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Available online at www.sciencedirect.com ScienceDirect Geochimica et Cosmochimica Acta 143 (2014) 189–206 www.elsevier.com/locate/gca Origin of carbonatites in the South Qinling orogen: Implications for crustal recycling and timing of collision between the South and North China Blocks Cheng Xu a,⇑, Anton R. Chakhmouradian b, Rex N. Taylor c, Jindrich Kynicky d, Wenbo Li a, Wenlei Song a, Ian R. Fletcher e a Laboratory of Orogenic Belts and Crustal Evolution, School of Earth and Space Sciences, Peking University, China b Department of Geological Sciences, University of Manitoba, Canada c School of Ocean and Earth Science, University of Southampton, UK d Department of Geology and Pedology, Mendel University, Brno, Czech Republic e Department of Applied Geology, Curtin University, Australia Available online 13 April 2014 Abstract Most studies of compositional heterogeneities in the mantle, related to recycling of crustal sediments or delaminated subcontinental lithosphere, come from oceanic setting basalts. In this work, we present direct geochronological and geochemical evidence for the incorporation of recycled crustal materials in collision-related carbonatites of the South Qinling orogenic belt (SQ), which merges with the Lesser Qinling orogen (LQ) to separate the South and North China Blocks. The SQ carbonatites occur mainly as stock associated with syenites. The data presented here show that zircon from the syenites yields an age of 766 ± 25 Ma, which differs significantly from the age of primary monazite from the carbonatites (233.6 ± 1.7 Ma). The syenites contain lower initial 87Sr/86Sr and higher eNd values. This indicates that the carbonatites do not have genetically related with the silicate rocks, and were directly derived from a primary carbonate magma generated in the mantle. The carbonatites show a Sr–Nd isotopic signature similar to that of the chondritic uniform reservoir (CHUR), and parallel Sm–Nd model ages (TCHUR) of 190–300 Ma. However, the rocks have extremely variable Pb isotopic values straddling between the HIMU and EM1 mantle end-members. Most carbon and oxygen isotopic compositions of the SQ carbonatites plot outside the field for primary igneous carbonates. Their d13C shows higher value than a ‘normal’ mantle, which implies an incorporation of recycled inorganic carbon. The carbonatites were emplaced close to the Mianlue suture, and followed the closure of the Mianlue ocean and Triassic collision of the South and North China Blocks. However, direct melting of the subducted Mianlue oceanic crust characterized by high eNd and low (EM1-like) 206Pb/204Pb values cannot explain the CHUR-like Nd signature and the Pb isotopic trend toward HIMU in the SQ carbonatites. We conclude that their parental magma was derived from a source incorporating the Mianlue oceanic crust mixed with an asthenospheric (or deeper) material characterized by high Pb and low Nd isotopic values. This material represents a deep-seated Proterozoic carbonate component recycled via mantle convection or localized upwelling. Notably, this model cannot explain the isotopic compositions of the Late Triassic (209–221 Ma) carbonatites in the LQ, characterized by a mantle-derived d13C, but EM1-like Sr–Nd–Pb isotopic compositions. This signature is best explained in terms of delamination of the lower continental crust thickened during the collision of the South and North China Blocks, and partial incorporation of the delaminated material into the LQ mantle source. Modeling of the measured Sr–Nd–Pb isotopic variations suggests that the source of the LQ carbonatites could be produced by mixing of 80–85% of ⇑ Corresponding author. Tel.: +86 1062753894. E-mail address: [email protected] (C. Xu). http://dx.doi.org/10.1016/j.gca.2014.03.041 0016-7037/Ó 2014 Elsevier Ltd. All rights reserved. 190 C. Xu et al. / Geochimica et Cosmochimica Acta 143 (2014) 189–206 mantle material and 15–20% of delaminated lower continental crust. The emplacement of the SQ and LQ carbonatites marked a gradual transition from a compressional tectonic regime, brought about by the collision of the South and North China Blocks to intra-orogenic extension in the waning stages of the Triassic orogeny. Ó 2014 Elsevier Ltd. All rights reserved. 1. INTRODUCTION 2. GEOLOGICAL SETTING AND PETROGRAPHY Carbonatites are mantle-derived carbonate-rich igneous rocks found in extensional settings. Most of the presently known carbonatite occurrences are situated in anorogenic settings related to intracontinental rifting. Collision-related carbonatites are much rarer (Tilton et al., 1998; Hou et al., 2006; Chakhmouradian et al., 2008) and remain poorly understood. These rocks provide valuable information on the composition of the mantle because (1) their isotopic characteristics are generally inherited from the mantle source owing to the initially high concentrations of Sr and rare-earth elements (REE) in carbonatitic magmas (e.g., Nelson et al., 1988), and (2) carbonate melts have extremely low viscosities (Treiman, 1989) facilitating their rapid ascent to the surface. The importance of recycled sediments and other crustal materials as a source for carbonatitic magmatism is disputed (Barker, 1996; Hoernle et al., 2002; Halama et al., 2008; Bell and Simonetti, 2010). Previously published isotopic studies of young carbonatites (<200 Ma) identified a high-l material (HIMU), Enriched Mantle 1 (EM1) and Focal Zone (FOZO) as the major mantle components involved in the generation of carbonatitic magmas (Tilton and Bell, 1994; Simonetti et al., 1998; Bell and Tilton, 2001, 2002). At present, the prevailing view is that these magmas are derived from a lithospheric source affected by asthenospheric upwelling or plumes involving deeper parts of the mantle (e.g., Bell and Tilton, 2001; Tolstikhin et al., 2002). This conventional isotopic framework is not applicable to recently discovered young carbonatites in the Qinling orogenic belt of central China. The Qinling belt is the product of ocean closure and continental collision in the Triassic, and hosts several carbonatite intrusions of Triassic age. These rocks provide critical information about deep mantle sources in the collisional environment and the role of crust recycling in magma generation. Previous studies show that the carbonatites not associated with alkalic silicate rocks from the northernmost Qinling have less radiogenic Nd values than EM1 (Huang et al., 2009; Xu et al., 2011), and do not record simple mixing between the HIMU and Enriched Mantle components. However, the southernmost Qinling carbonatites examined in this study are associated with syenites and characterized by slightly depleted isotopic compositions. Their origin and genetic relationship with the syenites are unknown. Thus, our primary objective is to develop alternative models of carbonatite petrogenesis in collisional settings that will explain both the structural relations and unusual isotopic signatures of the Qinling rocks. In a broader context, correct interpretation of the geodynamic evolution of the Qinling belt gleaned from isotopic data is important in unraveling the tectonic history of East Asia. The Qinling zone is linked to the Dabie-Sulu ultrahigh-pressure metamorphic belt to the east (Fig. 1) and separates the North China Block from the South China Block. Closure of the Mianlue Ocean in the Triassic resulted in collision of the South China Block with the South Qinling belt (SQ) along the Mianlue suture, followed by amalgamation of the South and North China Blocks (Meng and Zhang, 1999; Ratschbacher et al., 2003). The southern margin of the North China Block was deformed in the process to produce the Lesser Qinling belt (LQ), which, together with the North Qinling belt (NQ), forms a passive continental margin (Xue et al., 1996). The Qinling orogen is bounded by the Lingbao–Lushan–Wuyang fault in the north and the