Tectonic evolution and continental crust growth of

Earth-Science Reviews 126 (2013) 178–205

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Earth-Science Reviews journal homepage: www.elsevier.com/locate/earscirev

Tectonic evolution and continental crust growth of Northern Xinjiang in northwestern China: Remnant ocean model Qin-Qin Xu a,b, Jian-Qing Ji a,⁎, Lei Zhao a,c, Jun-Feng Gong a, Jing Zhou a, Guo-Qi He a, Da-Lai Zhong b, Jin-Duo Wang d, Lee Griffiths e a

Key Laboratory of Orogenic Belts and Crustal Evolution, School of Earth and Space Sciences, Peking University, Beijing 100871, China State Key Laboratory of Lithospheric Evolution, Institute of Geology and Geophysics, Chinese Academy of Science, Beijing 100029, China Institute of Geology, Chinese Academy of Geological Sciences, Beijing 100037, China d Shengli Oil Field Company Ltd., China Petroleum and Chemical Corp., Dongying 257022, China e Ionflight LLC, 50 Terminal St., Suite 311, Charlestown, MA 02129, USA b c

a r t i c l e

i n f o

Article history: Received 22 November 2012 Accepted 7 August 2013 Available online 19 August 2013 Keywords: Ophiolite Geochemistry Remnant ocean Continental crust growth Northern Xinjiang

a b s t r a c t The Northern Xinjiang region is located in the southwestern part of the Central Asian Orogenic Belt (CAOB, also known as the Altaid Tectonic Collage). Despite extensive research aimed at understanding the crustal growth of the CAOB and the evolution of the Paleo-Asian Ocean, the tectonic evolution mechanism of continental crust growth in Northern Xinjiang remains controversial. The geology of Northern Xinjiang is characterized by widespread ophiolites, granitoids, intermediate–basic dikes. Most of the ophiolites were generated in the early Paleozoic. The ophiolites are widely spread around the Junggar Basin, but their distribution does not indicate a welldefined band. Their outcrops are generally related to various faults. The basic rocks widespread in Northern Xinjiang are grouped into two categories: (i) gabbros, diabases basalts of the ophiolites and (ii) basic dikes that intrude into the Paleozoic strata granite plutons. The basic rocks associated with the early Paleozoic ophiolites were reworked by later geothermal events with a peak 40Ar/39Ar age of 310–290 Ma. The basic dikes intruded into Paleozoic strata and granite plutons during the Carboniferous–Jurassic, displaying three peaks of emplacement at 260–250 Ma, 220 Ma, and 200–190 Ma. These two types of basic rocks and the documented Variscan magmatic rocks were derived from the same source. Their isotope geochemical characteristics and widespread distribution suggest that since the Paleozoic, a large geochemical province has existed in Northern Xinjiang with an affinity to mid-ocean ridge basalts (MORB) and ocean island basalts (OIB), which is related to a long-lived remnant ocean and the underlying early Paleozoic oceanic crust. The existence of remnant oceanic crust in Northern Xinjiang was confirmed by seismic, gravity and aeromagnetic data. Therefore, we propose the following remnant ocean model for the Paleozoic tectonic evolution of Northern Xinjiang. It consists of three stages: 1) oceanic crust formation and deposition of the overlying volcanic-sedimentary rocks during the early Paleozoic; 2) retaining of the remnant ocean with marine sediments deposited during the early stage of the late Paleozoic; and 3) widespread and pervasive emplacement of Variscan granites, intermediate–basic dikes, and their volcanic equivalents during the Late Carboniferous and Early Permian, and termination of marine sedimentation at the end of the Early Permian. The tectonic evolution of Northern Xinjiang has been in a state of intracontinental deformation since the Mesozoic. The Variscan granitoids and basic dikes of Northern Xinjiang originated from the partial melting of the remnant oceanic crust formed in the early Paleozoic. These Variscan intrusive rocks represent the production of continental crust transferred from the basic crust. The Phanerozoic continental growth of Northern Xinjiang was completed by mass transfer from the early Paleozoic remnant oceanic crust; this approach may considerably change our views of continental growth. © 2013 Elsevier B.V. All rights reserved.

Contents 1. 2. 3.

Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Geological setting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Distribution and ages of the ophiolites in Northern Xinjiang . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

⁎ Corresponding author at: School of Earth and Space Sciences, Peking University, No. 5 Yiheyuan Road, Haidian District, Beijing 100871, China. Tel.: +86 10 6275 3040. E-mail address: [email protected] (J.-Q. Ji). 0012-8252/$ – see front matter © 2013 Elsevier B.V. All rights reserved. http://dx.doi.org/10.1016/j.earscirev.2013.08.005

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3.1. Ophiolite distribution . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.2. Ages of ophiolites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4. Field occurrences and petrology of basic rocks in Northern Xinjiang . . . . . . . . . . . . . . 5. Whole-rock Sr, Nd, and Pb isotope geochemistry . . . . . . . . . . . . . . . . . . . . . . 5.1. Analytical method . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.2. Results . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6. Laser 40Ar/39Ar dating . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.1. Analytical method . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.2. Results . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7. Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.1. Sr, Nd, and Pb isotopic data for Variscan magmatic rocks in Northern Xinjiang—a summary 7.2. Characteristics of the source area . . . . . . . . . . . . . . . . . . . . . . . . . . 7.3. The basement nature of the Junggar Basin based on geophysical data . . . . . . . . . . 7.4. Previously proposed tectonic models and the remnant ocean model . . . . . . . . . . 7.5. Tectonic evolution and continental crust growth of Northern Xinjiang . . . . . . . . . 7.5.1. Tectonic evolution of Northern Xinjiang . . . . . . . . . . . . . . . . . . . 7.5.2. Continental crust growth of Northern Xinjiang . . . . . . . . . . . . . . . 8. Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Acknowledgments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

1. Introduction The growth and evolution of the continental crust has been an important subject of research and debate in the earth sciences (Jahn et al., 2000a, 2000b, 2002c; Hawkesworth and Kemp, 2006). It is widely agreed that the formation of the continental crust was essentially completed in the Precambrian and that the amount of new crust formed in the Phanerozoic is minor or insignificant (Reymer and Schubert, 1984; Armstrong, 1991; Stein and Hofmann, 1994; Taylor and McLennan, 1995; Albarède, 1998; Condie, 1998; Hawkesworth and Kemp, 2006). This hypothesis was supported in earlier Nd isotope studies of granitoids from several intensely studied classic orogenic belts, such as the Caledonides, the Hercynides, the Yangtze–Cathaysia of southeast China, and the Himalayas (Allègre and Othman, 1980; Liew and Hofmann, 1988; Patchett, 1992; Darbyshire and Shepherd, 1994; Chen and Jahn, 1998). However, the idea of negligible growth in the Phanerozoic was challenged by the presence of very large volumes of “juvenile” crust (unlike recycled Precambrian crust) in several orogenic belts, such as the western North American Cordilleras (DePaolo, 1981; Samson et al., 1989; Samson and Patchett, 1991; Samson et al., 1995; Whalen et al., 1996), the Lachlan and New England Fold belts of eastern Australia (McCulloch and Chappell, 1982; Hensel et al., 1985; Collins, 1996, 1998) and the Altaid Tectonic Collage (Sengör et al., 1993; Sengör and Natal'in, 1996). The Central Asian Orogenic Belt (CAOB; Mossakovsky et al., 1994; Jahn et al., 2000b; Windley et al., 2007), also known as the Altaid Tectonic Collage (Sengör et al., 1993; Sengör and Natal'in, 1996), is situated between the European craton to the west, the Siberian craton to the east, and the Tarim and North China cratons to the south (Fig. 1A). It is the largest Paleozoic accretionary orogen in the world and is generally thought to be related to the closure of the Paleo-Asian Ocean (the Paleo-Ocean defined by units located between the European, Siberian, Tarim, and North China cratons during the Neoproterozoic–late Paleozoic; Coleman, 1989; Zonenshain et al., 1990; Şengör et al., 1993; Mossakovsky et al., 1994; Dobretsov et al., 1995; Şengör and Natal'in, 1996; Buslov et al., 2001; Dobretsov et al., 2003, and references herein; Kovalenko et al., 2004; Windley et al., 2007), representing one of the most important sites of Phanerozoic crustal growth in the world (Sengör et al., 1993; Dobretsov et al., 1995; Kovalenko et al., 1996a, 1996b; Hu et al., 2000; Jahn et al., 2000a, 2000b, 2000c, 2004; Hong et al., 2004; Jahn, 2004; Kovalenko et al., 2004; Wilde et al., 2010). Northern Xinjiang, which occupies the northwest edge of China, is an important part of the CAOB (Fig. 1A). It has been the subject of

