MSS) in mantle rocks, have compositions suggesting equilibrium with a sulfide melt (Craig, 1973) and commonly are interpreted as residual from partial melting (Lorand, 1987) or as crystallisation products of sulfide liquids trapped during partial melting (Mitchell and Keays, 1981). Studies of these mantle sulfides may provide critical constraints on metasomatic processes and the evolution of the lithospheric mantle (Aulbach et al., 2004; Griffin et al., 2002, 2004; Wang et al., 2003a,b, 2009).

It has been proposed that the Cathaysia Block, as part of the larger South China Block, was probably associated with western Laurentia when the supercontinent Nena (Gower et al., 1990) or Columbia (Rogers and Santosh, 2002) formed in Paleoproterozoic time (Li et al., 2008; Rogers and Santosh, 2002; Zhao et al., 2004). Until the Neoproterozoic supercontinent Rodinia formed and before it broke up, the Cathaysia Block was still located between Laurentia, Australia and probably Siberia (Li et al., 1995, 1996b). Although there is more evidence to clarify the location of the Cathaysia Block during the Neoproterozoic, poor exposure of early Proterozoic basement rocks makes it difficult to reconstruct its relationship with other cratons and blocks during that time. As a counterpart of crustal rocks to form the tectonic plate, lithospheric mantle should accompany the processes to assemble and breakup of a supercontinent. Thus, the evolution of the lithospheric mantle beneath the Cathaysia Block may record a similar tectonic history to the overlying crust. In this study, we document and characterise major-element and Os isotopic compositions of mantle sulfides, describe an unusual suite of Co-rich sulfides, and present the noble gas compositions of the sulfide-bearing peridotites from the Penghu Islands, Taiwan. We propose a plume-related model for the formation of these sulfides and use their Os model ages to correlate them with known Proterozoic mantle plumes on supercontinents Nena/Columbia and Rodinia. 2. Geological background and sample description The South China Block comprises the Yangtze craton in the northwest and the Cathaysian foldbelts (Cathaysia), including the Cenozoic mountain ranges of Taiwan, in the southeast (Fig. 1). The Yangtze craton has a late Archean-Paleoproterozoic (P1.7 Ga) nucleus surrounded by late Mesoproterozoic orogenic belts; the Cathaysia Block is generally considered to be floored by Paleo- to Mesoproterozoic continental crust (Chen and Jahn, 1998) although our studies (Wang et al., 2009) and other recent

100°

Peiliao

110°

130°E

120°

MONGOLIA

Yangtze

40°

Kueipi

Makung

CHINA

30°

Minxi

10 km

Hainan 0

200 km

115°

ra St ch

aT ren

South China Sea

nil

Qilin

Penghu Islands

Ma

Tungchiyu

20°N

PHILIPPINE

it

Niutoushan

Ta i

0

TAIWAN

w an

Penghu Islands Cathaysia

23.5°N

Taiwan

ugh a Tro 25°

aw Okin

KuanhsiChutung Ryukyu Trench

Philippine Sea

20°

125°E

Fig. 1. Tectonomagmatic map of southeastern China. Black areas—late Cenozoic intraplate basalts; stars—localities of mantle xenoliths; solid star—locality of xenoliths in this study. Insets: main sample localities in the Penghu Islands and the regional map.

K.-L. Wang et al. / Journal of Asian Earth Sciences 37 (2010) 229–245

work (Xu et al., 2005, 2007; Yu et al., 2008) suggest the presence of Archean domains. Eastern China has experienced extension due to the late Mesozoic subduction of the Paleo-Pacific plate beneath Eurasia and/or the Indo-Eurasia collision since the early Paleogene. Along the continental margin, this extension led to the formation of early Tertiary basins, and subsidence of the Taiwan Straits accompanied the opening of the South China Sea (Chung et al., 1994). Widespread intraplate basaltic volcanism (65 Ma to Recent; Fig. 1) has occurred in these basins in conjunction with the regional extension and some of the volcanic rocks carry abundant mantlederived xenoliths. The mantle xenoliths investigated in this study come from two localities in the Penghu Islands in the Taiwan Strait: Kueipi (KP) village on the main island and Tungchiyu (TCY) islet southeast of

231

the main island (Fig. 1). The Penghu islands are made up of Miocene basaltic lavas (16–8 Ma; Chung et al., 1994), which can be divided into alkaline and tholeiitic suites. The xenolith suite is dominated by spinel peridotites with minor spinel pyroxenites. The KP xenoliths are dominantly spinel lherzolites from alkali basatic lavas (13–10 Ma; Lee, 1994), with calculated equilibration temperatures between 880 and 1040 °C (Wang et al., 2003b). Sulfides and rare amphiboles are found in the spinel lherzolites, indicating modal metasomatism. The modal sulfide content is from 0.1 to 0.3 vol%. The TCY xenoliths are mostly spinel lherzolites and pyroxenites with minor harzburgites, and are derived from a pyroclastic layer (13.2–11.2 Ma; Ho et al., 2000). Their equilibration temperatures lie in a narrow range from 950 to 1000 °C (Wang et al., 2003b). Sulfides and amphibole have been found in all sam-

Fig. 2. (a) Sulfides under the microscope (transmitted and reflected-light) and X-ray maps of element distribution in representative sulfide grains from each group (KP and TCY). Red circles mark locations of sulfides with X-ray maps in the thin sections. (b) Image under the microscope (transmitted light) showing triple junction position of interstitial sulfides in a TCY peridotite.

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Table 1 Reconstructed bulk major-element compositions (wt.%) of Penghu sulfides. Minerala

Sample no.

c

Fe

Co

Ni

Cu

S

O

Total

Me/Sb

Ni/(Ni + Fe)

