Metamorphic growth and recrystallization of zircons in extremely 18O-depleted rocks during eclogite-facies metamorphism: Evidence from U–Pb ages, trace elements, and O–Hf isotopes

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Geochimica et Cosmochimica Acta 75 (2011) 4877–4898 www.elsevier.com/locate/gca

Metamorphic growth and recrystallization of zircons in extremely 18 O-depleted rocks during eclogite-facies metamorphism: Evidence from U–Pb ages, trace elements, and O–Hf isotopes Yi-Xiang Chen a, Yong-Fei Zheng a,⇑, Ren-Xu Chen a, Shao-Bing Zhang a, Qiuli Li b, Mengning Dai c, Lu Chen d a

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 b State Key Laboratory of Lithospheric Evolution, Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing 100029, China c State Key Laboratory of Continental Dynamics, Department of Geology, Northwest University, Xi’an 710069, China d State Key Laboratory of Geological Processes and Mineral Resources, China University of Geosciences, Wuhan 430076, China Received 20 September 2010; accepted in revised form 1 June 2011; available online 16 June 2011

Abstract A combined in situ SIMS and LA-(MC)-ICPMS study of U–Pb ages, trace elements, O and Lu–Hf isotopes was conducted for zircon from eclogite-facies metamorphic rocks in the Sulu orogen. The two microbeam techniques sampled various depths of zircon domains, revealing different element and isotope relationships between residual magmatic cores and new metamorphic rims and thus the geochemical architecture of metamorphic zircons which otherwise cannot be recognized by the single microbeam technique. This enables discrimination of metamorphic growth from different subtypes of metamorphic recrystallization. Magmatic cores with U–Pb ages of 769 ± 9 Ma have positive d18O values of 0.1–10.1&, high Th/U and 176Lu/177Hf ratios, high REE contents, and steep MREE-HREE patterns with negative Eu anomalies. They are interpreted as crystallizing from positive d18O magmas during protolith emplacement. In contrast, newly grown domains have concordant U–Pb ages of 204 ± 4 to 252 ± 7 Ma and show negative d18O values of 10.0& to 2.2&, low Th/U and 176Lu/177Hf ratios, low REE contents, and flat HREE patterns with weak to no Eu anomalies. They are interpreted as growing from negative d18O fluids that were produced by metamorphic dehydration of high-T glacial-hydrothermally altered rocks during continental subductionzone metamorphism. Differences in d18O between different domains within single grains vary from 0.8& to 12.5&, suggesting different degrees of O isotope exchange between the positive d18O magmatic core and the negative d18O metamorphic fluid during the metamorphism. The magmatic zircons underwent three subtypes of metamorphic recrystallization, depending on their accessibility to negative d18O fluids. The zircons recrystallized in solid-state maintained positive d18O values, and REE and Lu–Hf isotopes of protolith zircon, but their U–Pb ages are lowered. The zircons recrystallized through dissolution exhibit negative d18O values similar to the metamorphic growths, almost completely reset U–Pb ages, and partially reset REE systems. The zircons recrystallized through replacement show variably negative d18O values, and partially reset REE, and U–Pb and Lu–Hf isotopic systems. Therefore, this study places robust constraints on the origin of metamorphic zircons in eclogite-facies rocks and provides a methodological framework for linking the different types of metamorphic zircons to petrological processes during continental collision. Ó 2011 Elsevier Ltd. All rights reserved.

⇑ Corresponding author. Tel./fax: +86 551 3603554.

E-mail address: [email protected] (Y.-F. Zheng). 0016-7037/$ - see front matter Ó 2011 Elsevier Ltd. All rights reserved. doi:10.1016/j.gca.2011.06.003

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1. INTRODUCTION Zircon is one of the widely used accessory minerals in dating metamorphic events, though it shows complex mineralogical and geochemical behaviors in high-grade metamorphic rocks (e.g., Geisler et al., 2007; Rubatto and Hermann, 2007; Xia et al., 2009; Chen et al., 2010). On one hand, zircon can grow by metamorphic reactions of protolith minerals under subsolidus conditions (Fraser et al., 1997; Pan, 1997; Bingen et al., 2001; Degeling et al., 2001), or crystallize from aqueous fluids (e.g., Liati and Gebauer, 1999; Rubatto and Hermann, 2003; Zheng et al., 2007a; Wu et al., 2009a; Xia et al., 2009; Chen et al., 2010) or hydrous melts (Vavra et al., 1996; Roberts and Finger, 1997; Keay et al., 2001; Rubatto, 2002; Rubatto et al., 2009; Xia et al., 2009; Liu et al., 2010). On the other hand, protolith zircons can be modified by metamorphic recrystallization to approach different extents of thermodynamic reequilibration, depending on the crystallinity of the zircons and their accessibility to metamorphic fluid/melt (e.g., Geisler et al., 2007; Rubatto and Hermann, 2007; Martin et al., 2008; Rubatto et al., 2008; Xia et al., 2009; Chen et al., 2010; Xia et al., 2010). The metamorphic recrystallization of protolith zircons can cause the fractionation of trace elements and the resetting of U–Pb and Lu–Hf isotope systems (e.g., Zheng et al., 2005; Wu et al., 2006; Zheng et al., 2006; Martin et al., 2008; Wu et al., 2008; Gerdes and Zeh, 2009; Wu et al., 2009b; Zeh et al., 2010). This process can proceed via the mineralogical mechanism of solid-state transformation, replacement alteration and/or dissolution reprecipitation (e.g., Hoskin and Black; 2000; Rubatto et al., 2008; Xia et al., 2009; Chen et al., 2010; Xia et al., 2010). However, the effect of this process on zircon O isotopes has rarely been studied (Valley et al., 1994; Booth et al., 2005; Martin et al., 2008) and thus poorly constrained relative to the U– Pb and Lu–Hf isotopes (Zheng et al., 2004, 2005, 2006; Xia et al., 2009; Chen et al., 2010; Xia et al., 2010). The rate of O diffusion in the crystalline zircon is very slow under anhydrous conditions, but it becomes much faster under hydrous conditions (Watson and Cherniak, 1997; Zheng and Fu, 1998). As a consequence, the solid-state recrystallization under eclogite- and granulite-facies conditions does not alter the O isotope signature of protolith magmatic zircons (Peck et al., 2003; Zheng et al., 2004; Page et al., 2007). Nor are their U–Pb and Lu–Hf isotope systems in the absence of fluid accessibility (e.g., Zheng et al., 2005, 2006; Chen et al., 2007a; Tang et al., 2008a; Xia et al., 2009; Chen et al., 2010; Xia et al., 2010). In contrast, zircon O isotopes would be partially or completely reset upon action of metamorphic fluid (Valley et al., 1994; Zheng et al., 2004; Booth et al., 2005; Martin et al., 2008), and so would zircon U–Pb and Lu–Hf isotopes (e.g., Zheng et al., 2005, 2006; Chen et al., 2007a; Gerdes and Zeh, 2009; Xia et al., 2009; Chen et al., 2010; Xia et al., 2010; Zeh et al., 2010). The metamorphic recrystallization can result in coupled variations between zircon U–Pb and O isotope systems (Booth et al., 2005; Martin et al., 2008), and between U–Pb and Lu–Hf isotope systems (Zheng et al., 2005; Wu et al., 2006; Zheng et al., 2006; Chen et al., 2007a; Gerdes and Zeh, 2009; Xia

et al., 2009; Chen et al., 2010; Xia et al., 2010; Zeh et al., 2010). The trace element distribution of metamorphic zircons is a powerful means to identify their genesis and thus vital for interpretation of their U–Pb ages for high-grade metamorphic rocks (e.g., Hoskin and Black; 2000; Rubatto, 2002; Whitehouse and Platt, 2003). The solid-state recrystallization of magmatic zircons is readily identified by the preservation of steep HREE patterns with negative Eu anomalies, whereas the metamorphic growth in the presence of garnet effect is characterized by flat HREE patterns without Eu anomalies (e.g., Rubatto and Hermann, 2007; Zheng et al., 2011). However, it becomes complicated for discrimination between replacement and dissolution recrystallized zircons because they sometimes show mixed REE patterns between magmatic protolith inheritance and metamorphic new growth (Xia et al., 2009; Chen et al., 2010; Xia et al., 2010). In this case, zircon O isotopes are expected to provide an additional dimension to the discrimination because the O isotopes are sensitive to fluid–mineral interaction and thus record the different processes of zircon growth and recrystallization. This requires a direct linkage of zircon O isotopes to the U–Pb and Lu–Hf isotopes as well as trace elements in the same property of domains. Zircons from ultrahigh-pressure (UHP) eclogite-facies metamorphic rocks in the Dabie–Sulu orogenic belt show variable degrees of 18O depletion, with negative d18O values as low as 10& to 6& (Rumble et al., 2002; Zheng et al., 2004; Tang et al., 2008a). Preliminary SIMS analyses revealed large core-rim d18O variations in some zircons but no statistical difference was observed between them (Chen et al., 2003). In contrast, laser fluorination analyses did not yield significant d18O differences between core and rim (Rumble et al., 2002; Zheng et al., 2004). This discrepancy probably results from bulk grain laser analyses with large sample sizes (usually 1.5–2.5 mg), which cannot resolve the intragrain d18O variations. Some negative d18O zircons were dated by U–Pb methods, yielding ages of 700–800 Ma (Rumble et al., 2002; Zheng et al., 2004; Tang et al., 2008a). Thus, these zircons were interpreted as crystallization in the Neoproterozoic from negative d18O magmas that were derived from remelting of high-T glacial-hydrothermally altered igneous rocks in a rifting zone. Then the negative d18O rocks would be transformed to the UHP eclogite-facies rocks during the Triassic continental subduction (Zheng et al., 2003, 2009). However, there are positive d18O values of 0.5–2.7& for magmatic zircons with U–Pb ages of 750–780 Ma but negative d18O values of 14.4& to 4.5& for coexisting rock-forming minerals from Neoproterozoic granites in the Dabie orogen (Zheng et al., 2007b). In particular, the negative d18O values as low as 14.4& to 10.0& for garnet from the granites provide a geochemical proxy for high-T hydrothermal alteration by glacial meltwater at 750 Ma. This finding has challenged the previous consensus about the origin of negative d18O zircons, implying that the negative d18O zircons can be crystallized either from negative d18O magmas in the Neoproterozoic or from negative d18O fluids during the Triassic subduction-zone metamorphism. Therefore, it remains to be resolved how the metamorphic zircons in the UHP

