He dating

Journal of Asian Earth Sciences 79 (2014) 741–758

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Thermal-tectonic history of the Baogutu porphyry Cu deposit, West Junggar as constrained from zircon U–Pb, biotite Ar/Ar and zircon/apatite (U–Th)/He dating Guang-Ming Li a,⇑, Ming-Jian Cao a,b, Ke-Zhang Qin a, Noreen J. Evans c,d, Brent I.A. McInnes d, Yong-Sheng Liu e a

Key Laboratory of Mineral Resources, Institute of Geology and Geophysics, Chinese Academy of Sciences, P.O. Box 9825, Beijing 100029, China Graduate University of Chinese Academy of Sciences, Beijing 100049, China c CSIRO Earth Science and Resource Engineering, 26 Dick Perry Ave., Kensington, WA 6151, Australia d John de Laeter Center for Isotope Research, Dept. Applied Geology/Applied Physics, Curtin University, Perth, WA 6945, Australia e State Key Laboratory of Geological Processes and Mineral Resources, Faculty of Earth Sciences, China University of Geosciences, Wuhan, China b

a r t i c l e

i n f o

Article history: Available online 22 June 2013 Keywords: Western Junggar Baogutu porphyry copper deposit Zircon U–Pb dating Ar/Ar dating Zircon (U–Th)/He dating Apatite (U–Th)/He dating Thermal-tectonic history

a b s t r a c t Understanding postmineralization tectonic movements in porphyry deposits, is critical to interpreting the complete thermal-tectonic history. This study reports new zircon U–Pb ages, hydrothermal biotite 39 Ar/40Ar age, and zircon and apatite (U–Th)/He ages from the Baogutu porphyry copper deposit, which, in conjunction with pre-existing geochronology and thermochronology and inverse modeling simulations, constrain the thermal-tectonic history of the deposit. Zircon LA-ICP-MS U–Pb concordia ages indicate a Late Carboniferous age of 320.1 ± 2.2 Ma for the diorite complex and 309.8 ± 2.2 Ma for the mineralized granodiorite porphyry. Hydrothermal biotite selected from a quartz–biotite–chalcopyrite vein yields a plateau age of 311.0 ± 1.8 Ma which agrees with the age of the granodiorite porphyry and a previously reported molybdenite Re–Os age, and suggests that hydrothermal fluid circulated about 310 m.y. ago. Diorite and granodiorite porphyry yield weighted mean zircon (U–Th)/He ages of 200.6 ± 5.7 Ma and 241.1 ± 8.1 Ma, respectively, with ages ranging from 221.0 Ma to 174.3 Ma and 225.9 to 261.6 Ma. Weighted mean diorite and granodiorite porphyry apatite (U–Th)/He ages of 87.4 ± 2.3 Ma and 120.0 ± 4.2 Ma were obtained with ages ranging from 68.9 Ma to 100.8 Ma, and from 91.0 Ma to 152.0 Ma, respectively. The wide range of zircon (U–Th)/He ages may be due to the combined effects of U and Th zonation and radiation damage, and radiation damage effect may also account for the wide range of apatite (U–Th)/He ages. The combined effects of depth and cooling due to meteoric water circulation contribute to an older (U–Th)/He age for the granodiorite porphyry, relative to the diorite. A five-episode cooling rate history for the diorite can be deduced by inverse model simulation: fast cooling—moderate fast cooling—relatively slow cooling—fast cooling again—very slow cooling. The thermal and tectonic history of the wall rock indicates that it suffered significant far-field effects from the Qiangtang–Eurasia and Lhasa–Qiangtang collision, however, no visible thermal effect from the India–Asia collision is observed. Ó 2013 Elsevier Ltd. All rights reserved.

1. Introduction The Central Asian Orogenic Belt (CAOB) (Mossakovsky et al., 1993), also known as the Altaids (Sengör, 1993; Sengör and Natal’in, 1996), extends from the Urals to the northwest Pacific and from the Siberian and Baltica cratons to the Sino-Korean and Tarim cratons. The CAOB is a Paleozoic accretionary orogen, composed of several Precambrian and Paleozoic units that amalgamated be-

⇑ Corresponding author. Tel.: +86 10 82998187; fax: +86 10 62010846. E-mail address: [email protected] (G.-M. Li). 1367-9120/$ - see front matter Ó 2013 Elsevier Ltd. All rights reserved. http://dx.doi.org/10.1016/j.jseaes.2013.05.026

tween Siberia, Baltica and Tarim. After the final Late Carboniferous–Early Permian amalgamation, the CAOB was subject to several episodes of Mesozoic and Cenozoic deformation due to successive Eurasian collision–accretion events which led to the CAOB being one of the world’s largest and most active intracontinental orogen (Molnar and Tapponnier, 1975; De Grave et al., 2007; Jolivet et al., 2007; Vassallo et al., 2007). The deformation of the Late Miocene and Pliocene even propagated into the Altai Mountains (Yuan et al., 2006b; Zhang et al., 2008), the Gobi–Altay region (Vassallo et al., 2007), and the Baikal rift region (Jolivet et al., 2009).

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Fig. 1. (a) Location of the study area in the Central Asian Orogenic Belt (modified after Jahn et al. (2004)). (b) Geological map of the West Junggar (modified after Feng et al. (1989) and XBGMR (1993)). Age data for ophiolite rocks are from Feng et al. (1989), Kwon et al. (1989), Zhang and Huang (1992), Buckman and Aitchison (2001), Zheng et al. (2005), Xu et al. (2006), Gu et al. (2009), Liu et al. (2009a), Chen and Guo (2010), Zhang and Guo (2010) and Chen and Zhu (2011). Age data for granite rocks are from Li et al. (2000), Chen and Jahn (2004), Chen and Arakawa (2005), Han et al. (2006), Su et al. (2006), Yuan et al. (2006a), Zhou et al. (2008), Geng et al. (2009) and Chen et al. (2010). Age data for gold and copper deposits are from Li et al. (2000) and Wang and Zhu (2007). Age data for Baogutu diorite–grandiorite plutons and porphyries are from this study.

Western Junggar is bounded by the Altai orogen to the north, the Tienshan orogen to the south, the Kazakhstan plate to the west and the Junggar basin to the east (Fig. 1a). The economic potential of western Junggar is significant, with more than 300 gold deposits and occurrences discovered (Shen et al., 1996). Moreover, the newly discovered Baogutu porphyry copper deposit is the first large porphyry deposit found in western Junggar. However, the thermal-tectonic evolution of the region is still poorly known with only very limited thermochronology studies (Li et al., 2010). For porphyry deposits, understanding postmineralization thermal-tectonic history is critical in order to comprehend the complete cooling and exhumation history of the deposit and to determine the regional tectonic evolution. Therefore, studies elucidating the full thermal-tectonic history of a porphyry copper deposit are necessary. With the advent of (U–Th)/He radiometric methods, particularly zircon and apatite (U–Th)/He thermochronometers which record the cooling history of rock samples between 160 °C to 200 °C (Reiners et al., 2004) and 40° to 85 °C (Wolf et al., 1998), respectively, we have the capability to link existing low-temperature thermochronometric methods such as the apatite fission track method (closure temperature 120 °C; Gallagher et al., 1998; Ketcham et al., 1999), and biotite 40Ar/39Ar method (closure temperature 280°C to 348 °C; Harrison et al., 1985; Grove and Harrison, 1996). However, successful application

of (U–Th)/He thermochronology for ore deposit studies is still very scarce (McInnes et al., 1999, 2005a,b; Danišík et al., 2010). In this study, systematical geochronology and multi-method thermochronology methods are applied to the Baogutu porphyry copper deposit. The objectives of the present work are to investigate the thermal-tectonic history of the Baogutu porphyry copper deposit, and provide some insight into the uplift history of the western Junggar. 2. Tectonic setting It is generally accepted that the western Junggar orogenic belt was accreted to the Kazakhstan plate during the Paleozoic. Three major Paleozoic NE–SW-trending fault systems are recognized in the western Junggar. From north to south they are the Barleik, Mayila and Dalabute faults, all dipping to the NW. At their southwestern extremities these faults curve sharply westwards to merge with the right-lateral Dzhungarian fault that cross-cuts the former three faults (Zhang et al., 1993). Allen and Vincent (1997) indicated that the Paleozoic fault zones have been reactivated by Mesozoic– Cenozoic collision events, and Mesozoic events have reactivated a larger number of fault zones than have been affected by the Cenozoic India–Asia collision. In addition, the far-field effects caused by collision–accretion events are also recorded by several stages of

G.-M. Li et al. / Journal of Asian Earth Sciences 79 (2014) 741–758

coarse, clastic sediment precipitation in the Junggar foreland basins (Hendrix et al., 1992; Allen et al., 1993; Vincent and Allen, 2001). Li et al. (2010) identified Late Cretaceous reactivation in western Junggar batholiths based on apatite fission track analyses. Thus, Mesozoic–Cenozoic events have played a significant effect on reactivation of western Junggar. Ordovician to Early Permian stratum is developed in the western Junggar (Feng et al., 1989). Divided by the Hueshengtaolege valley, the western Junggar is divided into northern and southern parts (Fig. 1b). Ordovician sedimentary rocks are restricted to the Tangbale area and small occurrences in the Shaburt Mountains (Fig. 1b). The main outcrops of Silurian rocks are in the south where Silurian flysch-type deposits and pillow lavas are interlayered with radiolarian chert in the Mayila area. Devonian stratum consists of molasse including coarse clastic sands and gravels, and flysch-like sediments in the south part, while in the north, Devonian rocks are mainly andesitic lavas and terrestrial clastic sandstones. The Carboniferous sedimentary rocks in the north and south parts of the western Junggar are not significantly different mainly comprising flysch-type sediments deposited under shallow marine and littoral conditions. The Early Permian rocks are mainly continental molasse with minor volcanic. Ophiolites are scattered across the western Junggar region in a series of discontinuous belts, generally striking NE–SW. From northwest to southeast these ophiolites include Barleik, Mayile, Tangbale, Dalabute, Hongguleleng and Karamay (Fig. 1b). Numerous studies have been focused on the geochronology of six ophiolites: for example, Dalabute ophiolite were dated from Middle Devonian to Late Carboniferous with ages ranging from 395 Ma to 302 Ma (Zhang and Huang, 1992; Gu et al., 2009; Liu et al., 2009a; Chen and Guo, 2010), and Karamay ophiolite from Early Devonian to Early Carboniferous with ages of 414–332 Ma (Xu et al., 2006). Carboniferous to Early Permian intrusions are widely distributed in western Junggar (Fig. 1b). Recently, Chen et al. (2010) obtained the first Late Silurian–Early Devonian pluton ages (422– 405 Ma) in the Xiemistai and Saier Mountains. In the southern part of west Junggar, two types of granitoids are recognized: large Atype granite batholiths including Miaoergou, Akbastao, Hatu, Tiechanggou, Karamay and Hongshan alkali-feldspar granites with the ages of Late Carboniferous–Early Permian age (Li et al., 2000; Chen and Jahn, 2004; Han et al., 2006), and small I-type granodiorite stocks including more than twenty stocks with age of Late Carboniferous (Shen et al., 2009; Shen et al., 2010). The stock V hosts the Cu and Au mineralization of Baogutu deposit which is the second largest porphyry copper deposit in Xinjiang, after Tuwu–Yandong located in east Tienshan.