Ni/Co

27.21 28.46 30.82 28.68 30.23

0.34 0.25 0.39 0.39 0.4

33.31 31.31 36 37.42 36.32

6.43 7.12 0.14 0.08 0.2

33.1 33.16 32.83 33 32.98

0.37 0.49 0.36 0.37 0.44

100.8 100.8 100.5 99.95 100.6

1.12 1.12 1.14 1.12 1.14

0.54 0.51 0.53 0.55 0.53

98 126 93 95 90

Fe-MSS (Pt nug.) Fe-MSS Fe-MSS

40.77 44.09 48.84

0.33 0.17 0.16

20.82 11.39 11.79

2.38 6.83 1.97

35.15 35.24 35.8

0.55 0.86 0.76

100 98.59 99.33

1.03 0.99 0.99

0.33 0.2 0.19

62 65 73

s2 s5

Fe-MSS Fe-MSS

32.29 49.86

0.31 0.18

23.29 14.36

7.07 0.12

30.86 34.79

4.05 0.56

97.87 99.87

1.14 1.05

0.41 0.21

76 78

KPH9810 KPH9812 KPH9815/1

s1 s3 s4e s5e

Fe-MSS ISS Fe-MSS Fe-MSS

48.29 29.74 49.75 50.97

0.25 0.18 0.13 0.1

12.11 16.51 12.34 12.35

3.68 19.12 0.07 0.88

35.83 33.21 31.95 34.28

0.63 0.36 3.61 0.9

100.8 99.13 97.84 99.5

1.01 1.08 1.1 1.07

0.19 0.35 0.19 0.19

48 90 99 121

KPH9816/1

s1e s2e s3

Fe-MSS Fe-MSS Fe-MSS

39.71 40.94 39.58

0.37 0.25 0.34

23.88 21.03 20.72

0.9 3.26 3.79

34.17 34.33 34.09

0.43 0.3 1.4

99.45 100.1 99.93

1.07 1.07 1.06

0.36 0.33 0.33

64 83 60

KPH9816/2

s1 s2 2s1 2s2

PN Fe-MSS ISS Fe-MSS

29.94 42.63 28.28 41.49

0.44 0.37 0.29 0.33

36 21.81 24.47 20.26

0.25 0.02 12.74 2.1

32.58 35.15 33.24 34.78

0.33 0.49 0.55 0.3

99.54 100.5 99.56 99.26

1.14 1.04 1.09 1.04

0.53 0.33 0.45 0.32

82 59 85 61

KPH9816/3

s1e s2 s3 s4e

Fe-MSS PN Fe-MSS ISS

30.47 27.33 45.77 29.09

0.31 0.25 0.24 0.2

29.51 34.76 17.3 23.94

6.28 4.54 0.01 13.1

32.39 30.72 35.77 33.02

0.66 1.13 0.85 0.42

99.63 98.73 99.94 99.77

1.14 1.21 1.01 1.11

0.48 0.55 0.26 0.44

95 137 73 119

KPH9822

s1 s2 s3 s4

Fe-MSS ISS Fe-MSS PN

41.99 29.96 41.22 30.91

0.41 0.17 0.3 0.38

21.29 18.27 18.66 35.28

0.19 17.09 4.15 0.21

34.74 32.44 34.08 32.43

0.42 1.54 1.26 0.46

99.05 99.48 99.68 99.66

1.04 1.11 1.06 1.15

0.33 0.37 0.3 0.52

52 110 62 94

KPH9826

s1 s2 s3 s4 s5 s6 2s1

ISS ISS Fe-MSS Fe-MSS CB ISS ISS

32.57 21.89 39.49 27.6 31.03 39.96 29.84

0.2 0.14 0.41 0.25 0.12 0.18 0.35

23.53 27.15 25.43 31.68 12.43 14.21 19.38

9.85 20.87 0.07 7.17 22.37 8.79 16.89

32.45 30.55 34.71 33 32.87 34.68 33.54

1.75 0.24 0.28 0.66 1.41 1.96 0.4

100.4 100.9 100.4 100.4 100.2 99.79 100.4

1.13 1.24 1.06 1.12 1.09 1.02 1.08

0.41 0.54 0.38 0.52 0.28 0.25 0.38

115 197 62 128 107 78 56

KP0201

s1 s2e s3 s4 s5 s6 s7

PN PN ISS ISS Fe-MSS ISS ISS

25.43 32.19 29.68 30.87 39.91 30.77 32.12

0.34 0.49 0.22 0.32 0.28 0.27 0.23

38.61 32.86 26.71 21.87 19.53 22.34 23.2

2.26 0.57 10.04 12.44 4.51 12.12 11.53

31.59 32.98 32.77 32.93 34.07 32.14 33.46

2.19 0.67 0.37 1.47 1.19 2.03 0.35

100.4 99.75 99.78 99.9 99.5 99.68 100.9

1.17 1.12 1.12 1.1 1.06 1.13 1.11

0.59 0.49 0.46 0.4 0.32 0.41 0.41

115 67 121 68 69 84 100

KP0202-1

s1 s2e s3

ISS ISS ISS

30.11 29.76 30.08

0.24 0.18 0.23

18.88 18.44 26.51

13.6 17.4 8.37

31.29 33.05 33.03

3.4 0.29 0.3

97.54 99.11 98.51

1.11 1.09 1.09

0.37 0.37 0.46

79 105 117

KP0202-2

s1 s2 s3

ISS ISS ISS

30.29 32.79 31.08

0.15 0.16 0.29

15.92 15.44 22.32

19.96 18.26 12.88

33.14 31.08 32.99

1.16 2.17 0.76

1.09 1.18 1.11

0.33 0.31 0.41

106 99 78

KP0203

s1 s2e s3e

Fe-MSS Fe-MSS ISS

42.94 30.15 25.35

0.21 0.22 0.21

14.2 26.53 32.27

6.1 7.46 10.02

34.18 32.15 31.41

1.55 1.51 0.41

99.17 98.02 99.66

1.04 1.11 1.19

0.24 0.46 0.55

67 118 157

KP0203

s4e s5

ISS Fe-MSS

31.23 32.42

0.25 0.18

23.51 21.97

7.89 10.51

31.94 33.36

2.2 1.09

97.03 99.54

1.09 1.08

0.42 0.39

92 120

KP0214-8

s1 s2 s3

ISS CP PN

28.55 19.54 27.24

0.17 0.08 0.3

21.43 8.93 38.42

16.31 41.23 0.04

32.71 28.79 32.75

0.54 0.29 0.31

99.7 98.86 99.05

1.11 1.29 1.12

0.42 0.3 0.57

128 117 130

KP0214-11

s1 s2 s3 s4 s5 s6e

Fe-MSS ISS ISS ISS ISS Ni-MSS

38.88 29.76 36.4 29.96 28.38 26.88

0.36 0.34 0.21 0.36 0.18 0.25

17.85 20.37 17.7 24.68 20.63 33.74

5.47 14.23 11.03 8.24 16.17 5.39

33.47 32.78 33.86 31.29 33.45 33.06

2.06 1.52 0.93 1.28 0.33 0.33

98.09 99 100.1 95.81 99.13 99.66

1.05 1.09 1.07 1.12 1.07 1.11

0.3 0.39 0.32 0.44 0.41 0.54

50 60 84 69 113 133

KP0214-16

s1 s2 s1e s2 s3 s4e

Ni-MSS ISS ISS Fe-MSS Fe-MSS ISS

29.68 30.49 34.97 31.74 32.75 33.47

0.28 0.29 0.21 0.23 0.32 0.22

32.16 26.96 19.23 26.32 24.77 17.97

4.64 8.13 11.85 7.47 7.33 13.88

32.99 33.04 33.07 33.3 31.88 32.62

0.27 0.37 0.36 0.44 2.19 1.26

100 99.27 99.7 99.5 99.24 99.43

1.12 1.1 1.11 1.1 1.14 1.11

0.51 0.46 0.34 0.44 0.42 0.34

116 94 91 113 77 83

KPH9801/1-3

ts1 ts2 s2 s3 s4

Ni-MSS Ni-MSS PN PN PN

KPH9801/2-1 Ol.-websterited

s1 s2e s3

KPH9803

KP0215

100.6 99.9 100.3

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K.-L. Wang et al. / Journal of Asian Earth Sciences 37 (2010) 229–245 Table 1 (continued) Minerala

Fe

Ni

Cu

S

O

Total

Me/Sb

s5e s6e s7

Fe-MSS ISS ISS

40 31.02 34.3

Ni/(Ni + Fe)

Ni/Co

0.28 0.27 0.34

22.5 20.39 23.22

2.98 12.81 7.82

33.82 31.43 33.37

0.5 2.97 0.46

100.1 98.9 99.51

1.09 1.13 1.09

0.35 0.38 0.39

80 75 69

TC0223 Sp.-pyroxenite

s1 s3e s4 s6 s8 s14 s16 2s3 2s5 2s7 2s8 2s9

Fe-MSS Fe-MSS Ni-MSS Fe-MSS Fe-MSS Fe-MSS Ni-MSS Fe-MSS Fe-MSS Fe-MSS Ni-MSS Ni-MSS

38.77 34.82 21.49 45.55 30.91 40.02 25.8 30.99 27.8 32.64 9.49 20.01

0.29 0.33 2.13 0.3 1.89 0.63 2.55 0.73 5.47 0.56 7.13 5.18

22.86 27.25 38.97 15.64 26.79 20.97 31.47 30.17 26.15 28.91 39.84 33.18

0.55 0.61 0.66 0.55 0.55 0.34 0.55 0.66 0.51 0.56 0.48 0.42

36.93 36.66 35.99 37.35 38.57 38.01 38.22 36.42 37.32 37.07 39.09 37.96

0.36 0.33 0.45 0.47 0.71 0.46 1.3 0.39 1.15 0.35 2.01 2.31

99.77 100 99.7 99.86 99.42 100.4 99.89 99.35 98.4 100.1 98.04 99.05

0.92 0.96 0.97 0.92 0.8 0.92 0.88 0.96 0.9 0.95 0.8 0.86

0.36 0.43 0.63 0.25 0.45 0.33 0.54 0.48 0.47 0.46 0.8 0.61

80 84 18 51 14 33 12 42 5 51 6 6

TC0225

s1 s2i s13 s15 s23 2s1 2s2

Ni-MSS Ni-MSS Ni-MSS ML Ni-MSS Ni-MSS Ni-MSS

1.71 5.02 3.67 2.4 17.25 1.76 1.27

3.87 5.68 4.69 4.82 2.44 1.9 6.77

52.88 54.28 54.5 53.16 42.63 59.25 52.23

0.54 2.04 2.36 1.75 2.24 1.08 0.18

30.49 32.2 34.12 34.15 33.74 34.48 35.78

6.92 0.69 0.43 1.86 0.5 0.33 1.94

96.41 99.9 99.77 98.15 98.8 98.8 98.18

1.06 1.14 1.04 1 1.06 1.01 0.93

0.97 0.91 0.93 0.95 0.7 0.97 0.98

14 10 12 11 17 31 8

TC0242/2 Harzburgite

s1 s2 s3 s4 s6 s7 s8 s10 s11 s12 s13 s16

ISS Ni-MSS Ni-MSS Ni-MSS Ni-MSS Ni-MSS ML ML ML Ni-MSS ML Ni-MSS

20.85 6.23 18.6 3.12 5.06 5.97 1.84 2.43 1.74 12.03 0.74 3.46

1.05 3.23 1.96 4.11 4.7 2.03 2.34 2.47 3.32 4.85 1.31 0.79

27.17 49.11 39.55 57.37 51.69 50.72 60.09 57.56 59.32 46.44 61 59.4

13.98 4.45 3.31 0.61 2.79 5.75 0.98 1.59 0.37 0.89 0.38 0.8

32.98 33.84 33.35 33.94 31.3 33.31 34.06 33.96 34.45 33.08 34.37 34.61

1.96 1 1.4 0.31 3.21 0.78 0.27 0.69 0.36 0.6 0.45 0.41

97.99 97.87 98.16 99.46 98.76 98.57 99.59 98.7 99.57 97.89 98.24 99.47

1.05 1.02 1.05 1.05 1.12 1.06 1.05 1.03 1.03 1.07 1.01 1.02

0.55 0.88 0.67 0.95 0.91 0.89 0.97 0.96 0.97 0.79 0.99 0.94

26 15 20 14 11 25 26 23 18 10 47 75

TC0247

s3 s5

Ni-MSS Ni-MSS

5.1 2.72

4.18 5.27

49 53.1

2.74 1.47

38.01 36.68

0.59 0.25

99.63 99.49

0.88 0.93

0.9 0.95

12 10

TC0247

s6 s7 s10

Ni-MSS Ni-MSS Ni-MSS

4.71 7.53 23.24

11.47 4.75 1.38

37.25 41.01 35.1

3.38 7.96 3.5

34.45 34.86 35.85

4.05 1.01 0.36

97.8 97.48 99.48

0.9 0.95 0.98

0.88 0.84 0.59

3 9 26

TC0248

ts1 ts5 2s1 s2 s3 s5 s6i s7 s10 s11 s12 s13 s16 s17 s18e s23 s24 s29 s30 s31 s33 s34 s36 s39 ss12 ss13

Ni-MSS Ni-MSS Ni-MSS Ni-MSS Ni-MSS Ni-MSS Ni-MSS ISS ML Ni-MSS Ni-MSS ISS Ni-MSS Ni-MSS Ni-MSS Ni-MSS Ni-MSS Ni-MSS Ni-MSS Ni-MSS Ni-MSS Ni-MSS Ni-MSS Ni-MSS ML ML

16.54 5.28 4.72 4.01 16.93 7.99 11.75 22.14 1.45 12.65 9.02 13.31 6.41 3.91 3.09 9.5 17.48 12.53 11.34 8.4 6.9 6.96 6.15 14.03 0.88 1.47