Mianlue–Bashan–Xiangguang fault in the south. Two suture zones incorporating ophiolitic mélange, the Shangdan suture in the north and the Mianlue suture in the south, are well documented (Meng and Zhang, 1999; Ratschbacher et al., 2003). The orogen is divided into the aforementioned SQ, NQ and LQ terranes, which are separated by the two sutures and several major faults (Fig. 1). The LQ is characterized by highly deformed, high-grade metamorphic basement rocks of the Mesoarchean Taihua group and granite-greenstone assemblages of the Neoarchean Dengfeng group (Kröner et al., 1988). The carbonatites intruded the Taihua group, which comprises amphibolite- to granulite-facies metamorphic rocks discordantly overlain by Mesoproterozoic volcanics. In the SQ, the Proterozoic basement consists of the Paleoproterozoic Douling group comprising paragneiss and granitic gneiss, Paleoproterozoic- to Neoproterozoic Wudang and Yunxi groups composed of alkali basalt, alkali trachyte, dacite, rhyolite and pyroclastic rocks, and the Neoproterozoic Yaolinghe group made up of tholeiitic basalt, spilitic diabase, alkali trachyte and spilite (Huang, 1993). Many occurrences of carbonatites have now been identified in this part of China; these rocks were emplaced into the deformed southern margin of the North China Block in the LQ and into the northern margin of the South China Block in the SQ. The SQ carbonatites are tectonically confined to the southern margin of the Mianlue suture. They were emplaced into syenitic rocks as stock and minor dikes, the largest of which covers 0.3 km2 at the current erosion level. The syenites were altered and impregnated by numerous calcite-bearing veinlets. The rock is composed predominantly of alkali feldspar and plagioclase (Fig. 2a); minor and accessory phases include biotite, amphibole, clinopyroxene, quartz, sulfides, calcite, fluorapatite, barite, REE minerals, rutile and zircon. The SQ carbonatites are composed predominantly of medium- to fine-grained calcite. Minor and accessory phases include biotite, feldspar, fluorapatite, ilmenite, pyrite, barite, fluorite, C. Xu et al. / Geochimica et Cosmochimica Acta 143 (2014) 189–206 191 Fig. 1. The location and geological setting of the carbonatites in the South Qinling (SQ) and Lesser Qinling orogenic belts (LQ-S in Shanxi, LQ-H in Henan; modified after Xu et al., 2010b, 2011). quartz, monazite, bastnäsite, synchysite, ferrocolumbite and rutile. Monazite is an early-crystallising mineral in the carbonatites and occurs as ovoid grains up to 200 lm across. The crystal does not exhibit patchy zoning (Fig. 2c), and contains little compositional variation and constant (La/Nd)CN ratio of 3 (Xu et al., 2010a). The ages of the SQ carbonatites and syenites have not been well constrained prior to the present work. In the LQ, the carbonatites are situated in the central part of the Shanxi Province (LQ-S) and in the western part of the Henan Province (LQH). They occur as dikes containing abundant sulfide mineralization including molybdenite, and are not associated with alkalic silicate rocks. Rhenium-Os measurements on molybdenite from LQ-S and LQ-H samples gave ages of 221.5 Ma (Du et al., 2004) and 209.5 Ma (Huang et al., 2009), respectively. The detailed geological and petrological information on LQ carbonatites was reported by Xu et al. (2007, 2011) and Huang et al. (2009). 3. ANALYTICAL METHODS Because the SQ carbonatite lacks zircon or baddeleyite, monazite was chosen for a radiochronological study of the rock. Its age was determined by U–Th–Pb methods using a sensitive high-resolution ion microprobe (SHRIMP) at the Curtin University, Australia. The mineral was separated from fresh samples using standard techniques, cast into 25 mm resin mounts together with grains of 425-Ma reference material 44069 (Aleinikoff et al., 2006) and polished to expose flat sections. The primary beam diameter was 25 lm, and primary O2 beam current was 1.5 nA. A mass resolution (M/DM at 1% of peak height) was 5000, and each analysis consisted of six scans of the mass spectrum. Data were acquired in a single analytical session, following the procedures described in Foster et al. (2000) and Fletcher et al. (2010). The primary data ratios were 206Pb+/270[UO2]+ and 208Pb+/264[ThO2]+, both of which were subjected to 1-dimensional calibrations (Fletcher et al., 2010) with standard 44069 as the primary age reference. Additional standard materials French, z2234 and 2908 provided reference U, Th and Y concentrations. Corrections for matrix effects in Pb/U and Pb/Th were made subsequently, in spreadsheet templates, using matrix correction factors from Fletcher et al. (2010). Data from French, z2234 and z2908 provided first-order confirmation of the correction factors for this session. 192 C. Xu et al. / Geochimica et Cosmochimica Acta 143 (2014) 189–206 a b c Fig. 2. Photomicrographs of the South Qinling syenites (a) and carbonatites (b, c). Kfs, K-feldspar; Ab, albite; Aug, augite; Bt, biotite; Ap, fluorapatite; Cal, calcite; Mn, monazite. Zircon crystals were extracted from representative syenite samples using conventional magnetic and heavy-liquid separation techniques, and then mounted in epoxy and polished. The crystals were examined using cathodoluminescence (CL) imaging to reveal their internal structure. Age determination was done using laser-ablation inductivelycoupled-plasma mass-spectrometry (LA-ICPMS). For LA-ICPMS measurements, we used an Agilent 7700 mass-spectrometer coupled to a 193-nm ArF excimer laser at the Institute of Geochemistry, Chinese Academy of Sciences (CAS). The laser beam was focused on the sample with a fluency of 10 J/cm2 and a spot of 32 lm in diameter at a repetition rate of 5 Hz for 40 s. Helium was used as a carrier gas to transport the ablated aerosol to the massspectrometer. Zircon Nancy 91500 was used as an external calibration standard to correct for instrumental mass bias and elemental fractionation. Zircon standards GJ-1 and Plešovice were employed for quality control. The Pb content of zircon was externally calibrated against NIST 610 with Si as an internal standard, whereas other trace elements were measured with Zr as an internal standard. Raw-data reduction was performed off-line using the ICPMSDataCal software (Liu et al., 2010). The trace-element composition of syenite samples was determined by solution ICPMS (ELAN DRC-e) at the Institute of Geochemistry, CAS. The analytical protocol for trace elements was described in detail in Liang et al. (2000). Replicate analyses (including well-characterised standards) indicated that the accuracy of trace element measurements was better than 10%. The Rb, Sr, Sm, Nd, U, Th and Pb concentrations of carbonatites and their constituent minerals were analyzed by solution ICPMS (Thermo Fisher Scientific X-Series II) in the School of Ocean and Earth Science at the University of Southampton (UK). The instrument was calibrated using five international rock standards, BIR-1, JGb-1, JB-3, BHVO-2 and JB-1a, which were run before, during, and after each group of analyses. The data processing procedure included linear drift correction, internal (matrix) correction, REE and Ba interference corrections, blank subtraction, calibration with international standards, and a dilution correction. In-run precision was routinely in the 1–5% range, and accuracy relative to the reference materials 5%. In situ LA-ICPMS determinations of Rb, Sr, U, Th and Pb concentrations of calcite were performed in polished sections at the Australian National University (ANU) using an analytical methodology similar to that described below for the Pb isotopic measurements. The calculation of element concentrations was like that reported by Eggins et al. (1997), and detection limits were determined using the approach of Longerich et al. (1996). Analytical precision was <5% at the ppm level. Carbon and oxygen isotopic compositions of calcite from the SQ carbonatites and associated