3. Geology of Baogutu deposit and samples The stock V is a dioritic complex with an area of about 0.6 km2 (Fig. 2), located within the eastern limb of the Xibeikulasi syncline. It is emplaced into the Lower Carboniferous Baogutu group and Xibeikulasi group and occurs at the intersection of the north–south and northeast–southwest trending faults. Lower Carboniferous strata contain the Xibeikulasi group, Baogutu group and Tailegula group from bottom to top in the Baogutu area with total exposed thickness ranging from 3300 m to 7000 m for the Xibeikulasi group, 2000 m to 5000 m for the Baogutu group, and thickness of 5000 m for the Tailegula group (XBGMR, 1993). These three stratigraphic units are similar being mainly composed of tuffaceous siltstone, silty tuff and felsic tuff with intercalations of pebbly greywacke, and lenses of limestone, marl and bioclastic limestone. Former studies indicate Early Carboniferous age (345–328 Ma) for these three stratigraphic units based on zircon U–Pb dating (An

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and Zhu, 2009; Tong et al., 2009; Guo et al., 2010). Shen et al. (2010) obtained 325 Ma for the diorite. After the diorite emplacement, later near east–west trending granodiorite porphyry intruded into the diorite and the stratum of Baogutu group (Fig. 2). This was accompanied by disseminated sulfides (pyrrhotite, chalcopyrite and pyrite) forming in the porphyry. Magmatic breccia cemented by fragments of matrix and rock-forming minerals with weak mineralization occurs in the upper and lower part of the porphyry in many drill holes. The shape of the ore-body (determined by 0.2% Cu contours) is that of a thick plate with dipping northwest, which is similar to that of the granodiorite porphyry which the ore-body spatially overlaps. In addition, the granodiorite porphyry displays relatively strong biotite alteration relative to the diorite. Therefore, we propose that the porphyry is responsible for copper and gold mineralization. In addition, Tang et al. (2010) obtained a zircon U–Pb age of 310 ± 3 Ma on porphyry which is in agreement with the mineralization age of 310 Ma for molydbenite Re–Os and hydrothermal biotite Ar/Ar (Song et al., 2007; Shen et al. 2012). The consistency of these ages also supports the direct relationship between the porphyry and mineralization. The Baogutu porphyry copper deposit contains 63  104 t Cu averaging 0.28 wt%, 14 t Au averaging 0.1 ppm, 1.8  104 t Mo averaging 0.011 wt% and 390 t Ag averaging 1.8 ppm (Zhang et al., 2009, unpublished data). Based on our detailed observations, we have divided mineralization at Baogutu into five paragenetic stages from early to late. Stage I: Ca–Na silicate alteration with the assemblage of actinolite + albite + sphene ± ilmenite ± magnetite. Stage II: extensive potassic alteration which produces dominant disseminated biotite and secondary vein-type K-feldspar. Stage III: quartz–sulfide mineralization with quartz–arsenopyrite, quartz–pyrrhotite–chalcopyrite, quartz–chalcopyrite, quartz–pyrite and quartz–molybdenite vein assemblages. Stage IV: propylitization with the assemblage of chlorite + calcite ± epidote. Stage V: secondary Ca–Al silicate alteration with the assemblage of prehnite ± epistilbite ± laumounite. Among these five stages, biotite alteration and propylitization are extensively developed, which leads to moderate biotite alteration in the core and propylitization on the periphery (Fig. 2). Unlike a typical porphyry Cu deposit, the mineralized porphyry and associated wall rock diorite contain a predominance of disseminated Cu–Mo–Au mineralization together with lesser amounts of vein-type mineralization. In this study, zircon and apatite samples used for U–Pb geochronology and (U–Th)/He thermochronology dating were selected from relatively fresh diorite (ZK003-233(473)) and granodiorite porphyry (BCK2-1) which were collected from drill hole ZK003 (at 473 m depth) and from the surface, respectively (Fig. 2). Diorite ZK003-233(473) is fine-grained, dark-grey in color, massive with subhedral granular texture, and composed of plagioclase (60%), K-feldspar (10%), quartz (4%), biotite (6%), hornblende (15%), and minor accessory minerals (e.g., sphene, apatite, zircon) and sulfides. The granodiorite porphyry BCK2-1 is fine-grained, porphyritic in texture, light-grey in color. The phenocrysts are made up of plagioclase (50%), K-feldspar (15%), quartz (10%), biotite (15%), hornblende (5%), and minor accessory minerals (e.g., sphene, apatite, zircon) and sulfides. In addition, hydrothermal biotite (ZK202176) selected from a quartz–biotite–chalcopyrite vein was used for 40 Ar/39Ar dating to obtain the timing of hydrothermal fluid circulation. This vein is continuous with a width of about 2.5 cm and shows a week alteration halo. Biotite is fine-grained and is found in banding parallel to the walls, whereas chalcopyrite is usually distributed throughout the center of this vein. Based on these characteristics, the quartz–biotite–chalcopyrite vein could be classified as a B type vein according to Gustafson and Hunt (1975), which is the product of main stage of mineralization. Details of these three samples such as sample depth, location, lithology, dating method and mineral assemblage are listed in Table 1.

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Fig. 2. Geological map of the Baogutu porphyry copper deposit showing the porphyry intrusion modified after Shen and Jin (1993). Inset figure is the cross section along EW line 0 showing lithologies and copper ore bodies after Zhang et al. (2006).

Table 1 Summary of location and lithology, analytical methods and mineral assemblage details of samples from Baogutu porphyry copper deposit. Sample

Depth (m)

Location

Lithology

Dating method

Mineral assemblage (in vol.%)

ZK003-233 (473) BCK2-1

473

Diorite

401

Zr U/Pb, Zr and Ap (U–Th)/He Zr U/Pb, Zr and Ap (U–Th)/He Bt Ar/Ar

Pl (60%), Kfs (10%), Qtz (4%), Bt (6%), Hbl (15%), and others (5%)

ZK202-176

45°280 20.100 N 84°320 09.200 E 45°280 19.400 N 84°320 18.900 E 45°280 23.300 N 84°320 23.200 E

0

Granodiorite porphyry Qtz–Bt–Ccp vein

Phenocrysts: Pl (50%), Kfs (15%),Qtz (10%), Bt (15%), Hbl (5%), and others (5%); and matrix: Pl, Qtz and Kfs Continuous vein with width about 2.5 cm, Bt occupy banding parallel to the walls, Ccp distribute in the center

Note the abbreviations for minerals: Ap, Apatite; Bt, Biotite; Ccp, Chalcopyrite; Hbl, Hornblende; Kfs, K-feldspar; Pl, Plagioclase; Qtz, Quartz; Zr, Zricon.

4. Analytical methods Mineral concentrates of zircon, apatite and biotite were obtained by routine crushing, sieving, heavy liquid, magnetic separation and careful handpicking. A representative number (100–200 per sample) of zircon and apatite grains were selected, mounted in epoxy blocks, carefully polished and carbon coated for cathodoluminescence (CL) images. 4.1. U–Pb dating Cathodoluminescence (CL) images (at 15 kV) were obtained using a CAMECA SX-50 microprobe at the Institute of Geology and Geophysics, Chinese Academy of Sciences (IGGCAS) in Beijing,

in order to characterize the internal structures of the zircon and apatite. Both optical photomicrograph and CL images were taken as a guide to selection of spots for zircon U–Pb dating. U–Pb dating was performed by laser ablation inductively coupled plasma mass spectrometry (LA–ICP–MS) at the State Key Laboratory of Geological Processes and Mineral Resources, China University of Geosciences, Wuhan. The laser ablation system was a GeoLas 2005, which was equipped with a 193 nm ArF excimer laser and a homogenizing optical imaging system coupled to an Agilent 7500a ICP-MS. Nitrogen was added into the central gas flow (Ar + He) of the Ar plasma to decrease the detection limit and improve precision (Hu et al., 2008). Off-line selection and integration of background and analytical signals, time-drift correction and quantitative calibration for trace element analyses and U–Pb dating

G.-M. Li et al. / Journal of Asian Earth Sciences 79 (2014) 741–758

were performed by ICPMSDataCal (Liu et al., 2008, 2010). Detailed operating conditions for the laser ablation system and the ICP-MS instrument and data reduction are the same as in Liu et al. (2008, 2010). Concordia diagrams and weighted mean calculations were made using Isoplot/Ex_ver3 (Ludwig, 2003). Individual analyses and weighted average of ages were reported with 1r uncertainty. Representative CL images of zircon with analytical numbers, U–Pb ages of diorite and granodiorite porphyry are shown in Fig. 3. 4.2. Ar/Ar dating Biotite separates were ultrasonically cleaned three times in deionized water and three times in acetone, each for 30 min, and then dried in preparation for irradiation. All aliquots samples were

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wrapped in aluminum foil and stacked in quartz vials with the international standard Brione muscovite (18.7 ± 0.09 Ma) (Hall et al., 1984). Neutron irradiation was carried out in position B4 of 49-2 Nuclear Reactor (49-2 NR) at the Beijing Atomic Energy Research Institute, with a flux of 6.5  1012 n (cm2 s)1 for 24 h, yielding J values of 0.00646. 40Ar/39Ar measurements were performed at IGGCAS in Beijing. The irradiated sample was transferred into a high-vacuum ‘‘Christmas tree’’ for 24 h at 250 °C, and was then dropped into the Ta crucible of an automated double-vacuum resistance furnace. Biotite was incrementally heated from 700 °C to 1300 °C at a pace of 30 °C or 50 °C with each step for 10 min. Isotopic measurement were made on the MM5400 mass spectrometer at IGGCAS. Detailed procedures for 40Ar/39Ar analyses follow Wang et al.