1.49 0.76 6.02 4.15 1.72 1.55 2.3 0.27 1.02 1.75 3.08 1.66 4.43 1.55 7.67 1.04 2.1 1.47 0.83 1.3 1.74 0.91 4.84 1.57 0.89 1.37

45.58 57.05 53.59 54.57 42.39 49.4 47.85 36.98 61.37 46.81 47.01 31.31 48.27 55.12 50.63 52.74 43.18 35.66 51.17 54.64 46.67 54.86 51.63 45.85 61.6 59.68

1.37 1.4 0.29 1.05 3.13 4.96 2.74 6.49 0.59 2.67 2.03 14.45 4.54 2.81 0.58 1.79 1.61 13.31 1.75 1.07 7.01 2.91 1.27 4.39 0.03 0.06

33.94 34.34 34.45 34.08 32.44 34.99 33.69 33.39 34.63 33.55 32.52 32.3 34.97 34.95 35.7 34.06 35.32 33.84 34.22 34.26 33.21 34.24 34.99 33.97 34.17 35.03

0.48 0.32 1.32 1.08 2.92 0.48 0.56 0.38 0.4 1.33 3.4 3.16 0.69 0.38 1 0.27 0.75 2.5 0.64 0.38 3.44 0.5 0.81 0.39 0.45 0.26

99.4 99.16 100.4 98.94 99.54 99.37 98.89 99.65 99.46 98.76 97.05 96.19 99.32 98.72 98.68 99.41 100.5 99.3 99.96 100.1 98.98 100.4 99.69 100.2 98.02 97.87

1.06 1.03 1.03 1.02 1.09 1 1.05 1.09 1.01 1.04 1.03 1.02 0.99 0.99 0.95 1.05 1.01 1.01 1.05 1.05 1.02 1.05 1 1.06 1.01 0.98

0.72 0.91 0.92 0.93 0.7 0.85 0.79 0.61 0.98 0.78 0.83 0.69 0.88 0.93 0.94 0.84 0.7 0.73 0.81 0.86 0.87 0.88 0.89 0.76 0.99 0.97

31 75 9 13 25 32 21 137 60 27 15 19 11 36 7 51 21 24 61 42 27 61 11 29 69 44

TC0276/1 Harzburgite

s2 s3 s4 s7 s8 s9 s10

Ni-MSS Ni-MSS Ni-MSS Ni-MSS ML Ni-MSS Ni-MSS

16.01 4.16 6.47 6.62 1.62 19.24 9.16

2.11 0.58 1.47 2.75 3.46 1.38 1.58

45.47 54.54 54.52 50.23 57.49 43.43 52.15

1.17 4.36 1.15 4.01 1.53 1.82 1.27

32.65 33.91 34.35 34.07 33.59 32.82 33.63

0.54 0.44 0.82 1.25 0.99 0.58 0.66

97.93 97.99 98.78 98.93 98.68 99.26 98.45

1.1 1.02 1.02 1.02 1.04 1.11 1.05

0.73 0.93 0.89 0.88 0.97 0.68 0.84

22 94 37 18 17 32 33

Sample no.

Co

(continued on next page)

234

K.-L. Wang et al. / Journal of Asian Earth Sciences 37 (2010) 229–245

Table 1 (continued) Minerala

Sample no. s11 s12 a b c d

Ni-MSS ML

Fe 4.46 1.58

Co 6.38 1.24

Ni 51.24 60.39

Cu 1.85 1.17

S

O

33.01 34.53

0.85 0.8

Total 97.78 99.72

Me/Sb

Ni/(Ni + Fe)

Ni/Co

1.06 1.02

0.92 0.97

8 49

MSS: monosulfide solid solution; ISS: intermediate solid solution; PN: pentlandite; CP: chalcopyrite; CB: cubanite; ML: millerite. Me/S: ratio of total metal/sulfur. Ni-MSS:Ni/(Ni + Fe) > 0.5. Lithology of host mantle rocks indicated under sample name; samples without indication are all sp.-lherzolite.

ples, reflecting extensive metasomatism. Sulfide modal abundance is between 0.6 and 1 vol%. Compared with the fresh KP lherzolite samples, the TCY lherzolites and pyroxenites experienced minor alteration with carbonate replacement of some silicate minerals. However, the original textures and mineral outlines are commonly preserved (Ho et al., 2006). The major-element data for these spinel lherzolites suggest that some of the lithospheric mantle beneath this region is refractory, and similar to that of typical Proterozoic lithospheric mantle (e.g. Griffin et al., 1999; Wang et al., 2003a,b, 2009). 3. Analytical methods Sulfides P50 lm across in polished thin sections (200 lm thick) and blocks of peridotites were chosen for analysis. All analyses were done at the Geochemical Analysis Unit (GAU) at GEMOC, Department of Earth and Planetary Sciences, Macquarie University, Sydney. Major-element contents were analysed using a CAMECA Camebax SX50 electron microprobe (EMP). Sulfide analyses were performed with an accelerating voltage of 20 kV, a beam current of 20 nA and a beam size of 63 lm. Counting times were 20 s on the peak and 10 s on each side of the background. Bulk sulfide compositions were reconstructed by combining EMP analyses of the individual low-temperature phases with their modal abundances. Modes were calculated from X-ray element distribution maps (Fig. 2a) using an image analysis program (NIH Image 1.62). A traverse of analyses across the grain was averaged and this was taken as the representative composition of those samples in which no distinct exsolved phases were visible in the X-ray maps. The reconstructed major-element data for the sulfides are reported in Table 1. In situ Re–Os isotope compositions of sulfide grains were analysed using a Merchantek LUV266 laser microprobe (LAM) with a modified ablation cell, attached to a Nu Plasma multiple collector (MC) ICPMS. All ablations were carried out using He as the carrier gas, which was blended with Ar in a 30 mL mixing chamber prior to introduction into the plasma, to enhance sensitivity and reduce elemental fractionation. Typical laser operating conditions are 5-Hz frequency, beam energy of 1–2 mJ/pulse and 50–80 lm spot size. The analytical procedures have been described in detail by Pearson et al. (2002) (see www.es.mq.edu.au/GEMOC). A dry aerosol of Ir, produced by a CETAC MCN6000 desolvating nebuliser, was bled into the gas line between the ablation cell and the ICPMS to provide a mass-bias correction. The total Ir signal is around 1.5–2.5 V during the analyses reported in this study. Pearson et al. (2002) have demonstrated that the relative fractionation coefficients for Os, Re, and Ir are internally consistent and that the Ir mass-bias correction procedure provides a significant improvement in precision relative to the use of the Os isotopes of the sample itself. Masses 188–194 were measured in Faraday cups, and masses 185 and 187 were measured in ETP ion counters. The ion counters were calibrated initially against the Faraday cups and one another using a two-cycle analysis of an Os standard solution, rather than the sequential analysis of Ir + Os and Re + Ir solutions used by Pearson et al. (2002). During ablation runs, a synthetic NiS bead with 200 ppm Os and Pt (PGE-A) was analysed between samples, to monitor and correct any drift in the ion counters. These corrections typically were less than 1% over a long day’s analytical session.

The overlap of 187Re on 187Os was corrected by measuring the Re peak and using 187Re/185Re = 1.6741, which has been determined by Pearson et al. (2002) and is identical to the IUPAC value (1.6740; Gramlich et al., 1973). Most 187Re/188Os ratios of sulfides in this study (0.017–1.078; Table 3) are well below the threshold (1.18) for which a reliable correction can be carried out (Pearson et al., 2002). The data were collected using the Nu Plasma time-resolved software, which allows the selection of the most stable intervals of the signal for integration. The selected interval is divided into 40 replicates to calculate the standard deviation and standard error. The precision and accuracy of the method are discussed in detail by Pearson et al. (2002). Under ideal circumstances (i.e. sulfides 50 microns in diameter, and containing P40 ppm Os), an internal precision of 0.1–0.3% (2SE) is routinely obtained for 187 Os/188Os; for smaller grains or lower Os contents (<5–10 ppm), an internal precision of 1–2% is routine. The external reproducibility of 187Os/188Os for the PGE-A standard over three years (April 2001 to June 2004), covering the period of the analyses presented here, is ±0.00030 (2sd; Table 2), and the mean value of 187Os/188Os is indistinguishable from that derived by TIMS analysis (0.10645 ± 0.00002; Pearson et al., 2002). Repeated analysis of a 25 ppb Os standard solution JM Os GEMOC (JMC standard solution prepared at GEMOC) using the same collector configuration gives a better external reproducibility (Table 2). The Re–Os data for the sulfides are reported in Table 3. The Os isotope compositions of these sulfides can be recast as model ages to reveal the history of mantle depletion. Two types of model age are presented here (Table 3). TMA model ages represent the time of separation of the sulfide from a chondritic mantle reservoir calculated using the measured Re/Os ratio of the sulfide (Shirey and Walker, 1998). Re-depletion model ages (TRD) assume complete removal of Re (zero Re/Os) during melting, and are minimum ages for separation of the sulfide from a chondritic reservoir (Shirey and Walker, 1998). If some Re is introduced to the sulfide after the initial melting event, the TMA model ages will be artificially old, and the true age will lie somewhere between the two model age estimates. The calculated model ages are dependant on the model reservoir selected to represent Os isotope evolution in the upper mantle. The differences between evolution models yield an uncertainty on the accuracy of any age of 0.4 Gyr, but do not change the relative positions of age peaks (Pearson et al., 2007). A recent empirical study of Os–Ir alloys from ophiolite sequences with well-constrained ages (Shi et al., 2007) supports an Enstatite Chondrite Reservoir (ECR), as the best model for the Os isotope evolution of the upper mantle; this is intermediate between the lower primitive upper mantle (PUM; Shirey and Walker, 1998) and higher carbonaceous-chondrite values. In this study, we use the model ages calculated from the ECR mantle model, with present-day 187Os/188Os = 0.1281 and 187Re/188Os = 0.421 (Walker et al., 2002a,b) to interpret the evolution of the lithospheric mantle. The model ages estimated using PUM (Shirey and Walker, 1998) are also listed in Table 3 for comparison. Noble gas analyses of selected olivine separates from KP and TCY lherzolites were undertaken by vacuum crushing experiments at Osaka University (OU) and the Australian National University (ANU). Olivine separates were hand-picked from five coarsely 185

192

a

JM Os

April 2001 to May 2004

JM Os (Far + IC)

April 2001 to May 2004

JM Os

Oct 2004 to Dec 2008

PGE-A

April 2001 to June 2004

GEMOC Re–Os

Nov 2002 to June 2004

PGE-A

Oct 2004 to Dec 2008

GEMOC Re–Os

Oct 2004 to Dec 2008

The

192

Average 2sd Average 2sd Average 2sd Average 2sd Average 2sd Average 2sd Average 2sd

(n = 54) (n = 70)

Os/188Os

1se

3.08040 0.00222 see notea

(n = 65) (n = 465) (n = 70) (n = 536) (n = 201)

Os/188Os was used to correct mass-bias for these analyses.