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extensive studies that examined the crustal growth of the CAOB and the evolution of the Paleo-Asian Ocean (Coleman, 1989; Feng et al., 1989; Xiao and Tang, 1991; Xiao et al., 1992; Sengör et al., 1993; He et al., 1994b; Sengör and Natal'in, 1996; Hu et al., 2000; Jahn et al., 2000b; Liu, 2002; Wang et al., 2003; Buckman and Aitchison, 2004; Xiao et al., 2008). However, the mechanism of continental crust growth and tectonic evolution in Northern Xinjiang remains controversial and is currently under debate, with models of a continuous single subduction– accretion processes (Sengör et al., 1993; Sengör and Natal'in, 1996; Wang et al., 2003), collision of various terranes with multiple subduction systems (Didenko et al., 1994; Mossakovsky et al., 1994; Buchan et al., 2001; Buslov et al., 2001; Badarch et al., 2002; Windley et al., 2002), an accretionary wedge (Xiao et al., 2008, 2009; Zhang, 2009), or mid-ocean ridge subduction (Liu et al., 2007, 2009; Geng et al., 2009; Yin et al., 2010). These models can be adapted to certain aspects of the tectonics of Northern Xinjiang. However, they are not supported by the following lines of geological facts: (1) the ophiolites are widespread around the Junggar Basin but their distribution does not indicate a band pattern (Fig. 1B); (2) the exposure of ophiolites is generally related to different faults and exhumation depths (Fig. 1B), and the strata in the fault zones are deformed by subvertical cleavages (Fig. 2A, B, and D); (3) outside the fault zones, the Paleozoic strata lie almost horizontally or tilt gently (Fig. 2B and C). This study focuses primarily on basic rocks of different ages and occurrences in Northern Xinjiang. Based on geology, geochemistry, and geochronology observations, we present a new model for the tectonic evolution of Northern Xinjiang, which can better explain the continental crust growth of Northern Xinjiang. 2. Geological setting Northern Xinjiang is located in the southwestern region of the CAOB (Fig. 1A) and includes (from north to south) the Altay, the Junggar Basin, and the Tian Shan tectonic domains (Fig. 3). According to the geological data (Feng, 1985; Feng et al., 1989; BGMRXUAR, 1993), no Precambrian strata has been clearly identified north of Tian Shan (Fig. 3B), and the oldest units are the Ordovician formations intermediate to basic volcanic rocks and flysch-type deposits. The Ordovician units are exposed mainly in the southwestern areas of the Junggar Basin and the Altay region. The main outcrops of Silurian rocks in the west are around the Tangbale and Mayile ophiolites, and in the east around the Kalamaili ophiolite (Figs. 1B and 3B). The Silurian strata are mainly Early-Silurian flysch-type deposits with ophiolitic

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Fig. 1. (A) Northern Xinjiang in the Central Asian Orogenic Belt (modified from Sengör et al., 1993; Jahn et al., 2000c). SCB — South Caspian Basin. The approximate location of Fig. 1B is indicated. The distribution of remnant ocean basins are from Aplonov (1995) (for West Siberia); Brunet et al. (2003) (for SCB); He et al. (2004) (for the west of Junggar). (B) Simplified geological map of Northern Xinjiang showing the major ophiolites and their ages (modified from Ma, 2002). Age data taken from a: Xiao et al. (1992), b: Xu et al. (2012), c: He et al. (2007), d: Zhu and Xu (2006), e: Jian et al. (2005), f: Jian et al. (2003), g: Xiao et al. (2006b), h: He et al. (2001), and i: this study.


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Fig. 2. Photographs showing the outcrops of Paleozoic strata in the study area. (A) Carboniferous strata deformed by a fault and subvertical cleavages (location: N45°48′03″, E85°05′43″) with the original bedding plane visible. (B) Carboniferous strata, lying almost horizontally outside the fault zone, but deformed in the fault zone (location: N45°47′52″, E85°07′18″). (C) Devonian strata, tilting gently outside the fault zone (location: N45°38′45″, E82°46′29″). (D) Devonian strata deformed by fault and subvertical cleavages (location: N45°37′54″, E82°46′11″).

Fig. 3. (A) Geographic position of Northern Xinjiang in East Asia (modified from Zhang et al., 2013); (B) Geological map of Northern Xinjiang showing sampling localities (modified from BGMRXUAR, 1993). F1 — Chepaizi fault; F2 — South Junggar fault; F3 — Kalamaili Fault; F4 — North Junggar fault; F5 — Karamay fault. The Junggar Basin is bound by F1–F5 (Geology Department, Chinese Academy of Sciences and Xinjiang Petroleum Administration Bureau, 1989; Zhou, 1994).