Fig. 3. Representative cathodoluminescence (CL) images of zircons with analytical numbers, U–Pb ages, CL images of apatite grains, microscopy images of zircon and apatite with analytical numbers, measured dimension and (U–Th)/He ages of diorite and granodiorite porphyry from the Baogutu porphyry copper deposit from top to bottom.

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(2006). CaF2 and K2SO4 crystals were fused to calculate Ca and K correction factors: [36Ar/37Ar]Ca = 0.000261 ± 0.000014; [39Ar/37Ar]Ca = 0.000724 ± 0.000028 and [40Ar/39Ar]K = 0.000880 ± 0.000023. The decay constants used in the age calculation were quoted from Steiger and Jäger (1977). Mass discrimination was monitored using an online air pipette from which multiple measurements were made before and after each experiment. The mean over this period was 0.99977 ± 0.00022 per amu and the uncertainty of this value was propagated into the age calculation (Wang et al., 2006). Plateau ages were determined from 3 or more contiguous steps, comprising 50% of the 39Ar released, revealing concordant ages at the 95% confidence level. Uncertainties on all data for this study are reported at the 95% confidence level. The raw data were processed using the ArArCALC (Koppers, 2002). 4.3. (U–Th)/He dating 4.3.1. Apatite Grains for (U–Th)/He thermochronology were selected under the microscope in order to avoid grains with cracks or U- and Th-rich mineral/fluid/gas inclusions that may contribute excess helium to the analysis. Grain measurements were taken for the calculation of an alpha correction factor (Ft; Farley et al., 1996) and every effort was made to select grains with a diameter larger than 70 lm in order to maximum helium gas values and minimize the Ft correction. Images of selected grains were recorded digitally and grain measurements were taken. Helium was thermally extracted from single shards, loaded into platinum micro-crucibles and heated using a 1064 nm Nd–YAG laser. 4He abundances were determined by isotope dilution using a pure 3He spike, calibrated daily against an independent 4He standard tank. The uncertainty in the sample 4He measurement is <1%. For degassed apatite, the U and Th content was determined by isotope dilution. 25 ll of a 50% (by volume; approximately 7 M) HNO3 solution containing approximately 15 ppb 235U and 5 ppb 230 Th was added to each sample. The apatite was digested in the spiked acid for at least 12 h to allow the spike and sample isotopes to equilibrate. Standard solutions containing the same spike amounts as the samples, in addition to 25 ll of a standard solution containing 25 ppb U and Th, were treated identically, as were a series of unspiked reagent blanks (25 ll of a 50% HNO3 solution). 250 ll of MilliQ water was added prior to analysis on an Agilent 7500CS mass spectrometer (at the University of Western Australia, TSW Analytical). U and Th isotope ratios were measured to a precision of <2%. Overall the apatite (U–Th)/He (AHe) thermochronology method at CSIRO has a precision of 3.6% for apatite, based on multiple age determinations (n = 40) of Durango standard over the analysis period which produce an average age of 30.8 ± 1.1 (2r) Ma. 4.3.2. Zircon Zircon grains were selected using the same criteria described for apatite, above. Measured and photographed grains were loaded in niobium microvials. Helium was thermally extracted from single crystals heated to >1200 °C using a 1064 nm Nd–YAG laser. Uranium and Th were also determined using isotope dilution inductively coupled mass spectrometry. Samples were removed from the laser chamber and transferred to Parr pressure dissolution vessels where they were spiked with 25 ll of a solution containing 15 ppb 235U and 15 ppb 230Th and digested at 240 °C for 40 h in 350 ll of HF. 80 ppb U and Th standard solutions containing the same spike amounts as samples were treated similarly, as were a series of unspiked reagent blanks. After the 40 h dissolution, samples were removed from the pressure vessels and allowed to dry for 2 days. 300 ll of HCl was added to each vial and a second bombing for 24 h at 200 °C ensured dissolution of fluoride salts.

Final solutions were diluted to 10% acidity for analysis on an Agilent 7500CS mass spectrometer (at the University of Western Australia, TSW Analytical). For single crystals digested in small volumes (0.3–0.5 ml), U and Th isotope ratios were measured to a precision of <2%. The zircon (U–Th)/He (ZHe) thermochronology method at CSIRO has an overall precision of <6% and more detailed description of U and Th analysis can be found in Evans et al. (2005). Representative microscopy images of zircon and apatite with analytical numbers, (U–Th)/He ages of diorite and granodiorite porphyry are shown in Fig. 3. 5. Results 5.1. Zircon U/Pb ages Zircon U–Pb dating results are presented in Table 2 and Concordia diagrams and weighted ages are presented in Fig. 4. Diorite yielded a zircon U–Pb Concordia age and weighted age of 320.1 ± 2.2 Ma and 320.5 ± 1.6 Ma, respectively (sample ZK003233(473), Figs. 4a and b). This consistent zircon U–Pb age suggests that the diorite was emplaced in Early to Late Carboniferous. The granodiorite porphyry of sample BCK2-1 displays a somewhat younger age, with a Concordia age of 309.8 ± 2.2 Ma and weighted age of 310.1 ± 1.6 Ma (Figs. 4c and d), corresponding to Late Carboniferous. 5.2. Biotite Ar/Ar age Hydrothermal biotite Ar/Ar analysis results are presented in Table 3, and age spectra and inverse isochrones are illustrated in Fig. 5. Hydrothermal biotite shows relatively well-resolved plateaus comprising 48.5% of the total 39Ar released, and presents consistent plateau and inverse isochron ages of 311.0 ± 1.8 Ma and 309.8 ± 1.9 Ma, respectively. An initial 40Ar/36Ar ratio of 303.9 shows an excess Ar component, compared with the atmospheric value (295.5). Erroneously old ages in low-temperature steps (700–1030 °C) account for 50% of the total released 39Ar. The monotonically decreasing low-temperature discordant sections (Fig. 5) were probably caused by low-temperature alteration and hydration (Lo et al., 1994; Wang et al., 2009), which is consistent with the observation of an excess Ar component. Wang et al. (2009) shows that samples with an increasing degree of alteration have increasing excess initial Ar component suggesting that some of our primary hydrothermal biotite transformed into hydrated secondary minerals (i.e., chlorite). 5.3. (U–Th)/He ages ZHe and AHe dating results are presented in Table 4 and radial plots are presented in Fig. 6. We analyzed 4 zircons grains from ZK003-233(473) (diorite). The four ZHe ages range from 221.0 to 174.3 Ma and the entire population has a weighted mean age of 200.6 ± 5.7 Ma (Fig. 6a, Table 4). We analyzed 5 zircons grains from BCK2-1 (granodiorite porphyry). The ZHe ages range from 225.9 to 261.6 Ma with a weighted mean age of 241.1 ± 8.1 Ma (Fig. 6b, Table 4). Six apatite crystals from ZK003-233(473) (diorite) yielded ages of 68.9–100.8 Ma, with a weighted mean age of 87.4 ± 2.3 Ma (Fig. 6a, Table 4). Seven apatite crystals from BCK21 (granodiorite porphyry) yielded ages of 91.0–152.0 Ma and a weighted mean age of 120.0 ± 4.2 Ma (Fig. 6b, Table 4). For both samples, AHe ages are younger than ZHe ages (Fig. 6), which is consistent with lower closure temperature for the AHe system. Besides, both ZHe and AHe ages show wide distribution. Additionally, both ZHe and AHe ages of BCK2-1 granodiorite porphyry are several tens of millions of years older than those of ZK003-233(473) diorite.