3.07995 0.00047 3.08036 0.00587 3.07986 0.00386 3.08137 0.00544 3.08193 0.00594

0.00005 0.00116 0.00044 0.00080 0.00057

190

Os/188Os

1.98302 0.00094 1.98314 0.00178 1.98282 0.00060 1.98175 0.00322 1.98186 0.00241 1.98246 0.00309 1.98264 0.00370

1se

0.00003 0.00062 0.00028 0.00053 0.00036

189

Os/188Os

1.21947 0.00025 1.21934 0.00135 1.21944 0.00018 1.21935 0.00202 1.21985 0.00121 1.21980 0.00208 1.21943 0.00314

1se

0.00002 0.00039 0.00017 0.00039 0.00026

187

Os/188Os

0.18365 0.00018 0.18349 0.00059 0.18369 0.00006 0.10640 0.00030 0.18359 0.00067 0.10634 0.00054 0.18363 0.00075

1se

0.00001 0.00007 0.00017 0.00008 0.00016

186

Os/188Os

0.12185 0.00310 0.12022 0.00207 0.12181 0.00189

1se

187

Re/188Os

0.00001

0.00009 0.00023 0.00013 0.00045 0.44978 0.10649 0.00014 0.00031 0.41118 0.12153

1se

0.00001 0.00003 0.00444 0.00005 0.00455

K.-L. Wang et al. / Journal of Asian Earth Sciences 37 (2010) 229–245

Table 2 Reference material and standard reproductivity of solution and in situ Os isotopic analyses.

235

236

Samplea

192

Os/188O

±2se

190

Kueipi (KP) peridotites KPH9801/2-1-s2e 3.0769 KPH9803-S5 3.0699 KPH9810-S1 3.0900 KPH9810-S-S1 3.0633 KPH9816/3-1-s1e 3.0987 KPH9816/3-1-S1-2 3.1036 KPH9816/3-1-s4e 3.0874 KPH9822-S0 3.0733 KPH9822-S4 3.0928 KPH9822-S4-2 3.0942 KPH9822-s-s1e 3.0612 KPH9822-S-S3 3.0592 KPH9822-s-s5e 3.0947 KPH9826-s1e 3.0894 KPH9826-s2e 3.0601 KPH9826-S4 3.0869 KP0201-s2e 3.1043 3.1196 KP0201-S4 KP0201-S5 3.0923 KP0201-S7 3.0815 KP0201-S-S4 3.0780 KP0201-s-s12 3.0556 KP0214-8-S3 3.1083 KP0214-11-S0 3.0922 KP0214-11-S3-1 3.1018 KP0214-16-S0-1 3.0933 KP0215-s1e 3.1440 KP0215-s4e 3.0594 KP0215-s5e 3.0836 KP0215-s6e 3.0508 KP0215-S8 3.0745 KP0215-S-S3 3.0825

0.0180 0.0200 0.0240 0.0280 0.0620 0.0440 0.0540 0.0166 0.0280 0.0220 0.0280 0.0380 0.0280 0.0360 0.0144 0.1480 0.1040 0.0880 0.0130 0.0060 0.0064 0.0220 0.1020 0.0320 0.0640 0.0420 0.1080 0.0380 0.0196 0.0340 0.0176 0.0060

Tungchiyu (TCY) peridotites TC0225-S11 3.0698 TC0225-S23 3.0265 TC0225-S-S24

3.0485

Os/188Os

±2se

189

1.9920 1.9685 1.9795 1.9843 1.9733 1.9895 1.9748 1.9739 1.9850 1.9858 1.9726 1.9556 1.9839 1.9844 1.9764 1.9880 1.9337 1.9419 1.9877 1.9722 1.9781 1.9759 2.0272 1.9769 1.9768 1.9650 1.9824 1.9542 1.9851 1.9521 1.9667 1.9849

0.0174 0.0150 0.0132 0.0300 0.0500 0.0300 0.0360 0.0102 0.0162 0.0088 0.0158 0.0198 0.0220 0.0192 0.0100 0.0980 0.0500 0.0620 0.0100 0.0040 0.0046 0.0060 0.0560 0.0180 0.0720 0.0150 0.0800 0.0320 0.0084 0.0260 0.0170 0.0038

0.0054 0.0860

1.9709 1.9578

0.0360

1.9616

Os/188Os

±2se

187

Os/188Os

±2se

187

1.2190 1.2251 1.2172 1.2209 1.2631 1.2370 1.2184 1.2127 1.2163 1.2219 1.2060 1.2159 1.2180 1.2318 1.2221 1.1715 1.1965 1.1875 1.2279 1.2113 1.2184 1.2128 1.2033 1.2337 1.2092 1.2104 1.2081 1.2049 1.2158 1.2046 1.2071 1.2176

0.0078 0.0114 0.0120 0.0260 0.0240 0.0196 0.0200 0.0086 0.0140 0.0046 0.0150 0.0168 0.0110 0.0200 0.0076 0.0920 0.0540 0.0540 0.0072 0.0034 0.0038 0.0090 0.0420 0.0200 0.0360 0.0086 0.0420 0.0184 0.0056 0.0166 0.0134 0.0030

0.1148 0.1245 0.1228 0.1153 0.1160 0.1194 0.1061 0.1187 0.1091 0.1182 0.1201 0.1218 0.1229 0.1190 0.1239 0.1327 0.1202 0.1256 0.1123 0.1210 0.1284 0.1481 0.1138 0.1212 0.1270 0.1155 0.1249 0.1216 0.1167 0.1195 0.1216 0.1167

0.0028 0.0020 0.0028 0.0036 0.0036 0.0016 0.0030 0.0030 0.0015 0.0014 0.0022 0.0013 0.0028 0.0022 0.0009 0.0064 0.0050 0.0046 0.0014 0.0028 0.0011 0.0020 0.0030 0.0026 0.0036 0.0036 0.0032 0.0020 0.0009 0.0013 0.0012 0.0013

0.0040 0.1280

1.2145 1.2446

0.0030 0.0300

0.1182 0.1184

0.0280

1.1953

0.0130

0.1276

Re/188Os

cOsb

±2se

Os/Pt

0.2294 0.2120 0.1417 0.1657 0.2168 0.1395 0.2951 0.7094 0.7240 0.1176 0.2178 0.0787 0.4337 0.1509 0.2931 0.5683 0.7777 0.5252 1.0776 0.8974 0.2798 0.5527 0.1025 0.4320 0.7201 0.1526 0.4560 0.2838 0.2111 0.3364 0.4089 0.5923

0.0068 0.0158 0.0030 0.0260 0.0048 0.0034 0.0220 0.0072 0.0138 0.0044 0.0280 0.0032 0.0094 0.0062 0.0076 0.0780 0.0220 0.0182 0.0320 0.0280 0.0044 0.0420 0.0040 0.0400 0.0138 0.0058 0.0240 0.0144 0.0120 0.0102 0.0048 0.0440

0.2 24.9 13 0.2 1.2 0.8 4.2 240 2.1 1.3 2.6 18 0.3 0.5 12 0.4 13 8.6 0.6 0.6 6.4 1.1 0.7 1.0 0.4 0.1 0.6 1.2 2.2 0.4 4.5 6.1

10.6 3.0 4.4 10.2 9.6 7.0 17 7.5 15 7.9 6.5 5.1 4.3 7.3 3.5 3.4 6.4 2.2 13 5.8 0.0 15 11 5.6 1.1 10.0 2.7 5.3 9.1 6.9 5.3 9.1

0.0010 0.0028

0.0171 0.1121

0.0002 0.0114

6.2 0.4

7.9 7.7

0.0022

0.1342

0.0062

1.1

0.6

TRDb (Ga)

2se

TMAb (Ga)

2se

TRDc (Ga)

TMAc (Ga)

1.9 0.5 0.8 1.8 1.7 1.2 3.1 1.3 2.6 1.4 1.1 0.9 0.7 1.3 0.6 0.7 1.1 0.4 2.2 1.0 0.0 2.9 2.0 1.0 0.2 1.8 0.5 0.9 1.6 1.2 0.9 1.6

0.4 0.3 0.4 0.5 0.5 0.2 0.7 0.4 0.2 0.2 0.3 0.2 0.7 0.3 0.1 1.0 0.7 0.7 0.2 0.4 0.1 1.0 0.4 3.6 0.5 0.5 0.6 0.3 0.2 0.2 0.4 0.5

4.0 1.0 1.1 2.9 3.4 1.8 9.7 2.0 3.9 1.9 2.3 1.1 31.9 2.0 2.0 1.9 1.3 1.5 1.5 0.9 0.1 8.5 2.6 59.5 0.2 2.7 5.8 2.8 3.2 5.8 25.9 4.1

0.8 0.6 0.6 0.8 1.0 0.3 1.3 0.6 0.3 0.3 0.6 0.2 18.5 0.5 0.4 2.6 0.9 2.7 0.1 0.4 0.5 0.9 0.6 26.3 0.7 0.8 5.8 0.9 0.2 0.9 5.6 0.5

1.8 0.4 0.6 1.7 1.6 1.1 3.1 1.3 2.6 1.3 1.0 0.8 0.6 1.2 0.5 0.8 1.0 0.2 2.2 0.9 0.2 3.2 1.9 0.9 0.0 1.7 0.3 0.8 1.5 1.1 0.8 1.5

4.1 0.8 1.0 2.9 3.5 1.7 10.7 1.6 3.4 1.8 2.2 1.0 8.3 1.9 1.7 2.0 1.1 0.1 1.3 0.7 0.7 7.9 2.6 13.0 0.0 2.7 2.4 2.7 3.1 6.5 90.0 3.3

1.4

0.1 0.4

1.4 1.8

0.1 0.5

1.3 1.3

1.4 1.7

0.3

0.1

0.5

0.1

-0.1

1.4 0.1

K.-L. Wang et al. / Journal of Asian Earth Sciences 37 (2010) 229–245

Table 3 In situ Re–Os isotope compositions of Penghu sulfides from Taiwan.