Metamorphic growth and recrystallization of zircons in negative d18O rocks

metamorphic rocks acquired their negative d18O values and whether different zircon domains can record the different mineralogical processes between protolith emplacement and subduction-zone metamorphism. The occurrence of negative d18O zircons in regional metamorphic rocks is a wonder on Earth. To understand their origin has important implications for the chemical geodynamics of continental subduction zones. Thus, it is intriguing to determine whether the negative d18O zircons crystallized from negative d18O magmas before subduction or grew from negative d18O metamorphic fluids during subduction. In the latter case, it would be linked to metamorphic dehydration of glacial-hydrothermally altered igneous rocks having negative d18O values. This thus requires discrimination between the zircons of dissolution recrystallization and metamorphic growth. The advances in the SIMS zircon O and U–Pb isotope analyses as well as the LA-(MC)-ICPMS analyses of trace elements, U–Pb and Lu–Hf isotopes provide us with an excellent opportunity to resolve these issues. This paper presents a combined study by these two microbeam techniques on metamorphic zircons from UHP eclogite and granitic gneiss in the Sulu orogen. The results allow direct recognition of geochemical layering in the zircons of core-rim structure. 2. GEOLOGICAL SETTING AND SAMPLES The Dabie–Sulu orogenic belt formed by the Triassic subduction of the South China Block beneath the North China Block (Cong, 1996; Zheng et al., 2003; Ernst et al., 2007). The Sulu orogen is considered as its eastern part, with an offset of 500 km to the northeast due to movement along the Tan-Lu fault. This orogen is bounded by the Jiashan-Xiangshui fault to the south and the Yantai-Qingdao-Wulian fault to the north (inset in Fig. 1), and segmented into a number of slices by several NE–SW-trending faults subparallel to the Tan-Lu fault (Xu et al., 2006; Tang et al., 2008b). According to available petrological and geochemical data, the Sulu orogen can be divided into high pressure (HP) and UHP metamorphic zones, both are unconformably overlain by Jurassic clastic strata and Cretaceous volcanoclastic cover, and intruded by Mesozoic granites (Zhang et al., 1995; Liu et al., 2004; Zhang et al., 2010). The HP zone consists mainly of quartz–mica schist, kyanite–mica–quartz schist, paragneiss, orthogneiss, marble and rare blueschist. The UHP zone consists mainly of amphibolite-facies orthogneiss and paragneiss, with minor garnet peridotite, eclogite, kyanite–topaz-bearing quartzite and marble. Eclogite mainly occurs as blocks or lenses in granitic gneiss, with some being enclosed by marble and garnet peridotite. The ultramafic bodies, meter- to kilometer in size, occur sporadically as layers or blocks throughout the UHP zone. Coesite and its pseudomorph were identified as inclusions in major minerals from eclogites (e.g., Liou and Zhang, 1996; Zhang et al., 2005) and in zircons from the country rocks (e.g., Liu et al., 2004, 2006, 2009, 2010). These observations demonstrate the eclogites and their country rocks together experienced UHP eclogite-facies metamorphism.

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The UHP rocks in the Sulu orogen underwent peak metamorphism at 3.0–4.5 GPa and 700–850 °C (e.g., Zhang et al., 1995, 2005, 2008). The UHP metamorphism occurred at 240 to 225 Ma, with a duration of 15 ± 2 Ma (Zheng et al., 2009). The majority of the UHP rocks have igneous protoliths of bimodal composition, formed probably in a continental rift zone in the middle Neoproterozoic (Zheng, 2008). Negative d18O values, as low as 10& to 6& (Yui et al., 1995; Zheng et al., 1996; Yui et al., 1997; Rumble and Yui, 1998; Zheng et al., 1998), were reported for UHP eclogite and quartz schist at Qinglongshan in the Donghai area (Fig. 1). The negative d18O values have been interpreted to indicate that premetamorphic protoliths were strongly affected by 18O-depleted meteoric water at high temperatures. TIMS and SIMS U–Pb dating for negative d18O zircons gave ages of 700–800 Ma (Rumble et al., 2002; Zheng et al., 2004), suggesting that the negative d18O signature was acquired in the middle Neoproterozoic. The four samples used in this study were taken from Qinglongshan, a hill located 12 km east of Donghai county (Fig. 1). Five rock types, including eclogite, orthogneiss, paragneiss, schist and quartzite, are exposed in this small area (Zhang et al., 2005). Two granitic gneisses (99QL07 and 99QL16), 350 m apart, were sampled along an E–W trending highway roadcut that traverses the southern side of Qinglongshan. Another granitic gneiss (00QL27) was sampled in the southern hillside of Qinglongshan. One eclogite (00QL16) was sampled from an outcrop at the top of Qinglongshan (Fig. 2 of Zheng et al., 1998). Granitic gneiss 99QL07 consists of K-feldspar, plagioclase, quartz, muscovite, and biotite, with minor garnet, titanite and zircon. Garnet, commonly associated with biotite and muscovite, is mostly skeletal, and only a few grains show subhedral morphology. Biotite and muscovite define the gneissosity as seen in thin section and hand specimen. Granitic gneiss 99QL16 consists of K-feldspar, plagioclase, quartz, muscovite, and biotite, with minor allanite, epidote, zircon and a little garnet. Allanite, with significant zoning and rimmed by epidote, along micas, defines the gneissosity of the sample. Granitic gneiss 00QL27 consists of K-feldspar, plagioclase, quartz, muscovite, and biotite, with minor epidote, garnet and zircon. Garnet in this sample shows a subhedral morphology. Eclogite 00QL16 consists of garnet, omphacite, phengite, epidote, quartz, kyanite, and rutile. Polycrystalline quartz after coesite with radial cracks was found in kyanite, epidote, and omphacite in this eclogite sample and the other eclogites from the same outcrop (Zhang et al., 1995, 2005). 3. ANALYTICAL METHODS Zircons were carefully studied for external morphology and internal structure with the binocular microscope and by cathodoluminescence (CL) imaging. Zircon O and U–Pb isotopes were firstly measured by CAMECA IMS1280 ion microprobe at State Key Laboratory of Lithospheric evolution in Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing. Then the same mounts were simultaneously measured for U–Pb isotopes and trace elements as well as Lu–Hf isotopes by the combined

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Fig. 1. Simplified geological map in the Donghai area, southwestern Sulu orogen (modified after Liu et al., 2004). Lithological units: (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- and biotite-bearing granitic gneiss; (10) epidote- and biotite-bearing granitic gneiss; (11) metamorphic rocks of supracrustal origin, including paragneiss, kyanite- and jadeite-bearing quartzite, and marble; (12) eclogite and ultramafic rocks; (13) ductile shear zone or fault; (14) sample site. Abbreviations: CCSD-PP1, CCSDPP2, and CCSD-MH are the Chinese Continental Scientific Drilling (CCSD) pre-pilot holes 1 and 2 and main hole, respectively; NCB, North China Block; SCB, South China Block; WQYF, Wulian-Qingdao-Yantai fault; JXF, Jiashan-Xiangshui fault.

LA-(MC)-ICPMS technique at State Key Laboratory of Continental Dynamics in Northwest University, Xi’an or only for U–Pb isotopes and trace elements by the LA-ICPMS method at State Key Laboratory of Geological Processes and Mineral Resources in China University of Geosciences, Wuhan. Details of the analytical procedures and instruments are presented online in Appendix I. 4. RESULTS Zircon U–Pb isotopes were analyzed for the eclogite and granitic gneisses by the both SIMS and LA-ICPMS meth-

ods, and the original data are listed in electronic Tables A1 and A2. Table A3 lists SIMS analytical data of zircon O isotopes, Table A4 lists LA-ICPMS analytical data of zircon trace elements, and Table A5 lists LA-MC-ICPMS analytical data of zircon Lu–Hf isotopes. All zircons are described and interpreted following the conventions of Corfu et al. (2003), Wu and Zheng (2004) and Zheng et al. (2007a). In doing so, a term of mineragraphy is used to outline the external morphology and internal structure of accessory minerals. These can be observed under the microscope and CL images, leading to possible distinction between different properties of domains with the core–

Metamorphic growth and recrystallization of zircons in negative d18O rocks

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Fig. 2. Representative zircon CL images for eclogite and granitic gneisses from Qinglongshan in the Sulu orogen. Red and yellow circles denote the SIMS spots with both O and U–Pb isotope data and with only O isotope data, respectively. Numbers in yellow and red denote the d18O value (in &) and the 206Pb/238U age (in Ma), respectively. The scale bar is 100 lm. Note that zircons from gneiss 99QL07 and 00QL27 have not only relict cores with middle Neoproterozoic U–Pb ages and positive d18O values but also mantle/rim domains with Triassic U–Pb ages and negative d18O values. In contrast, zircon grains from gneiss 99QL16 have not only patchy-zoned cores of pre-Triassic U–Pb ages and variably negative d18O values but also homogeneous rims of Triassic U–Pb ages and negative d18O values. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

mantle–rim structure. Furthermore, the residual core is defined for domains having relative blurring of oscillatory zones in CL images, and the mantle or rim is defined for metamorphic domains surrounding the cores. A detailed description of individual samples for zircon mineragraphy, U–Pb ages, O isotopes, REE and Lu–Hf isotopes is presented in Appendix II and illustrated in Figs. 2–8. Table 1 presents a summary of zircon U–Pb ages, O–Hf isotopes and trace elements analyzed by the both SIMS and LA(MC)-ICPMS methods for the eclogite and granitic gneisses. Metamorphic growths occur as either single grains or mantle/rims, so that the distinction between the domains of magmatic and metamorphic origins is made in terms of the mineragraphical and geochemical parameters. The results indicate that the metamorphic rims are thin domains

surrounding the magmatic cores that underwent different degrees of metamorphic recrystallization. The genetic classification of metamorphic zircons follows the generalization of Xia et al. (2009) and Chen et al. (2010) based on zircon mineragraphy, U–Pb ages, REE contents and patterns, and Lu–Hf isotopes. This includes three subtypes of metamorphic recrystallization (solid-state, replacement and dissolution) and two types of metamorphic growth (from aqueous fluid and hydrous melt). Fig. 2 presents some representative CL images for a few zircons with the notations of d18O value and apparent 206Pb/238U age. Four types of REE patterns are defined in Fig. 6 for different zircon domains with respect to their growth and recrystallization during the eclogite-facies metamorphism. Table 2 presents a summary of representative features for the four

