Oxygen isotope geochemistry of ultrahigh-pressure metamorphic rocks from 200–4000 m core samples of the Chinese Continental Scientific Drilling

Chemical Geology 242 (2007) 51 – 75 www.elsevier.com/locate/chemgeo

Oxygen isotope geochemistry of ultrahigh-pressure metamorphic rocks from 200–4000 m core samples of the Chinese Continental Scientific Drilling Ren-Xu Chen, Yong-Fei Zheng ⁎, Bing Gong, Zi-Fu Zhao, Tian-Shan Gao, Bin Chen, Yuan-Bao Wu CAS Key Laboratory of Crust–Mantle Materials and Environments, School of Earth and Space Sciences, University of Science and Technology of China, Hefei 230026, China Received 23 August 2006; received in revised form 16 February 2007; accepted 3 March 2007 Editor: D. Rickard

Abstract Oxygen isotope studies were carried out for ultrahigh-pressure metamorphic rocks from the main hole of the Chinese Continental Scientific Drilling in the Sulu orogen. The samples of interest include various types of lithology (mainly eclogite and gneiss) in a depth of 200 to 4000 m, containing five continuous core segments between contrasting lithologies. The results show a large variation in δ18O value from −10.41 to 9.63‰ for constituent minerals. Distinct 18O depletions are observed in the frequently alternated layers of eclogite and gneiss, in which the variations of δ18O values are gradual, regardless of lithologies. Both equilibrium and disequilibrium O isotope fractionations are observed between quartz and the other minerals. Special attention was paid to the relationship between distance, petrography and δ18O value of adjacent samples. The results show O isotope heterogeneities between the different and same lithologies on scales of about 20 to 50 cm, corresponding to the maximum scales of fluid mobility during the continental collision. Amphibolite-facies retrograde metamorphism during exhumation caused mineral reactions and O isotope disequilibria between some of the minerals. Considerable changes occur in δ18O and petrography at the contact between eclogite and gneiss, suggesting that the contact between different lithologies is the most favorable place for fluid activity. Despite the widespread retrogression, retrograde fluid was internally buffered in the stable isotope compositions. The retrograde fluid is of deuteric origin and thus was derived from the decompression exsolution of structural hydroxyl. Although locally external fluids became available along fault zones and lithological layers, it still acted within the exhumed slabs with internal origins. Fluid flow also took place after the amphibolite-facies retrogression, but it only affects the O isotope compositions of feldspar and mica. Premetamorphic protoliths are deduced to have heterogeneous δ18O values due to varying degrees of meteoric water–rock interaction before continental collision. Minimum depth of the 18O depletion is up to 3300 m. Together with areal 18O depletion in surface outcropped rocks along the Dabie–Sulu orogenic belt, at least 66,000 km3 of supracrustal rocks were interacted with meteoric water along the northern margin of the South China Block during the Neoproterozoic. © 2007 Elsevier B.V. All rights reserved. Keywords: O isotope; Petrography; Eclogite; Gneiss; Aqueous fluid; Continental collision; UHP metamorphism; CCSD; Sulu orogen

⁎ Corresponding author. E-mail address: [email protected] (Y.-F. Zheng). 0009-2541/$ - see front matter © 2007 Elsevier B.V. All rights reserved. doi:10.1016/j.chemgeo.2007.03.008

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1. Introduction The Dabie–Sulu ultrahigh-pressure (UHP) metamorphic belt in east-central China is the largest known UHP terrane in the world. Since findings of coesite and diamond in metamorphic rocks from this region (e.g., Okay et al., 1989; Wang et al., 1989; Xu et al., 1992), it has attracted extensive interest from the geoscience community (cf. Ernst and Liou, 1999; Jahn et al., 2003; Rumble et al., 2003; Zheng et al., 2003, and references therein). In order to understand the fluid regimes which can greatly assist in the development of formation and exhumation of UHP metamorphic terranes, many studies have been made on stable isotope geochemistry of UHP metamorphic rocks. Finding of anomalously low δ18O values of −11 to −4‰ in the UHP eclogites and gneisses indicates exchange of meteoric water with their protoliths before the Triassic subduction (e.g., Yui et al., 1995; Zheng et al., 1996; Baker et al., 1997; Rumble and Yui, 1998; Zheng et al., 1998, 1999; Fu et al., 1999; Zheng et al., 2001; Xiao et al., 2002; Fu et al., 2002, 2003). Preservation of such low δ18O signatures indicates very low fluid mobility during the bulk processes of subduction and exhumation of the continental crust (Fu et al., 2001; Zheng et al., 2003). These stable isotope studies have been successfully used to characterize chemical geodynamics and fluid regime during continental subduction and exhumation. As demonstrated by Zheng et al. (2003), continent–continent collision is not only characterized by fast subduction and fast exhumation of cold slabs with short-lived residence at mantle depths, but also its fluid regime is attributed to redistribution of deuteric water that attests significant activity during the exhumation with major derivation from decompression exsolution of structural hydroxyl in nominally anhydrous minerals. It has been long recognized that fluids play important roles in geochemical cycling during subduction of young and hot oceanic crust (e.g., Touret, 2001; Miller et al., 2002; Schmidt and Poli, 2003). Compared to the oceanic crust, the continental crust that was subducted to mantle depths and suffered the UHP eclogite-facies metamorphism is relatively poor in water (Liou et al., 1997; Fu et al., 2001; Zheng et al., 2003). Nevertheless, substantial quantities of H2O can still be transported to depths greater than 100 km by minor amounts of hydroxyl-bearing minerals such as epidote, phengite, zoisite, talc, topaz and lawsonite (e.g., Liou et al., 1995; Zhang et al., 1995, 2000, 2002; Schmidt and Poli, 2003; Li et al., 2004), and by nominally anhydrous minerals such as garnet, pyroxene and rutile under UHP conditions (e.g., Zhang et al., 2001; Su et al., 2002; Katayama and Nakashima, 2003; Su et al.,

2004; Zhang et al., 2004; Xia et al., 2005; Katayama et al., 2006; Chen et al., 2007). During exhumation, due to an abrupt decrease in pressure, hydroxyl-bearing mineral decomposition and hydroxyl exsolution would release significant amounts of aqueous fluid, resulting in amphibolite-facies retrogression, formation of quartz veins in eclogites and even syn-exhumation magmatism (Zheng et al., 2003; Li et al., 2004). These petrological and geochemical studies of surface samples have not only provided important constraints on fluid activity during pre- through syn- to post-UHP metamorphic phases, but also put forward a lot of questions: (1) how did water/rock interaction take place at vertical depths within an UHP terrane? (2) how large volumes of UHP rock have the 18O-depleted signature? (3) how continuous do mineral δ18O values change at the transition between different UHP lithologies? (4) on which scale can mineral O isotopes record fluid activity during the continental collision? In fact, there is still limited knowledge about fluid behavior during exhumation, especially about the quantitative scales of fluid activity in different UHP lithologies. The Chinese Continental Scientific Drilling (CCSD) in the Dabie–Sulu UHP metamorphic belt provides an advantageous and rare chance for collecting UHP rock samples continuously from the subsurface (Xu et al., 1998). Major goals of the CCSD include: (1) revealing of the crustal structure of convergent plate boundaries, (2) providing constraints on crust–mantle interactions and mantle behavior during deep subduction of continental crust, and (3) investigating the variation of fluid composition with depth, the flow patterns and cycling mechanisms, as well as changes of the fluid regime in time. O isotope variations through vertical sections from 100 to 3000 m for CCSD main hole have been investigated by Xiao et al. (2006) and Zhao et al. (2007), and the results show that meteoric water/rock interaction has reached depths of at least 2700 m and the O isotope shows different behavior on different scales. This paper presents a combined study of petrographic observation and O isotope analysis on selected samples from 200 to 4000 m for the CCSD main hole, with particular attention to the relationship between distance, petrography and δ18O value of adjacent samples. Five continuous core segments consisting of different UHP lithologies were investigated comprehensively in order to provide geochemical constraints on fluid behavior during the continental collision. Compared with the previous studies, the present study is not only an important account of oxygen isotope geochemistry from the lower part of the CCSD core, but also the first demonstration of the magnitude of the ancient hydrothermal system that produced

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the famous premetamorphic 18O depletion in the Dabie– Sulu orogenic belt. 2. Geological setting and samples UHP metamorphic rocks outcropping in the Dabie– Sulu orogenic belt represent exhumed products of the Triassic subduction of the South China Block beneath the North China Block (e.g., Cong, 1996; Liou et al., 1996; Li et al., 1999; Zheng et al., 2005). The eastern part of the orogenic belt was displaced approximately 500 km of leftlateral strike-slip by the NE-trending Tan-Lu fault to the

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northeast from the Dabie orogen to form the Sulu orogen (Fig. 1). The latter is separated from the North China Block to northwest by the Wulian–Qingdao–Yantai Fault (WQYF), and the South China Block to southeast by the Jiashan–Xiangshui Fault (JXF). The Sulu orogen can be divided into fault-bounded UHP and HP metamorphic zones (Xu et al., 2006). The UHP zone consists mainly of eclogite, orthogneiss, paragneiss, amphibolite, garnet peridotite, pyroxenite, and marble. The eclogite is generally enclosed within gneissic rocks and marble as layers and blocks. Although amphibolite-facies overprinting is significant, evidence

Fig. 1. Sketch map of geology in the Sulu orogen and the Donghai area, showing major lithotectonic units and the locations of outcrops and prepilot drillhole of CCSD-PP1, CCSD-PP2 and the CCSD main hole (modified after Liu et al., 2001). (1) Quaternary; (2) Tertiary basalt; (3) Cretaceous basin; (4) Cretaceous granite; (5) aegirine-bearing granitic gneiss; (6) amphibole-bearing granitic gneiss; (7) garnet-bearing granitic gneiss; (8) biotite-bearing granitic gneiss; (9) amphibole-bearing and biotite-bearing granitic gneiss; (10) epidote-bearing and biotite-bearing granitic gneiss; (11) supracrustal rocks, including paragneisses, kyanite-bearing and jadeite-bearing quartzite, and marble; (12) eclogite and ultramafic rocks; (13) ductile shear zone or fault; (14) drilling hole. NCB—North China Block; SCB—South China Block; WQYF—Wulian–Qingdao–Yantai fault; JXF—Jiashan–Xiangshui fault.