Table 2 LA–ICPMS zircon U–Pb isotope data for diorite and granodiorite porphyry from Baogutu porphyry Cu deposit in western Junggar. Spot

Element(ppm) Pb

Th

Th/U

Isotopic ratios 207

U

Pb/206Pb

Age (Ma) ±1r

207

Pb/235U

±1r

206

Pb/238U

±1r

207

Pb/206Pb

±1r

207

Pb/235U

±1r

206

Pb/238U

±1r

80.6 110 67.8 111 132 117 155 107 136 134 127 95.6 154 81.4 237 232 163 118 99.3 135 114 89.8 318

0.78 0.85 0.92 0.99 0.80 0.82 0.86 1.02 0.74 0.82 0.76 0.55 1.01 0.89 0.84 0.49 1.90 0.99 1.14 1.65 0.75 1.04 1.23

0.05406 0.05251 0.05450 0.05764 0.05220 0.05872 0.05230 0.05260 0.05691 0.05847 0.05870 0.05954 0.05755 0.06353 0.05282 0.05177 0.04890 0.05598 0.05403 0.06036 0.05388 0.05866 0.05437

0.00245 0.00253 0.00264 0.00296 0.00178 0.00221 0.00194 0.00233 0.00232 0.00354 0.00230 0.00262 0.00188 0.00297 0.00239 0.00223 0.00271 0.00294 0.00282 0.00300 0.00318 0.00362 0.00197

0.37530 0.36142 0.37753 0.40864 0.36701 0.41047 0.36363 0.36668 0.40199 0.41155 0.41071 0.42449 0.40370 0.44028 0.36350 0.36421 0.34123 0.39192 0.38556 0.41610 0.37618 0.40871 0.38143

0.01681 0.01671 0.01784 0.02214 0.01263 0.01545 0.01326 0.01609 0.01670 0.02401 0.01586 0.02054 0.01299 0.02239 0.01551 0.01554 0.01821 0.01985 0.01969 0.02005 0.02258 0.02581 0.01414

0.05095 0.05090 0.05098 0.05104 0.05103 0.05100 0.05081 0.05103 0.05094 0.05112 0.05101 0.05099 0.05096 0.05003 0.05073 0.05105 0.05107 0.05106 0.05274 0.05069 0.05083 0.05099 0.05092

0.00067 0.00058 0.00075 0.00058 0.00050 0.00064 0.00052 0.00069 0.00058 0.00066 0.00060 0.00061 0.00057 0.00079 0.00063 0.00060 0.00063 0.00076 0.00092 0.00074 0.00072 0.00080 0.00057

372 309 391 517 295 567 298 322 487 546 567 587 522 728 320 276 143 450 372 617 365 554 387

69 109 109 118 78 81 88 100 91 133 81 96 72 94 104 100 162 117 119 107 133 135 81

324 313 325 348 317 349 315 317 343 350 349 359 344 370 315 315 298 336 331 353 324 348 328

12 12 13 16 9 11 10 12 12 17 11 15 9 16 12 12 14 14 14 14 17 19 10

320 320 321 321 321 321 319 321 320 321 321 321 320 315 319 321 321 321 331 319 320 321 320

4 4 5 4 3 4 3 4 4 4 4 4 3 5 4 4 4 5 6 5 4 5 3

Granodiorite porphyry (BCK2-1) 1 6.75 62.6 2 5.71 43.3 3 6.68 51.9 4 10.3 105 5 7.11 55.8 6 6.45 44.3 7 6.31 64.6 8 8.29 78.7 9 6.53 38.3 10 6.65 65.9 11 9.02 98.5 12 7.45 72.9 13 7.71 55.7 14 4.59 40.4 15 7.13 52.2 16 6.46 41.7 17 11.1 107

108 92.3 110 163 115 103 94.8 126 108 101 133 113 123 71.1 115 103 173

0.58 0.47 0.47 0.64 0.48 0.43 0.68 0.63 0.35 0.66 0.74 0.65 0.45 0.57 0.46 0.41 0.62

0.05271 0.05545 0.05592 0.05417 0.05609 0.05982 0.05464 0.05139 0.05210 0.05276 0.05428 0.05883 0.05484 0.05914 0.05656 0.05952 0.05278

0.00219 0.00224 0.00197 0.00176 0.00201 0.00270 0.00214 0.00172 0.00211 0.00238 0.00189 0.00256 0.00209 0.00308 0.00187 0.00259 0.00188

0.35734 0.37396 0.36671 0.36128 0.38015 0.41101 0.36795 0.34863 0.35294 0.35348 0.36802 0.39853 0.37116 0.39624 0.38225 0.41051 0.35811

0.01522 0.01459 0.01229 0.01150 0.01373 0.01856 0.01414 0.01158 0.01451 0.01533 0.01328 0.01729 0.01402 0.01938 0.01275 0.01626 0.01288

0.04946 0.04940 0.04813 0.04864 0.04931 0.05049 0.04953 0.04931 0.04931 0.04930 0.04933 0.04931 0.04928 0.04939 0.04929 0.05110 0.04934

0.00062 0.00055 0.00048 0.00040 0.00057 0.00071 0.00062 0.00041 0.00055 0.00066 0.00051 0.00052 0.00048 0.00073 0.00056 0.00073 0.00051

317 432 450 389 457 598 398 257 300 317 383 561 406 572 476 587 320

94 89 78 74 112 103 89 78 97 102 78 94 85 115 74 96 81

310 323 317 313 327 350 318 304 307 307 318 341 321 339 329 349 311

11 11 9 9 10 13 10 9 11 12 10 13 10 14 9 12 10

311 311 303 306 310 318 312 310 310 310 310 310 310 311 310 321 310

4 3 3 2 4 4 4 3 3 4 3 3 3 4 3 4 3

G.-M. Li et al. / Journal of Asian Earth Sciences 79 (2014) 741–758

Diorite (ZK003-233(473)) 1 5.67 62.8 2 7.86 94.0 3 4.95 62.5 4 8.39 110 5 9.30 105 6 8.46 96.5 7 11.1 133 8 7.79 109 9 9.52 100 10 9.62 110 11 8.77 96.2 12 6.43 52.5 13 11.6 155 14 5.75 72.4 15 17.5 198 16 15.3 114 17 14.9 308 18 9.01 116 19 8.19 113 20 11.8 222 21 8.12 85.6 22 6.95 93.0 23 25.0 391

747

748

G.-M. Li et al. / Journal of Asian Earth Sciences 79 (2014) 741–758

Fig. 4. Zircon U–Pb concordia diagrams and weighted ages of the diorite and granodiorite porphyry from the Baogutu porphyry copper deposit.

6. Discussion 6.1. Interpretation of intra- and inter-sample variation in single crystal (U–Th)/He age Zircon and apatite from both samples show a range of ZHe ages ranging from 221.0 to 174.3 Ma for diorite and from 261.6 to 225.9 Ma for granodiorite porphyry, and AHe ages ranging from

68.9 to 100.8 Ma for diorite and from 91.0 to 152.0 Ma for granodiorite porphyry that far exceeds analytical uncertainty (Fig. 6, Table 4). Many factors have been proposed to explain the intrasample variation in single crystal ages including: the presence of U- and Th-rich inclusions or fluid inclusions (e.g., Lippolt et al., 1994), He implantation from the surrounding matrix (e.g., Spencer et al., 2004; Danišík et al., 2010), variation in crystal size (e.g., Reiners and Farley, 2001), U and Th zonation (e.g., Farley et al., 1996;

Table 3 The 40Ar/39Ar analysis results.

a

Temperature (°C)

36

37

38

39

40

700 750 800 840 920 960 1000 1030 1060 1090 1120 1160 1200 1300

0.000696 0.000947 0.001415 0.001972 0.001219 0.001372 0.001041 0.000502 0.000296 0.000169 0.000087 0.000084 0.000077 0.000053

0.000066 0.000026 0.000437 0.000701 0.000170 0.000055 0.000253 0.000318 0.000291 0.000255 0.000340 0.000525 0.000948 0.001250

0.000127 0.000193 0.000316 0.000467 0.000365 0.000416 0.000518 0.000704 0.000672 0.000539 0.000233 0.000068 0.000035 0.000016

0.000343 0.000899 0.002233 0.004354 0.005784 0.006541 0.012925 0.023716 0.023721 0.019659 0.008771 0.002236 0.000745 0.000296

0.221514 0.316052 0.491698 0.726148 0.549078 0.635181 0.728729 0.861790 0.778266 0.619806 0.280078 0.090273 0.047931 0.024250

Ar[Ca]b

Ar[Cl]b

Ar[k]b

Ar[r]b

Age (Ma)

2r (Ma)

40

472.16 418.77 349.68 348.77 346.27 370.02 345.53 321.22 311.76 310.39 310.53 313.22 358.5 319.72

141.74 72.19 43.09 30.91 14.61 14.32 5.95 2.35 2.03 1.95 2.17 5.37 13.31 45.55

7.14 11.47 15.00 19.74 34.42 36.19 57.81 82.81 88.77 91.94 90.82 72.51 52.69 36.43

Age of the fluence monitor Bern 4 M is 18.7 ± 0.1 Ma (Hall et al., 1984). Ar[a], 37Ar[Ca], 38Ar[Cl], 39Ar[k], 40Ar[r] denote the atmospheric 36Ar, 37Ar derived from Cl, K/Ca = [39Ar[k]/37Ar[Ca]] * 0.56 (Wang et al., 2006).

b 36 c

Ar[a]b

39

Ar[r]b (%)

Ar derived from k, and radiogenic

Ar[r]b (%)

40

40

K/Cac

2r

0.31 0.80 1.99 3.88 5.15 5.83 11.52 21.13 21.14 17.52 7.82 1.99 0.66 0.26

3.03 20.44 2.96 3.60 19.74 68.67 29.59 43.29 47.27 44.75 14.96 2.47 0.46 0.14

16.77 280.49 2.36 1.81 40.43 434.82 41.03 47.48 56.80 61.23 15.39 1.75 0.17 0.04

Ar in mV, respectively.

749

G.-M. Li et al. / Journal of Asian Earth Sciences 79 (2014) 741–758

Fig. 5. The 40Ar/39Ar age spectra and inverse isochron diagram. The solid squares denote the steps used for calculating the isochrones. Inset figure is the photograph of quartz–biotite–chalcopyrite vein from which biotite is selected for Ar/Ar dating.

Meesters and Dunai, 2002; Danišík et al., 2010), radiation damage (e.g., Hurley, 1952; Nasdala et al., 2004; Shuster et al., 2006; Shuster and Farley, 2009).