veinlets in the syenites were measured at the Institute of Geochemistry, CAS, using a continuous-flow isotope ratio mass spectrometer (IsoPrime). The results were expressed in per mil (&) using the conventional d-notation, i.e., d = (R1/R2 1) 1000, where R1 is the 13C/12C or 18O/16O ratio in the sample and R2 is the corresponding isotope ratios in the standards C. Xu et al. / Geochimica et Cosmochimica Acta 143 (2014) 189–206 (V-PDB for C and V-SMOW for O). Reproducibilities of d13C and d18O values, measured as one standard deviation for the carefully selected reference materials (GBW04405, GBW04406, GBW04417), were 0.15& and 0.2&, respectively. The Sr and Nd isotopic compositions of the selected samples from the SQ carbonatites were analyzed by a seven-collector VG Sector 54 mass-spectrometer with a separable-filament source in the School of Ocean and Earth Science, University of Southampton. Strontium and Nd isotopic ratios were determined as the average of >100 ratios by measuring ion intensities in multidynamic collection mode. Mass fractionation corrections for Sr and Nd isotopic ratios were normalized to 86Sr/88Sr = 0.1194 and 146 Nd/144Nd = 0.7219, respectively. Total blanks were <1.1 ng for Sr and <<0.2 ng for Nd, which represent <0.05% of the mass of both elements in the measured fraction. Repeated measurements for the Nd standard JMC321 and Sr standard NBS987 yielded 143Nd/144Nd = 0.511125 ± 11 (2r, n = 45) and 87Sr/86Sr = 0.710252 ± 11 (2r, n = 169), respectively. These values are similar to the commonly accepted reference values (e.g., Jochum et al., 2005). Strontium isotopic composition of calcites from the SQ carbonatites and veinlets in syenites was analyzed by LAMC-ICPMS using a Thermo Neptune instrument at the ANU. Two calcite crystals from the same hand specimen were embedded in a polished epoxy mount with their cores exposed. The spot diameter used for the measurement was 178 lm for the calcite and 233 lm for the Tridacna standard. The average 87Sr/86Sr ratio obtained for the standard was 0.70913, i.e., close to the accepted in-house value of 0.70917. Lead fractions were separated and purified using HBr and anion exchange resin and then analyzed by MC-ICPMS (Thermo Neptune) in the School of Ocean and Earth Science at the University of Southampton. Doublespiking (Southampton–Brest–Lead 207–204 spike, SBL-74) was used to correct for instrumental mass bias on the measurements. The Pb isotopic composition of the samples was obtained from the natural and mixture runs by iterative calculation adopting an exponential mass-bias correction. Repeated measurements for Pb standard NBS981 gave 206 207 Pb/204Pb = 16.9404 ± 25, Pb/204Pb = 15.4973 ± 20 and 208Pb/204Pb = 36.7173 ± 65 (2r, n = 32). Total procedural blanks were Pb < 50 pg, which is considered negligible relative to the sample loads and analyzed concentrations (>300 ng Pb). The Pb isotopic composition of calcites from the SQ carbonatites was also measured in polished thin sections using a HP7500 Agilent LA-ICPMS at the ANU. The diameter of the ablation spot was 54 lm. The NIST 610 glass was used as a calibration standard for all samples, with 43Ca as an internal standard for quantitative analysis. In-run signal intensities for 206Pb, 207Pb and 208Pb were monitored during the measurement to make sure that the laser beam stayed within the grain chosen for analysis and did not penetrate inclusions. The 204Pb content could not be quantified because of systematic Hg contamination. The 207Pb/206Pb and 208Pb/206Pb ratios in the calcites were calculated so that Pb concentration ratios estimated from 206 Pb, 207Pb and 208Pb measurements were multiples of 193 0.9098 (207Pb/206Pb value in the NIST 610 calibration standard) and 2.169 (208Pb/206Pb value in NIST 610: Jochum et al., 2005), respectively. 4. RESULTS The SHRIMP and LA-ICPMS U–Th–Pb data for monazite and zircon grains from the SQ carbonatites and syenites, respectively, are given in Tables 1 and 2 and shown in Fig. 3. Trace elements and C–O–Sr–Nd–Pb isotopic values for the examined syenites and carbonatites are listed in Tables 3–6 and summarized in Figs. 4–6. 4.1. U–Th–Pb geochronology Monazite is widely used as a U–Th–Pb geochronometer to determine the magmatism timing (e.g., Harrison et al., 1995). Experiment shows that a 10 lm-sized monazite grain would have a Pb closure temperature in excess of 900 °C (Cherniak et al., 2004). Carbonatite magmas contain abundant volatiles, which result in low crystallization temperature (Jago and Gittins, 1991). The carbonatite lava flows in the crater of the volcano Oldoinya Lengai (Tanzania) show extremely low eruption temperature of 544–590 °C (Jago and Gittins, 1991). Therefore, the U–Th–Pb system in monazite is a valuable chronometer for dating carbonatite magmatism. The common Pb contents of the analyzed monazite samples from the SQ carbonatites are quite variable; all analyses with >2% common Pb (either 208Pb or 206 Pb) were excluded from further consideration, even though these discarded data are generally consistent with the retained analyses, implying that the chosen 2% cut-off limit is sufficiently conservative. Only three 208Pb/232Th measurements gave “high” common Pb (2.2–3.3%) and were rejected. Using the Isoplot software (Ludwig, 2003), the 16 analyzed grains gave a weighted mean age of 235.5 ± 2.1 Ma with MSWD = 0.91. A similar assessment of the 206Pb/238U data gave a weighted mean age of 231.5 ± 2.6 Ma with MSWD = 1.2. Since the 208Pb/232Th and 206Pb/238U data were calibrated independently, they can be combined, giving an age of 233.6 ± 1.7 Ma with MSWD = 1.3 for the main phase of monazite crystallization (Fig. 3a). The remaining two monazite grains appear to have been partially recrystallized between 220 and 190 Ma, but the outliers show too much scatter to indicate a distinct datable post-emplacement recrystallization (growth?) episode other than the main 235 Ma result. The younger monazite contains slightly higher U content and has a lower average Th/U ratio relative to the primary generation, possibly indicating a crustal input. According to the Isoplot results for several monomineralic fractions (biotite, ferrocolumbite, pyrite, fluorapatite and monazite) from the SQ carbonatites and whole-rock measurements, the 206Pb/204Pb vs. 207Pb/204Pb (“Pb–Pb”) isochron gave a combined age of 262 ± 25 Ma with MSWD = 357 (Fig. 3b). This value is (within error) in general agreement with the SHRIMP monazite age. Zircons from the syenites are euhedral oscillatory-zoned crystals >70 lm in length (Fig. 3c). Twenty-four LAICPMS analyses of the zircon crystals show variable Th 194 C. Xu et al. / Geochimica et Cosmochimica Acta 143 (2014) 189–206 Table 1 SHRIMP U–Pb and Th–Pb data from monazites in South Qinling carbonatites, central China. Analysis U (ppm) Th (ppm) Self-consistent Th–Pb group, in MY-01.4 173 4236 MY-01.12 145 2175 MY-01.26 76 2789 MY-01.19 99 1776 MY-01.15 87 2895 MY-01.21 249 1524 MY-01.20 43 1082 MY-01.23 111 2795 MY-01.9 101 2049 MY-01.24 135 2010 MY-01.25 64 2839 MY-01.7 94 2064 MY-01.11 136 1007 MY-01.2 114 2396 MY-01.6 39 1392 MY-01.10 49 3636 Th/U 4f206 (%) 206 Pb/ U ±s 238 Pb/232Th sequence 25 1.86 0.0355 15 0.24 0.0367 37 1.63 0.0359 18 2.60 0.0362 33 3.11 0.0349 6.1 0.07 0.0373 25 1.36 0.0370 25 6.97 0.0368 20 1.51 0.0371 15 1.55 0.0373 45 6.13 0.0378 22 2.59 0.0364 7.4 0.33 0.0363 21 1.87 0.0367 36 1.00 0.0365 75 2.84 0.0379 t[206Pb/ 238 U] (Ma) ±s 4g208 (%) 208 Pb/ Th ±s t[208Pb/ 232 Th] (Ma) ±s 232 208 0.0006 0.0006 0.0007 0.0007 0.0007 0.0006 0.0009 0.0008 0.0007 0.0007 0.0009 0.0007 0.0006 0.0007 0.0009 0.0010 224.6 232.3 227.1 228.9 221.0 236.4 234.1 232.8 234.6 236.3 239.3 230.4 229.8 232.1 231.1 240.1 3.8 3.9 4.5 4.3 4.5 3.6 5.4 5.1 4.4 4.0 5.6 4.6 3.9 4.4 5.6 5.9 0.49 0.11 0.29 0.95 0.59 0.07 0.36 1.90 0.49 0.69 0.96 0.76 0.28 0.57 0.18 0.26 0.0114 0.0114 0.0115 0.0116 0.0116 0.0116 0.0117 0.0117 0.0117 0.0118 0.0118 0.0119 0.0119 0.0119 0.0120 0.0121 