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for UHP metamorphism has been recognized in coesitebearing eclogite, schist, ultramafic rocks, and some gneisses (Zhang et al., 1995; Liou and Zhang, 1996; Ye et al., 2000; Liu et al., 2002, 2004a,b, 2005; Yang et al., 2003). Peak metamorphic conditions are defined by coesite-bearing eclogites with temperature of 750 to 850 °C and pressure of 2.5 to 4.5 GPa (e.g. Liou et al., 1998; Liu et al., 2004c). The HP zone is mainly composed of quartz–mica schist, chloritoid–kyanite– mica–quartz schist, marble, and rare blueschist (Zhang et al., 1995; Liu et al., 2004c). The main hole (MH) of Chinese Continental Scientific Drilling Project (CCSD), with a depth up to 5158.8 m, is located at the southwestern part of the Sulu orogen (N34°25′, E118°40′), about 17 km southwest of Donghai City (Fig. 1). This area is underlain mainly by paragneiss, orthogneiss and supercrustal rocks mostly with Neoproterozoic protolith ages (Liu et al., 2004a,b; Zheng et al., 2004), which were intruded by Cretaceous granite and unconformably overlain by Quaternary sedimentary cover (Fig. 1). Finding of coesite in the

gneiss and eclogite in this area, as well as finding of micro-diamond in the eclogite from Maobei (Xu et al., 2003, 2005) demonstrate that the high-grade rocks in this area were subducted to mantle depths to undergo UHP metamorphism. Metamorphic rocks recovered from CCSD-MH are various types of gneiss, eclogite, amphibolite, marble and peridotite. From 0 to 100 m, only rock cuttings were collected. The section from 100 to 2050 m has a core recovery of ∼80%. Eclogite and orthogneiss are the principal lithological types; other rocks include paragneiss, ultramafic rock and rarely, schist and quartzite. The section from 2050 to 4000 m is mainly composed of orthogneiss and paragneiss, with minor amounts of eclogite and amphibolite occurring as interlayers (Fig. 2). Coesite was found as inclusions within zircons from both gneiss and eclogite (Liu et al., 2001, 2002, 2004a,b, 2005, Zhang et al., 2006a), demonstrating that both eclogite and gneiss were subjected to in situ UHP metamorphism. P–T estimates of 675 to 815 °C and 3.1 to 4.4 GPa were obtained for eclogite at depths of 0 to

Fig. 2. Lithological profile of CCSD main hole from depths of 0 to 4000 m (revised after Liu et al., 2005).

R.-X. Chen et al. / Chemical Geology 242 (2007) 51–75

2050 m from CCSD-MH (Zhang et al., 2006b). SHRIMP U–Pb dating for coesite-bearing domains of zircons from the CCSD gneisses yields ages of 226 to 234 Ma for the UHP metamorphic event and about 750 to 780 Ma for the host metaigneous protolith (Liu and Xu, 2004; Liu et al., 2004a,b). These dates are consistent with known results from multi-method geochronological studies of surfaceexposed UHP metamorphic rocks along the Dabie–Sulu orogenic belt (e.g. Ames et al., 1996; Hacker et al., 1998; Li et al., 1999, 2000; Hacker et al., 2000; Zheng et al., 2003, 2004; Li et al., 2004). A combined study of petrography and geochronology for the CCSD-MH eclogite also reveals occurrence of a HP eclogite-facies recrystallization phase at 216 ± 3 Ma (Zhao et al., 2006). After general inspection of the CCSD-MH cores, we selected 87 samples for petrographic and O isotope investigations. The sampling depths range from 220 to 4000 m (Fig. 2). They include 51 samples from five continuous core segments, with particular attention to the transitions between eclogite, amphibolite and gneiss. The other 36 samples are sporadic at different depths. The five continuous core segments are composed of interlayered amphibolite and eclogite, gneiss and eclogite, eclogite and schist, gneiss, amphibolite and eclogite, and eclogite and amphibolite, respectively. They are from depths of (1) 1637.23 to 1640.55 m, (2) 1921.64 to 1924.03 m, (3) 2710.21 to 2711.66 m, (4) 3296.90 to 3300.20 m, and (5) 3585.10 to 3586.42 m, respectively, named as the first, second, third, fourth and fifth core segments. Lithologies, mineral assemblages, and depths of samples are summarized in Table 1. A general description of petrology for the CCSD-MH cores (especially for 0 to 2050 m) have been presented by Liu et al. (2004a, 2005), Su et al. (2005) and Zhang et al. (2006a,b). Fig. 3 illustrates the relationship between different lithologies in the fourth core segment. 3. Analytical methods Minerals were separated for isotopic analyses. After crushing, a shaking bed was used to separate heavy minerals, followed by magnetic separation using a magnetic separator, and then purified by hand picking under binocular microscopes. All mineral separates used for O isotope analyses are grained in sizes of 40 to 60 μm. Oxygen isotope analysis was accomplished for mineral separates by the laser fluorination techniques at University of Science and Technology of China in Hefei. A 25 W CO2 laser MIR-10 was used: its emission wavelength is 10.6 μm, which lies in the infrared domain, and can thus be well absorbed by O-bearing compounds; its power can be adjusted in the range of 0 to 100%, which is very important during sample analysis;

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and the diameter of the laser beam can be adjusted from 100 to 1820 μm. Minerals weighing about 1.5 to 2.0 mg were reacted with BrF5 under vacuum conditions. Obtained O2 was directly transferred to a mass spectrometer for O isotope ratio measurement. 18O/16O ratios were measured in a Finnigan Delta XP mass spectrometer and reported in the δ18O notation relative to the VSMOW standard. A number of replicate analyses gave the reproducibility better than ± 0.1‰ (Zheng et al., 2002). Three international standards and two national standard of China were used during the laser fluorination analyses: δ18O = 5.8‰ for UWG-2 garnet (Valley et al., 1995); δ18O = 5.2‰ for SCO-1 olivine (Eiler et al., 1996); δ18O = 10.0‰ for 91500 zircon (Zheng et al., 2004); δ18O = 11.1‰ for the National Standard of China GBW04409 quartz (Zheng et al., 1998); δ18O = − 1.7‰ for the National Standard of China GBW04410 quartz. Oxygen isotope composition of calcite was analyzed by Gasbench-MS online technique at University of Science and Technology of China in Hefei. This is an integrated procedure by using a Finnigan Gasbench II and a Finnigan MAT-253 mass spectrometer (MS) via open split in a Finnigan Conflo III interface. CO2 from calcite was extracted by reaction with phosphoric acid at 72 °C in the Gasbench. Then obtained CO2 was analyzed in the mass spectrometer in a continuous flow mode. The National Standard of China GBW04416 with δ18O = 18.96‰ was used as the reference material. The results were reported in the δ18O notation relative to the VSMOW standard. The reproducibility of replicate analyses is better than ± 0.2‰ for δ18O values. Assuming preservation of O isotope equilibration at the scale of sample measurement, apparent temperatures of quartz– mineral pairs are calculated with 2σ uncertainties using the fractionation equations calibrated by Zheng (1991, 1993a,b). As illustrated by Zheng et al. (1998, 1999), similar temperatures can be obtained if calibrations of Matthews (1994) are applied. 4. Results 4.1. Petrography Lithologies and mineral assemblages for the core samples are shown in Table 1 and Fig 4. Mineral composition of the eclogite from CCSD-MH can vary from sample to sample. Petrographic studies indicate four types of eclogite according to the content of minerals (Table 1): quartz-rich eclogite (quartz N 10 vol.%), rutilerich eclogite (rutile in 5–10 vol.%), phengite- and/or kyanite-rich eclogite (phengite + kyanite N 10 vol%) and normal eclogite. Similar subdivision was also made for

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Table 1 Modal abundance (%) of minerals in UHP rocks from CCSD-MH in the Sulu orogen Rock type

Depth (m)

Qz

Gt

Omp

Sym

02–648.9 02–652 02–700 02–704 02–738.7 02–739.5 02–927 02–928 04A-1 04A-2 04A-3 04A-4 04A-5 04A-6 04A-7 04A-8 04A-9 04A-10 04A-11 04A-12 04A-13 04C-1A 04C-2A 04C-3A 04C-5A 04C-6A 04C-8A 04C-9A 04C-10A 04C-12A 04C-13A 04B-1A 04B-2A 04B-3A 04B-4A 04B-5A 04B-6aA 04B-6bA 04B-6cA 04B-7aA 04B-7bA 04B-7cA 05B-1aA

Eclogite Eclogite Eclogite Eclogite Gneiss Gneiss Gneiss Gneiss Eclogite Eclogite Gneiss Eclogite Eclogite Gneiss Gneiss Gneiss Schist Eclogite Eclogite Eclogite Gneiss Gneiss Gneiss Gneiss Amphibolite Amphibolite Eclogite Eclogite Eclogite Eclogite Amphibolite Amphibolite Amphibolite Amphibolite Amphibolite Amphibolite Eclogite Eclogite Eclogite Eclogite Eclogite Eclogite Eclogite

648.9 652.0 700.0 704.0 738.7 739.5 927.0 928.0 222.15 347.65 367.00 572.00 1067.00 1170.30 1262.39 1390.60 1508.21 1644.70 1645.65 1756.03 1795.00 1637.23 1637.47 1637.83 1638.05 1638.18 1638.67 1638.85 1639.02 1640.15 1640.55 1921.64 1921.79 1922.30 1922.40 1922.89 1923.09 1923.29 1923.46 1923.63 1923.83 1924.03 2710.21

10–15

35 30 40 60

10 60–70 40–50 40

20–30 b5 b5

5–10 T 45–50 20–25 35–40 25 5 15–20 20–30 5–10 30–35 25–40 30–40 10–15 T T T 30–40 40 40 30 10–15 b2 b5 5–10 5 b5 5

T 30 50 20–30 30–40

30–40 20–25

5–10 10–15

5 55 40 40 40 30 40 25

Mus

10–15 5 1–2

5–10 b5 35

50–60 5–10

60 30–40 60

T

5 5–10 1–2 b5

Amp

10, Chl: b 5 10–20, Chl:5 20 10

10–15, Chl:4

T 30–40

35 30 20 30

T T 1–2

5–10 10–20 10–15 5–10

b5

Other

Ap

Zr

Vein

Ttn(T)

T T T

T T T T

Cc Fel

T T T T T T T T T T T T T T T T T T T T T T T T T T T T T T T T T T T T T T

T T T T T T T T T T T T T T T T T T T T T T T T T T T T T T T T T T T T T T

b2 b2

T

Ttn(T)

25 T 5

T T

8 2–3

T

10–25

3–5 10–15, Chl:5 5–10 30–40

Ky(T)