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Table 1 Summary of zircon U–Pb ages, O–Hf isotopes and trace elements analyzed by both SIMS and LA-(MC)-ICPMS for UHP eclogite and granitic gneisses at Qinglongshan in the Sulu orogen. SIMS analyses Spota

206

Pb/238U age (Ma)

LA-(MC)-ICPMS analyses Th/U

d18O (&)

Eclogite 1.1c 1.2c 2.1n 3.1c 4.1c 4.2c 5.1n 7.1n 11.1n 12.1n 15.1n 17.1c 18.1n

00QL16 432 ± 8

0.38

235 ± 5 443 ± 7 450 ± 7

0.03 0.53 0.64

212 ± 4 226 ± 4 235 ± 7 219 ± 4 212 ± 7 330 ± 5 245 ± 4

0.04 0.04 0.03 0.04 0.04 0.20 0.03

7.1 7.6 2.5 6.1 7.6 8.4 2.4 2.6 2.2 2.3 3.7 6.2 2.5

Granitic 1.1c 1.2 2.1c 3.1 3.2 4.1c 4.2 4.3 5.1 5.2 6.1c 6.2 7.1c 7.2c 7.3 8.1 8.2c 9.1 m 9.2c 10.1c 10.2 10.3 11.1 11.2c 12.1 13.1 13.2c 13.3

gneiss 99QL07 758 ± 14 1.34 249 ± 7 0.54 752 ± 12 0.90 518 ± 23 1.15 217 ± 4 0.16 799 ± 13 1.36 766 ± 12 1.11 216 ± 4 0.24 714 ± 11 0.82 219 ± 4 0.38 774 ± 12 1.05 784 ± 13 0.59 658 ± 11 2.77 718 ± 11 2.41 219 ± 4 0.38 214 ± 4 0.55 735 ± 12 1.28 420 ± 7 0.95 695 ± 11 1.71 616 ± 10 0.85 358 ± 8 0.99 234 ± 4 0.49 216 ± 3 0.22 692 ± 11 0.63 221 ± 4 0.52 222 ± 4 0.57 722 ± 11 0.59 665 ± 15 0.95

2.0 8.1 2.8 4.2 9.2 3.7 2.8 8.3 4.2 9.2 0.1 0.9 5.8 2.2 9.4 8.6 1.8 8.4 2.0 2.2 0.7 4.7 8.4 0.6 9.1 8.8 2.7 0.4

Granitic 1.1c 1.2r 1.3c 2.1c 2.2r 3.1r 3.2c 4.1r 4.2 4.3c 5.1 5.2 5.3

gneiss 99QL16 476 ± 10 1.02 215 ± 3 0.03 363 ± 6 0.90 426 ± 7 0.91 214 ± 3 0.05 223 ± 3 0.03 375 ± 6 0.69 216 ± 3 0.02 229 ± 4 0.96 502 ± 9 0.82 528 ± 8 0.80 437 ± 7 0.96 456 ± 8 1.10

1.2 9.8 3.6 4.2 9.6 9.0 8.9 9.2 9.9 1.6 0.4 2.8 4.8

Spota

206

Pb/238U age (Ma)

Th/U

1c

440 ± 4

2c 3n 7n 12n 13n 17n 20c 18n

176

Lu/177Hf

b

Eu/Eu* c

HREE (ppm)

(Yb/Gd)Nc

0.59

0.31

740

36.2

405 ± 4 207 ± 4 218 ± 5 247 ± 9 229 ± 7 243 ± 7 371 ± 3 239 ± 9

1.29 0.05 0.06 0.05 0.05 0.09 0.48 0.06

0.62 2.31 1.96 1.02 0.56 1.59 0.23 0.92

752 16.1 18.1 19.5 14.5 12.5 1773 13.1

19.8 9.67 9.86 6.32 7.18 16.5 18.6 15.4

1c 2 3c

663 ± 9 578 ± 8 678 ± 6

1.29 1.32 1.94

0.002286 0.000575 0.004168

1.9 1.6 4.2

0.43 0.42 0.48

1267 415 2112

15.7 28.8 19.8

7 8c 9 10 11 12 14c 15

214 ± 3 765 ± 8 655 ± 7 648 ± 7 250 ± 4 247 ± 4 707 ± 6 621 ± 5

0.21 1.06 1.06 1.04 0.31 0.31 0.63 0.63

0.000169 0.002040 0.002023 0.002142 0.001396 0.000280 0.003392 0.001054

2.4 0.1 0.5 1.1 1.1 1.3 7.8 0.2

0.80 0.30 0.31 0.30 0.53 0.96 0.16 0.15

121 1038 1088 1263 200 186 1696 720

28.8 22.3 22.1 21.4 40.1 33.8 33.9 29.3

19c 20 17 18c 35 m 36c 32c 33

671 ± 8 242 ± 4 526 ± 7 692 ± 8 471 ± 9 599 ± 9 611 ± 6 624 ± 9

2.48 0.12 2.48 1.49 1.20 1.51 1.27 0.92

0.005020 0.000341 0.001302 0.002632 0.000798 0.002376 0.001961 0.000935

4.9 2.8 0.8 4.6 8.9 0.5 0.2 0.8

0.35 0.66 0.53 0.38 0.51 0.50 0.26 0.32

2696 192 942 1337 572 910 1026 669

15.8 25.3 23.3 20.0 27.6 19.9 29.7 32.0

29 30c 31

520 ± 6 649 ± 7 639 ± 8

0.72 0.71 1.73

0.000823 0.002464 0.000779

2.7 4.5 7.9

0.33 0.21 0.47

515 1618 846

36.6

1c 2 3c 4c 5

487 ± 5 501 ± 13 521 ± 6 512 ± 5 455 ± 7

0.73 0.74 0.80 1.63 1.54

0.001789 0.001202 0.000806 0.002408

6.4 5.0 4.6 11.8

0.43 0.56 0.52 0.43 0.41

869 557 437 1343 1341

34.3 25.0 31.5 15.8 15.9

6c 10 11 12c 7 8 9

403 ± 6 260 ± 7 535 ± 10 577 ± 6 432 ± 12 465 ± 4 414 ± 5

1.32 0.12 1.46 1.49 0.79 1.43 1.16

0.001784 0.000864 0.002017 0.002338 0.001182 0.004876 0.002875

2.6 2.7 3.1 2.5 7.4 7.5 9.2

0.17 0.58 0.50 0.52 0.58 0.14 0.21

1136 341 930 1271 486 2480 1321

14.8 13.2 15.6 12.9 20.0 13.0 14.1

eHf(t)

31.0 17.7

Metamorphic growth and recrystallization of zircons in negative d18O rocks Table 1 (continued ) SIMS analyses a

LA-(MC)-ICPMS analyses

Spot

206

Pb/ U age (Ma)

Th/U

d O (&)

Spota

206

Pb/238U age (Ma)

Th/U

176

eHf(t)

6.1r 6.2c 7.1r 8.1r 9.1r 10.1r 11.1 11.2 11.3r 12.1r 12.2r 13.1r 13.2c

222 ± 3 592 ± 9 212 ± 3 214 ± 5 224 ± 4 219 ± 3 585 ± 14 627 ± 10 211 ± 3 226 ± 4 219 ± 4 216 ± 4 398 ± 6

0.02 1.02 0.04 0.02 0.02 0.02 0.95 0.92 0.04 0.03 0.05 0.02 1.01

9.5 8.4 10.0 9.6 9.8 9.8 1.5 0.5 9.8 9.6 9.4 9.4 4.1

13 14c 18 19 25 26r 28 29

416 ± 7 413 ± 7 462 ± 7 457 ± 7 350 ± 8 234 ± 5 479 ± 6 411 ± 7

0.63 1.13 0.74 0.86 0.81 0.45 0.98 0.78

0.000837 0.001870 0.000905 0.000980 0.000893 0.000650 0.001717 0.000867

34

579 ± 17

0.32

37r

231 ± 6

0.09

8.9 3.1 2.6 2.7 9.2 8.8 3.6 1.5 8.8 1.1 9.5 2.8 0.7 4.0 8.7 8.8 0.4 2.8 8.7 2.6 9.8 2.7 3.5 1.1 0.4 2.4 6.4

1 2c 3 4c 5 6 7m 8c 9m 10c 11 12c 13 14 15 16 17r 18c 19r 20c 21 m 22c 23 24c 25 m 26c 27

538 ± 7 647 ± 9 602 ± 9 562 ± 11 361 ± 17 221 ± 7 321 ± 6 729 ± 11 220 ± 5 599 ± 7 215 ± 4 682 ± 8 589 ± 6 247 ± 5 553 ± 10 637 ± 8 240 ± 9 733 ± 9 226 ± 6 704 ± 7 304 ± 7 764 ± 7 318 ± 5 532 ± 9 225 ± 6 728 ± 9 376 ± 6

0.81 0.72 0.66 0.64 0.37 0.44 0.45 0.58 0.43 0.86 0.38 1.35 0.91 0.30 0.45 0.88 0.37 0.99 0.36 0.99 0.47 1.39 0.36 0.72 0.53 0.80 0.45

Granitic 1.1r 1.2c 2.1 3.1c 4.1 5.1 6.1 m 6.2c 7.1 m 7.2c 8.1 9.1c 10.1 10.2 11.1 12.1 13.1r 13.1c 14.1r 14.1c 15.1 m 15.2c 16.1 16.2c 17.1 m 17.2c 18.1 a b c