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Fig. 3 (continued).

debris, unconformably overlying the Tangbale ophiolite. Early- and Late-Silurian flysch sediments are interlaid by Middle-Silurian pillow lavas and radiolarian cherts at Mayile. In the east, the undivided Silurian sequences, distributed along the Kalamaili ophiolite, are composed of gray green sandstones, siltstones, and sandy limestone. The strata in this region are dominated by Devonian to Carboniferous sediments (Fig. 3B). The widespread Devonian rocks consist of mainly marine mudstone, radiolarian cherts, limestone, volcanoclastic rocks, and andesitic volcanic rocks. The Carboniferous sedimentary rocks are mainly flysch-type sediments deposited under shallow marine conditions with minor interbedded pyroclastic materials. The Permian rocks in Northern Xinjiang are composed of terrestrial volcanic and volcanoclastic rocks with minor marine sediments. The Triassic–Quaternary units consist of alluvial, fluvial, and lacustrine deposits. Importantly, the geology of Northern Xinjiang is characterized by widespread ophiolites, granitoids, and intermediate–basic dikes (Figs. 1B and 3B). Geochronological data suggest that the ophiolites were formed in the Paleozoic, with ages ranging from 561 Ma to 332 Ma (Feng et al., 1989; Kwon et al., 1989; Xiao et al., 1992; Zhang and Huang, 1992; Xiao et al., 2006b; Huang et al., 1997; He et al., 2001; Jian et al., 2003, 2005; Xu et al., 2006; Zhu and Xu, 2006; He et al., 2007; Gu et al., 2009), indicating that the oceanic crust formation began during the Cambrian. The granitoids in this region are generally

accepted as post-collisional granites originating either from depleted mantle (Han et al., 1997, 1999; Chen and Arakawa, 2005) or from Paleozoic oceanic crust (e.g., Coleman, 1989; Feng et al., 1989; Hu et al., 2000). The emplacement times for these granitoids are between the Middle–Late Carboniferous and the Permian (Hopson et al., 1989; Han et al., 1997; Chen and Jahn, 2004; Han et al., 2006; Su et al., 2006a, 2006b; Briggs et al., 2007; Long et al., 2008; Zhou et al., 2008b; Wang et al., 2009; Chen et al., 2010; Han et al., 2010b). They are characterized by low initial 87Sr/86Sr ratios, positive εNd, and young Sm–Nd model ages (TDM), which is in sharp contrast with the coeval granitoids emplaced in the European Caledonides and Hercynides (Jahn et al., 2000a, 2000b, 2000c; and references therein), indicating Phanerozoic crust growth in the CAOB (Han et al., 1997, 1998, 1999; Hu et al., 2000; Jahn et al., 2000a, 2000b, 2000c; Chen and Jahn, 2004; Chen and Arakawa, 2005; Han et al., 2006; Chen et al., 2010; Tong et al., 2010). The intermediate– basic dikes were emplaced from the Carboniferous until the Jurassic (Li et al., 2004; Xu et al., 2008; Zhou et al., 2008a; Zhang et al., 2009; Yin et al., 2010), and are also considered to be related to the crustal growth during the Phanerozoic (Li t al., 2004; Han et al., 2006; Xu et al., 2008). The Tian Shan mountain range extends roughly east–west and separates the Junggar Basin in the north from the Tarim Basin to the south (Fig. 3A). It is a late Paleozoic orogen that was strongly modified by large-scale strike–slip faulting in an intracontinental setting during


the Late Permian and Early Triassic (Shu et al., 1998; Laurent-Charvet et al., 2002, 2003; Li et al., 2002; Wang et al., 2002; Natal'in and Şengör, 2005; Van der Voo et al., 2006). The present-day Tian Shan is believed to have been shaped by Cenozoic uplift in response to the India–Asia collision (e.g., Avouac et al., 1993; Hendrix et al., 1994; Abdrakhmatov et al., 1996; Yin et al., 1998; Bullen et al., 2001, 2003; Charreau et al., 2005, 2006; Huang et al., 2006; De Grave et al., 2007; Zhang et al., 2007; Sun et al., 2009; Wang et al., 2010). The Altay range, extending in the southeast–northwest direction (Fig. 3A), is also a late Paleozoic orogen (He et al., 2004; Yuan et al., 2004; Xu et al., 2005) that has been undergoing multi-stage uplift and denudation since the Mesozoic (Liu et al., 2002; Yuan et al., 2004; Bao et al., 2005; Guo et al., 2006; Yuan et al., 2006; Liu et al., 2012). The present-day topography and relief of the Altay began forming 25–15 Ma (Liu et al., 2012) or 8–6 Ma (Xu, 2011) as a result of the farfield effects of the India–Asia collision (Xu, 2011; Liu et al., 2012). The Junggar Basin is a large petroliferous basin bound by faults F1–F5 (Fig. 3B; Geology Department, Chinese Academy of Sciences and Xinjiang Petroleum Administration Bureau, 1989; Zhou, 1994). There is no agreement concerning its origin, nature, and age (Hsü, 1988; Carroll et al., 1990; Zhao, 1992a; Cai et al., 2000). Hsü (1988) suggested that the Junggar was a Carboniferous back-arc basin separated from the Paleo-Tethys Ocean to the south by a late Paleozoic frontal arc and filled