Ttn(T) Ttn(T) 1–2 T 1–2

T

2 5, Chl:3 Chl b 2 Chl b 3 5–10 b5

1 5 b5 b3 3 T 1–2 4 5 2–5 b2

Ilm

5 10–15

10, Chl:4

40 40–50

5–10 20–25 10–15 15–20 5–10, Chl b 3 5–10

50–60 60 50 40–50 60 40–50

5 b2 5

T b2 10 b5

Rt 2 4 5 T

5–10

T

50–60 10–20 30–40 20–30

30–35

Bt

12 T 5

30 10 20 10–15 15–20 15–20 40 50–55 40–50 40–50 50 50 b2

30–40 20–25 5

5–10 5–15 b5 27 5–10 10–20 20–25 30 30–40

b5 10–15 20 10–20

Ep

5

30–40

2 b5 b3 b5

Mt

5–10

25 20 30–40 5 20 40 30

Kfs

30 30 20–30 30–35

5 45 50 b2 30–35 20

Pl

8, Chl:2

5

T 2 T 2 2–3 2 2–5 1–2 T 5 1–2 1 2 2 2 2–3 2–5

2–3 T T

1–2 1–2 2 Ky(T)

T T

Ky(2) Ky(2) Ttn(T)

Cc, Fel Fel Cc, Fel

Cc Cc Cc Cc, Fel Cc, Fel Cc, Fel Cc, Fel Cc Fel Cc, Fel Fel Fel Fel Fel Cc, Fel Fel Fel

R.-X. Chen et al. / Chemical Geology 242 (2007) 51–75

Sample

Eclogite Eclogite Eclogite Eclogite Eclogite Eclogite Schist Schist Schist Schist Gneiss Gneiss Gneiss Gneiss Amphibolite Amphibolite Amphibolite Amphibolite Eclogite Eclogite Eclogite Eclogite Gneiss Gneiss Amphibolite Amphibolite Amphibolite Amphibolite Amphibolite Eclogite Gneiss Gneiss Eclogite Amphibolite

2710.41 2710.53 2710.61 2710.74 2710.84 2711.07 2711.32 2711.46 2711.54 2711.66 3296.90 3297.02 3297.20 3297.42 3297.57 3297.72 3297.85 3298.00 3299.73 3299.80 3300.00 3300.20 3585.10 3585.23 3585.66 3585.86 3586.07 3586.21 3586.42 2007.49 2076.75 2210.95 2271.10 2377.30

05A-10A 05A-9A 05A-8A 05A-7A 05A-6A 05A-5 05A-4A 05A-3A 05A-2A 05A-1A

Eclogite Gneiss Amphibolite Gneiss Gneiss Gneiss Amphibolite Gneiss Amphibolite Gneiss

2768.33 2820.97 2888.58 2997.46 3204.84 3553.80 3772.45 3918.39 3963.23 3999.75

10–15 5–10 5–10 b5 5 5–15 20–30 25 30 30 40 10–20 35 40 b5

20–30 20–25 15–20 10–20 20 10–15

b2 b2 T T b3 b2

30–35 40–45 50–55 40–45 30–40 40–50

20 25 10–20 20 25 30 25 30 10–15 20 b 15 25

T

b5 T b5 b2 T 30–40 35 10 5–10

20–30 10–15 30–40 20–30

b5 2–5 1–2 5–10

b5 b2 b5 b2

5 T 1

35

2–5

30 20

b2

T 5 b5 T 5 5–10 5–10 5

40 5 5 b5 5 b5

5 40–50 10–20 30–35 30–40 20 5–10 40–45 5 35

40

25 20–30

2–5 1–5

20–30

T

T 5–10

5

b2

T 2

T

40 60 50–60 80 70

10

2–3

20

2–5

20

5, Chl:3

1–2 5

2, Chl:10–15 10–15 5, Chl:2 15, Chl:10 20–30, Chl:5 5–10 2, Chl:b2 1–2, Chl:T T, Chl:1–2

25 5 35

1–2 T 1–2

b5 5 2–5 b5 5

2–5 2 4 2–4

Ttn(T)

1–2 2 2

Ttn(T) Ttn(T)

Cc(15) Cc(10) Cc(10) 2 2 2

Cc(5)

Ttn(T) 20 1–5

5 10–15

5

Ttn(T) Ttn(T)

2

T 10–20 10–15 30–35 20–30 20–25 30 30 40 25

2 T T T

Chl b 5 5–10, Chl:4

T

40

1 5 5 2–3 2 b2

Cc(2)

T 20–30 10–15

20–30 50

70–80 70 65 50

5–10 Chl: b 5 Chl:5–10 20 20–30 15 10–15

5–10 20 25

60

5, Chl:3 5 5–10 5–10, Chl:4 10–15, Chl:5 10–15, Chl:6

b5

60 70–80 40–50 50–60 25 20 30

35

2 5 b5 15–20 15–20 10 20–30 20–30 20–25 30–40

Cc(5)

b 10 b5 5 5–10 45 40 5–10 20

10, Chl:3 10–15 5 1 2–5 15, Chl:5 10 10–15, Chl:7 18, Chl:5 5–10 T(Chl)

b5

b2

40

30–40 b2 40

Ttn(T)

T T T T T T T T T T T T T T T T T T T T T T T T 2 T T T T T T T T T

T T T T T T T T T T T T T T T T T T T T T T T T T T T T T T T T T T

T T T T T T T

T T T T T T T T T T

T T

Cc, Fel Cc, Fel Cc, Fel Fel Fel Cc Cc Cc Cc Cc, Fel Cc b 5 Cc Cc, Fel Fel Cc Cc, Fel Cc Cc Cc Cc Cc Fel

Fel Cc, Fel Fel Fel

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05B-1aC 05B-1b 05B-1cA 05B-2A 05B-3 05B-4A 05B-5A 05B-6 05B-7 05B-8 05C-1A 05C-2 05C-3 05C-4A 05C-4C 05C-5A 05C-6A 05C-7A 05C-8 05C-9 05C-10 05C-11A 05D-1A 05D-2A 05D-5A 05D-6A 05D-7A 05D-8A 05D-8C 05A-15A 05A-14 05A-13A 05A-12A 05A-11A

Cc Cc Cc

Abbreviations: Qz—quartz, Pl—plagioclase, Kfs—potassic feldspar, Gt—garnet, Omp—omphacite, Mus—muscovite, Ep—epidote, Mt—magnetite, Bt—biotite, Amp—amphibole, Rt—rutile, Ky—kyanite, Chl—chlorite, Ilm—ilmenite, Ttn—titanite, Cc—calcite, Sym—symplectite, Ap—apatite, Zr—zircon, Fel—felsic vein, T—trace.

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R.-X. Chen et al. / Chemical Geology 242 (2007) 51–75

Fig. 3. Photo showing the different lithologies for the fourth core segment 05C.

CCSD-MH core samples (Su et al., 2005; Zhang et al., 2006a,b). General paragenesis of the eclogites is garnet + omphacite + rutile ± phengite ± kyanite ± quartz ± plagioclase ± biotite. The minerals show variable degrees of pervasively retrograded texture (Fig. 4a, b and c). Rutile occurs in all the investigated eclogites; in some samples, it is rimmed by ilmenite and/or titanite. Omphacite is generally replaced by amphibole and sodic plagioclase symplectite, and garnet by amphibole and plagioclase symplectite (Fig. 4b). Felsic and calcite veins were observed in some of the samples (Fig. 4c). The gneiss mainly consists of quartz + plagioclase ± biotite ± epidote ± K-feldspar ± phengite ± rutile ± amphibole ± garnet ± magnetite (Fig. 4d and e). A few of the gneisses are rich in biotite and amphibole but poor in felsic minerals relative to the most gneisses. They are renamed as the schist, but discussed geochemically within the gneiss category. In some samples, biotite is partially replaced by chlorite. Major mineral assemblages record amphibolite-facies overprinting. But occurrences of relict garnet in some gneisses and mineral inclusions (e.g., coesite) in zircon from the lithology indicates that it was subjected to UHP metamorphism together with the enclosed eclogite (Liu et al., 2004a,b, 2005). Felsic and calcite veins are found in some samples (Fig. 4f). The amphibolite is generally composed of amphibole, quartz, plagioclase, apatite, epidote, biotite and rutile (Fig. 4g), with abundant symplectite after the

primary omphacite and large amounts of residual garnet in some of the samples. Thus it may be a product of retrograded eclogite. Felsic and calcite veins are found in some samples (Fig. 4h). The five continuous core segments are composed of different lithologies (Figs. 3 and 4). Contacts between the different lithologies are sharp (Fig. 3), indicating that they are original bedding prior to the UHP metamorphism. Previous studies of tectonics (Xu et al., 2006), O isotopes (Xiao et al., 2006; Zhao et al., 2007), petrology and geochemistry (e.g., Zhang et al., 2006b), and geochronology (Liu et al., 2004a,b) have also demonstrated that the eclogite and gneiss were subjected to in situ UHP metamorphism. Table 1 presents a detailed outline of lithologies and mineral assemblages for the five core segments. The first core segment 04C and the fourth core segment 05C consist of a transition from gneiss, amphibolite to eclogite. Degree of retrograde metamorphism increases with decreasing distance from the contact between eclogite and gneiss. This is indicated by not only the amphibolite layer between eclogite and gneiss, but also the amount of symplectite after omphacite and garnet (Table 1). For the third core segment 05B, there is no amphibolite between eclogite and schist, but the eclogites adjacent to the schist have larger extent of retrogression. This is indicated by the observation that more amphibole grains in symplectite after omphacite occurs in the eclogite adjacent to gneiss. For the fifth core segment 05D that consists

R.-X. Chen et al. / Chemical Geology 242 (2007) 51–75

of a transition between gneiss and amphibolite, a felsic vein of about 3 cm length occurs as an interlayer between them. Calcite was commonly observed to occur as matrix and/or vein in the gneiss and amphibolite.

59

4.2. Oxygen isotopes The O isotope compositions of mineral separates from eclogite, amphibolite, and gneiss are listed in Table 2. The

Fig. 4. Microphotographs of eclogite, gneiss and amphibolite from the CCSD main hole. (a) Eclogite (02–652, 652.00 m), containing garnet, omphacite, rutile. (b) Eclogite (05A-10A, 2768.33 m), containing garnet, omphacite, amphibole, quartz, rutile. Most of the omphacite was replaced by symplectite, and garnet is commonly rimmed by symplectite. (c) Eclogite (04A-11, 1645.65 m), consisting of garnet, omphacite, with minor phengite and rutile. Felsic and calcite vein was found in the sample. (d) Gneiss (02–927, 927.00 m), consisting of quartz, plagioclase, garnet, biotite, epidote, muscovite and K-feldspar. Biotite is partially replaced by chlorite. (e) Gneiss (04A-8, 1390.60 m), containing quartz, plagioclase, K-feldspar, biotite, with minor magnetite. (f) Gneiss (02–738.7, 738.7 m), consisting of quartz, plagioclase and biotite. Calcite vein was found in the sample. (g) Amphibolite (05A-8A, 2888.58 m), containing amphibole, plagioclase, quartz, epidote and biotite. Biotite was partially replaced by chlorite. (h) Amphibolite (05A-2A, 3963.23 m), consisting of amphibole, K-feldspar and quartz. Calcite vein was found cutting through the amphibole. Mineral abbreviations: Gt = garnet, Omp = omphacite, Phg = phengite, Sym = symplectite, Qz = quartz, Rt = rutile, Ilm = ilmenite, Ttn = titanite, Am = amphibole, Bt = biotite, Pl = plagioclase, Kfs = K-feldspar, Ep = Epidote, Mus = muscovite, Mt = magnetite, Cc = Calcite.