238

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gneiss 00QL27 213 ± 3 0.42 658 ± 9 0.72 765 ± 11 0.60 802 ± 11 1.13 218 ± 3 0.42 221 ± 3 0.42 773 ± 11 0.57 691 ± 10 0.48 217 ± 3 0.39 649 ± 9 0.71 223 ± 3 0.45 750 ± 11 1.24 698 ± 10 0.69 215 ± 3 0.27 230 ± 4 0.21 223 ± 3 0.32 226 ± 3 0.24 750 ± 11 0.93 219 ± 4 0.37 721 ± 10 0.70 332 ± 5 0.66 776 ± 11 1.00 518 ± 17 0.91 771 ± 11 1.12 260 ± 4 0.44 752 ± 11 0.70 613 ± 9 1.01

18

Lu/177Hf

b

Eu/Eu* c

HREE (ppm)

(Yb/Gd)Nc

2.8 3.5 9.1 5.6 6.3 3.0 7.2 6.3

0.57 0.52 0.45 0.52 0.52 0.65 0.46 0.56

356 813 499 598 509 247 896 464

27.9 14.1 19.5 18.4 18.7 63.4 16.0 22.3

0.000604

2.3

0.64

257

12.6

0.000826

0.7

0.87

369

313

0.48 0.52 0.27 0.42 0.34 0.52 0.59 0.30 0.51 0.41 0.69 0.38 0.41 0.42 0.36 0.42 0.58 0.52 0.60 0.43 0.47 0.42 0.39 0.39 0.66 0.41 0.50

1510 1021 840 1494 437 207 422 666 102 2899 201 2719 1936 426 533 1965 434 2752 122 2388 453 2650 291 3084 205 1722 1192

13.4 13.4 17.8 13.9 33.1 22.1 48.7 20.1 9.20 11.5 19.8 10.8 11.3 40.9 32.8 11.0 49.6 11.0 12.2 10.8 12.5 11.4 26.9 12.9 19.5 11.8 21.2

Postfixes c, m, r, and n denote zircon core, mantle, rim and no zoning domains, respectively. Initial Hf isotope ratios are calculated at t = 780 Ma. p Eu/Eu* = EuN/ (Sm  Gd)N, where N denotes the normalization to the chondrite values of Sun and McDonough (1989).

types of zircon domains. Collectively, the zircons from the eclogite and granitic gneisses experienced metamorphic modification of different degrees. The integrated SIMS and LA-ICPMS study enables us to: (1) compare the element and isotope relationships between different depths of metamorphic zircons, (2) understand the genesis of negative d18O zircons, and (3) discriminate between the different subtypes of metamorphic recrystallization. To correlate the zircon U–Pb ages, O isotopes, trace elements and Lu–Hf isotopes, it is critical to ensure that the different types of microbeam analyses were actually made on the same property of domains (Cavosie et al., 2006; Kemp et al., 2010). There is generally no problem to corre-

late the SIMS U–Pb and O isotope data because the two types of analyses have comparable spot sizes (25 lm) and depths (5 lm). However, caution must be taken in correlating the SIMS data with the LA-(MC)-ICPMS data because the laser ablation has a larger spot size (44 lm) and a greater depth (20 lm). In order to examine whether the two methods have sampled the same property of zircon domains, we compared the Th/U ratios and 206Pb/238U ages for individual analyses (Fig. 9). Generally, the two methods give consistent results for most residual cores in all samples and for the homogeneous grains in eclogite 00QL16. Therefore, all the analyses for the eclogite by the two methods are considered to

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Fig. 3. Concordia diagrams of zircon U–Pb isotope data analyzed by SIMS and LA-ICPMS, respectively, for eclogite and granitic gneisses at Qinglongshan in the Sulu orogen. Panels: (a and b) eclogite 00QL16; (c and d) gneiss 99QL07; (e and f) gneiss 99QL16; and (g and h) gneiss 00QL27. Note panels (a, c, e and g) are the SIMS results, and panels (b, d, f and h) are the LA-ICPMS results. The d18O values are shown for certain grains along with the SIMS U–Pb data.

sample the same property of domains (Fig. 9a and b). However, there are remarkable departures for the three granitic gneisses, where most of the laser analyses yield significantly

higher Th/U ratios and 206Pb/238U ages than the SIMS analyses (Fig. 9c–h). Such departures appear to result from the difference in sampling depths. While the SIMS analyses

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Fig. 4. Histograms of zircon O isotope data for (a) eclogite and (b–d) granitic gneisses at Qinglongshan in the Sulu orogen. Black, blue, and red colors denote the domains with pre-Triassic 206Pb/238U ages, concordant Triassic U–Pb ages, and no U–Pb ages, respectively. Note the U– Pb ages indicated here are all SIMS analytical results. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

have only sampled the surface to 5 lm depth, the laser analyses have ablated the subsurface to 20 lm depth. In this regard, the laser analyses have often sampled the residual cores beneath the metamorphic rims. Therefore, caution must be taken when interpreting the two methods of analytical data from the same spots. For granitic gneiss 99QL07 (Fig. 9c and d), there are six domains where the SIMS spots were overlapped by the laser spots. The two methods of analysis yielded significantly different results. For five of the six domains, the SIMS analysis gave low Th/U ratios of 0.22–0.55 and concordant Triassic U–Pb ages of 214 ± 4 to 249 ± 7 Ma (Table A1), whereas the laser analysis yielded high Th/U ratios of 0.72–2.48 and discordant pre-Triassic U–Pb ages of 520 ± 6 to 648 ± 7 Ma (Table A2). For the remaining domain, the SIMS analysis (#5.1) gave a high Th/U ratio of 0.82 and a pre-Triassic U–Pb age of 714 ± 11 Ma (Table A1), whereas the laser analysis (#11) yielded a relatively low Th/U ratio of 0.31 and a Triassic U–Pb age of 250 ± 4 Ma (Table A2). The other analyses are considered to have sampled the same property of domains by the both methods (Fig. 9c and d).

For granitic gneiss 99QL16 (Fig. 9e and f), there are large differences in seven rim domains, where the SIMS analysis gave low Th/U ratios of 0.02 to 0.05 (except #4.2 with a high Th/U ratio of 0.96) and concordant Triassic U–Pb ages of 214 ± 3 to 229 ± 4 Ma, whereas the laser analysis yielded high Th/U ratios of 0.12 to 1.46 and discordant pre-Triassic U–Pb ages of 260 ± 7 to 535 ± 10 Ma (Fig. 9e and f). The other analyses are considered to have sampled the same property of domains by the both methods (Fig. 9e and f). For granitic gneiss 00QL27 (Fig. 9g and h), there are large differences in five mantle/rim domains. For four of them, the SIMS analyses gave relatively low Th/U ratios of 0.21–0.42 and concordant Triassic U–Pb ages of 213 ± 3 to 230 ± 4 Ma. In contrast, the laser analyses yielded high Th/U ratios of 0.37–0.88 and discordant preTriassic U–Pb ages of 361 ± 17 to 637 ± 8 Ma. The SIMS analysis of a remaining mantle domain (#17.1 m) gave a relatively high Th/U ratio of 0.44 and a pre-Triassic U– Pb age of 260 ± 4 Ma, whereas the laser analysis yielded a comparable Th/U ratio of 0.44 but a Triassic U–Pb age of 225 ± 6 Ma. For one mantle domain (#6.1 m), the SIMS

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Fig. 5. Th/U ratios and d18O values vs. SIMS U–Pb ages for zircons from eclogite and granitic gneisses at Qinglongshan in the Sulu orogen. Domains with concordant Triassic U–Pb ages are indicated as filled circles, and other domains are denoted as filled triangles. Panels: (a and b) eclogite 00QL16; (c and d) gneiss 99QL07; (e and f) gneiss 99QL16; and (g and h) gneiss 00QL27.

analysis gave a relatively high Th/U ratio of 0.57 and a concordant U–Pb age of 773 ± 11 Ma, whereas the laser analysis gave a comparable Th/U ratio of 0.45 but a discordant U–Pb age of 321 ± 6 Ma. The other analyses are considered to have sampled the same property of domains by the both methods. The differences in Th/U ratios and 206Pb/238U ages between the SIMS and laser analyses indicate that many zircons consist of varying proportions of residual core and

metamorphic product (new growth or recrystallized domain). This reveals the geochemical architecture of metamorphic zircons which otherwise cannot be recognized by the single microbeam technique. In order to correlate the SIMS O isotopes with the laser REE and Lu–Hf isotopes, we prefer to obtain the data of interest from the same property of domains. Therefore, Table 1 only includes zircon domains that were analyzed for the SIMS O isotopes together with the SIMS U–Pb ages or the LA-(MC)-ICPMS

Metamorphic growth and recrystallization of zircons in negative d18O rocks

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Fig. 6. Chondrite-normalized REE patterns for zircons from eclogite and granitic gneisses at Qinglongshan in the Sulu orogen. Normalization values are from Sun and McDonough (1989). Panels: (a) eclogite 00QL16; (b–d) gneiss samples 99QL07, 99QL16 and 00QL27, respectively. Note the division of types is based on integrated analyses of zircon U–Pb ages, trace elements, Lu–Hf isotopes, and/or O isotopes. Type III domains corresponds to metamorphic growth from aqueous fluids; Types I, II and IV are assigned to solid-state, dissolution and replacement recrystallized domains, respectively.