The most common U- and Th-rich inclusions in apatite are zircon and monazite, which survive the nitric acid U–Th dissolution technique but contribute He to the analysis. Unlike apatite, the effect of U- and Th-rich minerals inclusions in zircon is minimized since natural zircon is highly enriched in actinides (U and Th), and total dissolution during chemical treatment. Fluid inclusions may contain a significant radiogenic 4He component, which could be ‘‘parentless’’, or contribute ‘‘excess He’’ (Ballentine et al., 2002; Ballentine and Burnard, 2002). While no detectable mineral or fluid inclusions were observed in the selected zircon and apatite and no extremely old ages were found. If zircon and apatite grains were surrounded by mineral phases with high U and/or Th contents, extraneous 4He will be implanted into these grains (Spencer et al., 2004). However, the perfect positive correlation between zircon/apatite He concentration and effective uranium concentration suggests there was no significant ‘‘parentless’’ 4He in these grains (Fig. 7a and b) and so the presence of undetected inclusions and He implantation from the surrounding matrix are not of major concern. Larger zircon and apatite crystals have a greater effective diffusion dimension and thus higher He retentivity than smaller grains (Farley, 2000). Larger grains will exhibit a higher closure temperature and yield older ages than smaller grains experiencing the same thermal history. A slightly negative correlation is seen between ZHe and equivalent spherical radius (three times of volume-to-surface ratio; Hourigan et al., 2005; Meesters and Dunai, 2002) in the diorite and granodiorite porphyry zircon grains

Table 4 Zircon and apatite (U–Th)/He ages for diorite and granodiorite porphyry from Baogutu porphyry copper deposit. U (ppm)

1r

Th (ppm)

1r

Th/U

eUa (ppm)

He (nmol/g)

Ftb

Raw date (Ma)

Corr datec (Ma)

2rd

325.0 406.3 465.1 233.0

8.9 11.1 12.9 6.4

195.2 184.5 222.5 169.9

7.22 6.81 8.18 6.23

0.60 0.45 0.48 0.73

370.9 449.7 517.4 272.9

226.9 352.8 358.7 204.1

0.64 0.65 0.64 0.66

112.4 143.6 127.1 137.0

174.3 221.0 199.1 208.0 200.6

5.1 6.2 5.7 6.0 5.7

BCK2-1 granodiorite porphyry 137662-1 2.4 223.9 137662-2 1.8 212.1 137662-3 2.4 218.8 137662-4 3.0 356.5 137662-5 2.6 257.9 Average

6.2 5.8 6.0 9.9 7.1

122.8 103.2 156.0 139.2 111.0

4.53 3.78 5.72 5.17 4.10

0.55 0.49 0.71 0.39 0.43

252.8 236.3 255.4 389.2 284.0

251.4 235.6 229.0 359.4 272.5

0.72 0.69 0.73 0.75 0.73

181.5 181.6 163.9 168.9 175.2

251.6 261.6 225.9 226.4 240.0 241.1

8.9 7.2 7.9 8.1 8.5 8.1

Mass (lg) Zircon ZK003-233(473) diorite 137661-1 1.1 137661-2 1.1 137661-3 0.9 137661-4 1.4 Average

Apatite ZK003-233(473) diorite 137683-1 8.7 137683-3 5.1 137683-4 3.0 137683-5 2.8 137683-6 3.2 137683-8 4.3 Average

8.61 10.98 14.70 9.29 24.63 17.29

0.36 0.45 0.59 0.36 0.98 0.68

16.81 17.56 14.90 24.07 28.00 21.35

0.66 0.68 0.57 0.91 1.08 0.84

1.95 1.60 1.01 2.59 1.14 1.23

12.6 15.1 18.2 14.9 31.2 22.3

3.7 5.7 5.5 4.6 11.9 8.3

0.79 0.75 0.68 0.69 0.70 0.72

54.30 70.20 56.28 57.25 70.66 68.97

68.9 94.2 82.6 82.8 100.8 95.3 87.4

2.1 2.8 2.3 2.4 2.3 2.2 2.3

BCK2-1 granodiorite porphyry 137684-1 1.7 5.57 137684-2 2.5 4.56 137684-3 5.3 3.27 137684-4 2.3 2.66 137684-5 2.7 2.56 137684-6 3.8 1.15 137684-7 1.9 5.57 Average

0.22 0.18 0.13 0.10 0.10 0.05 0.22

18.99 20.67 11.64 7.17 10.39 5.11 18.67

0.71 0.77 0.43 0.29 0.38 0.20 0.73

3.41 4.54 3.56 2.69 4.07 4.45 3.35

10.0 9.4 6.0 4.3 5.0 2.3 10.0

5.0 5.0 2.5 1.8 1.7 0.8 4.6

0.63 0.65 0.74 0.66 0.68 0.70 0.62

92.43 98.73 76.60 74.85 63.76 64.14 86.12

146.7 152.0 103.7 113.9 93.4 91.0 139.1 120.0

3.9 3.9 3.1 3.6 2.7 7.7 4.7 4.2

a eU – effective uranium concentration that weighs U and Th for their alpha productivity and is computed as [U] + 0.235[Th]. Mass (and resultant U and Th concentrations) are based on microscopy grain measurements and assumed characteristic grain morphologies (see Reiners, 2005). b Ft – alpha ejection correction of Farley et al. (1996), assuming homogeneous eU distribution (no eU zonation). c Corr date includes Ft correction. d 2r errors propagated from U, Th and He measurement uncertainties and grain-length measurement uncertainty.

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G.-M. Li et al. / Journal of Asian Earth Sciences 79 (2014) 741–758

Fig. 6. (a) Comparison of zircon and apatite (U–Th)/He ages of diorite displayed on a radial plot (Galbraith, 1988; Vermeesch, 2009). (b) Comparison of zircon and apatite (U– Th)/He ages of granodiorite porphyry displayed on a radial plot (Galbraith, 1988; Vermeesch, 2009). Each population is represented by a average age and uncertainty calculated as a geometric mean and standard deviation, respectively, from single grain (U–Th)/He ages and associated analytical errors.

(Fig. 7c and d), which suggests the age variation can not be explained by the grain size effect. 6.1.1. Interpretation of intra-sample variation in single zircon crystal (U–Th)/He age A uniform concentration of U and Th is typically assumed when correcting for alpha ejection. However, U and Th are typically heterogeneously distributed in zircon, commonly in self-similar oscillatory zonation patterns over a range of scales (Fowler et al., 2002). Our zircon grains show typically oscillatory zoning in CL images (Fig. 3) and although actinide profiling was not performed on the grains, it can be assumed that some degree of compositional zoning was present. As a consequence, the assumption of uniform distribution of U and Th was likely inaccurate, leading to a commiserate inaccuracy in ZHe age calculation which had been investigated by many studies (e.g., Tagami et al., 2003; Hourigan et al., 2005). Radiation damage caused during natural radioactive decay will produce structural changes reflected in physical properties (e.g. Radlinski et al., 2003) and hence a decrease in chemical resistance. Highly radiation-damaged zircon shows enhanced susceptibility to loss of He (Hurley, 1952; Nasdala et al., 2004) and Pb (Rizvanova et al., 2000) that often significantly affects ZHe and U–Th–Pb ages, leading to ages that deviate from the true geologic age and a wide grain-to-grain age variation. Several representive studies suggest that He diffusion from zircon may not be significantly affected until an unusually high dosage is reached (2–4  1018 a/g; Nasdala et al., 2004; Reiners et al., 2004; Reiners, 2005). The zircon in this study had a crystalization age of about 300 Ma and U concentration of less than 400 ppm, leading to an a dose of almost 5  1017 a/g, far lower than that required to impact age determination. However, apparent negative correlation between ZHe age and eU in zircon grains, especially for zircon samples from granodiorite porphyry (Fig. 7e), suggest that radiation damage may play a role for the wide range of ZHe ages. Thus, we consider that the combined effects of U and Th zonation and radiation damage may account for the majority of ZHe variation in this study. 6.1.2. Interpretation of intra-sample variation in single apatite crystal (U–Th)/He age Some recent studies have shown that U–Th zonation does produce significant age bias in apatite (U–Th)/He dating (Farley et al.,

1996; Farley, 2002; Meesters and Dunai, 2002; Boyce and Hodges, 2005; Danišík et al., 2010). CL images of typical apatite from the diorite and granodiorite porphyry do not show complicated zoning (Fig. 3), but heterogeneity in U and Th distribution was not assessed. Highly radiation-damaged zircon and titanite rapidly lose He (Hurley, 1952), while Shuster et al. (2006) and Shuster and Farley (2009) found that radiation damage accumulation in apatite crystals impedes He diffusion, based on natural and artificial radiation damage studies. The effect of radiation damage may produce tens of degrees of variation in the closure temperature of apatite and lead to higher eU apatite undergoing less fractional He loss and yielding older ages than in lower eU apatite grain for the same time–temperature path (Flowers et al., 2009). This effect is most pronounced when thermal histories are characterized by slow cooling and prolonged residence in the He partial retention zone (Flowers et al., 2009). Many researches have suggested that ageeU correlations in apatite grains can be explained by the effects of radiation damage on He retentivity, and demonstrate the utility of this effect for thermal history interpretation (Green et al., 2006; Flowers et al., 2007; Flowers et al., 2008; Ault et al., 2009; Flowers and Kelley, 2011). Although both samples show relatively low U and Th content with eU ranging from 12.6 to 31.2 ppm for diorite and from 2.3 to 10.0 ppm for granodiorite porphyry, both samples show significant positive correlation between AHe age and eU (Fig. 7f). Therefore, radiation damage effect may be the main cause of variation in AHe age. If this is the case, our results suggest that even several ppm of U and Th in apatite could produce significant radiation damage under slow cooling and prolonged residence in He partial retention zone (see Section 6.3). 6.1.3. Interpretation of inter-sample variation in single zircon and apatite crystal (U–Th)/He age We use the mean value of corrected ages as the representative value of ZHe and AHe for each sample. The ages of ZHe and AHe are 200.6 Ma and 87.4 Ma for the diorite and 241.1 Ma and 120.0 Ma for the granodiorite porphyry. Thus, the differences of ZHe and AHe age between diorite and granodiorite porphyry are 40.5 Ma, and 32.6 Ma, respectively. Two possible reasons can be used to explain the large difference of samples: the effect of depth and cooling due to meteoric water. Diorite, collected from a drill hole, is 473 m deeper than the granodiorite porphyry collected from the surface. Many studies