60

Table 2 Oxygen isotope compositions of minerals in UHP metamorphic rocks from CCSD-MH (δ18O in ‰) Rock type

Depth (m)

02–648.9 02–652 02–700 02–704 02–738.7 02–739.5 02–927 02–928 04A-1 04A-2 04A-3 04A-4 04A-5 04A-6 04A-7 04A-8 04A-9 04A-10 04A-11 04A-12 04A-13 04C-1A 04C-2A 04C-3A 04C-5A 04C-6A 04C-8A 04C-9A 04C-10A 04C-12A 04C-13A 04B-1A 04B-2A 04B-3A 04B-4A 04B-5A 04B-6aA 04B-6bA 04B-6cA 04B-7aA 04B-7bA 04B-7cA 05B-1aA

Eclogite Eclogite Eclogite Eclogite Gneiss Gneiss Gneiss Gneiss Eclogite Eclogite Gneiss Eclogite Eclogite Gneiss Gneiss Gneiss Schist Eclogite Eclogite Eclogite Gneiss Gneiss Gneiss Gneiss Amphibolite Amphibolite Eclogite Eclogite Eclogite Eclogite Amphibolite Amphibolite Amphibolite Amphibolite Amphibolite Amphibolite Eclogite Eclogite Eclogite Eclogite Eclogite Eclogite Eclogite

648.9 652.0 700.0 704.0 738.7 739.5 927.0 928.0 222.15 347.65 367.00 572.00 1067.00 1170.30 1262.39 1390.60 1508.21 1644.70 1645.65 1756.03 1795.00 1637.23 1637.47 1637.83 1638.05 1638.18 1638.67 1638.85 1639.02 1640.15 1640.55 1921.64 1921.79 1922.30 1922.40 1922.89 1923.09 1923.29 1923.46 1923.63 1923.83 1924.03 2710.21

Qz

Pl

Kfs

6.08

1.91 2.96 0.12 −0.94 3.55 8.66 2.17 −0.45 1.57 −2.69 7.39

Gt 2.16 4.23 1.81 3.23

2.91 − 0.33

− 0.02

Mus

Ep

Mt

2.90 3.57 2.21 3.02

−3.73 0.63 −2.63 −1.80 −1.09 −3.32

0.30 − 0.74

Omp

− 0.87 − 2.55

−1.66 −0.91 −4.37

8.47 9.39 9.30 9.43 9.10

8.11 7.80 7.33

5.32 5.09 5.66

7.27 8.06 7.58 7.55

−1.49 − 2.79

8.87 8.71 9.24

− 4.82

− 3.38

− 2.43

−4.22

− 7.55 − 10.41 − 1.10

− 3.11 − 7.83 0.23 0.65

− 0.65 1.44 − 0.19

8.34 8.53 2.05

− 0.37

5.39 5.34 5.18 5.84 5.75 5.55 − 1.46

1.01 1.17 0.71 0.77 1.06

5.90 5.43 5.99

6.42 6.68 6.41 6.19 5.65 6.40 − 1.29

WR 4.34 3.77 2.05 3.15 0.97 1.20 − 1.54 − 0.94 0.63 − 0.21 0.14 − 1.10 − 2.58 1.10 − 2.25 5.27 1.02 5.24 5.05 5.46 7.75 8.00 7.89 7.74

5.02

8.20

4.85 5.11 5.32 4.93 5.29 5.27 4.86 −1.21

− 2.79 (Ky)

0.24

4.55

8.75 8.20

− 4.63 − 8.40

0.67

4.78

6.71 6.51

− 3.94 − 6.04

1.24 1.28 1.04

3.44 5.36 5.36 5.73 5.67

Other

2.68 (Ttn) 1.01 (Ttn)

6.10 6.79 5.34 5.74 5.49 5.45 0.82

7.52 5.06 5.35 4.86 4.83

Chl

− 0.75

− 2.78 − 1.71 − 5.92 − 6.59

0.10

1.57 4.52 4.99 5.05

Rt 0.34

1.03 − 0.78 5.11

9.10

Hb

1.67

− 3.70 0.41 − 1.84

Bt

− 2.82

4.00 4.50 4.50 5.23 3.88 4.48

1.12 1.22 1.18 1.21 1.07 1.56 1.61 1.45 1.67 2.08 1.94 − 5.87

4.38 6.53 (Ky) 4.83 5.78 (Ky) 6.24 (Ky) 5.98 (Ky)

5.24 5.56 5.47 5.30 4.11 4.35 4.69 5.23 4.55 4.68 5.29 5.33 5.20 5.57 5.86 5.28 − 0.61

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Sample No.

Eclogite Eclogite Eclogite Eclogite Eclogite Eclogite Schist Schist Schist Schist Gneiss Gneiss Gneiss Gneiss Amphibolite Amphibolite Amphibolite Amphibolite Eclogite Eclogite Eclogite Eclogite Gneiss Gneiss Amphibolite Amphibolite Amphibolite Amphibolite Amphibolite Eclogite Gneiss Gneiss Eclogite Amphibolite Eclogite Gneiss Amphibolite Gneiss Gneiss Gneiss Amphibolite Gneiss Amphibolite Gneiss

2710.41 2710.53 2710.61 2710.74 2710.84 2711.07 2711.32 2711.46 2711.54 2711.66 3296.90 3297.02 3297.20 3297.42 3297.57 3297.72 3297.85 3298.00 3299.73 3299.80 3300.00 3300.20 3585.10 3585.23 3585.66 3585.86 3586.07 3586.21 3586.42 2007.49 2076.75 2210.95 2271.10 2377.30 2768.33 2820.97 2888.58 2997.46 3204.84 3553.80 3772.45 3918.39 3963.23 3999.75

1.99 1.69 2.51 3.77 2.78 2.64 2.78 2.11 1.87 1.74 8.12 8.75 7.77 7.41

1.31 − 0.68 0.54 1.79 1.67 − 0.37 0.77 1.14 0.59 5.61 6.18 5.19 5.19 4.54 3.69 3.81 4.21

2.87

0.33 −0.68

− 0.59 − 1.53 0.55

− 0.91 − 1.60 − 2.14 − 1.62

− 0.28

5.25 5.31 5.13 5.57

3.60 3.75

6.81

4.44 4.49

4.01 4.28 4.57 4.88

3.95

− 3.14 − 3.37 1.33 0.91 3.73 3.66

2.95 (Cc) 4.76 (Cc) 4.52 (Cc) 1.03

− 2.95 3.48 5.53

5.44

5.13

4.65

4.08

4.87

6.21

− 2.85 0.41 − 0.47 − 2.47

2.03 4.60

4.42 4.87

2.06 (Ttn) 1.09

2.14 − 3.62 1.74 2.48

− 1.43 1.89

− 1.98

1.89 4.16

4.20 (Cc)

5.69

5.80 5.93

6.59 5.43 6.92

7.40 (Cc) 1.97 1.61 1.78 1.55 1.99

1.03

2.92 1.93 2.36 2.90

3.28

− 6.30 − 5.35 − 5.75 − 3.07 − 3.39 − 5.60

5.65

4.65

3.59 − 3.26 2.95 3.85 5.31 6.40 6.75 5.64 7.16

2.18 3.18 3.11 3.09

3.79 2.97

4.26

5.17 5.22

−1.52 −1.73 −2.43 −2.25

5.93

5.59

8.92 7.93 7.48 9.00 3.10 8.24 6.05 − 0.19 5.57 5.55 9.63 9.22 9.52 8.54 9.28

− 3.22 − 3.42 − 3.10 − 3.35 − 3.66 − 3.26 − 3.32 − 3.16 − 3.63 − 3.45

5.47 5.31 4.08

5.34 5.69

3.54 5.67 5.03

−1.70

5.37

9.30 4.78 8.46 8.23

− 2.26 − 2.67 − 1.25 0.76 − 1.46 − 1.37

0.75 − 1.65 0.49

− 0.45 − 5.26 − 1.73 − 0.54 3.53 2.44 1.16 1.49 2.36

−4.53

4.40 3.89 3.33

− 3.51

5.20 (Cc) 0.02 0.68

3.06 (Ttn)

− 1.32 − 2.04 − 2.19 0.56 − 0.55 − 1.15 − 0.15 − 0.21 − 0.52 − 0.69 6.52 6.34 6.35 6.34 4.21 4.05 4.27 4.32 4.10 4.06 4.33 4.26 5.62 5.63 1.65 1.11 3.45 3.51 4.94 6.69 6.82 5.85 0.62 4.39 4.99 − 3.47 4.06 3.21 6.18 5.77 8.40 4.93 7.79

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05B-1aC 05B-1b 05B-1cA 05B-2A 05B-3 05B-4A 05B-5A 05B-6 05B-7 05B-8 05C-1A 05C-2 05C-3 05C-4A 05C-4C 05C-5A 05C-6A 05C-7A 05C-8 05C-9 05C-10 05C-11A 05D-1A 05D-2A 05D-5A 05D-6A 05D-7A 05D-8A 05D-8C 05A-15A 05A-14 05A-13A 05A-12A 05A-11A 05A-10A 05A-9A 05A-8A 05A-7A 05A-6A 05A-5 05A-4A 05A-3A 05A-2A 05A-1A

Abbreviations: Qz—quartz, Pl—plagioclase, Kfs—potassic feldspar, Gt—garnet, Omp—omphacite, Mus—muscovite, Ep—epidote, Mt—magnetite, Bt—biotite, Hb—hornblende, Rt—rutile, Ky—kyanite, Chl—chlorite, Ttn—titanite, Cc—calcite, WR—whole-rock (estimates δ18O value from modal abundance).