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with Carboniferous and Permian flysch. Carroll et al. (1990) pointed out that the Junggar Basin originated during the late Paleozoic as either a remnant early- to middle-Paleozoic ocean basin or a Middle–Carboniferous back-arc basin bound by emergent volcanic arcs south of the Siberian Craton. Zhao (1992a) suggested that the Junggar Basin was originally formed by an upper mantle uplift during the Early–Middle Carboniferous. Cai et al. (2000) proposed that the basin was initially rifted in the Early Permian followed by thermal cooling subsidence during the Late Permian. It then went through the cratonic basin stage from the Triassic to the Paleogene, and finally developed into a foreland basin. Despite the varying opinions and controversy surrounding the origins of the Junggar Basin, many researchers agree that the Junggar Basin is a superimposed basin that originated during the late Paleozoic (Wu, 1986; Zhao, 1992a; Kuang, 1993; Wang et al., 1999; Zhao et al., 2008b, 2010a). 3. Distribution and ages of the ophiolites in Northern Xinjiang 3.1. Ophiolite distribution Throughout the study area, the widespread ophiolites display the following distribution characteristics: (1) The ophiolites are widely distributed around the Junggar Basin, but not as a well-defined band (Figs. 1B and 4).

Fig. 4. Distribution of ophiolites on the southwest side of Junggar Basin (location shown in Fig. 1B).

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(2) The ophiolite suites are generally complete and usually consist of metamorphic peridotite (serpentinite), layered gabbros and dikes, basaltic lavas, and cherts, but certain areas display a lack of cherts or pillow lavas. The basaltic lavas have both layered and pillow structures. (3) The strata deposited on the ophiolites are early Paleozoic and dominated by Ordovician–Silurian volcanoclastic rocks, cherts, flysch, and limestone (Table 1; Fig. 4). The ophiolites and Ordovician–Silurian strata generally have distinct unconformable or fault contacts with the Devonian and Carboniferous strata (Table 1; Figs. 4–6). The Devonian and Carboniferous strata sometimes lie directly and evenly over the ophiolites where the early Paleozoic strata are absent. 4. The ophiolite outcrops are generally related to various faults formed in different times (Table 1; Figs. 4–6). For example, the outcrop of the Darbut ophiolite is clearly orientated along the Darbut fault (Fig. 5A) striking east–northeast. According to the relationship between the faults and strata and the deformation characteristics of the strata, this fault can be classified as a Neogene sinistral strike–slip fault (Xu et al., 2009). Moreover, the southwest part of the outcrop is affected by the nearly E–W-striking fault (Fig. 5B) formed in the Paleogene (Xu et al., 2009). Similarly, the Kalamaili ophiolite zone extends along the Kalamaili fault zone striking 290° (Fig. 6). The dragging of chert lenticles in the fault zone indicates that the Kalamaili fault zone is a dextral strike–slip belt (Zhao et al., 2012); seismic data and 40Ar/39Ar age show that this fault zone was formed 270–260 Ma (Lu, 2008; Zhao et al., 2012; Tu et al., 2013).

3.2. Ages of ophiolites According to previous studies, ophiolite ages in Northern Xinjiang range from 561 Ma to 332 Ma (Table 1; Feng et al., 1989; Kwon et al., 1989; Xiao et al., 1992; Zhang and Huang, 1992; Huang et al., 1997; He et al., 2001; Jian et al., 2003, 2005; Xiao et al., 2006b; Xu et al., 2006; Zhu and Xu, 2006; He et al., 2007; Gu et al., 2009). In the eastern part of the region, 13 zircons from the layered gabbro in Zhaheba yielded an emplacement age of 489 ± 4 Ma (Jian et al., 2003), and a plagiogranite in the Tuziquan ophiolite section of Zhaheba was dated at 503 ± 7 Ma (Xiao et al., 2006b). In the south of Zhaheba, microfossils obtained from the Kalamaili ophiolite indicated that the formation age of this ophiolitic mélange was Middle–Late Ordovician (He et al., 2001). In the western part of Northern Xinjiang, the plagioclase and sphene Pb–Pb ages of the Tangbale leucogabbro are around 523–508 Ma (Kwon et al., 1989; Xiao et al., 1992), and the Sm–Nd isochron age of the Tangbale ophiolite is between 489 and 447 Ma (Zhang and Huang, 1992). Radiolarian fossils of the Middle Ordovician were found in cherts from the Tangbale ophiolite, which stratigraphically intercalate with or overlie mafic lavas of the ophiolite (Feng et al., 1989; Xiao et al., 1992). In the north of the Tangbale belt, Silurian radiolarian fossils were found in the cherts of the Mayile ophiolite (Feng et al., 1989; Xiao et al., 1992; Zhang and Huang, 1992). However, the zircon U–Pb ages of Yushitasi gabbro, diorite, and granite in the ophiolite from southern Mayile reveal

Table 1 Compiled age data and related information for the ophiolites in Northern Xinjiang. No.