G.-M. Li et al. / Journal of Asian Earth Sciences 79 (2014) 741–758

751

Fig. 7. (a) Zircon He concentration versus eU-effective uranium concentration, (b) apatite He concentration versus eU-effective uranium concentration, (c) zircon (U–Th)/He ages versus Req-equivalent spherical radiu#, (d) apatite (U–Th)/He ages versus Req-equivalent spherical radius for diorite and granodiorite porphyry, (e) zircon (U–Th)/He ages versus eU-effective uranium concentration, (f) apatite (U–Th)/He ages versus eU-effective uranium concentration. #Req is calculated by three times of volume-tosurface ratio (Hourigan et al., 2005; Meesters and Dunai, 2002), eU weighs the decay of the two parents for a productivity and is computed as U + 0.235Th.

have shown decreasing fission track and (U–Th)/He ages with increasing depth (Coyle et al., 1997; Coyle and Wagner, 1998; House et al., 1999; Lorencak et al., 2004; Stockli and Farley, 2004; Söderlund et al., 2005; Wolfe and Stockli, 2010). Based on AHe-vs-depth trend from drill cores in Precambrian basement of southeast Sweden, Söderlund et al. (2005) calculated exhumation rate of 17 m/my. If depth is the main reason, we can calculate an exhumation rate of 11.7 m/my and cooling rate of 0.34 °C/my during 240–200 Ma and exhumation rate of 14.5 m/my and cooling rate of 0.36 °C/my during 120–87 Ma assuming a geothermal gradient of 25 °C/km. However, these rates of exhumation and cooling

are far less than the rates obtained from modeling all available chronology data (Fig. 10b). Thus, the effect of depth cannot be the major reason for the wide ZHe and AHe age difference between the diorite and granodiorite porphyry. No previously published crystalization ages of granitoid rocks or gold/copper mineralization ages yield Late-Permian or younger ages (Fig. 8), which suggests that no significant heating event occured in the region since 270 Ma and that a prolonged exhumation history is preserved in the dioritic complex. The only suitable mineral barometer for diorite and granodiorite porphyry is Al-in-hornblende (Hammarstrom and Zen, 1986;

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G.-M. Li et al. / Journal of Asian Earth Sciences 79 (2014) 741–758

Fig. 8. Age distribution of granitoid rocks and hydrothermal gold deposits in western Junggar. There are 55 age data for granitoid rocks sourced from Li et al. (2000), Chen and Jahn (2004), Chen and Arakawa (2005), Han et al. (2006), Su et al. (2006), Yuan et al. (2006), Tang (2008), Zhou et al. (2008), Geng et al. (2009), Chen et al. (2010) and this study, and our five unpublished zircon U–Pb ages. Ten mineralization ages for hydrothermal gold and copper deposits are from Li et al. (2000), Song et al. (2007), Wang and Zhu (2007), Liu et al. (2009b) and this study.

Hollister et al., 1987; Johnson and Rutherford, 1989; Schmidt, 1992). The necessary mineral assemblage (plagioclase + biotite + hornblende + alkali feldspar + quartz + sphene + iron–titanium oxide) is present in these rocks. Based on the Al-inhornblende barometer, hornblende compositions of diorite and granodiorite porphyry from Wei and Zhu (2010) and Dai et al. (2010), respectively (Table 5), give an average pressure of 2.46 kbar and 1.68 kbar, respectively, which is equivalent to 8.9 km and 6.1 km, respectively, assuming a lithostatic pressure of 3.64 km/ kbar. The depth of 8.9 km is consistent with the total thickness of the Baogutu group and Tailegula group ranging from 7000 m to 10,000 m. This depth implies that meteoric water will penetrate into the diorite and porphyries very slowly during later exhumation. However, due to the relatively high degree of fracturing and common occurrence of veinlets in the granodiorite porphyry, meteoric water can readily penetrate into this porphyry, leading to the more rapid cooling of this porphyry relative to the diorite, and causing argillic alteration which is very common in porphyry deposit (e.g. Khashgerel et al., 2009; Kodeˇra et al., 2010; Li et al., 2011; Palacios et al., 2011). The earlier cooling of the porphyry by circulation of meteoric water will generate older ZHe and AHe ages for the porphyry than that of diorite, because of earlier entrance into and passing through the He partial retention zone of zircon and then apatite.

Table 5 Representative EPMA data of chemical composition (wt%) and calculated formula of hornblende from the Baogutu deposit. Lithology

Diorite a

Granodiorite porphyry a

Sample Occurrence

D-1 Interstitial

D-2

D-3

D-4

D-5

P-1b Phenocryst

P-2b

P-3b

P-4b

P-5b

SiO2 TiO2 Al2O3 FeO MnO MgO CaO Na2O K2O Cr2O3 NiO H2O* Total

46.54 2.12 7.92 14.23 0.24 13.90 11.05 1.44 0.63 0.04 0.05 2.08 100.23

46.93 1.93 7.53 14.32 0.22 14.06 11.05 1.38 0.53 0.00 0.02 2.09 100.06

47.49 1.92 7.61 13.53 0.21 14.56 11.17 1.41 0.54 0.00 0.05 2.10 100.58

47.30 1.95 7.42 13.51 0.23 14.55 11.08 1.30 0.60 0.01 0.00 2.09 100.05

47.88 1.65 7.13 13.29 0.23 14.62 11.07 1.40 0.54 0.03 0.05 2.09 99.96

48.52 1.23 5.46 11.63 0.42 15.77 11.39 1.51 0.52 0.04 0.00 2.06 98.55

48.10 1.34 6.57 12.12 0.33 14.91 11.25 1.41 0.44 0.00 0.01 2.06 98.54

47.43 1.47 6.62 12.79 0.41 14.72 10.97 1.75 0.47 0.03 0.06 2.06 98.78

48.59 1.53 6.75 12.84 0.20 14.56 11.13 1.56 0.30 0.04 0.03 2.08 99.61

47.57 1.79 7.14 12.83 0.20 14.26 11.19 1.57 0.34 0.00 0.08 2.07 99.04

Structural formula calculated based on 23 oxygens Si 6.696 6.749 AlIV 1.304 1.251 VI Al 0.038 0.025 Ti 0.229 0.209 Cr 0.005 0.000 Fe3+ 0.880 0.921 Fe2+ 0.832 0.801 Mn 0.029 0.026 Mg 2.981 3.014 Ni 0.006 0.003 Ca 1.703 1.703 Na 0.403 0.384 K 0.115 0.098 OH* 2.000 2.000 Total 17.220 17.184 P (kbars) P1 P2 P3 P4 Average a

2.83 2.81 2.22 3.38 2.81

2.50 2.44 1.94 3.06 2.49

a

a

a

6.775 1.225 0.053 0.205 0.000 0.856 0.758 0.025 3.097 0.005 1.707 0.391 0.099 2.000 17.197

6.779 1.221 0.032 0.211 0.002 0.892 0.727 0.028 3.109 0.000 1.701 0.362 0.110 2.000 17.173

6.863 1.137 0.068 0.178 0.003 0.822 0.772 0.028 3.124 0.006 1.700 0.389 0.098 2.000 17.187

7.047 0.935 0.000 0.134 0.005 0.631 0.782 0.052 3.415 0.000 1.773 0.425 0.096 2.000 17.294

6.986 1.014 0.111 0.146 0.000 0.630 0.843 0.041 3.228 0.001 1.751 0.397 0.082 2.000 17.229

6.897 1.103 0.032 0.161 0.003 0.747 0.808 0.050 3.191 0.007 1.709 0.493 0.087 2.000 17.290

6.990 1.010 0.135 0.166 0.005 0.618 0.927 0.024 3.123 0.003 1.716 0.435 0.055 2.000 17.206

6.906 1.094 0.127 0.195 0.000 0.591 0.967 0.025 3.086 0.009 1.740 0.442 0.063 2.000 17.245

2.51 2.45 1.95 3.08 2.50

2.38 2.31 1.84 2.95 2.37

2.14 2.04 1.64 2.73 2.13

0.78 0.51 0.49 1.44 0.81

1.74 1.58 1.30 2.34 1.74

1.79 1.64 1.34 2.39 1.79

1.84 1.69 1.38 2.44 1.84

2.22 2.13 1.71 2.80 2.22

Data from Wei and Zhu (2010). Data from Dai et al. (2010). * Calculated value by Tindle and Webb (1994). The pressure values of P1, P2, P3 and P4 are calculated based on Hammarstrom and Zen (1986), Hollister et al. (1987), Johnson and Rutherford (1989) and Schmidt (1992), respectively. b

G.-M. Li et al. / Journal of Asian Earth Sciences 79 (2014) 741–758

Therefore, we ascribe the difference in ZHe and AHe ages between the diorite and granodiorite porphyry to the combined effects of burial depth and cooling due to meteoric water circulation. Because of more significant effect of meteoric water on the granodiorite porphyry, only the (U–Th)/He ages of diorite are used in the inverse modeling simulation, and the cooling and exhumation rates are calculated based on the inverse modeling simulation on the diorite sample. 6.2. Thermal history of the Baogutu porphyry copper deposit In order to further decipher the details of Baogutu thermal history, we model the prolonged exhumation history in the Baogutu deposit using diorite apatite and zircon (U–Th)/He data and existing apatite fission track data, assuming a lack of significant heating events since 270 Ma as described above. For the thermal history simulations described below, we used the thermochronology age to constrain the temperature–time conditions (Table 6). The geothermal gradient has been 25 °C/km since Early Jurassic, 29 °C/km during Early Permian to Late Triassic and 33 °C/km before Early Permian (Qiu et al., 2002), and we assume a mean surface temperature of 10 °C. Based on a depth of 473 m for ZK003-233(473), the temperature of the diorite sample is near 20 °C, which is utilized in the simulation. Our thermal history simulations begin at 320 Ma, the crystallization age of the diorite. The ages of mineralization including molybdenite Re–Os age and hydrothermal biotite Ar/Ar age are not utilized, because these ages are coeval with the crystallization age of the granodiorite porphyry, suggesting a direct genetic relationship. Moreover, molybdenite Re–Os and biotite Ar/Ar systems have closure temperatures of 500 °C (Suzuki et al., 1996) and 348 °C (Grove and Harrison, 1996), respectively and it is unlikely that thermal equilibrium was reached between the diorite and the high temperature hydrothermal fluid from which molybdenite and biotite precipitated. Liu et al. (2009b) reported a 297 Ma hydrothermal biotite K–Ar age (close temperature of 300 °C; McDougall and Harrison, 1999). We consider it likely that thermal equilibrium was reached between the diorite and the biotite K–Ar system and, therefore, several temperature–time constraints come from this study and existing thermochronology data: 300 °C and 297 Ma for hydrothermal biotite K–Ar age from Liu et al. (2009b) and the closure