3.0868

0.0198

1.9604

0.0124

1.2165

0.0092

0.1139

0.0022

0.0632

0.0010

0.3

TC0242/2-S11

3.1127

0.0460

1.8955

0.0300

1.2647

0.0180

0.1157

0.0020

0.2528

0.0036

1.1

11 9.9

TC0242/2-S13 TC0242/2-S-S5 TC0242/2-S-S9 TC0242/2-S-S19 TC0247-S-15 TC0247-nS-6 TC0247-nS-8 TC0248-S6 TC0248-S11 TC0248-S12

3.1166 3.1052 3.0939 3.0790 3.0859 3.0404 3.0688 3.0699 3.0704 3.0391

0.1120 0.0840 0.0054 0.0136 0.0520 0.0460 0.0156 0.0240 0.0154 0.0440

1.9122 1.9842 1.9870 1.9828 1.9772 1.9543 1.9862 1.9830 1.9717 1.9749

0.0660 0.0660 0.0054 0.0106 0.0400 0.0400 0.0112 0.0188 0.0138 0.0300

1.2621 1.2509 1.2213 1.2181 1.2198 1.2185 1.2263 1.2207 1.2252 1.1956

0.0420 0.0440 0.0038 0.0082 0.0340 0.0260 0.0086 0.0096 0.0062 0.0220

0.1128 0.1161 0.1163 0.1145 0.1072 0.1120 0.1148 0.1187 0.1223 0.1224

0.0034 0.0026 0.0005 0.0017 0.0022 0.0017 0.0010 0.0011 0.0009 0.0028

0.2770 0.2351 0.0428 0.0330 0.8049 0.1892 0.0176 0.2172 0.1909 0.2149

0.0084 0.0188 0.0011 0.0005 0.0138 0.0050 0.0004 0.0048 0.0046 0.0052

0.5 1.5 0.6 1.7 1.5 1.4 2.2 0.6 0.5 1.0

12 9.6 9.4 10.8 16 13 10.6 7.5 4.7 4.7

TC0248-S18e

3.0715

0.0580

1.9649

0.0360

1.2374

0.0240

0.1186

0.0028

0.2590

0.0042

0.6

7.6

Tungchiyu (TCY) pyroxenite TC0223-2S11 3.0385 TC0223-S-S9-1 3.0403 TC0223-S-S10-1e 3.0181 TC0223-S-S11 3.1056 TC0223-S-S12 3.0483 TC0223-S-S18 3.0730 TC0223-S-S19 3.0877 TC0223-S-S20 3.0756 TC0223-S-S25 3.0929 TC0223-S-S75 3.0901 TC0223-S-S77-3 3.0748 TC0223-S-S81 3.1146 TC0223-S-S102 3.0611 3.1030 TC0223-S-S114 TC0223-S-S115 3.0493 TC0223-S-S130e 3.0437 TC0223-S-S132 3.0764

0.0420 0.0840 0.0440 0.0920 0.0280 0.0460 0.0580 0.0520 0.0260 0.0200 0.0440 0.0280 0.0580 0.0200 0.0280 0.0420 0.0260

1.9478 1.9394 1.9669 1.9921 1.9773 1.9832 1.9729 1.9886 1.9623 1.9799 1.9686 2.0015 1.9994 1.9936 1.9634 1.9637 1.9968

0.0240 0.0580 0.0220 0.0880 0.0170 0.0380 0.0580 0.0320 0.0220 0.0180 0.0260 0.0166 0.0540 0.0240 0.0240 0.0320 0.0164

1.2161 1.2075 1.1904 1.2438 1.2080 1.2132 1.2198 1.2078 1.2104 1.2204 1.2179 1.2107 1.2194 1.2296 1.2020 1.2103 1.2326

0.0280 0.0400 0.0220 0.0500 0.0132 0.0240 0.0320 0.0220 0.0102 0.0154 0.0172 0.0186 0.0520 0.0220 0.0220 0.0220 0.0164

0.1247 0.1215 0.1222 0.1278 0.1222 0.1234 0.1257 0.1249 0.1214 0.1232 0.1218 0.1251 0.1328 0.1287 0.1279 0.1174 0.1258

0.0020 0.0020 0.0020 0.0026 0.0005 0.0016 0.0022 0.0016 0.0011 0.0014 0.0014 0.0012 0.0032 0.0011 0.0014 0.0020 0.0016

0.3363 0.1566 0.0924 0.1113 0.0581 0.1101 0.0605 0.0786 0.0733 0.1312 0.0902 0.0948 0.1044 0.0213 0.1136 0.2979 0.1623

0.0034 0.0172 0.0144 0.0034 0.0010 0.0026 0.0022 0.0017 0.0011 0.0044 0.0022 0.0019 0.0034 0.0005 0.0022 0.0054 0.0080

0.5 0.9 0.9 0.5 0.7 1.3 5.7 7.6 1.7 1.0 0.6 0.4 0.7 0.6 0.9 1.0 0.5

2.9 5.4 4.8 0.4 4.8 3.9 2.1 2.7 5.5 4.0 5.1 2.5 3.4 0.2 0.4 8.6 2.0

2.0

0.3

1.9

2.3

0.3

2.3 4.3

0.4

1.7 2.1 1.7 1.7 1.9 2.9 2.3 1.9 1.3 0.8

0.7

1.7

4.4

0.8

0.5 0.4 0.07 0.2 0.3 0.2 0.1 0.2 0.1 0.4

6.1 3.8 1.8 2.1 3.4 4.0 2.0 2.7 1.5 1.6

1.4 0.8 0.08 0.3 0.4 0.4 0.2 0.3 0.2 0.8

2.1 1.6 1.6 1.8 2.9 2.2 1.8 1.2 0.7 0.7

6.5 3.8 1.8 2.0 3.0 4.1 1.9 2.6 1.3 1.5

1.3

0.4

3.4

1.0

1.3

3.4

0.5 0.9 0.8 0.0 0.8 0.7 0.3 0.5 1.0 0.7 0.9 0.4 0.7 0.1 0.0 1.5 0.3

0.3 0.3 0.3 0.4 0.08 0.2 0.3 0.2 0.2 0.2 0.2 0.2 0.5 0.1 0.2 0.3 0.2

2.4 1.5 1.1 0.1 1.0 0.9 0.4 0.6 1.2 1.0 1.1 0.5 0.9 0.1 0.0 5.0 0.5

1.4 0.4 0.4 0.5 0.09 0.3 0.4 0.3 0.2 0.3 0.3 0.2 0.6 0.2 0.3 0.9 0.4

0.4 0.8 0.7 0.1 0.7 0.5 0.2 0.3 0.8 0.6 0.8 0.3 0.9 0.2 0.1 1.4 0.2

2.1 1.3 0.9 0.2 0.8 0.7 0.2 0.4 1.0 0.8 1.0 0.4 1.2 0.3 0.2 5.3 0.3

Sample labelled ‘‘-s2” represents different sulfide grains in a single thin section. Calculated at basalt eruption age of 13 Ma, using measured 187Re/188Os and decay constent of 1.66E 11 (Shen et al., 1996). 187Os/188Os and 187Re/188Os of Enstatite Chondrite Reservoir (ECR) are 0.1281 and 0.421 (Walker et al., 2002b). c Calculated using Primitive Upper Mantle (PUM) whose 187Os/188Os and 187Re/188Os are 0.127 and 0.40186 (Shirey and Walker, 1998). Other parameters are the same as calculation using ECR. In TCY sulfides, TRD ages underlined and in Italic are for Co-rich samples; samples in Italic are with superchondritic Ni/Co ratios. b

K.-L. Wang et al. / Journal of Asian Earth Sciences 37 (2010) 229–245

a

TC0242/2-S6

237

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K.-L. Wang et al. / Journal of Asian Earth Sciences 37 (2010) 229–245

Table 4 He, Ne and Ar compositions in olivine separates from Penghu peridotites, Taiwan by vacuum crushing experiments. KP peridotites a

TCY peridotites TC0242-2b

TC0275b

TC0225a

1.943 200

1.905 100

0.759 100

7.68 500

1.48 0.06 6.7 2.4

6.41 0.01 6.7 0.4

1.62 0.08 6.9 0.4

0.48 0.01 n. d.

15.48 0.98 11.1 0.6

n.d.

n.d.