0.0002 0.0002 0.0002 0.0002 0.0002 0.0002 0.0002 0.0002 0.0002 0.0002 0.0002 0.0003 0.0002 0.0002 0.0002 0.0002 229.1 229.8 230.9 232.3 233.7 233.9 234.6 234.8 235.5 236.8 237.4 238.7 238.8 239.8 241.3 242.7 4.2 4.0 4.2 4.4 4.5 3.7 3.9 4.4 4.4 4.2 4.0 5.0 4.0 4.1 5.0 4.0 Young outliers MY-01.1 358 MY-01.17 375 1929 2258 5.4 6.0 0.11 0.49 0.0302 0.0341 0.0005 0.0005 191.9 216.0 3.3 3.2 0.13 0.53 0.0094 0.0109 0.0002 0.0002 189.0 218.4 4.1 3.8 >2% common 208Pb MY-01.18 49 MY-01.13 300 MY-01.5 83 761 987 587 15 3.3 7.1 5.17 1.26 3.78 0.0361 0.0418 0.0353 0.0009 0.0047 0.0008 228.8 263.7 223.8 5.8 28.9 4.9 2.25 2.78 3.30 0.0115 0.0117 0.0117 0.0002 0.0004 0.0002 230.4 235.5 235.2 4.2 7.7 4.7 4f206 [4g208] is the proportion of 206Pb [208Pb] calculated to be common Pb on the basis of measured 204Pb and modeled common Pb composition (Stacey and Kramers, 1975) at the approximate sample age. All listed Pb isotope data are corrected for common Pb, based on measured 204Pb/206Pb. Listed uncertainties are 1r and include all components of statistical precision but not the uncertainties in the mean of the data for the primary reference material 44069 of Aleinikoff et al. (2006) or uncertainties in matrix corrections. and U concentrations (89–800 ppm and 112–507 ppm, respectively), with a Th/U ratio of 0.6–1.6. Owing to subsolidus re-equilibration and Pb loss, these analyses yielded a discordant U–Pb age of 766 ± 25 Ma (MSWD = 2), and only four samples gave a concordant age of 749 ± 11 Ma (MSWD = 1.6). 4.2. Trace element composition The SQ syenites have highly variable trace element abundances, particularly for Ba, Th, U, Nb, Ta, Pb Zr, Hf and REE. They are characterized by negative Ta, Sr, Hf, Sc and Cr anomalies in the primitive mantle-normalized trace element diagram (Fig. 4a). Positive Ba and Nb anomalies are observed in some samples. The chondrite-normalised REE patterns exhibit LREE enrichment with (La/ Yb)CN ratios ranging from 19 to 56, and negligible to small positive Eu anomalies (Eu/Eu* = 0.8–1.8; Fig. 4b). Trace element composition of the SQ carbonatites, reported by Xu et al. (2010a), shows enrichment in Sr and LREE, and negative Zr–Hf, Pb, Sc and Cr anomalies in Fig. 4a. The chondrite-normalised REE patterns of the carbonatites (Fig. 4b) show a steep negative slope and negligible Ce and Eu anomalies. Compared to the carbonatites, the syenites contain lower level of Sr, and greater enrichment in Rb and Ba. Both rock types have similar Sc, V, Ni, Cr and Ga contents. 4.3. C–O–Sr–Nd–Pb isotopic compositions The carbon and oxygen isotopic compositions of calcite separates from the SQ carbonatites show significant variation, ranging from 3.54& to 6.96& d13 CV-PDB and from 8.62& to 14.36& d18OV-SMOW, respectively (Fig. 5). Most of these data plot outside of the field of primary, unaltered carbonatites identified by Taylor et al. (1967). The d18OV-SMOW values of all SQ samples are higher than the normal range of mantle values, and some samples show higher d13CV-PDB values in comparison with the typical mantle range. The Sr and Nd isotopic results for the SQ carbonatites and their constituent minerals (calcite, biotite, pyrite, ferrocolumbite, fluorapatite and monazite) show low level of radiogenic Sr [(87Sr/86Sr)i 0.7036–0.7040] and little variation in eNd values (from +0.6 to 1.1) approaching to the Chondritic Uniform Reservoir (CHUR in Fig. 6a). The model age (TCHUR, Sm–Nd) is mainly in the range of 0.19–0.30 Ga (Table 5). All analyzed carbonatite samples show a wide range of initial Pb isotopic values. Some agecorrected ratios may be associated with large uncertainties for the high U, Th contents and 238U/204Pb, 232Th/204Pb ratios (Table 5). Not uncommonly, the applied corrections yield unreasonably low or high ratios, resulting in initial Pb isotopic ratios outside of the ‘normal’ mantle range. The primary carbonatitic calcite, analyzed in situ by Table 2 LA-ICPMS U–Pb data from zircons in the South Qinling syenites. Th (ppm) U (ppm) Th/U 207 Pb/206Pb ±1r 207 Pb/235U ±1r 206 Pb/238U ±1r q 207 Pb/206Pb Age (Ma) ±1r 207 Pb/235U Age (Ma) ±1r 206 Pb/238U Age (Ma) ±1r 66.3 23.2 16.6 48.6 35.1 31.4 23.5 49.0 26.4 77.5 60.0 29.4 27.4 44.9 33.5 58.0 35.8 24.8 27.1 51.9 32.3 62.1 93.1 23.5 352 88.9 108 235 181 156 95.8 322 130 485 457 158 141 215 137 369 190 107 125 273 144 345 800 125 376 160 112 289 209 204 153 283 164 448 370 188 178 272 220 381 214 162 175 332 217 390 507 146 0.93 0.56 0.96 0.81 0.86 0.76 0.62 1.14 0.79 1.08 1.23 0.84 0.79 0.79 0.62 0.97 0.89 0.66 0.72 0.82 0.66 0.89 1.58 0.86 0.0645 0.0645 0.0683 0.0646 0.0648 0.0657 0.0661 0.0652 0.0668 0.0682 0.0674 0.0679 0.0657 0.0635 0.0651 0.0668 0.0655 0.0648 0.0668 0.0667 0.0648 0.0640 0.0649 0.0683 0.0009 0.0013 0.0015 0.0009 0.0010 0.0010 0.0010 0.0011 0.0011 0.0008 0.0010 0.0013 0.0013 0.0010 0.0010 0.0010 0.0010 0.0010 0.0012 0.0009 0.0010 0.0009 0.0009 0.0014 1.1309 1.0050 1.0476 1.0935 1.0862 1.0281 1.0621 1.0504 1.0767 1.1122 1.0581 1.0929 1.0705 1.1018 1.0598 1.0412 1.0995 1.0375 1.0855 1.0381 1.0264 1.0203 1.0273 1.0826 0.0182 0.0220 0.0244 0.0148 0.0172 0.0156 0.0153 0.0177 0.0167 0.0126 0.0179 0.0213 0.0212 0.0172 0.0165 0.0157 0.0175 0.0151 0.0192 0.0135 0.0166 0.0138 0.0141 0.0199 0.1260 0.1123 0.1109 0.1220 0.1211 0.1127 0.1159 0.1163 0.1168 0.1177 0.1129 0.1167 0.1180 0.1252 0.1170 0.1125 0.1209 0.1161 0.1181 0.1128 0.1144 0.1150 0.1140 0.1158 0.0012 0.0011 0.0013 0.0008 0.0010 0.0008 0.0008 0.0009 0.0008 0.0008 0.0008 0.0011 0.0010 0.0010 0.0008 0.0010 0.0009 0.0008 0.0011 0.0007 0.0008 0.0006 0.0007 0.0009 0.5877 0.4319 0.5036 0.4877 0.4978 0.4546 0.4733 0.4476 0.4652 0.5734 0.4402 0.4804 0.4402 0.5189 0.4491 0.5798 0.4429 0.4990 0.5331 0.4664 0.4406 0.4182 0.4329 0.4427 767 767 880 761 769 798 809 789 831 876 850 865 794 724 789 833 791 766 833 828 769 743 769 880 31 39 46 29 228 30 25 37 33 25 32 36 44 33 33 169 33 27 41 28 28 25 28 43 768 706 728 750 747 718 735 729 742 759 733 750 739 754 734 725 753 723 746 723 717 714 718 745 9 11 12 7 8 8 8 9 8 6 9 10 10 8 8 8 8 8 9 7 8 7 7 10 765 686 678 742 737 688 707 710 712 717 690 712 719 761 713 687 736 708 720 689 699 702 696 707 7 6 8 5 5 4 5 5 5 4 5 6 6 6 5 6 5 5 6 4 5 4 4 5 C. Xu et al. / Geochimica et Cosmochimica Acta 143 (2014) 189–206 MY@1 MY@2 MY@3 MY@4 MY@5 MY@6 MY@7 MY@8 MY@9 MY@10 MY@11 MY@12 MY@13 MY@14 MY@15 MY@16 MY@17 MY@18 MY@19 MY@20 MY@21 MY@22 MY@23 MY@24 Pb (ppm) 195 196 C. Xu et al. / Geochimica et Cosmochimica Acta 143 (2014) 189–206 a b c Fig. 3. Weighted average plots of 206Pb/238U (black line) and 208Pb/232Th (grey line) ages derived from monazite SHRIMP data (a), and whole-rock and mineral Pb–Pb isochron age (b) for the South Qinling carbonatites, and LA-ICPMS analytical U–Pb data for zircons in the South Qinling syenites with cathodoluminescence images of representative zircons (c). In b, Bt = biotite, Mn = monazite, Ct = ferrocolumbite, Ap = fluorapatite, Py = pyrite; the rest of the data are whole-rock samples. The large MSWD is related to small measurement errors produced by the double spike method stemming from an increase in the error of regression when the isochron does not pass through the error range for each sample; MSWD would be significantly lower if the errors for each point were magnified. These data point to either an underestimated U/Pb ratio, or a geological phenomenon disturbing the primary isochron. LA-ICPMS, contains tens of ppm Pb, whereas its U and Th abundances are at or below their detection limit (Table 6), suggesting that a radiogenic contribution to the Pb budget of this mineral is negligible. Hence, the Pb budget of the primary calcite is an accurate measure of the composition of its parental magma, which in turn, reflects the composition of its mantle source(s). The examined calcite samples show a wide range of 207Pb/206Pb (0.74–0.88) and 208Pb/206Pb (1.79–2.15) ratios (Table 6), forming a nearly linear array between the EM1 and HIMU end-members (Fig. 6b) similar to the