Location

Coordinates

Strata deposited on the ophiolites & contact relationship

Trend of structure line

Age

1

Tangbale

45°05′–45°20′N 83°00′–83°33′E

60°, 315°

2

A'lashankou-Aibi Lake

45°10′–45°20′N 82°45′–83°00′E

Pb–Pb; 508 Ma (Xiao et al., 1992), 523 Ma (Kwon et al., 1989); Sm–Nd; 489–447 Ma (Zhang and Huang, 1992); Radiolaria; Middle Ordovician (Feng et al., 1989; Xiao et al., 1992) Zircon U–Pb; 531 Ma (Jian et al., 2005)

3

Mayile

45°20′–45°30′N 82°30′–83°00′E

4

Kufu

45°20′–45°45′N 83°00′–83°20′E

5

Southwest of Darbut

45°40′–46°00′N 84°05′–85°00′E

Ordovician sandstone, shale, sand limestone, and volcaniclastic rocks; Silurian flysch deposits, tuff, and andesitic volcanic rocks; fault and unconformable contact Ordovician tuff, tuff sandstone, phyllite, and sandstone; fault and unconformable contact with Middle Devonian Silurian tuff, tuff sandstone, and sandstone; fault and unconformable contact with Middle Devonian and Early Carboniferous Silurian tuff, tuff sandstone, muddy sandstone, and chert; fault and unconformable contact with Middle Devonian and Early Carboniferous Devonian spilitic pillow lava, chert, and limestone/Ordovician chert; fault and unconformable contact with Early Carboniferous

6

Northeast of Darbut

46°00′–46°08′N 85°00′–85°10′E

7

Baijiantan-Baikouquan

45°40′–45°50′N 85°00′–85°10′E

8

Tarbgatay

46°50′–47°00′N 83°15′–83°35′E

9

Hongguleleng

46°40′–46°50′N 85°15′–86°30′E

10

Zhaheba

46°05′–46°35′N 89°10′–90°00′E

11

Tuziquan

46°20′–46°35′N 89°00′–89°15′E

12

Kalamaili

44°40′–45°15′N 89°10′–90°55′E

Devonian spilitic pillow lava, chert, and limestone; fault and unconformable contact with Carboniferous Ordovician chert; unconformable contact with Carboniferous–Permian Early Ordovician limestone, chert, and tuff; fault contact with Late Devonian and Early Carboniferous Ordovician tuff, limestone and tuff sandstone; fault contact Ordovician chert and carbonate; fault and unconformable contact with Middle Devonian Ordovician chert and carbonate; fault and unconformable contact with Middle Devonian Ordovician calcareous sandstone, flysch, and volcanic rocks; fault and unconformable contact with Devonian– Carboniferous

315°, 340°

80°, 290°

20°, 45°, 80°

60°, 80°

60°

60°

290°

45°, 290° 315°, 330°

320°

290°

Radiolaria; Silurian (Feng et al., 1989; Xiao et al., 1992; Zhang and Huang, 1992); Zircon U–Pb; 415 Ma (Jian et al., 2005) Early–Middle Cambrian (Xu et al., 2012) Radiolaria; Silurian (Feng et al., 1989; Xiao et al., 1992; Zhang and Huang, 1992); Zircon U–Pb; 415 Ma (Jian et al., 2005) Sm–Nd; 395Ma (Zhang and Huang, 1992); Zircon U–Pb; 391 Ma (Gu et al., 2009); Microfossils; Middle–Late Ordovican (He et al., 2007) Conodonts; Ordovician (Samygin et al., 1997) Sm–Nd; 385 Ma (Zhang and Huang, 1992); Zircon U–Pb; 391 Ma (Gu et al., 2009) Microfossils; Middle–Early Ordovician (He et al., 2007); Zircon U–Pb; 414–332 Ma (Xu et al., 2006) Zircon U–Pb; 478 Ma (Zhu and Xu, 2006)

Zircon U–Pb; 475 Ma (Jian et al., 2005); Sm–Nd; 444 Ma (Zhang and Huang, 1992) Zircon U–Pb; 489 Ma (Jian et al., 2003, 2005); Sm–Nd; 479 Ma (Liu and Zhang, 1993) Zircon U–Pb; 503 Ma (Xiao et al., 2006b); Sm–Nd; 561 Ma (Huang et al., 1997); Radiolaria; Ordovician (Li, 1991) Microfossils; Middle–Late Ordovician (He et al., 2001)