2.8 0.8 12.4 1.7 0.041 0.006

2.3 0.7 10.3 1.3 0.035 0.005

0.7 0.5 12.3 4.3 0.051 0.019

2.0 0.2 10.1 2.4 0.033 0.009

0.65 0.02 550 6

1.53 0.02 517 7

1.73 0.03 844 17

0.41 0.05 496 27

5.70 0.08 322 5

3.20 0.11 381 2

KP0202

KP0202

Weight (g) strokes

9.82 500

2.001 200

[4He] (109) ± 3 He/4He (R/RA) ±

1.88 0.12 4.6 6.0 )

)

[22Ne] (10 ± 20 Ne/22Ne ± 21 Ne/22Ne ± [36Ar] (10 ± 40 Ar/36Ar ±

12

10

b

KP0214-20

b

Gas amounts are in cm3STP. Quoted errors: 1r; n.d.: not determined owing to low gas amounts. Crushing yields of grain sizes less than 150 lm were typically 50–70%, except for sample KP0202a (<10%). a Measurements undertaken at ANU. b At Osaka Univ.

crushed lherzolites. About 2 g of olivine separates of samples KP0202, KP214-20, TC242-2 and TC0275) and 10 g of two additional samples (KP0202-duplicate and TC0225) were used at the OU and ANU, respectively. Both institutes are equipped with a VG 5400 noble gas mass spectrometer, an air actuated-type vacuum crusher and a similar gas handling system (Matsumoto et al., 1998, 2001). Typical procedural blanks both at the OU and ANU were (1–5)  10 11, (4–8)  10 12 and (2–4)  10 9 cm3STP for 4He, 20Ne and 40Ar, respectively. The noble gas results are given in Table 4. 4. Results and discussion

and eruption of the host basalt magmas (cf. Lorand, 1987; Griffin et al., 2002; Aulbach et al., 2004). Nearly all sulfides from the KP lherzolites consist of interfingered Fe-rich monosulfide solid solutions (MSS) and pentlandite, commonly with an outer rim of chalcopyrite. Sulfides from the TCY lherzolites and pyroxenite have a Fe-rich MSS core, surrounded by a Ni-rich MSS mantle and a distinctive Co-rich rim. They sometimes contain pentlandite and millerite, but fewer TCY sulfides (compared with KP sulfides) have Cu-rich phases such as chalcopyrite, chalcopyrrhotite and bornite. The reconstructed bulk compositions of most KP sulfides are MSS and Cu-rich sulfides, whereas the bulk compositions of most TCY sulfides are Ni (Co)rich MSS and millerite (Wang et al., 2009).

4.1. Geochemical characteristics and origin of Co-rich sulfides 4.1.1. Petrographic characteristics of sulfide assemblages Some sulfides with spherical (or rarely elongated) habits are enclosed in olivine, opx or cpx, but most are anhedral interstitial grains (Fig. 2a). Some intergranular sulfides occur in veinlets along rims of silicate grains or infiltrating cracks in them. Some of the KP sulfides enclosed in silicate minerals occur as tiny (<20 lm) droplets coexisting with CO2 fluid inclusions along secondary inclusion planes in cpx or opx. In contrast, most of the TCY sulfides are interstitial and few are enclosed. In altered TCY lherzolites and pyroxenites, most interstitial sulfides occur at triple junctions, surrounded by serpentinised or carbonatised pseudomorphs of silicate mineral grains (Fig. 2b). The occurrences of sulfides in interstitial veinlets or cracks, as droplets along secondary fluid inclusion planes in silicate minerals, and at triple junctions of silicate grains, suggest a metasomatic origin (Lorand et al., 2003). X-ray element distribution maps show that most sulfides consist of several phases, with different major-element compositions (e.g., Ni, Fe, Cu and Co; Fig. 2a). Mantle sulfides never preserve high-temperature phase relations; they recrystallize to low-temperature assemblages such as pyrrhotite + pentlandite + chalcopyrite ± pyrite, independently of the cooling rate (Ballhaus et al., 2001 and references therein). Based on the rounded or polygonal shapes of the sulfide grains, their internal structure and the compositions of coexisting phases, these multi-phase intergrowths are interpreted as low-temperature assemblages that exsolved from high-temperature solid solutions upon cooling during ascent

4.1.2. Unique Co-rich sulfides In contrast to the KP sulfides with Fe–Cu-rich, S-deficient characteristics (Fe = 19.5–51.0 wt.%; Ni/(Ni + Fe) = 0.19–0.59; Cu = 0.1– 22.4 wt.%), sulfides from the TCY locality are Ni-rich and some (37 out of 118 grains) have remarkably high Co contents (Ni = 27.2–61.6 wt.%; Co = 0.3–11.5 wt.% (Fig. 3a); Ni/(Ni + Fe) = 0.55– 0.99). Generally the KP sulfides have slightly higher metal/sulfur (Me/S = 0.99–1.29) ratios than the TCY sulfides (0.80–1.12). Sulfides in the TCY pyroxenite have the lowest Me/S (0.80–0.97). This indicates that the TCY sulfides, especially those in pyroxenite, formed under sulfur-saturated conditions. Wang et al. (2009) demonstrated that in the porphyroclastic KP peridotites, the Cu-rich sulfides can be explained by percolation of differentiated melts enriched in Cu and Pd with high Re/Os, which partly modified the original residual MSS and also crystallized sulfides under low melt/rock ratios. On the other hand, in the equigranular, extensively metasomatised TCY peridotites, most Ni and S-rich sulfides reflect extensive interaction between pre-existing sulfides and percolating melts at high melt/rock ratios. The melt was not as differentiated as the one(s) that affected the KP peridotites. A few residual MSS and low-Cu sulfides survived this metasomatism and retain their ancient depletion signatures. The high Co TCY sulfides have subchondritic Ni/Co ratios (<21; 5–20, median = 12) similar to Ni–Co-rich sulfides found in xenoliths from kimberlites (Aulbach et al., 2004) and in the diamond inclusion (Davies et al., 2004) at Lac de Gras, Slave Craton. In contrast, the Ni/Co ratios of all KP sulfides and other sulfides from TCY

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K.-L. Wang et al. / Journal of Asian Earth Sciences 37 (2010) 229–245

a

8

b8

Co (wt.%) 6

6

4

4

2

2

0 28

30

32

34

36

38

0 28

40

O (wt.%)

30

32

c

34

36

38

40

S (wt.%)

S (wt.%) 8

d

O (wt.%)

8

Co (wt.%)

6

6

4

4

2

2

0

0 0

2

4

6

8

0

10

20

Cu (wt.%)

Co (wt.%) Kueipi (KP) sulfides Tungchiyu (TCY) sulfides (peridotite) Tungchiyu (TCY) sulfides (pyroxenite) Fig. 3. Major-element variations in reconstructed Penghu sulfides.

are superchondritic (>21; 48–157, median = 83). As such high Co contents have only been reported for few other occurrences (Aulbach et al., 2004; Davies et al., 2004), it could be suggested that the high Co bulk composition may be an artefact or due to secondary alteration. The O content may be used as a potential indicator of the extent of desulfurization of primary sulfides (that is, the extent of the secondary alteration). There is no correlation between the oxygen and Co contents of these TCY sulfides (Fig. 3c). Some rare TCY sulfides with oxygen contents above 10 wt% do show a negative correlation between oxygen and sulfur contents. These sulfides appear to have been severely altered and were excluded from discussion in this study. Cu-rich sulfides most likely result from fractionation of sulfide melts. However, the Co contents of the TCY sulfides are not correlated with Cu contents (Fig. 3d), so the Co-rich nature of some TCY sulfides is probably not to due to such melt fractionation. A recent study reported that Ni–Co-rich sulfides occur in serpentinised ocean-floor peridotites (Klein and Bach, 2009). However, such serpentinisation leads to reducing conditions, S loss and formation of S-deficient Ni–Co-rich sulfides which is contrary to formation of the TCY sulfides under sulfur-saturated condition. In the Fe–S–(Ni + Co) tetrahedral plot (Fig. 4), Ni– Co-rich sulfides in serpentinised peridotites apparently distribute at the lower-temperature phases such as pentlandite, heazlewoodite, awaruite and pyrite, distinct from the TCY Co-rich sulfides locating along the higher-temperature (>900 °C) MSS field. Noted that the Co-rich sulfides have higher sulfur content than that of Ni–Co-rich sulfides in serpentinised peridotites except for pyrites

(Fig. 4). Moreover, some TCY Co-rich sulfides still preserve primary Os isotopic signatures (see discussion later in Section 4.2) which could not be explained by the secondary serpentinisation. Therefore, the Co-rich nature of these sulfides is interpreted to be a primary feature, rather than a result of fractionation, contamination or alteration. The interstitial position of these sulfides makes it unlikely that their high Co contents are derived from the host silicate minerals (Co contents are 140–175 ppm in olivine; <70 ppm in orthopyroxene and clinopyroxene; 240–330 ppm in spinel; Table 5). The major silicate phases of sulfide-bearing lherzolites from all Penghu xenolith localities have similar Co contents, yet the Co-rich sulfides are confined to the TCY lherzolites and pyroxenite. It appears unlikely that only TCY sulfides could be affected by the silicate minerals. Even if the Ni and Co contents of TCY sulfides are modified by decomposition of silicate minerals due to alteration, the subchondritic Ni/Co ratios of some TCY sulfides are still not consistent with Ni and Co distributions as reflected in their partition coefficients (Table 6). 4.1.3. Petrogenetic evidence from partition coefficients of Ni and Co In order to generate subchondritic Ni/Co, Co must concentrate at least as strongly as Ni in the sulfides. Although partition coefficient data for Co between metal solid/metal liquid are quite limited, it has been demonstrated that metal solid/metal liquid partition coefficients of the siderophile elements including Co and Ni depend strongly on the S content of the system (Li and Agee,

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K.-L. Wang et al. / Journal of Asian Earth Sciences 37 (2010) 229–245

Pyrite

MSS (1100°C)

MSS (1000°C)

MSS (900°C)

Polydymite

Pyrrhotite Pentlandite Heazlewoodite

Awaruite

(Co-rich)

Sulfides

Fig. 4. Bulk sulfide compositions plotted in the Fe–S–(Ni + Co) tetrahedral plot. The 900 °C, 1000 °C and 1100 °C MSS phase are after Kullerud et al. (1969).