so-called East African Carbonatite Line (EACL). However, combined with Sr–Nd–Pb isotopic data, the SQ carbonatites form an anomalous isotope pattern (Fig. 6c, d), which does not follow the HIMU–EM1 mixing line. Note that, whereas the Sr and Nd isotopic values for the Bulk Silicate Earth reservoir (BSE) can be estimated with reasonable accuracy for 235 Ma, the ratios for DMM, HIMU, EM1 and EM2 end-members cannot be constrained with the same degree of confidence because of the absence of ocean-island volcanics older than 190 Ma. Thus, we cannot exactly evaluate the isotopic data for the sources in terms of specific mantle endmembers. The Sr–Nd isotopic compositions of the SQ syenites have strongly positive eNd (t) value of 9.6–10.3 and low radiogenic initial 87Sr/86Sr of 0.7011–0.7034. The Sm–Nd model age of TCHUR does not fit the syenites with quite low value of 0.03–0.11 Ga. The Sr–Nd–Pb isotopic compositions of carbonatites from the LQ areas, reported by Huang et al. (2009) and Xu et al. (2011), approach, and trend toward less radiogenic Sr and Nd values than, the EM1 end-member (Fig. 6a). Class et al. (2009) argued that EM1 and EM2 components with narrowly defined isotopic compositions may not exist, and individual island or volcano with “enriched mantle affinity” may instead form a trend toward its own unique end-member. The available data for the SQ and LQ carbonatites plot as clearly distinct clusters (Fig. 6a, c and d), rather than forming a mixing pattern involving the HIMU and EM1 end-members. In contrast, Cenozoic carbonatites from northwestern Pakistan and the Chinese Panxi region in the Indo-Asian collision zone conform to the well-established C. Xu et al. / Geochimica et Cosmochimica Acta 143 (2014) 189–206 197 Table 3 Trace element compositions (ppm) of the South Qinling syenites (MY-S17–S20) and carbonatites. Sc V Cr Ni Ga Rb Sr Y Zr Nb Ta Ba La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Hf Pb Th U MY-S17* MY-S18 MY-S19 MY-S20 Carbonatites n = 5* 4.67 190 4.39 22.0 25.0 113 702 25.2 65.7 1151 2.95 15169 154 360 29.5 97.2 12.7 3.59 12.1 1.06 4.47 0.84 2.40 0.31 1.94 0.29 1.93 25.2 10.9 2.24 0.67 101 3.29 11.7 28.7 123 648 86.8 390 387 9.80 1550 462 800 70.9 253 37.1 12.2 43.0 4.58 20.3 3.81 9.67 1.17 7.27 0.92 4.41 37.2 52.5 11.7 4.18 142 1.17 4.74 19.1 31.4 520 18.3 79.6 242 1.39 2520 147 306 32.0 111 14.9 3.77 13.5 1.06 3.83 0.75 2.22 0.27 1.77 0.25 1.26 11.0 5.07 1.27 2.68 40.7 0 3.34 15.9 65.6 1120 16.8 84.6 77.1 1.07 18400 42.2 79.7 9.03 36.0 7.57 5.0 9.44 0.70 3.43 0.70 1.86 0.24 1.48 0.20 2.29 11.7 5.11 0.45 2.33 61.0 2.85 9.44 17.9 11.4 5675 113 62.1 240 5.04 859 503 880 67.2 235 35.8 10.6 33.9 4.08 20.8 3.94 10.7 1.38 8.14 1.15 0.68 21.7 21.2 9.23 Data of Rb, Sr, Ba, Zr, Hf, Nb, Ta, Pb, Th, U and REE from * samples were reported by Xu et al. (2010a). Table 4 C–O isotopic compositions of calcites from the South Qinling carbonatites and calcite veinlets (MY-17) in the syenite. Sample MY-0 MY-01 MY-1 MY-4 MY-5 MY-6 MY-7 MY-07 MY-8 MY-10 MY-11 MY-12 MY-13 MY-14 MY-15 MY-16 MY-17 d13CPDB 5.81 4.40 6.96 4.76 3.63 5.30 5.62 6.46 3.78 5.68 4.35 3.54 3.79 4.20 4.49 4.29 4.14 (&) d18OSMOW 11.04 11.60 9.22 12.21 14.36 12.04 10.92 8.62 10.51 12.49 11.87 11.69 12.43 12.23 11.0 11.50 10.94 (&) HIMU–EM1 and EM1–EM2 trends, respectively, and show isotopic compositions different with respect to the SQ and LQ rocks. 5. DISCUSSION 5.1. Are the SQ carbonatites and syenites genetically related? The SQ carbonate rocks were emplaced intrusively into the syenites. In common with calcite carbonatites from other localities (e.g., Hornig-Kjarsgaard, 1998; Xu et al., 2007; Chakhmouradian et al., 2008, among many others), the calcite and fluorapatite in the examined carbonate rocks are enrichment in Sr and REE (Xu et al., 2010a). The Y/Ho ratio (average 28.6) of the whole rocks resembles the primary mantle value (28.9; McDonough and Sun, 1995). The rocks contain mantle-derived Sr–Nd isotopes. Thus, our field observations and geochemical data support the interpretation that the examined carbonate rocks are bona fide carbonatites rather than marble rafts metasomatised by syenite intrusions. Models for carbonatite genesis have been debated and essentially consider two options (e.g., Gittins, 1989): (1) that carbonatites evolved as a secondary magma during differentiation of a mantle-derived silicate parental melt; or (2) that carbonatites were derived directly from a mantlederived primary carbonate magma. The first option has dominated because of close spatial relationship between carbonatites and alkalic silicate rocks. Although the SQ carbonatites are intimately associated with the syenites, the radiochronological data obtained in the present work imply that the two rock types are not derived from a common parental magma. Zircon crystals from the syenites define a U–Pb age of 760 Ma, and clearly different from 198 Table 5 Sr–Nd–Pb isotopic compositions of carbonatites and their constituent minerals and syenites (MY-S17–S19) from the South Qinling orogen. Type Rb (ppm) Sr (ppm) 87 MY-0 6.96 0.36 6162 9631 0.19 239 0.35 36.5 0.75 4.06 my-7s my-8s my-9s MY-13 WR Ap WR WR Bt Ap WR Mn Ct WR WR Bt Ct WR Py Py WR MY-15 WR MY-16 WR MY-17 CV MY-S17 MY-S18 MY-S19 Samples MY-0s MY-01 MY-5 MY-5s MY-7 MY-0 MY-0s MY-01 MY-5 MY-5s MY-7 87 (87Sr/86Sr)i Sm (ppm) Nd (ppm) 147 143 eNd (t) TCHUR (Ga) 0.003 <0.001 0.703650 ± 5 0.703586 ± 6 0.70364 0.70359 3088 625 9800 4009 293 205 <0.001 1.090 <0.001 0.026 0.007 0.057 0.703651 ± 21 0.707534 ± 4 0.703610 ± 18 0.704118 ± 9 0.704050 ± 12 0.704235 ± 7 0.70365 0.70389 0.70361 0.70403 0.70403 0.70405 27.1 14.5 4.13 3724 5338 291 0.021 0.008 0.041 0.703773 ± 5 0.703808 ± 9 0.704189 ± 8 0.70370 0.70378 0.70405 72.1 156 1620 n=9 2327 n=3 3410 n=2 2842 n=6 702 648 520 0.006 0.005 <0.001 0.703685 ± 7 0.703780 ± 5 0.70392* ± 23 (n = 2) 0.70366 0.70376 0.70392 453 867 106 910 163 843 256 54410 770 161 288 397 311 203 122 105 278 0.0817 0.1032 0.0981 0.0738 0.0835 0.1026 0.0912 0.0721 0.0919 0.0992 0.0928 0.0961 0.1003 0.0965 0.0729 0.0732 0.0912 0.512422 ± 4 0.512480 ± 4 0.512446 ± 4 0.512394 ± 3 0.512434 ± 5 0.512478 ± 3 0.512452 ± 3 0.512409 ± 3 0.512434 ± 4 0.512462 ± 4 0.512483 ± 4 0.512474 ± 5 0.512464 ± 7 0.512481 ± 4 0.512415 ± 4 0.512418 ± 4 0.512507 ± 6 0.76 0.28 0.79 1.07 0.58 0.30 0.45 0.73 0.83 0.51 0.09 0.18 0.50 0.06 0.63 0.58 0.61 0.29 0.26 0.30 0.30 0.27 0.26 0.27 0.28 0.30 0.28 0.23 0.25 0.28 0.24 0.28 0.27 0.19 <0.001 0.70361* ± 2 (n = 2) 0.70361 35.7 245 0.0881 0.512487 ± 6 0.31 0.21 <0.001 0.70381* ± 28 (n = 2) 0.70381 33.8 224 0.0913 0.512482 ± 5 0.12 0.23 <0.001 0.70373* ± 1 (n = 2) 0.70373 WR WR WR 0.16 0.29 0.01 n=9 0.01 n=3 bdl n=2 0.01 n=6 113 123 31.4 61.2 148 17.2 111 22.5 143 38.5 6485 117 26.4 44.2 63.1 51.6 32.4 14.7 12.7 41.9 0.460 0.549 0.172 0.705963 ± 14 0.708671 ± 14 0.705284 ± 27 0.70107 0.70271 0.70341 12.7 37.1 14.9 97.2 253 111 0.0790 0.0886 0.0812 0.512577 ± 21 0.512614 ± 9 0.512552 ± 20 10.3 10.1 9.6 0.08 0.03 0.11 Type U (ppm) Th (ppm) 238 235 232 206 207 208 204 204 204 204 204 204 204 (206Pb/ Pb)i (207Pb/ 204 Pb)i (208Pb/ 204 Pb)i (207Pb/ 206 Pb)i (208Pb/ 206 Pb)i 74.5 172 41.1 4.63 650 229 60.1 431 472 91.7 320 3909 0.54 1.25 0.30 0.03 4.71 1.66 0.44 3.13 3.42 0.67 2.32 28.3 28.0 129 7.43 17.3 294 44.2 12.9 1.46 36.5 44.1 21.5 217 25.034 ± 1 30.632 ± 3 35.418 ± 3 18.118 ± 1 33.367 ± 2 24.074 ± 1 23.813 ± 1 48.960 ± 3 39.396 ± 3 24.984 ± 1 34.603 ± 5 107.32 ± 24 15.906 ± 1 16.167 ± 2 16.385 ± 1 15.527 ± 1 16.317 ± 1 15.869 ± 1 15.854 ± 1 17.138 ± 1 16.662 ± 2 15.917 ± 1 16.418 ± 2 20.130 ± 1 38.381 ± 3 41.540 ± 6 43.110 ± 5 44.352 ± 1 43.223 ± 4 38.247 ± 1 38.491 ± 3 208.403 ± 1 38.642 ± 5 39.403 ± 3 39.111 ± 1 40.382 ± 13 22.270 24.279 >30 17.946 <10 15.596 21.580 >30 21.872 21.574 22.722 <0 15.765 15.844 16.307 15.518 <15.4 15.438 15.740 >16 15.771 15.744 15.815 <15 38.054 40.029 43.023 44.150 39.786 37.731 38.340 >50 38.215 38.887 38.860 37.846 0.708 0.653 1.709 1.649 0.865 