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62

R.-X. Chen et al. / Chemical Geology 242 (2007) 51–75

results show large variations in O isotope ratios (Figs. 5 and 6). The eclogite has δ18O values of −3.32 to 7.52‰ for garnet, −2.55 to 5.84‰ for omphacite, −0.45 to 9.30‰ for quartz and −2.79 to 6.81‰ for muscovite. The amphibolite has δ18O values of −4.53 to 5.23‰ for amphibole, −3.26 to 6.75‰ for plagioclase and −0.19 to 9.22‰ for quartz. δ18O values of minerals from the gneiss vary from −2.69 to 9.63‰ for quartz, −0.74 to 8.11‰ for plagioclase and −7.83 to 3.53 for biotite. Based on mineral δ18O values and modal abundance, whole-rock δ18O values are estimated to be about −2.58 to 5.86‰, −3.47 to 5.77‰ and −2.25 to 8.40‰ for the eclogite, amphibolite and gneiss, respectively (Table 2). All the investigated samples can be divided into two groups according to their bulk δ18O values: 18O-depleted rocks (δ18O as low as − 3.7‰ for garnet) and normal δ18O rocks with δ18O N + 5.6‰. An O isotope profile from 0 to 4000 m for CCSD-MH is illustrated in Fig. 5. It appears that the O isotope compositions vary with spatial distribution, and mineral δ18O values are significantly heterogeneous. In particular, abrupt changes in δ18O occur on scales of about 100 m. 18O depletion is observed down to 1600 m, from 2300 to 2700 m, about 2900 m, and 3200 to 3300 m, respectively, regardless of lithology. In contrast, the samples from 1650 to 2300 m and 3300 to 4000 m have the normal δ18O values without 18O depletion. The lowest δ18O values occur at depths of 900 to 1100 m, 2500 to 2600 m and 2900 m, respectively, close to the orthogneiss layers. Figs. 7, 8, 9, 10 and 11 show five O isotope profiles of minerals in the eclogite, amphibolite and gneiss from the first, second, third, fourth and fifth core segments, respectively. It appears that the O isotope compositions vary with spatial distribution, the eclogite and/or gneiss from the first, second, fourth, fifth core segments have the normal or slightly reduced δ18O values, whereas those from the third core segment show significantly reduced δ18O values. O isotope heterogeneity on different scales can be investigated by looking at the relationship between distance, petrography and δ18O value of adjacent samples. For the first core segment 04C, the amphibolite layer separates the eclogite from the gneiss. Because of the difficulty in recovering mineral separates from the amphibolite, only few minerals were available for the O isotope analysis. The results show that the eclogite has δ18O values of 4.83 to 5.35‰ for garnet, 5.36 to 5.73‰ for omphacite, 8.71 to 9.24‰ for quartz, 5.43 to 5.99‰ for muscovite, and 0.71 to 1.17‰ for rutile. The gneiss has δ18O values of 9.30 to 9.43‰ for quartz, 7.33 to 8.11‰ for plagioclase, 7.55 to 8.06‰ for K-feldspar, 5.45 to 5.74‰ for muscovite, −0.65 to 1.44 for magnetite. As shown by Fig. 7, δ18O values for quartz, feldspar and

Fig. 5. Oxygen isotopic compositions vs. depths of CCSD-MH core samples in the Sulu orogen (for clarity not all measured values are presented in this figure).

mica are almost concordant with each other for gneisses 04C-1A to 04C-3A at depths of 1637.23 to 1637.83 m, and eclogites 04C-8A to 04C-12A at depths of 1638.67 to 1640.15 m are relatively homogeneous in their mineral δ18O values (Table 2). This suggests that the mineral δ18O values in the eclogite and gneiss are homogeneous on scales of 48 cm and 148 cm, respectively. The gneiss has similar δ18O values for quartz and muscovite to the eclogite, whereas the amphibolite has significantly low δ18O values for muscovite but high δ18O values for garnet. A high δ18O value of 7.52‰ for garnet occurs in amphibolite 04C-6A that is 49 cm away from adjacent eclogite 04C-8A. Magnetite from the gneiss shows variable δ18O values from −0.65 to 1.44‰ on scales of 24 to 36 cm. All of these suggest that the δ18O heterogeneity can occur on scales of 22 to 49 cm in this core segment. For the second core segment 04B, the eclogite has δ18O values of 4.86 to 5.32‰ for garnet, 5.18 to 5.84‰

R.-X. Chen et al. / Chemical Geology 242 (2007) 51–75

63

Fig. 6. Plots of the O isotope composition of quartz vs. coexisting minerals in eclogites and gneisses from the CCSD main hole. Fractionation lines at different temperatures are drawn by applying the theoretical calibrations of Zheng (1991, 1993a,b).

for omphacite, 5.65 to 6.68‰ for muscovite, and 1.45 to 2.08‰ for rutile. δ18O values for single minerals in the amphibolite range from 8.20 to 8.75‰ for K-feldspar,

3.88 to 5.23‰ for amphibole and 1.07 to 1.22‰ for rutile. Except for amphibole, the other mineral δ18O values are relatively homogeneous in both eclogite and amphibolite,

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R.-X. Chen et al. / Chemical Geology 242 (2007) 51–75

Fig. 7. Profiles of δ18O values in the first core segment 04C of CCSDMH.

Fig. 9. Profiles of δ18O values in the third core segment 05C of CCSD-MH.

respectively. The minerals from the eclogite show a similar δ18O variation except muscovite. The eclogite and amphibolite have similar δ18O values for feldspar. In contrast, the rutile from the amphibolite has lower δ18O values than those from the eclogite. Similarly, the amphibole from the amphibolite has lower δ18O values than omphacite from the eclogite, suggesting that aqueous fluid for amphibolite-facies retrogression was depleted in 18 O relative to the precursor omphacite. As shown by Fig. 8, amphibole from amphibolite 04B-3A at a depth of 1922.30 m has a δ18O value of 5.23‰, whereas amphibole from amphibolite 04B-4A at a depth of 1922.40 m has a δ18O value of 3.88‰ (Table 2). Amphibole from the adjacent amphibolites has higher or lower δ18O values of ∼4.5‰ (not only 04B-1A and 04B-2A at depths of 1921.64 and 1921.79 m, but also 04B-5A at depth of 1922.89 m). Thus a δ18O difference of 1.35‰ occurs in amphibole on scales of 20 to 51 cm. For the third core segment 05B, the eclogite has δ18O values of − 2.67 to 0.76‰ for garnet, − 1.70 to 0.33‰ for

omphacite, 1.69 to 3.77‰ for quartz, − 1.53 to 0.55‰ for muscovite, − 0.68 to 1.79‰ for plagioclase, − 3.66 to − 2.82‰ for biotite, and − 6.30 to − 3.07‰ for rutile. The schist has δ18O values of 1.74 to 2.78‰ for quartz, − 0.37 to 1.14‰ for plagioclase, − 2.14 to − 0.91‰ for epidote, − 2.43 to − 1.52‰ for amphibole, and − 3.63 to − 3.16‰ for biotite. As illustrated by Fig. 9, except for biotite, both eclogite and schist are significantly heterogeneous in mineral δ18O values. This is indicated by not only a large δ18O difference between eclogite and schist, but also variations in δ18O among eclogite and schist. For quartz, two samples of eclogite 05B-1aA and 05B-1aC away from the contact have relatively homogeneous δ18O values of 1.99 to 2.05‰; three samples of eclogite 05B-1b, 05B-1cA and 05B-2A show slight increases in δ18O value from 1.69 to 3.77‰; two samples of eclogite 05B-3 and 05B-4A and one sample of schist 05B-5A have homogeneous δ18O values of 2.51 to 2.78‰; three samples of schist 05B-6, 05B-7 and 05B-8 show relatively homogeneous δ18O values of

Fig. 8. Profiles of δ18O values in the second core segment 04B of CCSD-MH.

Fig. 10. Profiles of δ18O values in the fourth core segment 05D of CCSD-MH.

R.-X. Chen et al. / Chemical Geology 242 (2007) 51–75

1.74 to 2.11‰. As illustrated in Fig. 8, however, eclogites 05B-1b, 05B-1cA, 05B-2A and 05B-3 at depths of 2710.53 to 2710.84 m show variable δ18O values of − 2.67 to 0.76‰ for garnet on a scale of 31 cm (Table 2). The similar-scale δ18O variations also occur in omphacite, rutile, muscovite quartz and plagioclase. These δ18O variations occur in the interior of the eclogite that is away from the contact, suggesting a fracture control of δ18O heterogeneity. For the fourth core segment 05C, the amphibolite layer separates the eclogite from the gneiss. The gneiss has δ18O values of 7.41 to 8.75‰ for quartz, 5.19 to 6.18‰ for plagioclase, and 5.34 to 5.69‰ for K-feldspar. The amphibolite has δ18O values of 4.01 to 4.88‰ for amphibole, 3.69 to 4.54‰ for plagioclase, and 2.18 to 3.18‰ for biotite. The eclogite has δ18O values of 5.13 to 5.57‰ for garnet, 3.60 to 3.75‰ for omphacite, and 1.55 to 1.99‰ for rutile. As illustrated by Fig. 10, the gneiss exhibits a slight decrease in δ18O values for quartz, plagioclase and muscovite when moving toward the contact between gneiss and amphibolite at depths of 3296.90 to 3297.42 m. Similarly, the amphibolite also shows a progressive decrease in mineral δ18O values toward the contact between amphibolite and gneiss at depths of 3297.57 to 3298.00 m. Biotite from amphibolite 05C-4C at 3297.57 m has a δ18O value of 2.18‰, which is significantly lower than δ18O values of 3.09 to 3.18‰ for amphibolites 05C-5A, 05C-6A and 05C-7A at depths of 3297.72 to 3298.00 (Table 2). About 1.0‰ difference occurs in biotite on a scale of 15 m (Fig. 9). Both amphibole and plagioclase from the amphibolite have lower δ18O values than garnet from the eclogite, suggesting that aqueous fluid for amphibolite-facies retrogression was depleted in 18O relative to the precursor omphacite and garnet. The eclogites are relatively homogeneous in δ18O values for rutile, omphacite and garnet, but heterogeneous for biotite, muscovite and quartz on scales of 20 to 27 cm. For the fifth core segment 05D, a felsic vein separates the amphibolite from the gneiss. Two samples of the gneiss with a distance of 13 m (05D-1A and 05D-2A, respectively, at depths of 3585.10 m and 3585.23 m) have similar δ18O values of 8.46‰ and 8.23‰ for quartz and 4.44 ‰ and 4.49‰ for K-feldspar. Amphibole from amphibolites 05D-6A and 05D-7A (at depths of 3585.86 m and 3586.07 m) adjacent to a felsic vein has δ 18 O values of 0.91 to 1.33‰, significantly lower than 3.66 to 3.73‰ for amphibolites 05D-8A and 05D-8C (3586.21 m and 3586.42 m) away from the vein. About 2.33 to 2.82 ‰ differences in δ18O values occur in amphibole on a scale of 14 cm (Fig. 10). For biotite, amphibolite 05D-7A and 05D-8A have δ18O values of 1.93 to 2.36‰ lower than 05D-6A and