Table 5 Ni and Co concentrations (ppm) of major silicate minerals in Penghu host peridotites determined by LAM-ICPMS. Olivine

Orthopyroxene

Clinopyroxene

Spinel

KP peridotites

Ni Co

3000–3200 140–160

740–830 59–64

300–370 18–24

2200–3100 240–290

TCY peridotites

Ni Co

3000–3800 145–175

660–.790 56–61

170–370 12–23

1500–3100 275–330

1996; Li et al., 1996a). According to partition coefficients from experimental studies (Table 6), Ni and Co do not fractionate during the crystallization of sulfide melts or the separation of residual sulfide under different sulfur (S)–pressure (P)–temperature (T) conditions. Ni and Co have similar sulfide solid/sulfide liquid partition coefficients and these change covariantly with variation in S–P–T conditions. However, there is general consensus that partition coefficients of Ni and Co for metal/silicate melt and sulfide melt/ silicate melt both change with varying P–T conditions. Their sulfide melt/silicate melt partition coefficients decrease with increasing pressure (Ballhaus et al., 2001; Chabot et al., 2005; Gaetani and Grove, 1997; Li and Agee, 2001; Tschauner et al., 1999), and the decrease is greater for Ni than Co. At pressures >28 GPa (corresponding to conditions in the present-day lower mantle), the sulfide melt/silicate melt partition coefficient of Ni becomes lower than that of Co (Li and Agee, 1996, 2001; Tschauner et al., 1999). During core formation, carbon-bearing metal-rich sulfide melts, segregated from the magma ocean, would have drained to the core and some of them might have been trapped in the mantle on the way down (Arculus and Delano, 1981; Rama Murthy and Hall, 1972). A recent estimate of the magma ocean conditions gives 30–60 GPa, >2000°K and fO2 = 2.2 DIW (Chabot et al., 2005), suggesting that sulfide melts segregated under this condition may

have been enriched in Co relative to Ni. Such metal-rich sulfide melts might be trapped in silicate minerals that were subsequently transported to the upper mantle by a plume that originated in the deep mantle, entraining lower mantle materials. According to Ballhaus and Ellis (1996) and Jana and Walker (1997), liquid immiscibility could be induced by cooling during the ascent of a plume to form two immiscible melts: a S-, Ni- and Co-poor metal melt, and a Ni- and Co-rich sulfide melt. The latter would precipitate Ni–Corich sulfides with lower metal/sulfur ratios (Aulbach et al., 2004). Any phase changes during ascent are likely to be isochemical. Once the rising material reaches the shallow depths where melting starts, the high degree of melting in a plume and their low solidus temperatures will ensure that all sulfides melt completely, retaining the Ni–Co concentration relationship. Some sulfides could be initially enclosed in, then become interstitial to, recrystallising silicate minerals as a result of annealing, maintaining their unique Ni/ Co. In contrast, other sulfides may have reached equilibrium under upper mantle condition, which eliminates the Co-rich characteristics of the deep-seated sulfides. Similar subchondritic Ni/Co sulfides have been reported in xenoliths from kimberlites in the Slave Craton, Canada, and were proposed to have been carried from the lower mantle by a mantle plume (Aulbach et al., 2004). This deep origin is consistent with the presence in the same kimberlite

241

K.-L. Wang et al. / Journal of Asian Earth Sciences 37 (2010) 229–245 Table 6 Partition coefficients of Ni and Co compiled from literatures. Log fO2

T (K)

P (GPa)

DCo

DNi

References

MSS-melt na na

727–827 1193–1323

0.0001 0.0001

0.19–1.17 0.7–1.4

Li et al. (1996a) Ballhaus et al. (2001)

S-rich liquid/S-poor liquid Metal sat. Metal sat. Metal sat. Metal sat.

2173 2173 2173 1623–1843

5 5 5 2

0.56 1.02 1.96 18.71

1.25 2.86 6.67 21.25

Jana and Walker (1997) Jana and Walker (1997) Jana and Walker (1997) Ballhaus and Ellis (1996)

Metal or sulfide melt/silicate melt 7.9 1623 8 1623 0.94 1623 10.3 1623 IW-0.17 2273 IW-0.56 2473 IW-0.73 2623 IW-0.79 2373 IW-1.59 2273 IW-1 2500 IW-1.15 2773 IW-1.3 2500 IW-1.4 2300

0.0001 0.0001 0.0001 0.0001 10 10 10 25 10 43 20 20 20

17.7 24.3 62 164 20 22 23 17 55 17 25 33

510 580 1600 4400 98 92 104 39 224 19 37 40 116

Gaetani and Grove (1997) Gaetani and Grove (1997) Gaetani and Grove (1997) Gaetani and Grove (1997) Li and Agee (2001) Li and Agee (2001) Li and Agee (2001) Li and Agee (2001) Li and Agee (2001) Li and Agee (2001) Ohtani and Yurimoto (1996) Ohtani et al. (1997) Ohtani et al. (1997)

Metal melt/perovskite Metal sat. Metal sat. Metal sat. Metal sat. Metal sat. IW-2.7

1717–1927 1897–1967 1947 1697 1727 2200

25–31 35–39 43 59 78 27

17 11.6 7.6 3.4 1.09 72

13.0 10.4 7.3 5.3 3.4 119

Tschauner et al. (1999) Tschauner et al. (1999) Tschauner et al. (1999) Tschauner et al. (1999) Tschauner et al. (1999) Ohtani et al. (1997)

Metal melt/Mg-wuestite Metal sat. Metal sat. Metal sat. Metal sat. Metal sat. Metal sat. IW-1.3 IW-3.54 IW-1.33 IW-3.11 IW-1.15

1800 2200 2200 2400 2200 2500 2200 2200 2200 2200 2500

5 5 9 9 21 25 9 9 18 18 20

53

135 123 81 87 44 34 55 479 34 148

O’Neill et al. (1998) O’Neill et al. (1998) O’Neill et al. (1998) O’Neill et al. (1998) O’Neill et al. (1998) O’Neill et al. (1998) Gessmann et al. (1999) Gessmann et al. (1999) Gessmann et al. (1999) Gessmann et al. (1999) Ohtani and Yurimoto (1996)

of diamonds containing lower mantle inclusions (Davies et al., 1999). In addition, the strongly layered nature of the Slave Craton has been interpreted as due to plume subcretion of the lower layer (Griffin et al., 1999a,b). 4.2. Sulfide Os model ages and records of mantle plumes Although most sulfides from both KP and TCY peridotites are inferred to have formed by crystallization of sulfide melts and/or metasomatic modification of primary sulfides, rather than as residual MSS, the sulfides dominantly have subchondritic 187Os/188Os and 187Re/188Os. Wang et al. (2009) demonstrated that Os model ages derived from sulfides that preserved primary features, and those affected only by Re-enrichment, can provide meaningful chronological information on the SCLM. These Os model ages are marked in bold font in Table 3; they yield TRD age peaks of 1.9, 1.7–1.6, 1.4– 1.3 and 0.9–0.8 Ga (Fig. 5). Notably, the KP sulfides do not show the age peak at 1.4 Ga, which is obvious in the TCY sulfides (Fig. 6). The age data may record periods of significant melt extraction or metasomatism events. These model ages suggest that many of the sulfides have resided in the SCLM at least since late Paleo-Proterozoic time, and some may have formed as early as the late Archean. It has been shown that these sulfide Os model ages are contemporaneous with the major tectonothermal events recorded in the overlying Cathaysian crust (Wang et al., 2003a, 2009). The close

32 30 18 16 209 12.9 87.1 11 17

correspondence of the sulfide TRD age peaks with crustal U–Pb, Nd and Hf model ages indicates that the sulfide ages reflect the timing of lithosphere-scale events (such as melting and metasomatism) in the SCLM, related to tectonothermal events that affected the overlying crust. The age of the Cathaysia Block has been pushed back by recent studies on crustal rocks (Xu et al., 2005, 2007; Yu et al., 2008), and this study of mantle sulfides provides some supporting evidence from its SCLM counterpart. The Cathaysia Block appears to contain Archean lithosphere terrane(s) older than previously thought. The Os model age record, along with recent studies on the assembly and breakup of supercontinents (e.g. Li et al., 2008; Zhao et al., 2004; references therein), may better constrain the lithospheric evolution of the Cathaysia Block since the Proterozoic. The Cathaysia Block has been interpreted as probably associated with western Laurentia when the Nena/Columbia supercontinent formed (Li et al., 2008; Rogers and Santosh, 2002; Zhao et al., 2004). The global 1.9-Ga superplume event corresponds to the assembly of the supercontinent Nena/Columbia, one of the major episodes of crustal growth on Earth (Condie, 1998). A minor 1.7-Ga superplume event (Isley and Abbott, 1999), associated with the emplacement of the 1.76–1.55-Ga anorogenic rapakivi granite in SW Laurentia, preceded the breakup of Nena/Columbia, or caused its breakup (Windley, 1993). Subsequent emplacement of widespread Mesoproterozoic (1.5– 1.3 Ga) anorogenic magma across Laurentia, resulting from large–

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0.825 Ga plume in South China Block to breakup Rodinia 1.468 Ga Moyie plume in W Laurentia to breakup Nena/Columbia 1.7 Ga plume in SW Laurentia prior to breakup of Nena/Columbia 1.9 Ga global superplume and major crustal formation