185

Fig. 5. (A) Distribution of the Darbut ophiolite on the west side of Junggar Basin (location shown in Fig. 1B); (B) Cross-section of the Darbut ophiolite. The strata in the fault zones were seriously deformed by shear faulting and intense cleavages (right photograph). Consequently, the strata were broken into thin slices and it is difficult to get the original bedding plane. The cleavage planes are sub-parallel with the fault planes.

that this mélange might have been generated during the Early–Middle Cambrian (Xu et al., 2012). Zhu and Xu (2006) discovered an ophiolite mélange in the Tarbgatay Mountains, with a zircon SHRIMP U–Pb age of 478 Ma. The formation age of the Hongguleleng ophiolite was constrained within the range of 475–444 Ma (Zhang and Huang, 1992; Jian et al., 2005). Xu et al. (2006) discovered an ophiolitic mélange distributed along the northwest edge of the Junggar Basin, and the ages of 414 and 332 Ma were determined by zircon SHRIMP dating of the altered gabbro samples taken from a Baijiantan segment. Because this ophiolitic

mélange has experienced multiple deformations (Zhu and Xu, 2007), these two ages are controversial (He et al., 2007). Microfossils obtained from the Karamay ophiolite demonstrate that the formation age of this ophiolitic mélange was Middle–Late Ordovician (He et al., 2007). Radiolarians from cherts overlying the Darbut ophiolite were dated as Middle Devonian (Feng et al., 1989; Xiao et al., 1992). The whole-rock Sm–Nd isochron age of the gabbros was measured at 395 Ma (Zhang and Huang, 1992). However, this age was not completely accepted due to the uncertainty of the Sm–Nd isochron method used for determining

186


Fig. 6. (A) Distribution of the Kalamaili ophiolite on the east side of Junggar Basin (location shown in Fig. 1B); (B) Cross-section of the Kalamaili ophiolite.

the ophiolite formation age (Gu et al., 2009). The LA-ICP-MS zircon U–Pb dating of gabbros from the Darbut ophiolite gave an age of 391 Ma (Gu et al., 2009). Despite these controversies, the formation age of ophiolites in Northern Xinjiang has been determined generally as early Paleozoic except for the Darbut ophiolite (Table 1; Feng et al., 1989; Kwon et al.,

1989; Xiao et al., 1992; Zhang and Huang, 1992; Jian et al., 2005; Zhu and Xu, 2006; He et al., 2007; Gu et al., 2009). The reported ages seem to indicate that the Darbut ophiolite was formed in the Early–Middle Devonian (Feng et al., 1989; Xiao et al., 1992; Zhang and Huang, 1992; Gu et al., 2009). However, there are no complete late Paleozoic ophiolitic assemblages found in Darbut; therefore, the geochronology and the


187

Fig. 7. Field examples of basic rocks in Northern Xinjiang. (A) Basic dikes in Karamay on the west side of the Junggar Basin; (B) Basic dikes in Qiakuerte on the northeast side of the Junggar Basin; (C) Gabbro dikes emplaced into Late Carboniferous near Fuyun County on the southern margin of Altay; (D) Basic dikes in Toksun on the southern margin of Tian Shan; (E) The thick layer of basalts of Tangbale ophiolite on the southwest side of the Junggar Basin; (F) Basalts of ophiolite in Sujiquan on the east side of the Junggar Basin; (G) Weathered peridotites and weatherresistant diabases in A'lashankou on the southwest side of the Junggar Basin; (H) Weathered peridotites and weather-resistant basalts in Zhaheba on the east side of the Junggar Basin.

microfossils with Devonian ages cannot be used to determine the development and age of the ophiolite (Zhao, 2010). Otherwise, the early Paleozoic ophiolitic assemblages are complete (Zhao, 2010),

and Ordovician radiolarians (Shu et al., 2001) and conodonts (Samygin et al., 1997) have been found in the Darbut ophiolite. Moreover, the ophiolitic mélange with Ordovician conodonts is unconformably overlain

188


Fig. 8. Photomicrographs of intermediate–basic dikes (A–E) and basic rocks of ophiolites (F–H) in Northern Xinjiang. (A) Euhedral plagioclase in the diabase; (B) Diabase texture; (C) Semieuhedral clinopyroxene in the gabbro; (D) Gabbro texture; (E) Porphyritic texture; (F) Clinopyroxene and plagioclase in the diabase; (G) Diabase texture; (H) Gabbro texture. Pl — plagioclase; Cpx — clinopyroxene.