U–Pb ages, trace elements and Lu–Hf isotopes. On this basis, if not specially mentioned, the following comparison and correlation are made only on the same property of domains that have consistent U–Pb results between the SIMS and laser analyses (Fig. 10). 5. DISCUSSION 5.1. Eclogite Significant differences occur in d18O values, U–Pb ages and REE patterns between the residual cores and homogeneous grains in zircons from eclogite 00QL16 (Table 1 and Figs. 2–7). The cores exhibit positive d18O values of 6.1– 10.1&, moderately steep MREE-HREE patterns with negative Eu anomalies, high Th/U ratios (>0.20) and pre-Triassic U–Pb ages. These features indicate that the cores are of magmatic origin and crystallized from positive d18O magmas before the continental subduction. In contrast, the homogeneous grains have negative d18O values of 2.2& to 3.7&, flat MREE-HREE patterns with no Eu anomalies, low Th/U ratios (<0.04) and concordant Triassic U– Pb ages. These features demonstrate that the grains are of metamorphic origin and precipitated from negative d18O fluids during the Triassic eclogite-facies metamorphism. The intergrain d18O differences for the eclogite vary from 8.3& to 13.8& (Table A3), suggesting very limited O isotope exchange between the positive d18O cores and the negative d18O metamorphic fluids. In this regard, the cores would only experience metamorphic recrystallization in solid-state

(e.g., Xia et al., 2009; Chen et al., 2010). It appears that the metamorphic fluid was not accessible to the individual cores. 5.2. Granitic gneiss A few zircon grains have apparent core–mantle–rim or core–rim structures for the three samples of granitic gneiss (Fig. 2). Most of the residual cores show middle Neoproterozoic U–Pb ages, positive d18O values of 0.1–3.7&, and negative Eu anomalies in the REE patterns (Table 1 and Figs. 2–7). This demonstrates that the cores are of magmatic origin and crystallized from positive d18O magmas in the middle Neoproterozoic. In contrast, most of the mantle/rim domains exhibit Triassic U–Pb ages, negative d18O values of 10.0& to 8.7&, and weakly negative Eu anomalies in the REE patterns. This indicates that the domains are of metamorphic origin and formed either by growth from negative d18O fluids during the Triassic eclogite-facies metamorphism or by dissolution recrystallization in thermodynamic reequilibration with the metamorphic fluid in O and Pb isotopes (for further discussion see Section 7). The intragrain d18O differences vary from 0.8& to 12.5& (Table A3), suggesting variable degrees of O isotope exchange between the positive d18O cores and the negative d18O metamorphic fluids. The strong O isotope exchange also occurred during the metamorphic recrystallization, resulting in variably negative d18O values for some recrystallized zircons (Table 1). Many zircons show a transitional relationship between the d18O values, apparent U–Pb ages, REE patterns and/

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Table 2 Summary of U–Pb ages, REE contents, Lu–Hf and O isotopes for different types of metamorphic zircons in UHP eclogite and granitic gneisses at Qinglongshan in the Sulu orogen. Metamorphic recrystallization

Subtypea Solid-state recrystallization Replacement recrystallization Sampleb 00QL16 99QL07 00QL27 99QL07 99QL16 00QL27 Mineragraphy Euhedral, blurred oscillatory zoning Cloudy Anhedral, Patchy to no zoning patchy zoning zoning Th and U High Th, U contents High Th, U contents contents Th/U ratios High High High High (0.95) High (>0.3) High (0.30 to (>0.20) (>0.6) (>0.58) 0.88) 206 Pb/238U PreClose to discordia Pre-Triassic Variable; pre- Variable; ages (Ma) Triassic upper intercept Triassic pre-Triassic HREE High High High High (855) High (most Variable contents (>740) (>410) (>666) >440) (290–2899) (ppm) Eu anomalies Strong negative Eu anomalies Strong negative Eu anomalies (Eu/Eu*) 176 Lu/177Hf na High na Relatively Variable but na ratios (most high (0.0013) most >0.001 >0.001) d18O values Positive Positive Positive na Negative Negative (&) (6.1 to (0.1 to 3.7) (0.7 to 3.1) (8.9 to (6.4 to 10.1) 0.4) 1.1) Representative #1c, 2c, #1c, 3c, #2c, 3, 4c, #35 m #1c, 3c, 4c, #10c, 14, 23, domainsc 20c 8c, 14c 8c 12c 27

Metamorphic growth Dissolution recrystallization 00QL16 99QL07 99QL16 Homogeneous Homogeneous Patchy grain and no zoning zoning High U low High Th, U contents Th contents Low (0.03) Relatively high High (0.12 to 0.70) (0.12–0.56) Concordant Triassic or close to Triassic Triassic High (1160) Low (111–287) Low (most < 340) No Eu anormaly na

00QL27 No zoning

High U low Th contents High (0.36–0.53)

na

na

Negative (9.4 to 8.1)

#10n

#7, 12, 20

Negative (9.7 to 9.6) #17, 20

Low (<0.1)

Low (<0.1)

Concordant Triassic Low (<453)

No or weak negative Eu anomalies Low (<0.001)

00QL16 99QL16 Homogeneous and no zoning

Negative (9.8 to 8.7) #6, 9 m, 11, 19r

Low (<31)

Low (<370)

No Eu Weak negative anomalies Eu anomalies Low (most < 0.001)

Negative (3.7 to 2.2) #3n, 6n, 7n, 8n

Negative (10.0 to 9.2) #26r, 37r

a Discrimination of metamorphic growth from different subtypes of recrystallization. For zircon domains analyzed by both SIMS and LA-ICPMS, only these with consistent U–Pb ages are indicated for HREE contents, Eu anomalies, and 176Lu/177Hf isotopes. When data are not available, they are denoted as “na”. b 00QL16 is eclogite, 99QL07, 99QL16 and 00QL27 are granitic gneiss samples. c The spots of laser ablation are indicated, where the c, m, r, n denote zircon core, mantle, rim and no zoning domains.

Y.-X. Chen et al. / Geochimica et Cosmochimica Acta 75 (2011) 4877–4898

Typea

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Fig. 7. Th/U ratios, HREE contents and Eu/Eu* ratios vs. LA-ICPMS U–Pb ages for zircons from eclogite and granitic gneisses at Qinglongshan in the Sulu orogen. Typology of the domains is the same as in Fig. 6. Panels: (a–c) eclogite 00QL16; (d–f) gneiss 99QL07; (g–i) gneiss 99QL16; and (j–l) gneiss 00QL27.

Fig. 8. 176Lu/177Hf, 176Hf/177Hf ratios and eHf(t1) values vs. LA-ICPMS U–Pb ages for zircons from UHP granitic gneisses at Qinglongshan in the Sulu orogen. The initial Hf isotope ratios are calculated at t1 = 780 Ma. Typology of the domains is the same as those in Fig. 6. Panels: (a–c) gneiss 99QL07; and (d–f) gneiss 99QL16.

or Lu–Hf isotopes (Figs. 5–8). This can be caused by different degrees of metamorphic recrystallization via the mineralogical mechanisms of solid-state transformation, replacement alteration, or dissolution recrystallization (e.g., Hoskin and Black, 2000; Rubatto et al., 2008; Xia

et al., 2009; Chen et al., 2010). For samples 99QL07 and 00QL27, most cores exhibit positive but scattered d18O values, variable pre-Triassic U–Pb ages, and relatively high Th/U ratios. This indicates that the cores crystallized from positive d18O magmas, but experienced limited extents of

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Fig. 9. Comparison of U–Pb isotope data analyzed by the SIMS and LA-ICPMS methods. X-axis values are those from the SIMS results, and Y-axis values are those from the LA-ICPMS results. Panels: (a and b) eclogite 00QL16; (c and d) gneiss 99QL07; (e and f) gneiss 99QL16; and (g and h) gneiss 00QL27. Error bars for Th/U ratios obtained by SIMS and laser ablation are set as 15% and 10% (2r), respectively. Error bars for 206Pb/238U ages are in 1r.

Fig. 10. Correlations between U–Pb ages, Th/U ratios, HREE contents, 176Lu/177Hf ratios and d18O values for zircons from granitic gneisses at Qinglongshan in the Sulu orogen. Panels: (a–d) gneiss 99QL07; (e–h) gneiss 99QL16; and (i–k) gneiss 00QL27. The SIMS U–Pb ages and Th/U ratios are used, and only the domains that give consistent U–Pb results from the LA-ICPMS analyses are included (see Fig. 9). Domains that formed by solid-state recrystallization (Type I), replacement recrystallization (Type IV), dissolution recrystallization (Type II) and metamorphic growth (Type III) are abbreviated as SSR, RR, DR and MG, respectively.

Metamorphic growth and recrystallization of zircons in negative d18O rocks

modification by the metamorphic fluid. In contrast, most mantle/rim domains have negative d18O values, concordant Triassic U–Pb ages, and relatively low Th/U ratios. Thus, these domains have formed either as overgrowths, or by dissolution recrystallization under intensive action of negative d18O metamorphic fluids (for further discussion see Section 7). A few domains fall between the two end-members, suggesting that they underwent replacement recrystallization upon intermediate action of the metamorphic fluids. Zircons in sample 99QL16 exhibit amoeboid structure with patchy-zoned cores and irregular homogeneous rims (Fig. 2d–f). The rims are easily distinguished from the cores by CL images. They exhibit strong negative and uniform d18O values, concordant Triassic U–Pb ages and low Th/ U ratios. This indicates that the rims grew from negative d18O fluids during the Triassic metamorphism. In contrast, most cores show less negative and scattered d18O values, variable pre-Triassic U–Pb ages, and high Th/U ratios. This suggests that the metamorphic fluids were accessible to these cores, resulting in varying degrees of recrystallization (e.g., Xia et al., 2009; Chen et al., 2010). 5.3. Metamorphic and protolith ages For most domains with concordant Triassic U–Pb ages, their Th/U ratios, REE patterns and Lu–Hf isotopes indicate that these Triassic U–Pb ages represent timing of zircon growth from metamorphic fluids at eclogite-facies conditions. Thus they provide direct dating of the metamorphic event. These ages are also consistent with known dates of 215–245 Ma for eclogite-facies metamorphism in the Dabie–Sulu orogenic belt (Zheng et al., 2009; Liu and Liou, 2011). In detail, the U–Pb ages have a large variation from 204 ± 4 to 252 ± 7 Ma (Tables A1 and A2), possibly indicating different genesis for the metamorphic domains (Fig. 2). A few domains have older U–Pb ages of 245 Ma and thus could form by dissolution recrystallization of protolith zircons. The other domains exhibit two clusters at 230 and 220 Ma, respectively, consistent with metamorphic growth during the eclogite-facies metamorphism. It is noted that many of the U–Pb spot ages are actually concordant, rendering the possibility unlikely that the ages may result from Pb loss event from a single population. Alternatively, the ages are interpreted to indicate different age populations, so that these domains may have grown at given stages of prograde, peak and retrograde metamorphism (e.g., Liu et al., 2004; Wan et al., 2005; Liu et al., 2006; Wu et al., 2006; Liu et al., 2009; Zhao et al., 2006; Zheng et al., 2007a). Further studies are thus required to find independent constraints on zircon growth at the different stages during the continental subduction-zone metamorphism. The laser U–Pb isotope analyses for residual zircon cores in eclogite 00QL16 gave a discordia upper-intercept age of 722 ± 100 Ma (Fig. 3b). SIMS and laser U–Pb analyses for residual cores in granitic gneiss 99QL07 yielded discordia upper-intercept U–Pb ages of 799 ± 32 Ma (Fig. 3c) and 809 ± 29 Ma (Fig. 3d), respectively. Similar U–Pb ages of 871 ± 120 Ma (Fig. 3e) and 771 ± 68 Ma (Fig. 3f), 767 ± 28 Ma (Fig. 3g) and 796 ± 50 Ma (Fig. 3h) were ob-