2.460 0.990 0.729 2.419 1.777 0.721 0.730 0.70 1.747 1.803 1.710 WR Ap WR WR Bt Ap WR Mn Ct WR WR Bt 20.7 29.7 7.31 2.07 37.2 30.9 39.3 144 2,151 31.9 200 1080 Pb (ppm) 19.3 13.4 14.7 30.5 4.65 9.26 44.7 79.5 377 24.4 49.3 40.3 Rb/86Sr 7.52 21.7 1.28 7.46 16.3 5.78 8.18 0.47 161 14.8 13.0 58.0 Sr/86Sr ±s U/ Pb U/ Pb Th/ Pb Pb/ Pb ± s Pb/ Pb ± s Sm/144Nd Pb/ Pb ± s Nd/144Nd ±s C. Xu et al. / Geochimica et Cosmochimica Acta 143 (2014) 189–206 Samples Initial Sr, Nd and Pb isotopic ratios of the South Qinling carbonatites and syenites are calculated from Rb, Sr, Sm, Nd, U, Th and Pb contents measured by ICPMS, assuming an age of 235 Ma and 760 Ma, respectively. The eNd (t) values are calculated based on present-day (147Sm/143Nd)CHUR = 0.1967 and (143Nd/144Nd)CHUR = 0.512638. WR, whole rock; Bt, biotite; Mn, monazite; Ct, ferrocolumbite; Ap, fluorapatite; Py, pyrite; cal, calcite; CV, calcite veinlets in syenites associated with carbonatites. *Sr isotope was analyzed by calcites from carbonatites using Neptune LA-MCICPMS at Australian National University. The italic initial Pb isotopes are out of a ‘normal’ mantle field, so are not shown in Fig. 6. MY-7s MY-8s MY-9s Ct WR Py Py 960 21.3 62.8 121 160 54.0 12,540 62,080 72.8 4.71 7.17 6.93 515 27.3 0.32 0.12 3.74 0.20 0.00 0.00 40.4 6.23 0.04 0.01 42.759 ± 4 24.340 ± 6 19.324 ± 2 19.227 ± 2 16.835 ± 2 15.880 ± 1 15.597 ± 1 15.590 ± 1 38.718 ± 6 38.764 ± 13 38.135 ± 4 38.048 ± 5 23.641 23.328 19.312 19.222 15.862 15.829 15.596 15.590 38.246 38.691 38.134 38.048 0.671 0.679 0.808 0.811 1.618 1.659 1.975 1.979 C. Xu et al. / Geochimica et Cosmochimica Acta 143 (2014) 189–206 199 the SHRIMP U–Th–Pb age of monazite from the carbonatites (235 Ma). The syenites show lower initial 87Sr/86Sr and higher eNd (t) values than the carbonatites (Fig. 6a; Table 5). Note that some carbonatites and their mineral phases have quite lower 87Rb/86Sr ratios (<0.001; Table 5), and age-correction cannot affect their initial Sr isotopic composition. Both liquid immiscibility and crystal fractionation have been invoked to explain derivation of small-volume carbonate melts from a hybrid alkali-rich carbonate–silicate magma (Koster van Groos and Wyllie, 1963; Verhulst et al., 2000; Halama et al., 2005, among many others). Immiscibility experiments in compositionally diverse silicate–carbonate systems have demonstrated that Pb, Nb, Th, U and most of the REE partition preferentially into the silicate liquid, whereas Sr, Ba and F into the conjugate carbonate fraction (Jones et al., 1995; Veksler et al., 1998, 2012). This pattern of element partitioning is inconsistent with a much higher content of primary LREE-rich fluorapatite, fluorcarbonates and monazite in the SQ carbonatites relative to the syenites. Fractionation of feldspars and ferromagnesian minerals from a hybrid carbonate–silicate magma is a potential alternative mechanism of carbonatite petrogenesis. Clearly, this process should yield an evolved melt depleted in trace elements compatible in feldspars, biotite and clinopyroxene and enriched in incompatible elements retained in the liquid phase. The published experimental studies demonstrated that in silicate melts (including trachytic compositions), Cr, V and Ni are strongly compatible with respect to clinopyroxene and biotite (Schmidt et al., 1999; Fedele et al., 2009). In feldspars, Ga is compatible, whereas Rb is incompatible with DGa 3 DRb (e.g., Bédard, 2006; Macdonald et al., 2010). Hence, a hybrid magma evolving by fractionation of these minerals would be expected to produce derivative melts with significantly lower Ga/Rb ratio and Cr, V and Ni contents relative to the early crystallization products (i.e., cumulate syenites). At SQ, this is clearly not the case; the Ga/Rb ratio is higher in the carbonatites (Ga/Rb = 0.7– 8.8) relative to that in the syenites (0.2–0.6). Both rock types contain similar contents of compatible transition metals of Cr, V and Ni (Fig. 4a). Based on the above evidence, we conclude that, despite their intimate spatial association, the SQ carbonatites are not genetically related to the syenites. 5.2. Mantle sources, crustal recycling and carbonatite petrogenesis The extensive radiogenic isotopic data of carbonatites worldwide indicate that most young carbonatites (<200 Ma) have Nd, Sr, and Pb isotopic compositions similar to ocean-island basalts (OIB; Zindler and Hart, 1986; Hart, 1988), involving HIMU, EM1 and FOZO mantle components. The HIMU component present in the isotopic compositions of OIB has been interpreted as Proterozoic (1–2 Ga) oceanic crust recycled through subduction (Hofmann, 1997). Based on the similarity of isotopic compositions between OIB and carbonatites, derivation of the latter from melting of carbonated eclogite has been proposed (Nelson et al., 1988). Published experimental data 200 C. Xu et al. / Geochimica et Cosmochimica Acta 143 (2014) 189–206 Table 6 In situ Pb isotopic composition of calcites of carbonatites from the South Qinling orogen. Samples U (ppm) Pb (ppm) Th (ppm) (207Pb/206Pb)i ± s (208Pb/206Pb)i ± s MY-7 (n = 6) MY-10 (n = 8) MY-13 (n = 9) MY-14 (n = 14) MY-15 (n = 3) MY-16 (n = 2) MY-17 (n = 6) MY-18 (n = 12) 0.03 0.05 0.01 0.01 bdl bdl 0.12 0.01 12.1 50.3 13.3 9.99 12.8 9.19 43.0 63.9 0.04 0.01 0.02 0.01 bdl bdl 0.54 0.12 0.743 ± 22 0.772 ± 10 0.861 ± 16 0.881 ± 11 0.886 ± 2 0.750 ± 22 0.833 ± 8 0.837 ± 8 1.808 ± 83 1.908 ± 25 2.144 ± 31 2.164 ± 23 2.152 ± 16 1.792 ± 23 2.148 ± 9 2.097 ± 19 U, Pb and Th compositions were analyzed by LA-ICPMS. Measured Pb isotopic ratios of calcite are assumed to represent their initial ratios owing to negligible U and Th contents in the host minerals. a b Fig. 4. Primitive-mantle-normalised trace-element (a) and chondrite-normalised REE abundances (b) of the South Qinling syenites and carbonatites. The average carbonatite is from Xu et al. (2010a). Normalisation values for primitive mantle and chondrite are from McDonough and Sun (1995). Fig. 5. Carbon and oxygen isotopic compositions from the carbonatites in the South Qinling orogen, shown together with the field of primary, igneous carbonates (Taylor et al., 1967) and typical mantle composition (Ray et al., 1999). suggested that such an origin is feasible (Yaxley and Green, 1994; Hammouda, 2003; Dasgupta et al., 2004, 2005; Yaxley and Brey, 2004). The experimental work of Dasgupta et al. (2005) shows that compositions of eclogite-derived carbonate melts span the range of natural carbonatites from oceanic and continental settings. The isotopic Sr–Nd–Pb compositions of oceanic carbonatites off the African coast contain a HIMU-derived signature (Fig. 3); these compositions have been interpreted to record melting of Proterozoic subducted oceanic crust stored within the deep mantle (Hoernle et al., 2002). Petrologic evidence supports the notion that the carbon cycle involving subducted material may extend to lower-mantle depths (Walter et al., 2011). However, on the basis of experimental phase equilibria, Hammouda (2003) argued that carbonates will most likely be removed from subducted oceanic crust before it reaches a depth of 300 km. Dasgupta et al. (2004) and Tsuno and Dasgupta (2012) counter-argued that melting-induced release of carbonate material from the subducted crust will not occur until the slab reaches at least the base of the upper mantle and thus carbonates are likely to be recycled into the deep mantle. On the basis of both radiogenic and stable isotopic data from carbonatites worldwide, Bell and Simonetti (2010) interpreted carbonatitic melts as being derived from a sublithospheric source without any appreciable contribution from crustal materials. The latter authors also emphasized that carbonatites are virtually absent in subduction-related settings. There is a general consensus that collision between the South China Block and Qinling occurred in the Triassic along the Mianlue suture (Meng and Zhang, 1999; Zhang et al., 2002; Ratschbacher et