Table 6 Input parameters for HeFTy (version beta 2). Parameters

Option/value for apatite

Option/value for zircon

Calibrations Model precision Activation energy Do Alpha calculation Age to report Age alpha correction Segment parameters Maximum geothermal gradient

Other Best 29.23 kcal/mol 0.6071 cm2/N Ejection Corrected Ketcham in prep. 2G, gradual 25 °C/km after 200 Ma

Surface temperature Thermal history

10 °C No heating since 270 Ma 180 °C, 200 Ma 116 °C, 122 Ma 65 °C, 90 Ma 20 °C, 0 Ma

Reiners et al., 2004 Best 40.4 kcal/mol 0.46 cm2/N Ejection Corrected Ketcham in prep. 2G, gradual 25 °C/km after 180 Ma and 29 °C/km during 180–300 Ma, and 33 °C/ km before 300 Ma 10 °C No heating since 270 Ma

Constrain T and t

700 °C, 320 Ma 700 °C, 297 Ma 180 °C, 200 Ma 116 °C, 122 Ma 65 °C, 90 Ma 20 °C, 0 Ma

753

temperature from McDougall and Harrison (1999), 183 °C and 200 Ma for diorite zircon (U–Th)/He from this study and the closure temperature from Reiners et al. (2004), 116 °C and 122 Ma for apatite fission track from Li et al. (2010) and the closure temperature from Ketcham et al. (1999), 68 °C and 87.4 Ma for apatite (U–Th)/He from this study and the closure temperature from Farley (2000). We adopt 700 °C for the temperature of zircon crystallization. We used the new radiation damage accumulation and annealing model (Flowers et al., 2009), and posed 10,000 candidate temperature–time paths using the inverse modeling capabilities of HeFTy for the diorite sample (ZK003-233(473)) thermal history simulation (Ketcham, 2005). Details of the inverse modeling and statistically significant data fits are described by Ketcham (2005), and details of the input parameters for all simulations are described in the Table 5. The inverse modeling simulation results are shown in Fig. 9 which suggests a five episode cooling history for the diorite sample: (I) Very fast cooling from Late Carboniferous to Early Permian (320–290 Ma); (II) Moderate fast cooling from Early Permian to Late Triassic (290–200 Ma); (III) Relatively slow cooling from Early Jurassic to Early Cretaceous (200–136 Ma); (IV) Fast cooling from Early Cretaceous to Late Cretaceous (136–100 Ma); (V) Very slow cooling since Late Cretaceous (100 Ma). According to the Al-in-hornblende barometer, hornblende compositions from the diorite yield 2.46 kbar which is equivalent to 8.9 km, assuming a lithostatic pressure of 3.64 km/kbar (Table 5). Combined with the geothermal gradient of 33 °C/km (after Qiu et al., 2002), the temperature of the wall rock in the Late Carboniferous should be 294 °C which is much lower than the crystallization temperature of the diorite. We consider that the very fast cooling in stage I is related to the heat dissipation from the diorite into the wall rock due to the temperature contrast. However, we cannot verify whether or not the relative fast cooling in stage II was caused by reduction of heat dissipation between the diorite and the wall rock or by relative rapid exhumation of the diorite. The diorite sample yielded a zircon (U–Th)/He age of 200 Ma with a closure temperature of 183 °C. The prolonged 120 Ma of heat dissipation may imply temperature equilibrium between the diorite and the wall rock. Therefore, both should exhibit the same cooling and exhumation history since the Late Triassic. Relatively slow cooling stage in III indicates tectonic quiescence during 200–136 Ma, followed by tectonic reactivation during 136–100 Ma (stage IV) and then by very slow cooling (stage V) after 100 Ma in an episode of tectonic quiescence. In order to determine the cooling and exhumation history of the western Junggar basement, the depth of 8.9 km and temperature of 294 °C are used to represent the characters of the wall rock when the diorite pluton intruded. As mentioned, the tectonic history of the diorite pluton also represents that of the wall rock since Late Triassic. Based on the geochronology, thermochronology and inverse modeling results, the wall rock should pass through the following points of time, temperature, and depth: 320 Ma with 294 °C and 8.9 km, 200 Ma with 183 °C (zircon closure temperature) and 6.3 km (the corresponding depth of zircon closure temperature assuming geothermal gradient of 29 °C/km), 136 Ma with 136 °C (read from the apatite inverse modeling simulation) and 5.4 km (the corresponding depth of 136 °C assuming geothermal gradient of 25 °C/km), 100 Ma with 40 °C (read from the apatite inverse modeling simulation) and 1.6 km (the corresponding depth of 40 °C assuming geothermal gradient of 25 °C/km), and 0 Ma with 20 °C and 0.5 km (the present time, temperature and depth of the diorite sample). Fig. 10 displays the geochronology and thermochronology data obtained in this study and from previous work, and also the cooling and exhumation path (T–t and D–t) of the wall rock. The most prominent character of T–t and D–t is very fast cooling and exhumation for episode IV during 136–100 Ma and very slow cooling and exhumation for episode V after 100 Ma.

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Fig. 9. Inverse modeling results using the RDAAM in HeFTy for the diorite sample (ZK003-233(473) apatite and zircon). The mean age and mean U, Th values are used in the simulation. Black squares represent the imposed temperature–time constraint based on thermochronology ages. The good-fit point of solution is shown by the darker gray lines, and the acceptable-fit point of solution in lighter gray lines. The solid black line represents the best-fit thermal history. (a) Inverse modeling results of apatite. The constraints imposed on the thermal history are indicated by four black squares. 200 Ma with temperature of 180 °C represent the mean age of zircon (U–Th)/He and the closure temperature (Reiners et al., 2004). The constraints of 135–110 Ma and 95–135 °C come from apatite fission track age of Li et al. (2010) and assumed apatite partial annealing zone. AHe age of 120–80 Ma and temperature of 20–80 °C comes from this study and assumed apatite partial retention zone. 0 Ma and 20 °C are present time and presumed surface temperature. (b) Inverse modeling results of zircon. These constrains imposed on the thermal history are indicated by the Six black squares. 320 Ma with temperature of 700 °C represent the crystallization age and its temperature of zircon from this study. The 300 Ma with 300 °C are data of hydrothermal biotite K–Ar from Liu et al. (2009b). Other four constraints are the same as in (a). Five episodes can be distinguished. Episode I is the fast cooling stage due to heat dissipation from the diorite into the wall rock. This is followed by Episode II where slower cooling is due to eithor slower heat dissipation between the diorite and the wall rock or slow exhumation of the pluton. Episode III is a relatively slow cooling stage, followed by fast cooling stage IV, and tectonic quiescence or very slow cooling stage V.

6.3. Implication for the thermal-tectonic evolution of Western Junggar After its final amalgamation of Late Carboniferous–Early Permian, the CAOB was subjected to several episodes of reactivation due to intensive collision–accretion events successively in Eurasia which included the Late Triassic collision of Qiangtang terrane with south margin of Eurasia along the Jinsha suture (Dewey et al., 1988; Schwab et al., 2004), the Late Jurassic–Early Cretaceous collision of Lhasa terrane with Qiangtang along the Bangong–Nujiang suture (Dewey et al., 1988; Kapp et al., 2007), the Late Cretaceous collision of Kohistan–Ladakh arc and Karakoram terrane to Eurassia along south Kunlun terrane and the Shyok suture (Gaetani et al., 1993; Schwab et al., 2004), and the Cenozoic collision of India–Asia along the Indus–Yalu suture (Yin and Harrison, 2000; Kapp et al., 2007). The Junggar and Tarim basins, as foreland basins

of the ancestral Tienshan, preserved several stages of coarse, clastic sedimentation that recorded those successive, distinct orogenic episodes (Hendrix et al., 1992, 1994; Allen et al., 1993; Sobel and Dumitru, 1997; Hendrix, 2000; Vincent and Allen, 2001). Through an apatite fission-track study, Dumitru et al. (2001) identified three main cooling and exhumation periods in the Junggar and Tarim foreland basins: latest Paleozoic, Late Mesozoic and Late Cenozoic. Moreover, multi-method thermochronology on granitoids from Kyrgyz Tienshan also recorded those episodes of reactivation, especially Late Cenozoic cooling (Glorie et al., 2010; De Grave et al., 2011, 2012). In this study, the wall rock displayed a relatively slow cooling and exhumation evolution in our model with a cooling and exhumation rate of 0.92 °C/Ma and 21.6 m/my, respectively during Late Carboniferous to Late Triassic (Fig. 10b). Another probable evolu-

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recorded intensive tectonic events during Early Jurassic to Early Cretaceous (Fig. 10a). Stage IV exhibits another episode of intensified cooling and exhumation of Western Junggar during Early Cretaceous (2.67 °C/ Ma for cooling and 106.7 m/my for exhumation). The average cooling rate of episode IV is consistent with the results of Glorie et al. (2010) on Northern Kyrgyz Tienshan granitods. We interpret this reactivation as a far-field effect of the collision of the Lhasa with Qiangtang which is confirmed in other CAOB regions by many methods (Dumitru et al., 2001; Glorie et al., 2010; De Grave et al., 2011, 2012). Episode V is characterized by a very slow cooling and exhumation history after 100 Ma with a cooling rate of 0.2 °C/Ma and exhumation of 11.0 m/my (Fig. 10b) and may represent a stage of tectonic quiescence. The far-field effect of the India–Eurasia collision has caused widespread Miocene and Pliocene basement cooling in the CAOB as recognized in many other studies (Hendrix et al., 1994; Sobel and Dumitru, 1997; Dumitru et al., 2001; Bullen et al., 2003; Sobel and Strecker, 2003; Sobel et al., 2006; De Grave et al., 2007, 2011, 2012). For example, Yuan et al. (2006) and Vassallo et al. (2007) identified the effects of India–Asia collision in the Altai Mountains and in the Gobi-Altay range based on apatite fission track data. However, our new data and previous (albeit very limited) thermochronology data indicate no reactivation since Late Cretaceous in the western Junggar which may be the result of an increase in the strength of the Junggar basement and strain partitioned by the NW–SE dextral strike-slip Dzhungarian fault (Allen and Vincent, 1997). However, more detailed thermochronology studies on western Junggar are needed to verify this result.