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tained for residual cores in granitic gneisses 99QL16 and 00QL27. Despite large analytical uncertainties for the upper-intercept ages, they are basically consistent not only with concordant SIMS U–Pb ages of 752 ± 12 to 799 ± 13 Ma (n = 6) and 750 ± 11 to 802 ± 11 Ma (n = 8) for residual cores in samples 99QL07 and 00QL27 (Table A1), but also with concordant laser U–Pb ages of 749 ± 7 to 765 ± 8 Ma (n = 4) and 764 ± 7 Ma (n = 1) for residual cores in the two gneisses (Table A2). The large uncertainties for the upper-intercept U–Pb ages are possibly caused by distribution of data points mostly close to the lower intercepts. This in turn reflects intensive modification of protolith zircons by the fluid-assisted metamorphism (Zheng et al., 2004; Wu et al., 2006). It is known that the Tukey’s biweight mean in the ISOPLOT program of Ludwig (2003) is best-suited to a normal distribution that is slightly “contaminated” with points not belonging to the distribution (Hoaglin et al., 1983). Using this statistical method for the discordia upper-intercept and spot U–Pb ages of middle Neoproterozoic, it yields an age of 769 ± 9 Ma for the domains of magmatic origin. This age, probably the best approximation to the protolith ages, is consistent with SHRIMP zircon U–Pb ages of 764 ± 17 to 789 ± 23 Ma for residual cores in eclogite at Qinglongshan (Zheng et al., 2004). It also falls in the range of 750–780 Ma for protolith U–Pb ages of the metaigneous rocks in the Dabie–Sulu orogenic belt (Zheng et al., 2009). 6. ORIGIN OF NEGATIVE d18O ZIRCONS 6.1. Possible d18O values for magmatic cores Because of the sluggish O diffusivity in crystalline zircon under anhydrous conditions (Watson and Cherniak, 1997; Zheng and Fu, 1998), solid-state recrystallized domains are capable of preserving the primary d18O values of magmatic zircons (Cherniak and Watson, 2003; Zheng et al., 2004). This conclusion gains support from this study (Table 2), particularly from a relatively larger dataset of solid-state recrystallized domains in granitic gneisses 99QL07 and 00QL27. For these domains, their d18O values do not exhibit significant correlations with U–Pb ages, REE contents/ ratios, and 176Lu/177Hf ratios (Fig. 10). In addition, seven SIMS analyses on gneiss 99QL07 gave concordant U–Pb ages close to the protolith age with a similar range of d18O values (0.1– 3.7&) to the other analyses on this type of zircon (Fig. 10a). Similarly, seven SIMS analyses on gneiss 00QL27 with the concordant U–Pb ages close to the protolith age gave a limited range of d18O values (2.4– 2.8&). They are comparable to the other d18O analyses on this kind of zircon from this sample (Fig. 10i) or from gneiss 99QL07. Thus, the d18O values for solid-state recrystallized domains can be used to approximate those for the primary magmatic zircons. The d18O values for this kind of domain are 6.1–10.1& for eclogite 00QL16, 0.1–3.7& for gneiss 99QL07, and 0.7–3.1& for gneiss 00QL27 (Table 1). Compared to normal mantle zircon values of 5.3 ± 0.3& (Valley et al., 1998), the zircon d18O values for the eclogite are significantly higher but those for the gneisses are remarkably lower. This suggests that the

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magmatic zircons in the eclogite crystallized from high d18O magmas, whereas those in the granitic gneisses crystallized from low d18O magmas. As revealed by the conventional BrF5 method and the laser fluorination technique, the UHP metamorphic rocks from the Dabie–Sulu orogenic belt have a large d18O range from 10.9& to 8.5& for zircon (Rumble et al., 2002; Zheng et al., 2004; Tang et al., 2008b) and from 11.1& to 10.6& for common rock-forming minerals (Rumble and Yui, 1998; Zheng et al., 1998; Fu et al., 1999; Zheng et al., 1999; Zheng et al., 2003; Tang et al., 2008a,b). On the other hand, magmatic zircons from greenschist-facies metamorphosed granites in this region have d18O values of 1.1& to 7.6& and U–Pb ages of 750–780 Ma, and the rock-forming minerals show d18O values of 14.4& to 12.5& (Wu et al., 2007; Zheng et al., 2007b, 2008). In particular, U–Pb ages of 782 ± 3 Ma were obtained for magmatic zircons with positive d18O values of 0.50–2.68& (Zheng et al., 2007b), providing a direct dating of low d18O magmatism in the northern margin of the South China Block. In this regard, the low d18O values of 0.1–3.7& for magmatic cores with the Neoproterozoic U–Pb age in this study can be interpreted to crystallize from low d18O granitic magmas, which were formed by remelting of high-T hydrothermally altered rocks in a rifting zone (Zheng et al., 2004; Wu et al., 2007; Zheng et al., 2007b; Zheng et al., 2008). This process was also used to explain the origin of low d18O igneous rocks in other areas (Taylor, 1977; Hattori and Muehlenbachs, 1982; Taylor, 1986; Bindeman and Valley, 2000; Bindeman et al., 2001, 2008). As listed in Table A1, six SIMS analyses on gneiss 99QL07 gave concordant U–Pb ages of 752 ± 12 to 799 ± 13 Ma with a weighted mean of 772 ± 18 Ma (MSWD = 1.8), and eight SIMS analyses on gneiss 00QL27 yielded concordant U–Pb ages of 750 ± 11 to 802 ± 11 Ma with a weighted mean of 767 ± 14 Ma (MSWD = 2.5). Taken together, the Tukey’s biweight mean of the discordia upper-intercept and concordant U– Pb ages of middle Neoproterozoic for residual cores yields an age of 769 ± 9 Ma for the domains of magmatic origin. These ages are close to the U–Pb age of 782 ± 3 Ma for emplacement of low d18O magmas at the northern margin of the South China Block (Zheng et al., 2007b, 2008). Thus, the low d18O cores from samples 99QL07 and 00QL27 were crystallized from low d18O magmas in the middle Neoproterozoic. On the other hand, high d18O magmas can be derived from remelting of low-T hydrothermally altered or chemically weathered rocks, like the formation of high d18O S-type granites (e.g., O’Neil and Chappell, 1977). For the mafic protolith of eclogite, the low-T alteration of basaltic rocks is the best way to elevate their d18O values to the presently observed ones for the magmatic cores. In either case, protoliths of the metaigneous rocks would acquire their higher or lower d18O values than the normal mantle values in the middle Neoproterozoic. 6.2. Origin of negative d18O metamorphic growths The d18O values for metamorphically grown or dissolution recrystallized domains are variably negative, but rela-

tively uniform for individual samples (Figs. 5 and 10). This indicates that these zircon domains in single samples would form in O isotope equilibrium with the same negative d18O fluids during the Triassic metamorphism. In view of such a large dataset of SIMS zircon O and U–Pb isotopes (Table 1), it is assumed that those analyses without corresponding U–Pb date but with comparable d18O values for spots showing concordant Triassic U–Pb ages are also the domains that have equilibrated with metamorphic fluids for both O and Pb isotopes. In this regard, the domains selected for comparison include those possibly formed by dissolution recrystallization. As a consequence, d18O values for these domains are estimated to be 3.7& to 2.2& for eclogite 00QL16, 9.4& to 7.1& for gneiss 99QL07, 10.0& to 9.0& for gneiss 99QL16, and 9.8& to 8.7& for gneiss 00QL27. Although the selected domains have a range of negative d18O values, rather uniform d18O values are associated with the concordant Triassic U–Pb ages for individual samples (Table 1; Figs. 4, 5 and 10). However, the core and the mantle/rim exhibit significant differences in both U–Pb age and d18O value (Figs. 3–5), indicating that they grew from contrasting sources at different times. In other words, there was a negative d18O reservoir within the subducted continental slice, enabling the growth of negative d18O zircons during the Triassic metamorphism. Previous studies revealed d18O values as negative as 11.1& to 6.0& for rockforming minerals from the UHP metamorphic rocks (Yui et al., 1995; Zheng et al., 1996; Yui et al., 1997; Rumble and Yui, 1998; Zheng et al., 1998; Fu et al., 1999; Zheng et al., 1999; Zheng et al., 2003; Tang et al., 2008a,b) and as negative as 14.4& to 10.0& for greenschist-facies metamorphosed granites (Wu et al., 2007; Zheng et al., 2007b, 2008) in the Dabie–Sulu orogenic belt. The negative d18O values for granite minerals are interpreted as being derived from high-T glacial-hydrothermal alteration at 748 ± 3 Ma (Wu et al., 2007; Zheng et al., 2007b, 2008). This was based on the observations that identical U–Pb ages of 747 ± 4 and 749 ± 5 Ma were obtained for unzoned or convoluted zoned zircon domains that mostly cut the primary zonation of magmatic zircons and some of them have high common Pb concentrations of 5.45–61.00% (Zheng et al., 2007b). In this regard, the magmatic rocks would acquire their negative d18O values for rock-forming minerals that are susceptible to water–rock interaction at 750 Ma. During the Triassic eclogite-facies metamorphism, the 18O depleted minerals would be transformed into negative d18O metamorphic minerals and fluids. Metamorphic dehydration is evident during this process of continental subduction-zone metamorphism (Zheng et al., 2003; Chen et al., 2007b, 2007c; Zheng, 2009; Zheng et al., 2009; Chen et al., 2011), resulting in the growth of negative d18O zircons from the negative d18O metamorphic fluids. An analog to the protolith of Dabie–Sulu UHP metaigneous rocks is altered Cenozoic basalts at Krafla in Iceland, which also have negative d18O values of 12.7& to 8.5& for whole-rocks and minerals (Hattori and Muehlenbachs, 1982). The basalts were erupted in a rifting zone; hyaloclastites were formed during glaciation. Because Iceland is