65

05D-8C which have δ18O values of 2.90 to 2.92‰, suggesting heterogeneous δ18O distribution in biotite on a scale of 21 cm. For calcite, three samples of amphibolite 05D-6A, 05D-7A and 05D-8C away from the felsic vein have relatively homogeneous δ18O values of 4.20 to 4.76‰, and one sample of amphibolite 05D5A adjacent to the vein has a significantly lower δ18O value of 2.95‰. A δ18O difference of 1.81‰ occurs in calcite between amphibolites 05D-5A and 05D-6A at a distance of 20 m. Quartz and feldspar from the amphibolite have lower δ18O values than those from the gneiss (Fig. 11). As illustrated in Fig. 6, the mineral-pair O isotope fractionations are in both equilibrium and disequilibrium in the eclogite, amphibolite and gneiss. The temperature errors are estimated to be about 30 to 50 °C in terms of the analytical uncertainties in mineral δ18O determination and fractionation curve calibration. The mineral-pair temperatures span a wide range and some of them are too low to represent UHP metamorphic conditions (Table 3). Except for eclogites 04A-2 and 05C-11A, nevertheless, O isotope geothermometry for quartz–garnet pairs from the eclogite yields temperatures of 560 to 785 °C (Table 3). They are close to petrological temperatures of 675 to 815 °C responsible for eclogite-facies metamorphic conditions (Zhang et al., 2006b). This indicates attainment and preservation of O isotope equilibrium between quartz and garnet at eclogite-facies conditions. On the other hand, quartz–omphacite, quartz–phengite and quartz–rutile pairs from the eclogite yield temperatures of about 300 to 780 °C, 380 to 665 °C and 400 to 535 °C (Table 3), respectively (except eclogites 04A-2 and 05C-11A). These are differentially lower than the temperature of the eclogite-facies metamorphism, suggesting the effect of retrograde isotopic exchange during

Fig. 11. Profiles of δ18O values in the fifth core segment 05D of CCSD-MH.

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Sample No.

Rock type

Depth (m)

Qz–Gt

Qz–Omp

02–648.9 02–738.7 02–739.5 02–927 02–928 04A-1 04A-2 04A-3 04A-5 04A-6 04A-7 04A-8 04A-11 04C-1A 04C-2A 04C-3A 04C-5A 04C-9A 04C-10A 04C-12A 04B-7bA 05B-1aA 05B-1aC 05B-1b 05B-1cA 05B-2A

Eclogite Gneiss Gneiss Gneiss Gneiss Eclogite Eclogite Gneiss Eclogite Gneiss Gneiss Gneiss Eclogite Gneiss Gneiss Gneiss Amphibolite Eclogite Eclogite Eclogite Eclogite Eclogite Eclogite Eclogite Eclogite Eclogite

648.90 738.70 739.50 927.00 928.00 222.15 347.65 367.00 1067.00 1170.30 1262.39 1390.60 1645.65 1637.23 1637.47 1637.83 1638.05 1638.85 1639.02 1640.15 1923.83 2710.21 2710.41 2710.53 2710.61 2710.74

620

555

630

370

785 180 615 795

560 100

Qz–Pl

2105 1350

515 490 490 370 400

Qz–Mus

Qz–Rt

1100 – 495

505 590 500 625 505 480

515

395 1100 400 490 490

Qz–Ep

Qz–Chl

Qz–Ky

230

470 440 445 400 305

Qz–Mt

Qz–Ttn

460 540 535

580

405 435 450

440 190

650

415

470

460 450

580

490 440 340

440 590 460 420

Qz–Hb

470 435

410

720 630 460 680 630 560 720 720 580 565 640 770

1800

Qz–Bt

780 725 –

595

Qz–Kfs

660 430 410 390 130 525 480 480 540 470 665 380 490

425 420 530 440

165 410 405 420 405 510 420 400 470 400 480

590

R.-X. Chen et al. / Chemical Geology 242 (2007) 51–75

Table 3 Oxygen isotope temperature for quartz–mineral pairs from CCSD-MH (in °C) ⁎

580 610

510

800 290 690 1055 780 380 370 365

610 –

–

345 380 370 430 410 440

535 400 500 560 500 540

260

500 480 620

310

300

440 975

1015 330 280

135 155

600 785

575

350

500

670

430

360 510

750

630 755

560

100 100

560

155

385

505 485

340 425 690

460 455 420 500

390 295 360 570 170 330 340 300 455

406

420 400

420

460

420

380 415 330 180 240 200 320

calibrations of Zheng (1991, 1993a,b).

340 450 300 370 370 320 250 300 320

430

320

495

430 490

150

620 520

480 320 370 380 400

400

180 180

480 410 480

430

R.-X. Chen et al. / Chemical Geology 242 (2007) 51–75

05B-3 Eclogite 2710.84 05B-4A Eclogite 2711.07 05B-5A Schist 2711.32 05B-6 Schist 2711.46 05B-7 Schist 2711.54 05B-8 Schist 2711.66 05C-1A Gneiss 3296.90 05C-2 Gneiss 3297.02 05C-3 Gneiss 3297.20 05C-9 Eclogite 3299.80 05C-11A Eclogite 3300.20 05D-1A Gneiss 3585.10 05D-2A Gneiss 3585.23 05D-6A Amphibolite 3585.86 05A-15A Eclogite 2007.49 05A-14 Gneiss 2076.75 05A-13A Gneiss 2210.95 05A-12A Eclogite 2271.10 05A-11A Amphibolite 2377.30 05A-10A Eclogite 2768.33 05A-9A Gneiss 2820.97 05A-8A Amphibolite 2888.58 05A-7A Gneiss 2997.46 05A-6A Gneiss 3204.84 05A-5 Gneiss 3553.80 05A-4A Amphibolite 3772.45 05A-3A Gneiss 3918.39 05A-2A Amphibolite 3963.23 05A-1A Gneiss 3999.75 ⁎ Temperature calculations are based on the

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R.-X. Chen et al. / Chemical Geology 242 (2007) 51–75

exhumation of the UHP eclogite. Temperatures for the quartz–garnet, quartz–omphacite and quartz–rutile pairs from eclogite 04A-2 are very low, 180°C, 100°C and 190°C, respectively (Table 3). Eclogite 05C-11A exhibits very high temperatures of 1015 °C and 975 °C for quartz–iotite and quartz–rutile pairs, respectively (Table 3). In particular, the δ18O of quartz is lower than that of garnet. These indicate O isotope disequilibrium between quartz and other minerals in the two samples. While the disequilibrium in eclogite 04A-2 is obviously caused by anomalously high quartz δ18O values, the disequilibrium in eclogite 05C-11A is due to anomalously low quartz δ18O values. This implies that the quartz in the two eclogites may be formed or recrystallized at a later stage, with incorporation of external fluid δ18O signatures into this mineral. As depicted in Fig. 6e, h and Table 3, O isotope geothermometries for the quartz–amphibole and quartz– epidote pairs from the amphibolite yield temperatures of 370 to 505 °C and 320 to 490 °C, respectively. Similarly, O isotope geothermometries for the quartz–epidote, quartz–muscovite, quartz–amphibole pairs from the gneiss give temperatures of 320 to 500 °C, 390 to 620 °C and 380 to 560 °C. These isotopic temperatures are generally compatible with O isotope equilibrium under amphibolite-facies conditions. However, the higher isotopic temperatures of 560 to 755 °C are obtained for the quartz–garnet pairs (Fig. 6b and Table 3). This indicates that the O isotope composition of garnet was not significantly affected by amphibolitefacies retrogression during exhumation. The O isotope temperatures derived from the quartz–plagioclase, quartz–K-feldspar and quartz–biotite pairs are often unreasonably higher than 800 °C or lower than 400 °C, indicating retrograde resetting without isotopic reequilibration by a later event after the amphibolite-facies metamorphism. 5. Discussion 5.1. Effect of retrograde alteration As shown in Fig. 4 and Table 1, the eclogite from the different depths exhibits different degrees of retrograde metamorphism. In some of the eclogite samples, most omphacites were retrograded to symplectite, and amphibole and biotite became abundant. The widespread occurrence of amphibolite in CCSD-MH indicates hydration of the UHP rocks during exhumation. In addition, a lot of felsic and calcite veins, which are the record of fluid flow along fractures, are observed in some of the samples. These indicate that the eclogite and

gneiss were significantly affected by retrograde fluid during exhumation. These are consistent with previous studies of petrography indicating that eclogite and gneiss from CCSD-MH not only underwent the UHP metamorphism but also suffered amphibolite-facies retrogression (e.g., Liu et al., 2002; Su et al., 2005; Zhang et al., 2006b). Therefore, there would be isotopic exchange between minerals and retrograde fluid during the retrogression, which may change not only the O isotope composition of minerals but also the O isotope fractionations between coexisting minerals. As depicted in Fig. 6g, c, j and Table 3, the O isotope temperatures for the quartz–omphacite, quartz–phengite and quartz–rutile pairs from the eclogite are 300 to 780 °C, 380 to 665 °C and 400 to 535 °C, respectively. Similarly, the quartz–epidote, quartz–muscovite, quartz–amphibole pairs from the gneiss and amphibolite yield temperatures of 320 to 500 °C, 380 to 620 °C, 370 to 560 °C, respectively (Fig. 6h, c, e and Table 3). These are differentially lower than temperatures of the eclogitefacies metamorphism, suggesting the effect of retrograde isotope exchange subsequent to the eclogite-facies metamorphism. The quartz–feldspar and quartz–biotite pairs give wide temperature ranges and some of them are too high or too low to represent metamorphic conditions, pointing to O isotope disequilibria between quartz, feldspar and biotite. This is also caused by retrograde alteration. On the other hand, the quartz–garnet pairs are at high-T isotopic equilibrium, yielding temperatures of 560 to 785 °C for the eclogite and gneiss, consistent with the eclogite-facies conditions. The variable isotopic temperatures for the different mineral pairs were caused by differential isotope exchange between the minerals and retrograde fluid during exhumation (e.g., Rumble and Yui, 1998; Zheng et al., 1998, 1999; Fu et al., 1999; Zheng et al., 2002; Zhang et al., 2003). The O isotope disequilibrium between omphacite and garnet is also found for some of the eclogites from CCSD-MH. As depicted in Fig. 12, the O isotope fractionations between omphacite and garnet in the eclogite can be classified into two groups: (1) negative and zero fractionations of − 1.7 to 0‰, which clearly point to isotopic disequilibrium and thus yield isotopic temperatures unreasonably lower than 600 °C when paring omphacite with quartz; (2) positive fractionations of 0.2 to 0.9‰ which are at O isotopic equilibrium. As shown by Fig. 4 and Table 1, symplectites and retrograde minerals such as amphibole and biotite occur in some of the eclogites. Thus, the O isotope disequilibrium between omphacite and garnet is obviously caused by retrograde alteration (Zheng et al., 2003).