7. Conclusion

Fig. 10. (a) Tectonic events associated with the ages obtained for the western Junggar from multi-geochronology and thermochronology. Molybdenite Re–Os age is from Song et al. (2007), hydrothermal biotite K–Ar from Liu et al. (2009b), and apatite fission track ages from Li et al. (2010). (b) Schematic cooling and exhumation path (T–t and D–t) of the wall rock of diorite pluton from western Junggar through time as derived from the obtained multi-method ages and inverse modeling result. See Sections 6.2 and 6.3 for discussion. K–K, Kohistan–Ladakh and Karakoram.

tion path can also explain the whole history from Late Carboniferous to Late Triassic (Fig. 10b, dashed line); very slow cooling and exhumation from Late Carboniferous to Middle Triassic, followed by very fast cooling and exhumation history from Middle Triassic to Late Triassic. This evolution history is consistent with the diorite ZHe age (200 Ma) and the inverse modeling simulation which indicates rapid cooling and exhumation, and tectonic reactivation during Middle Triassic to Late Triassic (Fig. 10a). In fact, this very fast cooling and exhumation episode is contemporaneous with the collision of Qiangtang with south margin of Eurasia (Dewey et al., 1988; Schwab et al., 2004), which is verified in the structural geologic, sedimentary and thermochronology record in the CAOB (Hendrix et al., 1992; Allen et al., 1993; Dumitru et al., 2001; De Grave et al., 2011, 2012). Therefore, we suggest that the tectonic reactivation caused by the Qiangtang collision is recorded in western Junggar. Stage III shows a relatively slow cooling and exhumation history during Early Jurassic to Early Cretaceous with rates of 0.73 °C/Ma for cooling and 13.6 m/my for exhumation (Fig. 10b), which can be interpreted as an episode of tectonic quiescence. This tectonic quiescence is supported by a relative lack of previously

Using multiple chronology approach including the application of zircon U–Pb, hydrothermal biotite Ar/Ar, zircon (U–Th)/He and apatite (U–Th)/He dating and inverse modeling simulation, we were able to reconstruct the thermal-tectonic evolution of the Baogutu intrusive complex. (1) The LA-ICP-MS zircon U–Pb concordia ages obtained for diorite and granodiorite porphyry show ages of 320.1 ± 2.2 Ma and 309.8 ± 2.2 Ma, suggesting a Late Carboniferous crystallization age. (2) Hydrothermal biotite displays 311.0 ± 1.8 Ma plateau age and excess Ar component caused by low-temperature alteration and hydration that produced an erroneously old age in low-temperature steps. (3) The combined effects of U and Th zonation or heterogeneous distribution and radiation damage may account for the range of ZHe ages (221.0–174.3 Ma for diorite and 261.6– 225.9 Ma for granodiorite porphyry), and radiation damage effect may be the main reason for the range of AHe ages (68.9–100.8 Ma for diorite and 91.0–152.0 Ma for granodiorite porphyry). The greater cooling impact of meteoric water and shallower depth lead to older ZHe and AHe ages for the granodiorite porphyry, relative to the diorite. (4) The inverse modeling simulation based on apatite and zircon (U–Th)/He suggests a five-episode cooling history for the diorite: very fast cooling during Late Carboniferous to Early Permian (320–290 Ma, stage I), followed by moderately fast cooling during Early Permian to Late Triassic (290–200 Ma, stage II), relatively slow cooling from Early Jurassic to Early Cretaceous (200–136 Ma, stage III), fast cooling from Early Cretaceous to Late Cretaceous (136–100 Ma, stage IV), and finally very slow cooling since Late Cretaceous (100 Ma, stage V).

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(5) As suggested by the thermal and tectonic history, our results show significant far-field effects of the Qiangtang–Eurasia and Lhasa–Qiangtang collision, but not the effect of India– Asia collision.

Acknowledgements The authors would like to thank Dr. Wang Fei and Dr. Yang LieKun for their assistance in Ar/Ar analysis, and Mrs Yan Xin and Miss Yang SaiHong for their assistance in cathodoluminescence images at the State Key Laboratory of Lithospheric Evolution of the Institute of Geology and Geophysics, Chinese Academy of Sciences. Special thanks to Brad McDonald and Celia Mayers for their assistance with (U–Th)/He analysis. The manuscript was improved by thoughtful comments of Li JinXiang. In particular, we express our gratitude to Professor Chen Huayong and another reviewer, and also editor Bernd Lehmann for their stimulating comments and excellent suggestions for improvement of the manuscript. This research was financially supported by Ministry of Land and Resources Comprehensive Research for Typical Mineral Deposits (Grant Number 20089932). References Allen, M.B., Vincent, S.J., 1997. Fault reactivation in the Junggar region, northwest China: the role of basement structures during Mesozoic–Cenozoic compression. Journal of the Geological Society 154, 151–155. Allen, M.B., Windley, B., Zhang, C., Guo, J., 1993. Evolution of the Turfan basin, Chinese central Asia. Tectonics 12, 889–896. An, F., Zhu, Y.F., 2009. SHRIMP U–Pb zircon ages of tuff in Baogutu formation and their geological significance. Acta Petrologica Sinica 25, 1437–1445 (in Chinese with English abstract). Ault, A.K., Flowers, R.M., Bowring, S.A., 2009. Phanerozoic burial and unroofing history of the western Slave craton and Wopmay orogen from apatite (U–Th)/ He thermochronometry. Earth and Planetary Science Letters 284, 1–11. Ballentine, C.J., Burnard, P.G., 2002. Production, release and transport of noble gases in the continental crust. Reviews in Mineralogy and Geochemistry 47, 481–538. Ballentine, C.J., Burgess, R., Marty, B., 2002. Tracing fluid origin, transport and interaction in the crust. Reviews in Mineralogy and Geochemistry 47, 539–614. Boyce, J., Hodges, K., 2005. U and Th zoning in Cerro de Mercado (Durango, Mexico) fluorapatite: insights regarding the impact of recoil redistribution of radiogenic 4 He on (U–Th)/He thermochronology. Chemical Geology 219, 261–274. Buckman, S., Aitchison, J., 2001. Middle Ordovician (Llandeilan) radiolarians from West Junggar, Xinjiang, China. Micropaleontology 47, 359–367. Bullen, M., Burbank, D., Garver, J., 2003. Building the northern Tien Shan: integrated thermal, structural, and topographic constraints. The Journal of Geology 111, 149–165. Chen, B., Arakawa, Y., 2005. Elemental and Nd–Sr isotopic geochemistry of granitoids from the West Junggar foldbelt (NW China), with implications for Phanerozoic continental growth. Geochimica et Cosmochimica Acta 69, 1307– 1320. Chen, S., Guo, Z.J., 2010. Time constrains, tectonic setting of Dalabute ophiolitic complex and its significance for Late Paleozoic tectonic evolution in West Junggar. Acta Petrologica Sinica 26, 2336–2344 (in Chinese with English abstract). Chen, B., Jahn, B.M., 2004. Genesis of post-collisional granitoids and basement nature of the Junggar Terrane, NW China: Nd–Sr isotope and trace element evidence. Journal of Asian Earth Sciences 23, 691–703. Chen, B., Zhu, Y.F., 2011. Petrology, geochemistry and zircon U–Pb chronology of gabbro in Darbut ophiolitic mélange, Xinjiang. Acta Petrologica Sinica 27, 1746– 1758 (in Chinese with English abstract). Chen, J.F., Han, B.F., Ji, J.Q., Zhang, L., Xu, Z., He, G.Q., Wang, T., 2010. Zircon U–Pb ages and tectonic implications of Paleozoic plutons in northern West Junggar, North Xinjiang, China. Lithos 115, 137–152. Coyle, D.A., Wagner, G.A., 1998. Positioning the titanite fission-track partial annealing zone. Chemical Geology 149, 117–125. Coyle, D.A., Wagner, G.A., Hejl, E., Brown, R., Van den Haute, P., 1997. The Cretaceous and younger thermal history of the KTB site (Germany): apatite fission-track data from the Vorbohrung. Geologische Rundschau 86, 203–209. Dai, H.W., Shen, P., Shen, Y.C., Pan, H.D., Liu, T.B., Meng, L., Guan, W.N., 2010. Mineralogy of ore-bearing porphyries in the Baogutu copper–gold belt of West Junggar and its geological significance. Xinjiang Geology 28, 440–447 (in Chinese with English abstract). Danišík, M., Pfaff, K., Evans, N.J., Manoloukos, C., Staude, S., McDonald, B.J., Markl, G., 2010. Tectonothermal history of the Schwarzwald Ore District (Germany): an apatite triple dating approach. Chemical Geology 278, 58–69.

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