Metamorphic growth and recrystallization of zircons in negative d18O rocks

located in the polar area with cold climate, the negative d18O values for the altered basalts are certainly associated with either the meteoric water of cold climate or glacial meltwater. The protolith of Dabie–Sulu UHP metaigneous rocks was also emplacement in a rifting zone with the possible Kaigas continental glaciation (Zheng et al., 2007b, 2008). The difference is that it underwent the UHP eclogite-facies metamorphism during the Triassic continental subduction (Zheng et al., 2003, 2004, 2009). The negative d18O basalts of Iceland would become negative d18O eclogites if they could be subducted to mantle depths for UHP metamorphism at a later time. 7. DISTINCTION BETWEEN METAMORPHIC GROWTH AND RECRYSTALLIZATION OF ZIRCON To decipher the genesis of metamorphic zircons is of great interest and importance to interpretation of zircon U–Pb ages for various rocks (Wu and Zheng, 2004; Gerdes and Zeh, 2009). It is critical to recognize different subtypes of metamorphic recrystallization and distinguish them from metamorphic growth (Xia et al., 2009; Chen et al., 2010; Xia et al., 2010). Previous discriminations of metamorphic zircons were based on CL images, U–Pb ages, REE and other trace elements, and Lu–Hf isotopes (e.g., Vavra et al., 1996, 1999; Hoskin and Black, 2000; Zheng et al., 2004, 2005, 2006; Chen et al., 2007a; Geisler et al., 2007; Zheng et al., 2007a; Martin et al., 2008; Rubatto et al., 2008; Gerdes and Zeh, 2009; Rubatto et al., 2009; Xia et al., 2009; Chen et al., 2010; Xia et al., 2010). Generally, the newly grown domains at eclogite-facies conditions are characterized by metamorphic U–Pb ages, lowered HREE contents and weakly negative Eu anomalies, and low Th/U and 176 Lu/177Hf ratios. These domains have variable REE patterns depending on coexisting minerals during zircon growth, typically with flat MREE-HREE patterns due to the garnet effect (e.g., Rubatto and Hermann, 2007; Zheng, 2009; Liu and Liou, 2011). The metamorphic fluid probably plays a significant role during the new growth. In contrast, solid-state recrystallized domains are less affected by the metamorphic fluid and thus tend to maintain the primary features of magmatic zircon. They commonly show blurred oscillatory zoning, close to protolith but lowered U–Pb ages, variably steep MREE-HREE patterns with negative Eu anomalies, and high Th/U and 176Lu/177Hf ratios (e.g., Hoskin and Black, 2000; Xia et al., 2009; Chen et al., 2010). However, the metamorphic fluid has variable extents of accessibility to the replacement and dissolution recrystallized zircons. The dissolution reprecipitation involves the intensive action of metamorphic fluids. As a result, it causes almost complete resetting of their mineragraphy, U–Pb chronometric and REE systems. Thus, the reprecipitated products represent the mixtures between protolith zircon and metamorphic fluid. The replacement alteration causes similar effects, but it only results in intermediate resetting of the U–Pb and REE systems. Because of the infiltration of metamorphic fluid along fractures in protolith zircons, the occurrence of porphyritic textures is common in the products of replacement recrystallization (Xia et al., 2009; Chen et al., 2010).

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In plots of Th/U ratios and d18O values versus Pb/238U ages for zircons from the studied rocks (Fig. 5), it is feasible to apply the above generalization to categorization of metamorphic growth and recrystallization (Table 2). This is particularly so for eclogite 00QL16, which exhibits distinctive features for the different properties of domains (Fig. 5a and b). As a consequence, Types I and II domains are assigned to the solid-state and dissolution recrystallized zircons, respectively; Type III domains correspond to the metamorphic growth from aqueous fluids (Fig. 6). However, the difficulty is encountered in the three granitic gneisses when recognizing some replacement or dissolution recrystallized zircons because they all exhibit variably negative d18O values and pre-Triassic or concordant Triassic U–Pb ages. In this case, the correlations between zircon d18O values, U–Pb ages, REE and Lu–Hf isotopes can provide tight constraints on the subtype of metamorphic recrystallization (Fig. 10). For eclogite 00QL16, metamorphic growths can be readily distinguished from the recrystallized grains by their negative d18O values, low Th/U ratios (Fig. 5a and b), low REE contents, and flat MREE-HREE patterns with no Eu anomalies (Type III in Fig. 6a), typical feature for metamorphic growth from aqueous fluids (e.g., Rubatto, 2002; Xia et al., 2009; Chen et al., 2010). In contrast, residual cores are characterized by positive d18O values, high Th/ U ratios, and steep MREE-HREE patterns with moderate negative Eu anomalies (Type I; Fig. 6a). The blurred oscillatory zoning and the apparent U–Pb ages significantly lower than the protolith age indicate that the domains are of magmatic origin but underwent different extents of solidstate recrystallization (e.g., Hoskin and Black, 2000; Xia et al., 2009; Chen et al., 2010). One grain (#10n; Type II in Fig. 6a) without O isotope data exhibits a concordant Triassic U–Pb age, a low Th/U ratio and no Eu anomaly, features similar to the metamorphic growth (Fig. 7a–c). But its much higher contents of Th and U, and a steeper REE pattern (Tables A2 and A4) indicate its formation by dissolution recrystallization that completely reset the U–Pb system but has partially inherited the REE pattern of protolith zircon (e.g., Chen et al., 2010). For gneiss sample 99QL07, most residual cores of positive d18O values (0.1–3.7&) are characterized by blurred oscillatory zoning, close to protolith U–Pb ages, high REE contents, high Th/U and 176Lu/177Hf ratios, and steep REE patterns (Type I in Figs. 6b, 7 and 8). These domains are interpreted as a result of solid-state recrystallization (e.g., Hoskin and Black, 2000; Xia et al., 2009; Chen et al., 2010). In contrast, Type II domains of negative d18O values (Table 1) are characterized by low REE contents, moderately steep REE patterns (Fig. 6b) with weakly negative Eu anomalies (Fig. 7e and f), and generally low 176 Lu/177Hf ratios (<0.001) (Fig. 8a). All the Type II domains have discordant pre-Triassic to concordant Triassic U–Pb ages and high Th/U ratios of 0.12–0.70. A survey of the literature data for UHP gneisses in the Dabie–Sulu orogenic belt reveals that almost all metamorphically grown zircons have Th/U ratios lower than 0.1 (Zheng et al., 2003; Liu et al., 2004; Zheng et al., 2004; Zheng et al., 2005; Liu et al., 2006; Wu et al., 2006; Zheng et al., 206

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2006; Liu et al., 2008; Xia et al., 2009; Zhang et al., 2009; Zheng, 2009; Zheng et al., 2009; Chen et al., 2010; Liu et al., 2010; Xia et al., 2010) with just a few exceptions (e.g., Liu et al., 2009; Liu and Liou, 2011). Liu et al. (2009) found that coesite-bearing zircon domains have strongly positive Ce anomalies and low U contents (<30 ppm) and attributed their high Th/U ratios (>0.5) to metamorphic growth from aqueous fluids of high oxygen fugacity. Alternatively, these features can be produced by dissolution recrystallization in the accessibility of aqueous fluids. Type II domains in gneiss 99QL07 exhibit slightly lower U contents and much lower Th contents compared to the residual cores (Table A2), indicating that their lowered Th/U ratios are mainly caused by decreased Th contents. In this regard, these domains would form by the dissolution recrystallization rather than the metamorphic growth. This is consistent with their low HREE contents and steep MREE-HREE patterns despite the occurrence of some garnets in this sample (Fig. 6b). A Type IV domain (Fig. 6b) in gneiss 99QL07 is more complicated to interpret by dissolution recrystallization. This domain (#9.1 m) has a d18O value of 8.4& (Table 1) with consistent U–Pb ages from the both SIMS and laser analyses (Tables A1 and A2). The laser analysis (#35 m) on this domain yielded a high Th/U ratio (1.20) and a pre-Triassic U–Pb age (471 ± 9 Ma), a REE pattern similar to Type I zircons (Fig. 6b) and a low 176Lu/177Hf ratio (Fig. 8a). The negative d18O value and partially reset REE and U–Pb isotopes indicate that the Type IV domain may have formed by replacement recrystallization. In this regard, three domains (#5.1, #7.1c and #13.3) that have SIMS U–Pb ages close to the protolith age and variably negative d18O values (5.8& to 0.4&; Fig. 5d) also probably formed by replacement recrystallization. Zircons from gneiss 99QL16 underwent intensive modification by metamorphic fluid (Fig. 2d–f). The patchyzoned cores show high REE contents and Th/U ratios, discordant U–Pb ages (Fig. 7g–i), steep REE patterns (Type IV domains in Fig. 6c), and high 176Lu/177Hf ratios (Fig. 8d). Most analyses gave variably negative d18O values of 4.7& to 0.4& except two significantly lower values of 8.9& to 8.4& (Fig. 5f). A correlation occurs between the d18O value and 206Pb/238U age for these spots (Fig. 5f), indicating a coupled variation in the O and U– Pb isotopes due to different extents of action by the negative d18O fluid. These features suggest that the core domains underwent replacement recrystallization. On the other hand, the rims exhibit negative d18O values, low Th/U ratios and concordant Triassic U–Pb ages, which can be distinguished from the other domains by the SIMS U–Pb and O isotope data alone (Fig. 5e and f). These domains show low REE contents, steep MREE-HREE patterns with weak negative Eu anomalies (Type III; Fig. 6c), and low 176 Lu/177Hf ratios (Fig. 8d). They are interpreted as metamorphic growth from negative d18O metamorphic fluids. Their REE features, different from those described for the same type of zircons by Xia et al. (2009) and Chen et al. (2010), may reflect the difference in modal abundance of LREE-rich accessory minerals. The presence of allanite in