R.-X. Chen et al. / Chemical Geology 242 (2007) 51–75

Fig. 12. O isotope fractionations between omphacite and garnet versus δ18O values of garnet from the CCSD main hole. Reference lines of ▵ = 0‰ and 1‰ place the approximate limit of equilibrium fractionations between omphacite and garnet, whereas the samples beyond the two lines are out of isotopic equilibrium.

Although the UHP eclogite and gneiss from CCSDMH suffered the amphibolite-facies retrogression, the O isotope composition of less resistant minerals can provide a useful constraint on the origin of retrograde fluid. O isotope geothermometers of quartz–mineral pairs (e.g., muscovite, epidote, magnetite and rutile) gave temperatures compatible with amphibolite-facies conditions (Fig. 11), suggesting that the retrograde minerals would have approached and preserved O isotope reequilibration under the amphibolite-facies conditions. This is also confirmed by the O isotope geothermometries for the amphibolite samples, some of them are the product of almost complete retrogression of eclogite. The quartz– amphibole and quartz–epidote pairs from the amphibolite also yielded temperatures compatible with amphibolitefacies conditions (Fig. 6e, h and Table 3). Garnet from the amphibolite shows similar δ18O values to those from the eclogite for the second core segment. Except for feldspar and mica in some samples, the retrograde and refractory minerals from the eclogite and gneiss from the five core segments show similar δ18O variations (Figs. 7, 8, 9, 10 and 11), falling within the δ18O range of their protoliths. Thus there is no significant O isotope shift during the retrograde metamorphism. This suggests that the retrograde fluid was either in, or very close to, isotopic equilibrium with the minerals from the UHP metamorphic rocks. Therefore, the retrograde fluid would be internally buffered in their stable isotope compositions (Zheng et al., 1999, 2003). Some of the quartz–feldspar and quartz–biotite pairs gave too high or too low temperatures to be compatible with known metamorphic conditions (Fig. 6a, f, i and Table 3). Because rates of O diffusion in feldspar and mica are much faster than other minerals (Zheng and Fu, 1998), the

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two minerals are susceptible to later resetting after the amphibolite-facies overprinting. While the other minerals (e.g. epidote, magnetite, rutile, and amphibole) achieved and preserved O isotope reequilibration acquired under the amphibolite-facies conditions, the feldspar and mica underwent O isotope exchange with fluids at a stage after amphibolite-facies metamorphism. Calcite vein occurs across amphibole in some samples (Fig. 3h). This indicates that the calcite vein was formed at a stage after amphibole. Thus it may be a product of lower amphibolite-facies or a later event after the amphibolite-facies retrogression. The calcite–biotite and calcite–hornblende pairs gave O isotope temperatures of 450 to 1145 °C and 505 to 1590 °C, respectively (Table 4). These temperatures are not compatible with the lower amphibolite-facies metamorphism. Therefore, a later fluid activity is more possible. As shown by Fig. 3 and Table 1, in addition, biotite in some samples is partially replaced by chlorite. The quartz–chlorite pairs gave O isotope temperatures of 100 to 180 °C (Table 3), which is significantly lower than those for amphibolite-facies conditions. This also suggests that water–rock interaction occurred between the UHP rocks and the fluid at a later stage than the amphibolite-facies retrogression rather than during the lower amphibolitefacies retrogression. Therefore, fluid activity did occur within the UHP rocks after the amphibolite-facies retrogression. Postcollisional fluid flow after the Triassic orogenic cycle of continental collision has been observed in the Dabie orogen (Gao et al., 2006; Zheng et al., 2007a). Nevertheless, this later water–rock interaction may only affect the O isotope compositions of feldspar and mica. For the other minerals, when paired with quartz, they all gave temperatures compatible with eclogite-facies and/or amphibolite-facies metamorphism (Fig. 11 and Table 3). This difference is attributable to differential exchange of O isotopes in these minerals with fluid during exhumation because the rates of O isotope diffusion in feldspar and mica are much faster than the other minerals (Zheng and Fu, 1998). Table 4 Oxygen isotope temperature for calcite–mineral pairs from the CCSDMH (in °C) ⁎ Sample No.

Rock type

Depth (m)

Cc–Bt

Cc–Hb

05A-4A 05C-6A 05D-6A 05D-7A 05D-8C

Amphibolite Amphibolite Amphibolite Amphibolite Amphibolite

3772.45 3297.85 3585.86 3586.07 3586.42

690 450 930 725 1145

1390 640 530 505 1590

⁎ Temperature calculations are based on the calibrations of Zheng (1993b).

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The transitions between the different lithologies from CCSD-MH provide us a chance to test the effect of amphibolite-facies retrogression on mineral O isotope composition. For the first core segment 04C, the anomalously low δ18O value occurs in muscovite from the amphibolite (Fig. 7), resulting in a large O isotope shift relative to the adjacent eclogite and gneiss. For the second core segment 04B, rutile from the amphibolite has lower δ18O values than that from the adjacent eclogite (Fig. 8). Similarly, amphibole from the amphibolite has lower δ18O values than omphacite and garnet from the adjacent eclogite. For the fourth core segment 05C, amphibole and plagioclase from the amphibolite have lower δ18O values than garnet from the adjacent eclogite (Fig. 10). For the fifth core segment 05D, two amphibolites 05D-6A and 05D-7A close to the felsic vein have significantly lower δ18O values for amphibole than those away from the felsic vein (Fig. 11). All of these observations suggest that aqueous fluids for the amphibolite-facies retrogression were depleted in 18O relative to the precursor minerals. A lot of felsic and calcite veins were observed in the eclogite and gneiss from CCSD-MH (Fig. 3 and Table 1). However, no obvious correlation was observed between the veins with changes in O isotope compositions. The internally buffered stable isotope compositions suggest that the veins were formed by fluid derived locally from either the UHP rocks themselves or external sources which are either in or very close to O isotopic equilibrium with the host eclogite and gneiss. The formation of vein is obviously caused by channelized flow of the retrograde fluid. In this way of flow, the fluid is a focus of H2O and thus can move to a distant place along the flow path (e.g., fracture). Channelized flow produces zones in which the fluid/rock ratio is much higher. The fluid interacts with less rock per unit volume and retains more of its initial source composition. Therefore, if the fluid is locally derived, it will not result in the O isotope shift. If the amount of the external fluid is very small, it will not significantly change the O isotope composition of the host rocks. For the first and fourth core segments 04C and 05C, the amphibolite layer separates the eclogite from the gneiss. The O isotope compositions of the amphibolite are quite different from those for the gneiss and eclogite, indicating significant δ18O variations across the different lithologies. For the fifth core segment 05D, the felsic vein occurs between the gneiss and the amphibolite, suggesting fluid flow between the different lithologies. For the third core segment 05B, the eclogite adjacent to the schist appears to have larger extent of retrogression than that away from the contact between eclogite and

schist. All these correlations between O isotope and petrographic changes indicate that fluid flow and mineral reaction did take place at the contacts between the different lithologies. Such contacts are probably the most favorable place for fluid activity. With respect to the origin of the retrograde fluid, Zheng et al. (2003) advocate that it was derived from the decomposition of hydrous minerals, the decrepitation of primary fluid inclusions, and the exsolution of structural hydroxyls from nominally anhydrous minerals. In particular, the exsolution of structural hydroxyls has gained more and more supports from the measurement of hydroxyl contents in nominally hydrous minerals such as omphacite, garnet, rutile and jadeite (e.g., Zhang et al., 2001; Su et al., 2002; Katayama and Nakashima, 2003; Zhang et al., 2004; Su et al., 2004; Xia et al., 2005; Katayama et al., 2006; Chen et al., 2007; Sheng et al., 2007; Zhao et al., 2007). Furthermore, quantitative measurements of the concentration of total water by means of TC/EA–MS online method (Chen et al., 2007; Gong et al., 2007a,b) show not only the presence of molecular water in nominally anhydrous minerals from UHP metamorphic rocks, but also higher concentrations of water in gneissic minerals than in eclogitic minerals. Accumulation of structural hydroxyl and molecular water liberated from the nominally anhydrous minerals can form an important source of the low δ18O retrograde fluid, resulting in the disequilibrium O isotope fractionations between omphacite and garnet in some of the eclogites and the lower O isotopic temperatures for the retrograde minerals when paired with quartz. In addition, external fluids could be a possible origin of the retrograde fluid. It has been found at some surface outcrops that the eclogite adjacent to the gneiss shows high degrees of retrograde metamorphism as indicated by more abundant symplectites or even conversion to amphibolite, where the contacts between interlayers of amphibolite/gneiss and eclogite show gradations in mineralogy (Zhang et al., 2003). This suggests that the retrograde reactions of eclogite to amphibolite may take place during exhumation due to the infiltration of external fluids along fault zones or lithological layers. The fluid flow in this case is channelized one. The fluid is also focused water, so that it can bring external fluids into the rock to cause the O isotope shift. This may explain a very low δ18O value for muscovite from amphibolite 04C-5A (Fig. 7). Similarly, the fluid forming the felsic vein between the gneiss and amphibolite from the fifth core segment 05D is probably of external origin. This is suggested by the O isotope change between the amphibolite adjacent to and far away from the vein (Fig. 11). Although the retrograde fluid is external to the