al., 2003). This collision was C. Xu et al. / Geochimica et Cosmochimica Acta 143 (2014) 189–206 a b c d 201 Fig. 6. Comparison of the Sr–Nd–Pb isotopic compositions of carbonatites and seynites in the South Qinling (SQ) and Lesser Qinling (LQ) orogenic belts and Indo-Asian collision zone. The error bars for in situ Pb isotopic analyses of calcite from the SQ carbonatites reflect in-run precision with one standard deviation. Data for the LQ-H and LQ-S carbonatites are from Huang et al. (2009) and Xu et al. (2011), respectively. EACL stands for the East African Carbonatite Line of Bell and Tilton (2001); PaCar and PxCar for Cenozoic Pakistani and Panxi (China) carbonatites in the eastern Indo-Asian collision zone, respectively (Tilton et al., 1998; Xu et al., 2003; Hou et al., 2006). DMM, HIMU, EM1 and EM2 are the principal mantle end-member components, as defined by Hart (1988). preceded by the closure of the Mianlue ocean, which is manifested in the presence of ophiolite, ocean-island and island-arc volcanic packages along the southern boundary of the SQ. The belt of ophiolitic and tectonic mélange extends E–W for about 160 km (Zhang and Meng, 1995). Siliceous sedimentary rocks of possible deep water origin and Early Carboniferous radiolarian fauna (e.g., Albaillella sp.; Latentifistula cf. ruestae; Entactinia variospina) have also been recognized in the ophiolite (Feng et al., 1996). The whole-rock 40Ar/39Ar data for basalts and U–Pb zircon data for diabase from these ophiolites gave an age range from 345 to 264 Ma (Lai and Qin, 2010), showing that the Mianlue crust formed from at least the Early Carboniferous (crust older than 345 Ma may have been subducted) through the Permian. The important radiometric constraint on the onset of subduction is a U–Pb zircon age of 246 Ma, obtained by Qin et al. (2008) for andesite. The SQ carbonatites are located close to the Mianlue suture and were emplaced at 235 Ma, implying that their parental melt may be somehow linked to subduction of the Late Paleozoic Mianlue crust. Higher d18OV-SMOW and d13CV-PDB values were determined for the SQ carbonatites relative to the range of values typically observed in mantle-derived rocks, including fresh primary carbonatites (Fig. 5). Carbon isotopic composition of carbonatites is less susceptible to modification by fluid infiltration and other subsolidus processes and, hence, is a more robust petrogenetic indicator in comparison with d18OV-SMOW (Deines, 1989). The high d13CV-PDB values recorded in the SQ samples can potentially be accounted for by liquid immiscibility, fractional crystallization, assimilation of sediments, or source contamination (Ray et al., 1999). Although these carbonatites are spatially associated with mantle-derived syenites, the above study shows they do not have genetically related, ruling out immiscibility as the driving mechanism for the C-isotopic excursion. Fractional crystallization can be ruled out because this process would generate a concerted increase in d13CV-PDB and d18OV-SMOW values (Ray and Ramesh, 2000), which is not observed in our case (Fig. 5). Assimilation of sediments enriched in heavy C is possible, but is extremely unlikely to contribute significantly to the C–O isotopic budget because the SQ rocks show a conspicuous negative Pb anomaly in their primitive-mantle trace-element pattern (Fig. 4a). Thus, we conclude that crustal carbon recycled to the mantle via subduction of the Mianlue oceanic crust is the most likely source of the heavy-C enrichment in the SQ carbonatites. The origin of carbonatites in orogenic system is disputed. Tilton et al. (1998) suggested that Cenozoic carbonatites from northwestern Pakistan in the Indo-Asian collision zone were derived from a mantle source similar to that beneath the East African Rift, i.e., involving upwelling of deep sublithospheric material, and argued that no 202 C. Xu et al. / Geochimica et Cosmochimica Acta 143 (2014) 189–206 crustal contribution was involved in the generation of carbonatite parental magmas. Hou et al. (2006) proposed a petrogenetic model for the Himalayan collision-related carbonatites involving recycling of an old oceanic crust with pelagic or terrigenous component into a deep mantle source. Combined the Sr–Nd–Pb isotopic compositions, the effect of recycling subducted Proterozoic (1.5 Ga) oceanic crust with <10% of sediments into the mantle is modeled in Fig. 7. The results of this modeling show that the isotopic compositions define a HIMU–EM1–EM2 mixing trend with different percentages of sedimentary material. Hence, melting of ancient subducted oceanic crust with a small sedimentary component is an alternative explanation for the peculiar isotopic composition of continental carbonatites in the Indo-Asian collision zone. However, the model yields Sr–Nd–Pb isotopic compositions of Proterozoic oceanic crust plotting above the compositions of the SQ and LQ carbonatites (Fig. 7). The Paleozoic Mianlue oceanic crust is characterized by high eNd (8–11) but low (EM1-like) 206 Pb/204Pb ratios (16.90–17.25) (Xu et al., 2002), and is hence different from the Proterozoic crust used in the modeling. The initial isotopic compositions of the Mianlue crust and subducted sediments calculated at 235 Ma cannot explain the CHUR-like Nd and HIMU-like Pb isotopic signatures of the SQ carbonatites. Their formation requires an additional component with less radiogenic Nd and similarly radiogenic Pb isotopic characteristics with respect to the HIMU end-member. Because marine carbonates have low 147 Sm/144Nd (0.11) and high 238U/204Pb ratios (22; Hoernle et al., 2002), the hypothetical mantle component with a high 206Pb/204Pb but low 143Nd/144Nd ratios is likely to be derived from the recycling of Proterozoic carbonates. These carbonates could be subducted with Proterozoic oceanic crust through deep mantle over a period of 1–2 Ga before resurfacing in mantle plumes (e.g., Hofmann, 1997) or asthenospheric upwellings. Most of the previously published isotopic studies of carbonatites appear to favor plumes to explain both the geochemical characteristics and tectonic setting of these rocks (Dauphas and Marty, 1999; Bell and Tilton, 2001; Tolstikhin et al., 2002; and references therein). The Qinling orogen lacks any field evidence for plume activity, which is in contrast with the occurrence of voluminous basalts in the Emeishan igneous province in the western part of the South China Block, where the main stage of flood magmatism at 251–253 Ma (Lo et al., 2002) was associated with the arrival of a mantle plume (Chung and Jahn, 1995). In this work, we prefer a model where the magma that produced the SQ carbonatites originated by melting of the subducted Mianlue crust mixed with ascending mantle material characterized by high Pb and low Nd isotopic ratios and derived from a deep-seated Proterozoic carbonate reservoir via asthenospheric upwelling. In comparison with the SQ carbonatites, the LQ rocks have Sr–Nd–Pb isotopic compositions approaching, and trending toward less radiogenic Nd value than EM1. The two groups also differ somewhat in their trace-element composition. The LQ carbonatites are characterized by flat to Fig. 7. Modeled Sr–Nd–Pb isotopic variations for the carbonatite sources in orogenic belts corresponding to a mixture of recycled oceanic crust with sediments and of mantle-derived material with delaminated lower crust. The f(Nd, Pb) was calculated using the method described by Bell and Tilton (2001), i.e., f(Nd, Pb) = [(143Nd/144Nd)2 + (206Pb/204Pb)2]1/2 {sin[arctan(143Nd/144Nd/206Pb/204Pb) + 0.000064]}. (I) (old OC + sed.) recycled 1.5-Ga oceanic crust (OC; after Rehkämper and Hofmann, 1997) and marine sediments (sed.) at various mass ratios up to 10 wt.%; the average global subducted sediment composition is after Plank and Langmuir (1998); (II) (Ml OC + sed.) recycled Mianlue oceanic crust (MI OC; after Xu et al., 2002) incorporating up to 8 wt.