gneiss 99QL16 would probably contribute to the large REE fractionation during the growth of metamorphic zircon (Liu et al., 2009; Rubatto et al., 2009). The interpretation of Type II zircons in gneiss 99QL16 (Fig. 6c) is a little complex. The two SIMS analyses on this type of domains (#4.1r and #9.1r) gave concordant Triassic U–Pb ages, low Th/U ratios and negative d18O values (Fig. 5f), suggesting metamorphic growth from aqueous fluids. In contrast, a laser analysis for the former site (#10 in Table A2) yielded a relatively low Th/U ratio of 0.12, a late Permian U–Pb age of 260 ± 7 Ma, and a steep MREEHREE pattern similar to Type IV zircons (Fig. 6c). These features suggest that the laser ablation sampled the dissolution recrystallized domain. For the latter site, the laser analysis (#25) gave a high Th/U ratio of 0.81, a pre-Triassic U– Pb age of 350 ± 8 Ma, and a REE pattern similar to Type IV domains (Fig. 6c), indicating that the replacement recrystallized domains may have been sampled by laser ablation. These differences suggest that the SIMS and laser analyses can sample different depths of zircons for their elemental and isotopic compositions. The SIMS analysis of one site (#18.1) yielded a negative d18O value of 7.7& (Table A3); the laser analysis on this site (#30) gave a relatively low Th/U ratio of 0.47, a late Permian U–Pb age of 262 ± 9 Ma (Table A2), and REE pattern similar to Type III zircons (Fig. 6c). The negative d18O value and almost completely reset U–Pb and REE systems indicate that this domain may have formed by dissolution recrystallization, but the laser analysis incorporated some mixture of the residual core. The SIMS O isotope analyses on the other two sites (#15.2 and #16.3) gave negative d18O values of 9.7& and 9.6&; the laser analyses (#17 and #20) yielded high Th/U ratios and pre-Triassic U–Pb ages (Table A2; Fig. 7g), slightly lower REE contents, and steep MREE-HREE patterns similar to Type IV zircons (Fig. 6c). These two domains are interpreted as forming by replacement recrystallization (e.g., Xia et al., 2009; Chen et al., 2010). At another site, the SIMS O analysis (#20.2c) gave a negative d18O value of 4.9&, and the laser analysis (#36c; Table A2) yielded a high Th/U ratio of 0.87, a preTriassic U–Pb age of 340 ± 6 Ma, and a REE pattern similar to Type IV zircons (Fig. 6c), indicating its formation by replacement recrystallization (e.g., Xia et al., 2009; Chen et al., 2010). For gneiss 00QL27, most residual cores of positive d18O values of 0.7& to 3.1& are characterized by blurred oscillatory zoning, close to protolith U–Pb ages, high REE contents, high Th/U ratios, and steep MREE-HREE patterns (Type I in Fig. 6d; Figs. 7 and 8). These domains are interpreted as a result of solid-state recrystallization (e.g., Hoskin and Black, 2000; Xia et al., 2009; Chen et al., 2010). In contrast, Type II domains of negative d18O values (Fig. 10) are characterized by low REE contents, moderately steep MREE-HREE patterns (Fig. 6d) with weak negative Eu anomalies (Fig. 7k and l). All the Type II domains have discordant pre-Triassic to concordant Triassic U–Pb ages and high Th/U ratios of 0.36–0.53. Similar to the arguments for Type II zircons in gneiss 99QL07, these domains are interpreted as a result of dissolution recrystallization. A few zircon domains of flat HREE patterns may

Metamorphic growth and recrystallization of zircons in negative d18O rocks

indicate the profound garnet effect during metamorphic recrystallization at eclogite-facies (e.g., Rubatto, 2002). Type IV domains in gneiss 00QL27 exhibit variably negative d18O values (Fig. 10), variable REE contents, and moderately steep REE patterns (Fig. 6d) with negative Eu anomalies (Fig. 7k and l). The partially reequilibrated O and REE systems indicate their formation by replacement recrystallization. In summary, metamorphically grown zircons can be distinguished from the three subtypes of metamorphically recrystallized zircons by evaluating their trace elements, U–Pb, Lu–Hf and O isotopes (Table 2). Metamorphic growths (Type III zircons) have negative but relatively uniform d18O values (Figs. 5 and 10), recording the O isotope composition of metamorphic fluids from which they grew. These zircons exhibit low REE contents, low Th/ U and 176Lu/177Hf ratios, and weakly negative Eu anomalies (Table 2). Solid-state recrystallized zircons (Type I domains) have positive d18O values and lowered U–Pb ages with relative blurring of oscillatory zones, and largely maintain the primary O isotope and REE features (Fig. 10). Dissolution recrystallized zircons (Type II domains) exhibit negative d18O values that probably approached O isotope equilibrium with negative d18O fluids and almost complete resetting of U–Pb ages, but their REE and Lu–Hf isotopes were only partially equilibrated with the metamorphic fluids. These features are different from those found by Martin et al. (2008), who ascribed the preservation of REE and O isotopes in metamorphic domains to formation by dissolution recrystallization in a partially closed system. The replacement recrystallized domains (Type IV zircons) generally have negative and intermediate d18O values between the solid-state and dissolution recrystallized domains. They cannot be distinguished without O isotope data. Interestingly, the differential metamorphic effects on zircon REE, U–Pb, Lu–Hf, and O isotope systems are consistent with the diffusion data for the respective elements (Cherniak and Watson, 2003), indicating lattice control of the elemental and isotopic systematics in zircon during the metamorphic modification. While the replacement altered domains have large O isotope variations that clearly indicate open-system behaviors, their Hf isotopes do not show significant differences (Fig. 8). This seemingly contradictory result can be explained by considering the mass balance for the elements that exchanged by fluid–mineral interaction. As the aqueous fluid has a very low Hf content but a high O content, Hf in the replacement altered domains is mainly from the precursor zircon whereas the O content is from both the fluid and the precursor zircons depending on the fluid/mineral ratios. 8. CONCLUSIONS The combined SIMS and LA-(MC)-ICPMS analyses of zircon U–Pb, O–Hf isotopes and trace elements enable us to compare the element and isotope variations for different sampling depths of zircons. The results allow a direct link between the ion probe and laser U–Th–Pb data on the same property of domains, providing an insight into the geo-

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chemical architecture of metamorphic zircons. Although many zircons are actually different proportions of mixtures between residual cores and metamorphic overgrowths, it is evident from our combined study of U–Pb dating by the two microbeam techniques that the outmost domains sampled by SIMS may not have the same property as those sampled by LA-(MC)-ICPMS. Therefore, caution must be taken when interpreting the two methods of analytical data from the same spots. While the Neoproterozoic protolith U–Pb age and the Triassic metamorphic U–Pb age are consistent with previous studies, the present study reveals significant contrasts in O isotopes between the residual cores of magmatic origin and the new growths (individual grain or mantle/rim) of metamorphic origin. The Neoproterozoic cores show positive d18O values, whereas the Triassic growths exhibit negative d18O values. Intragrain d18O values from 0.8& to 12.5&, indicating variable degrees of O isotope exchange between the positive d18O magmatic zircons and the negative d18O metamorphic fluids during continental subduction-zone metamorphism. The residual magmatic cores commonly have Neoproterozoic U–Pb ages, high Th/U ratios, positive d18O values, high REE contents, and steep REE patterns with negative Eu anomalies. They are interpreted to have crystallized from positive d18O magmas but experienced metamorphic recrystallization in solid-state with very limited accessibility to metamorphic fluid. The positive d18O values for protolith zircons are either higher or lower than normal mantle zircon values. This is ascribed to their crystallization from high or low d18O magmas that were originally generated by remelting of either low-T or high-T hydrothermally altered rocks in the Neoproterozoic. In contrast, some mantle/rim domains have Triassic U–Pb ages, low Th/U ratios, negative d18O values, low REE contents, weak to no Eu anomalies, and variable REE patterns depending on the garnet effect. They are interpreted to have precipitated from negative d18O fluids that were produced by metamorphic dehydration of high-T glacial-hydrothermally altered rocks during the Triassic continental collision. The coupled O isotope variations with U–Pb ages, Th/ U ratios, REE patterns and/or Lu–Hf isotopes provide robust constraints on distinction between three subtypes of metamorphic recrystallization with respect to the accessibility of aqueous fluid. Solid-state recrystallized zircons still maintained the positive d18O values, and REE and Lu–Hf isotopes of magmatic zircon, but show some extents of lowering in U–Pb ages. Dissolution recrystallized zircons generally show negative d18O values similar to the metamorphic growths, almost complete resetting of U–Pb ages and partial reequilibration of REE systems, and they still have high Th/U and 176Lu/177Hf ratios. In contrast, replacement recrystallized domains usually exhibit variable negative d18O values, and partial resetting of REE, and U–Pb and Lu–Hf isotope systems. Therefore, recognition of the three subtypes of metamorphic recrystallization enables mineralogical elucidation of different metamorphic processes and geological interpretation of zircon U–Pb dates for continental subduction-zone metamorphic rocks.

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This study was supported by funds from the Chinese Ministry of Science and Technology (2009CB825004) and the Natural Science Foundation of China (40830318 and 40921002). Thanks are due to Xianhua Li, Yu Liu, Guoqiang Tang, Yafei Wang and Yuya Gao for their assistance with the SIMS analyses in Beijing, to Xiaoming Liu and Honglin Yuan for their assistance with the LA-MCICPMS analyses in Xi’an, and to Zhaochu Hu and Yongsheng Liu for their assistance with the LA-ICPMS analyses in Wuhan. CYX thanks Dr. Q.-X. Xia for her discussion on the distinction between different types of metamorphic zircon. Detailed and critical reviews by Prof. Simon Wilde and two anonymous referees greatly help improvement of the presentation. We are also grateful to Prof. Simon Wilde for his thorough proof-reading and polishing of the language, and to Dr. Yuri Amelin for his constructive comments and editorial handling of the manuscript.

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