R.-X. Chen et al. / Chemical Geology 242 (2007) 51–75

rock itself, it still acted within the exhumed slab and may come from the decompression exsolution of structural hydroxyl and molecular water from adjacent rock units. 5.2. Scales of fluid activity during continental collision The profile analysis of mineral O isotopes in the CCSD-MH core samples provides us a great advantage to test whether the O isotope heterogeneity occurs on small scales down to centimeters that are not able to be observed from the surface outcrops. In particular, mineral δ18O differences between adjacent samples can be correlated with their distance and petrography. The results can be used to place quantitative constraints on the scale of fluid mobility during the continental collision. For the first core segments, differences in garnet and magnetite δ18O values between adjacent samples occur on scale of 22 to 49 cm (Fig. 7). For the second core segments, differences in amphibole δ18O values between adjacent samples occur on scale of 20 to 51 cm (Fig. 8). For the third core segments, differences in garnet δ18O values between adjacent samples occur on a scale of 31 cm (Fig. 9). For the fourth core segments, differences in biotite δ18O values between adjacent samples occur on scales of 15 to 27 cm (Fig. 10). For the fifth core segments, differences in amphibole, biotite and calcite δ18O values between adjacent samples occur on scales of 14 to 21 cm (Fig. 11). Although homogeneous δ18O values occur locally on different scales up to 150 cm, the abrupt changes in δ18O are evident on scales of about 100 m along the vertical O isotope profile of 200 to 4000 m depths (Fig. 5). In particular, the O isotope heterogeneity is significant on scales of about 20 to 50 cm between the different lithologies and even within the same lithologies (Figs. 7–11). As discussed in the next section, the δ18O distribution in the eclogite and gneiss from CCSD-MH is primarily dictated by their protolith δ18O values. Thus the occurrence of δ18O heterogeneities on different scales indicate that the fluid mobility is very limited during the continental collision, depending on protolith lithologies and their contacts. They are generally smaller than the scales of 20 to 50 cm at thermodynamic equilibrium. This provides for the first time a quantitative estimate of fluid flow scales during the continental collision. It is the small scales of fluid activity that did not efficiently redistribute the O isotopes either within the same lithologies or among the different lithologies (e.g., Rumble et al., 2000; Zheng et al., 2003). The previous discussions have concluded that the metamorphic fluid affecting the eclogite and gneiss from the five CCSD-MH core segments is internally buffered

71

in stable isotope compositions. In particular, the retrograde fluid is depleted in 18O relative to the precursor minerals. Thus, the retrograde fluid in the UHP rocks is deuteric and thus principally derived from the decompression exsolution of structural hydroxyl. The retrograde fluids were directly derived from the UHP rocks themselves. Along the contacts between different lithologies and fractures in the interior of the same lithologies, fluid flow during the continental collision would basically occur in a channelized way. Even so, the scale of fluid flow is very limited, probably less than 20 to 50 cm provided that the physical flow rates of retrograde fluid resemble rates of chemical equilibrium towards O isotopic homogeneity. In this regard, accurate estimates of fluid flow scale involve the account of difference in rate between fluid flow and chemical equilibrium during fluid−rock interaction with respect to geochemical thermodynamics and kinetics. If the fluid filtration is faster than the isotopic homogenization, kinetic fractionation is predominant to result in significant δ18O gradients on small scales. In contrast, if the fluid filtration is slower than the isotopic homogenization, no isotopic gradient is expected to occur between adjacent lithologies. Because the metamorphic fluid is comprised of molecular water, it can intimately interact with the rock matrix. This leads to a strong potential for continuous reequilibration with the matrix materials, so that the rock matrix also exerts a control on the composition of metamorphic fluid. Consequently, the resultant isotope compositions are internally buffered and thus principally depend on the O isotopic composition of premetamorphic protoliths. 5.3. Inheritance of protolith compositions Metamorphic rocks generally have δ18O values of 6 to 18‰ (Hoefs, 2004), depending on the geochemical nature of protoliths (i.e. igneous or sedimentary origin). The UHP eclogites and associated gneisses from the surface outcrops along the Dabie–Sulu orogenic belt show a large variation in δ18O from − 11 to 10‰, with both equilibrium and disequilibrium fractionations of O isotopes between coexisting minerals (Zheng et al., 2003, and references therein). Preservation of the O isotope equilibrium fractionations among the minerals of the UHP eclogites and gneisses indicates that the UHP rocks acquired their unusually negative δ18O values down to − 11‰ by high-T meteoric–hydrothermal alteration (Rumble and Yui, 1998; Zheng et al., 1998, 1999, 2003). A combined study of zircon U–Pb dating and O isotope analysis demonstrates that the meteoric–hydrothermal circulation and local low δ18O

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magmatism occurred at middle Neoproterozoic (Rumble et al., 2002; Zheng et al., 2004), long time before the Triassic continental collision. The eclogite, amphibolite and gneiss in the selected core samples from CCSD-MH also show variable 18O depletion relative to the normal mantle-derived rocks (Table 2 and Fig. 5). O isotope equilibrium fractionations between quartz and garnet have been preserved, yielding reasonable temperature estimates responsible for eclogite-facies metamorphic conditions (Fig. 6b and Table 3). Therefore, the 18O depletion in the eclogite, amphibolite and gneiss from CCSD-MH is also ascribable to the high-T meteoric– hydrothermal alteration and low δ18O magmatism of their protolith before the continental subduction. Previous studies demonstrate that the different UHP lithologies from CCSD-MH were subjected to in situ UHP metamorphism (Xu, 2004; Liu et al., 2004a, 2004b; Zhang et al., 2006b). This study shows a significant variation in mineral δ18O values from CCSD-MH of 200 to 4000 m (Fig. 5). In particular, variable 18O depletions relative to the normal mantle occur at depths down to 3300 m, regardless of lithology. However, it does not mean that the ancient meteoric water–rock interaction would have penetrated into a depth of at least 3300 m because the continental collision (subduction and exhumation) has reset the original depths of premetamorphic lithologies. In other words, this depth cannot be considered as the minimum depth of 18O depletion because the UHP rocks from CCSD-MH are also part of subducted and exhumed slabs, with significant changes in horizontal and vertical directions. Thus it has nothing to do with original depths of their protoliths during the Neoproterozoic meteoric–hydrothermal alteration and low δ18O magma emplacement. Therefore, the “in situ” origin of the core sequence does not allow us to estimate depths of meteoric water penetration. Nevertheless, at least 3 to 4 km depths are required to bring about the 18O depletion of mafic and felsic igneous protoliths if no low 18 O magmas were involved. With respect to anatexis of hydrothermally altered low δ18O rocks in active rift zones, on the other hand, at least 5 km depth is necessary for generation of granitic melts under conditions of either H2O excess or thermal anomaly. In this regard, infiltration of meteoric water may be as deep as 8 to 9 km. Previous studies on surface samples demonstrate that the water–rock interaction in the Dabie–Sulu orogenic belt at the Neoproterozoic occurred in an area of over 20,000 km2 (Zheng et al., 2004). Considering 3300 m as the depth of 18O depletion, we can estimate that more than 66,000 km3 of supracrustal rocks were interacted with meteoric water during the Neoproterozoic. This implies

that a large volume of very 18O depleted meteoric water was involved in the protoliths of UHP metamorphic rocks. Furthermore, the lowest δ18O values occur at the site close to the thick orthogneiss layers (Fig. 5). Zircon U–Pb dating for the orthogneiss layers yields protolith ages of about 750 Ma (Liu et al., 2004b), consistent with the time of magma emplacement and meteoric water–rock interaction along the northern margin of the South China Block (Zheng et al., 2004, 2007b). Therefore, the low 18O data from the CCSD-MH core samples lend support to the conclusion that the heating engines driving the hydrothermal circulation of meteoric waters were continental rift magmatism in association with the Rodinia breakup and the Neoproterozoic deglaciation (Rumble et al., 2002; Zheng et al., 2004, 2007b). The vertical O isotope profile of 200 to 4000 m depths shows a large δ18O variation from − 3.5 to 8.4‰ for estimated whole-rock (Fig. 5), with abrupt changes in δ18O on scales of about 100 m. These indicate O isotope heterogeneities at different scales of depth. Differences in protolith δ18O values are considered as the primary cause for the observed δ18O heterogeneity. It appears that the Neoproterozoic water–rock interactions were heterogeneous in different lithologies of protolith, probably also limited to the scales of 20 to 50 cm. 6. Conclusions Mineral separates in UHP metamorphic rocks from CCSD-MH at depths of 200–4000 m show considerable changes in O isotope ratios. Eclogite has δ18O values of − 3.32 to 7.52‰ for garnet and − 2.55 to 5.84‰ for omphacite, amphibolite has δ18O values of − 4.53 to 5.23‰ for hornblende and − 3.26 to 6.75‰ for plagioclase, and gneiss has δ18O values of − 2.69 to 9.63‰ for quartz and −0.74 to 8.11‰ for plagioclase. Both equilibrium and disequilibrium O isotope fractionations occur between coexisting minerals, indicating the differential effects of retrograde alteration on UHP mineral assemblages. Considerable changes in mineral δ18O values occur at the contact between eclogite and gneiss, in concordance with the petrographic observations that the eclogite adjacent to the gneiss shows larger extent of retrogression and even occurrence of amphibolite. It appears that the contact between different lithologies is the most favorable place for fluid activity during continental collision. Amphibolite-facies retrogression during exhumation results in mineral reactions and O isotope disequilibria between some of the minerals, but retrograde fluid was internally buffered in the O isotope

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compositions, thus probably derived from the decompression exsolution of structural hydroxyl. Although locally external fluids can be a possible origin for retrograde fluid along fault zones or lithological layers, they still acted within the exhumed slab and thus may also be derived from the decompression exsolution of structural hydroxyl from adjacent rock units. Fluid flow also occurred after the amphibolite-facies retrogression, but it only influenced the O isotope compositions of feldspar and mica that are susceptible to retrograde O isotope exchange. Different scales of δ18O heterogeneity occur in the eclogite and gneiss from the vertical CCSD-MH profile in the depth range of 200 to 4000 m, including the five continuous core segments. These are ascribed to the primary δ18O distributions in their protoliths. Mineral O isotopic equilibrium was achieved on different scales, mostly on scales of millimeter to centimeter. While O isotopic heterogeneities occur among different lithologies and even within the same lithologies, abrupt changes in δ18O occur on a scale of about 100 m. A quantitative estimate of fluid activity scales was made by account of the relationship between distance, petrography and δ18 O value of adjacent samples. Homogeneous δ 18 O values are only observed on different scales less than 1.5 m, and heterogeneous δ18O distributions are limited on scales of about 20 to 50 cm. While heterogeneous δ18O distributions along the 200–4000 m depth profile are principally inherited from their protolith, microscale δ18O changes indicate fluid activity on the maximum scales of 20 to 50 cm during the continental collision. Preservation of O isotope equilibrium at eclogitefacies temperatures suggests that eclogite and gneiss protoliths acquired their low δ18O values by high-T meteoric water–rock interaction before the subduction. The minimum depth of 18O depletion is up to 3300 m for the CCSD-MH rocks. Together with an areal distribution of 18O depletion over 20,000 km2 on the surface outcrops, it is estimated that at least 66,000 km3 of supracrustal rocks were interacted with the meteoric water during the Neoproterozoic rift magmatism. Acknowledgments This study has been supported by funds from the National Ministry of Science and Technology (2003CB716501) and the Natural Science Foundation of China (40573011). Thanks are due to Benchao Ding and Xiangping Zha for their assistance with oxygen isotope analysis. Comments by Dr. D. Rickard and two anonymous reviewers help improve the presentation.

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