8

0.8-0.9

Penghu sulfides (n = 40)

Relative probability

1.6-1.7 1.3-1.4

6

1.9

4

Archean Hf model ages of zircons (2.5-3.4 Ga) Inherited Archean discordant U-Pb ages of zircons (2.5-2.7 Ga)

2

1

2

3

4

TRD (Ga) Fig. 5. Cumulative probability curve (Ludwig, 2000) of TRD model ages of all Penghu sulfides. Thick line represents the distribution pattern of ages. The histogram of ages is shown by thin grey lines. Corresponding mantle plume events are shown for comparison. See the text for details.

scale mantle upwelling (Hoffman, 1989) or the ascent of a mantle plume, further affected the Nena/Columbia supercontinent. This episode of anorogenic bimodal magmatism (mainly anorthositegabbro-peraluminous granite-rhyolite) was found in the Belt Basin region in western Laurentia and is known as the 1468 Ma Moyie event (Ernst and Buchan, 2001; Ernst et al., 2008). Li et al. (2008) proposed that the Cathaysia Block was an extension of Laurentia, since the former has similar crustal composition to what may be expected for the western source region of the Belt Basin (Ross et al., 1992). In Neoproterozoic time the supercontinent Rodinia, formed by the Grenvillian orogeny, started to breakup at the South China–Australia–Laurentia triple junction at 0.8 Ga, possibly due to a 825-Ma superplume as evidenced by the ages of synchronous dike swarms (Li et al., 2003). The ages of four distinct populations of sulfides in the SCLM beneath the Cathaysia Block apparently correlate with these superplume events (Fig. 5). Os model ages of the TCY Co-rich sulfides may further reveal the link between their formation and these superplume events. The TRD ages Co-rich sulfide grains which have been analysed by EMP for major-element concentrations and LAM-MC-ICPMS for Os isotope composition show can be divided into four episodes: 2.0, 1.7, 1.4 and 0.8 Ga (samples with underlined and Italic letters in Table 3). Among them, one low 187Re/188Os (0.0632) Co-rich sulfide has a TRD model age of 2.0 Ga, interpreted as its formation age. Considering that the uncertainty of these model ages is about 0.1–0.3 Ga (2r), the Os model age pattern of the TCY Co-rich sulfides suggests that the formation of the unique sulfides could be closely related to the corresponding superplume events. The Os model ages record the melting of mantle plumes at shallow depth to precipitate Corich sulfide melt; those that were then entrained by silicate minerals could retain their subchondritic Ni/Co. However, not all those sulfides with one of the four ages of proposed superplume events are supposed to be Co-rich. It is obvious that preservation of the unique Co-rich sulfide melt from the deep should be scarce as later re-melting, re-crystallization and metasomatism at shallow depth under open-system condition modify the distribution of Ni and Co in sulfides. Most Penghu sulfide grains are

interstitial to silicate minerals, making it unlikely that they could preserve the primary sulfide melt composition efficiently, although some might initially been enclosed and become interstitial due to grain-boundary migration during recrystallisation of the silicate minerals under closed-system condition. This may explain why only a small proportion of the TCY sulfides (less than one third irrespective of age group), show unique subchondritic Ni/Co ratios. There may be two reasons why no KP sulfides show Co-rich characteristics. First, few KP sulfides have low 187Re/188Os ratios (<0.07; Table 3) indicating a substantial influence by Re-addition metasomatism. If this metasomatism occurred at shallow depths it might obliterate the subchondritic Ni/Co ratios. Second, Os model age patterns of TCY Co-rich sulfides suggest a major formation event at 1.4-Ga plume event. However, there is no age peak at 1.4 Ga in the KP sulfides (Fig. 5); which suggests either the part of SCLM left untouched by the plume or later metasomatism eliminated the Os age record of 1.4 Ga under the KP domain. One might argue why Co-rich sulfides in other settings, where mantle plumes are generally considered to be important such as xenoliths in ocean island and oceanic plateau basalts, were not observed elsewhere and destruction of Co-rich sulfides during several billion years of tectonothermal evolution of lithosphere does not apply to young plume-related settings. The reasons might be: (1) only parts of mantle plumes originated from the lower mantle depth (Montelli et al., 2004). Those plumes originate from bottom of the upper mantle are unable to carry the Co-rich sulfides. (2) Studies on single sulfides in such young plume-related settings are limited. Whether there is Co-rich sulfides in such setting or not still remains uncertain and needs more further works to testify. 4.3. Noble gas ratios in the host peridotites Olivine separated from peridotites from the KP and TCY localities has very low noble gas concentrations (Table 4). The 40Ar/36Ar ratios of the olivine separates (322–844) show small excess 40Ar, relative to atmospheric argon with 40Ar/36Ar = 295.5. The 20 Ne/22Ne ratios in the samples appear to be higher than the atmo-

K.-L. Wang et al. / Journal of Asian Earth Sciences 37 (2010) 229–245

Relative probability

a

1.6

3

KP sulfides in peridotites (n = 14)

1.9

2

1

0 0

1

2

3

4

TRD (Ga)

b

4

Relative probability

1.3-1.4

1.6-1.7 TCY sulfides in peridotites (n = 13) 1.9

5. Conclusions

0.8 2

1

0

1

2

3

4

TRD (Ga)

Relative probability

from TCY locality) have 3He/4He of 4.6  6.7 R/RA, where R represents the measured 3He/4He ratio and RA the 3He/4He ratio of the atmosphere (1.4  10 6). One sample, TC0225 from the TCY locality, has a distinctively high 3He/4He ratio of 11 R/RA (Table 4). The 3He/4He ratios of 4.6  6.7R/RA in three samples are consistent with the value commonly found in old and metasomatised SCLM worldwide (Dunai and Porcelli, 2002). These are slightly lower than the MORB R/RA of 7–9 (Graham, 2002) and are interpreted as due to addition of metasomatic fluids enriched in radiogenic 4He to the SCLM, similar to the process generating HIMU basalts (Dunai and Porcelli, 2002). In contrast, the value of 11 R/RA is slightly higher than the typical MORB value. This relatively high 3He/4He can be interpreted as evidence for a primordial 3He-enriched component, which could have been transported by a mantle plume from the lower mantle. Similar helium isotope ratios (13 RA) have also been found in 250 Ma old Siberian flood basalts (Basu et al., 1995). Thus, the helium isotope composition found in sample TC0225 is consistent with the presence of mantle plume materials as suggested by the low-Ni/Co sulfide geochemistry and the sulfide Os model ages from the TCY lherzolites.

3

0

c

243

0.8-0.9

5

TCY sulfides in pyroxenite (n = 13)

4

3

2

1

0

0

1

2

3

A suite of unusual sulfides with subchondritic Ni/Co ratios (<21), together with other normal ones with superchondritic Ni/ Co ratios, occurs in lherzolites and pyroxenite from Penghu Islands, Taiwan, which locates at SE margin of the Cathaysia Block. Data on Ni–Co partition coefficients suggest a deep mantle origin for these sulfides, and their Os model ages record thermo-tectonic events corresponding the timing of known major mantle plumes. A plume origin is supported by limited noble gas data on separated olivines in their host peridotites. Such primary Co-rich sulfides have only been reported from two other occurrences, in mantle xenoliths entrained by kimberlites (Aulbach et al., 2004) and inclusions in diamonds (Davies et al., 2004), from the Slave Craton, Canada, where an association with lower mantle diamond inclusion phases supports a lower mantle origin. We thus propose that Co-rich sulfide melts may have been trapped in the lower mantle during core– mantle differentiation under high pressure (deep magma ocean) condition and then transported to the upper mantle by mantle plumes that entrained lower mantle materials. Age peaks in the spectrum of Os model ages (TRD) from the sulfides correspond to global superplume events related to the assembly of the Nena/Columbia supercontinent (1.9 Ga), the emplacement of anorogenic rapakivi granite in SW Laurentia prior to the breakup of Nena/Columbia (1.7 Ga), the 1468 Ma Moyie event in the breakup of Nena/Columbia and the 0.8 Ga breakup of Rodinia. The tight link between the Os model ages of these Corich sulfides and the mantle plume events manifest the Co-rich sulfide genetic relationship to these mantle plumes.

4

TRD (Ga) Fig. 6. Cumulative probability diagram of TRD model age patterns for (a) KP sulfides; (b) TCY sulfides in peridotites and (c) TCY sulfides in pyroxenite. Thick line represents the distribution pattern of ages. The histogram of ages is shown by thin grey lines. Note that peaks at 1.9, 1.6 and 0.8 Ga are recorded by both KP and TCY sulfides but the 1.4 Ga peak is only recorded by TCY sulfides. See the text for details.

spheric value of 9.8, thus indicating the existence of mantle neon in the samples. However, because of the large uncertainties in the neon isotopic ratios, the neon results do not allow the characterization of the sources of mantle neon (e.g. MORB-like or plume-related). On the other hand, the helium isotope compositions in the Penghu samples show more definitive results. Three lherzolite samples (KP0202 and KP214-20 from KP locality and TC0242-2

Acknowledgments We thank Chi-Yu Lee and Kung-suan Ho for help in the field work and Shiau-Huei Lin for sample preparation. We are grateful to Carol Lawson, Suzy Elhlou and Peter Wieland for assistance with the analytical facilities at GAU, Macquarie University in Sydney. Comments from reviewers, Jin-Hui Yang, Sonja Aulbach, and the editor Bor-Ming Jahn have improved the presentation of various aspects of this manuscript. Funding for this research was provided by the Australian Research Council (ARC), Macquarie University, the Ministry of Education and the National Science Council, Taiwan. The analytical data were obtained using instrumentation funded by ARC LIEF, and DEST, Systemic Infrastructure Grants, industry partners and Macquarie University. This is contribution

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