% of marine sediments (sed.; Plank and Langmuir, 1998); (III) (MM + LC) mixing of depleted mantle-derived material (MM; estimate from the Mianlue oceanic crust; Xu et al., 2002) with delaminated lower crust (LC; data from Gao et al., 1998a; Jahn et al., 1999; Lu et al., 2006) at various mass ratios up to 25 wt.%. Data for the Cenozoic Pakistan (PaCar) and Chinese Panxi carbonatites (PxCar) in the eastern Indo-Asian collision zone, and the LQ-H and LQ-S carbonatites in Lesser Qinling orogen are shown for comparison. See Fig. 6 for references. C. Xu et al. / Geochimica et Cosmochimica Acta 143 (2014) 189–206 weakly LREE-enriched chondrite-normalized patterns [(La/Yb)CN = 1.0–5.5; Xu et al., 2007], which is in marked contrast with published data for most other carbonatites [(La/Yb)CN > 20; Chakhmouradian et al., 2008). In addition, these rocks contain economic level of Mo, representing the only known molybdenite deposit associated with carbonatites worldwide (Xu et al., 2010b). Molybdenite is not found in the SQ carbonatites, which contain economic level of REE and are characterized by the normal pattern of LREE enrichment (Fig. 4b). Therefore, the recorded differences in trace-element and isotopic budget suggest that the SQ and LQ rocks were derived from compositionally distinct mantle sources. The LQ carbonatites contain mantle-derived d13CV-PDB ( 5.3% to 7%) and slightly variable d18OV-SMOW (7.6% to 9.5%) values (Xu et al., 2010b). Recycling of subducted Paleozoic Mianlue oceanic crust incorporating different proportions of a sedimentary component cannot explain the isotopic variation of the LQ carbonatites. It is noteworthy in this context that according to the available geophysical and geochemical evidence, lower crust is absent beneath the Qinling (Gao et al., 1998b). On the basis of Pb isotopic evidence for the LQ-S carbonatites, Xu et al. (2011) suggested that the observed crustal thinning resulted from underthrusting of the southern margin of the North China Block during its Triassic collision with the South China Block, followed by gravitational collapse of the Qinling orogen and delamination of its keel into the mantle. This type of geodynamic environment has been discussed previously by Arndt and Goldstein (1989), Kay and Mahlburg-Kay (1993), and Lustrino (2005). Mixing of the delaminated lower crust with mantle materials at different proportions could generate a carbonated source in the sub-LQ mantle showing less radiogenic Nd isotopic characteristics. The oldest basement rocks exposed in the LQ region are the Mesoarchean Taihua group comprising upper-amphibolite- to granulite-facies metamorphic assemblages. Lower-crust xenoliths from the central part of the North China Block are dominated by two-pyroxene and garnet-two-pyroxene mafic granulites (Fan and Liu, 1996), implying that these rocks were likely the principal component of the delaminated lower crust. Using the average chemical composition of granulite from the southern margin of the North China Block (Gao et al., 1998a,b; Jahn et al., 1999; Lu et al., 2006), we calculated that a maximum contribution of 15% and 20% of delaminated continental crust is sufficient to produce the unusual isotopic characteristics of the LQ-S and LQ-H carbonatites, respectively (Fig. 7). 5.3. Tectonic implications The Qinling orogen merges with the Kunlun and Qilian orogens to the west (Fig. 1) to compose one of the most prominent tectonic zones in central East Asia. However, the exact timing of collision between the South and North China Blocks along the Qinling orogen remains controversial; contrasting lines of evidence suggest that it either terminated in the Late Triassic (Wang et al., 2007; Qin et al., 2010; Shen et al., 2014) or began in the Late Triassic 203 and lasted into the Jurassic (Ames et al., 1993; Ratschbacher et al., 2003; Sun et al., 2002; Hacker et al., 2004). Emplacement of carbonatites in both the SQ and LQ places constraints on the regional tectonics following the closure of the Mianlue ocean. The co-existence of slightly depleted to enriched mantle sources under the Qinling belt reflects subduction and recycling of oceanic and, as we propose in the present work, also lower continental crustal materials into the mantle. Published mineral Sm–Nd and U–Pb ages indicate that ultrahigh-pressure metamorphism in the Dabie-Sulu terrane (Fig. 1) took place between 210 and 245 Ma, peaking at 230–240 Ma (Li et al., 2000; Ayers et al., 2002; Zheng et al., 2003). Paleomagnetic studies show that the Permian–Triassic collision between the North and South China Blocks was accompanied by a 70° clockwise rotation of the latter, causing the collision to progress from east to west (Zhao and Coe, 1987). This, in turn, produced extension in the Dabie-Sulu terrane, leading to rapid exhumation of ultrahigh-pressure metamorphic rocks, and synchronous compression in the Qinling, leading to subduction of the Mianlue oceanic crust in the Early to Middle Triassic. This was accompanied by Permian–Triassic plume activity at Emeishan in the western part of the South China Block (Chung and Jahn, 1995). The SQ carbonatites were emplaced in the Late Triassic (235 Ma) at the northwestern margin of the South China Block, and can be convincingly linked to the recycled Mianlue oceanic crust. We suggest that asthenospheric upwelling triggered decompressional melting of the recycled oceanic crust, producing the SQ carbonatites. Both asthenospheric activity and melting could be related to the transition from a transpressional to transtensional regime at the end of the Late Triassic. Underthrusting and thickening of the continental crust beneath the Qinling orogen caused its gravitational collapse and delamination of the lower crustal material which not only modified the isotopic budget of the mantle source of the LQ carbonatites, but also probably caused lithospheric extension (Jull and Kelemen, 2001). Extensional structures facilitated melting of the modified mantle source and the ascent and emplacement of carbonatitic magmas produced in the process. Therefore, our data conclusively demonstrate that intercontinental collision of the South and North China Blocks terminated in the Late Triassic. 6. CONCLUSIONS The carbonatites in the SQ orogenic belt at the northwest margin of the South China Block were emplaced at 233.6 ± 1.7 Ma, i.e., strongly younger than the associated syenites (766 ± 25 Ma). The initial Sr–Nd isotopic and trace element compositions do not support that both rock types have genetically related. The SQ carbonatites were directly generated by primary carbonate magmas in the mantle. The rocks exhibit slight depletion in radiogenic Sr, minor variation in eNd values close to the CHUR, but a wide range of initial Pb isotopes straddling between the EM1 and HIMU end-members. Most of their C–O isotopes plot outside of the field of typical primary igneous carbonates. Emplacement of the SQ carbonatites was preceded by 204 C. Xu et al. / Geochimica et Cosmochimica Acta 143 (2014) 189–206 the closure of the Paleozoic Mianlue ocean and collision between the South China Block and Qinling microplate along the Mianlue suture. Their isotopic characteristics are mostly consistent with the melting of subducted oceanic crust mixed with an ascending deeper mantle component characterized by high 206Pb/204Pb and low 143Nd/144Nd values and possibly derived from recycled Proterozoic carbonates. In contrast, the Late Triassic (209–221 Ma) LQ carbonatites at the south margin of the North China Block contain EM1-like Sr–Nd–Pb and mantle-derived C–O isotopic compositions, distinguishing them from the SQ carbonatites. The isotopic budget of these LQ rocks is best explained by delamination of lower continental crust beneath the Qinling following the Late-Triassic collision between the South and North China Blocks, and incorporation of 15–20% of lower crustal granulites into the subcratonic mantle. 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