189

Table 2 Sr, Nd, and Pb isotope analyses of basic rocks in Northern Xinjiang. Sample

SJQ-3

TBL-4

BJT-9

QIA-3-3

AL-8

TBL-3-1

TBL-3-3

BJT-1-3

0820SE-4

Location

44°52′32″N

45°06′32″N

45°43′41″N

46°29′11″N

45°09′11″N

45°06′39″N

45°06′39″N

45°43′43″N

45°56′57″N

90°30′07″E

83°23′10″E

85°07′07″E

89°07′02″E

83°10′27″E

83°21′32″E

83°21′32″E

85°07′13″E

90°11′45″E

Pb/206Pb 207 Pb/206Pb 208 Pb/206Pb 206 Pb/204Pb 207 Pb/204Pb 208 Pb/204Pb Rb (μg/g) Sr (μg/g) (87Rb/86Sr)ma (87Sr/86Sr)ma 2σ t (Ma) (87Sr/86Sr)t Sm (μg/g) Nd (μg/g) (147Sm/144Nd)ma (143Nd/144Nd)ma 2σ εNd(0) εNd(t) (143Nd/144Nd)t

0.05496 0.85319 2.08174 17.69405 15.49524 37.39013 2.72 573.60 0.013721 0.703556 0.000028 470 0.703464 5.00 19.92 0.151827 0.512903 0.000014 5.16 7.87 0.512435

0.05465 0.85300 2.08892 17.87312 15.58403 37.96413 24.18 607.10 0.115247 0.706703 0.000014 500 0.705882 4.04 15.67 0.155817 0.512857 0.000028 4.27 6.89 0.512346

0.05195 0.81112 2.01852 18.14478 15.55127 38.19371 17.97 268.79 0.193474 0.705857 0.000022 460 0.704589 4.22 14.03 0.181750 0.512969 0.00002 6.45 7.34 0.512421

0.05315 0.83319 2.06185 18.04629 15.63357 38.03206 48.20 732.40 0.190431 0.705871 0.000016 489 0.704544 3.18 11.61 0.165976 0.512793 0.000022 3.02 4.94 0.512261

1.30 146.78 0.025565 0.704154 0.000022 500 0.703972 3.46 11.49 0.182153 0.512831 0.00002 3.77 4.71 0.512235

0.05224 0.81460 2.00633 17.96380 15.52626 37.82718 0.96 185.78 0.014921 0.704919 0.000020 508 0.704811 1.82 6.76 0.162860 0.512808 0.000022 3.31 5.52 0.512266

0.05159 0.80659 1.99746 18.00152 15.55419 37.84025 12.70 211.60 0.173698 0.707011 0.000016 508 0.705753 1.69 5.96 0.171796 0.512859 0.00002 4.31 5.94 0.512287

0.05400 0.83955 2.06745 17.86225 15.51046 37.69766 2.43 74.92 0.093930 0.706188 0.000038 460 0.705573 1.80 5.78 0.188090 0.512895 0.00002 5.01 5.52 0.512328

1.51 284.40 0.015408 0.704539 0.000012 503 0.704428 1.48 5.11 0.174888 0.512812 0.000024 3.39 4.79 0.512235

Sample

XZ-5

BKQ-7

YM-4b

YM-8b

YM-12b

YM-16b

YM-24b

YM-31b

YM-39b

204

Location

45°46′11″N

45°59′42″N

49°40′12″N

45°40′30″N

45°40′26″N

45°40′26″N

45°40′23″N

45°41′20″N

45°39′47″N

84°26′32″E

85°17′37″E

84°43′12″E

84°42′11″E

84°41′24″E

84°41′24″E

84°38′06″E

84°33′47″E

84°50′17″E

Pb/206Pb Pb/206Pb 208 Pb/206Pb 206 Pb/204Pb 207 Pb/204Pb 208 Pb/204Pb Rb (μg/g) Sr (μg/g) (87Rb/86Sr)ma (87Sr/86Sr)ma 2σ t (Ma) (87Sr/86Sr)t Sm (μg/g) Nd (μg/g) (147Sm/144Nd)ma (143Nd/144Nd)ma 2σ εNd(0) εNd(t) (143Nd/144Nd)t

0.05541 0.86344 2.10514 17.83714 15.57076 37.87602 2.22 146.06 0.043888 0.704270 0.000020 460 0.703983 1.51 4.37 0.209008 0.513029 0.000024 7.64 6.92 0.512400

0.05538 0.86268 2.10944 18.00409 15.57336 38.04508 2.79 366.40 0.022018 0.705191 0.000024 460 0.705047 1.31 2.57 0.309662 0.513349 0.000042 13.87 7.24 0.512416

0.05270 0.81875 2.03803 17.96180 15.48330 37.96480 62.67 492.78 0.367900 0.705304 0.000018 250 0.703996 4.31 17.49 0.149200 0.512975 0.000011 6.57 8.09 0.512731

0.05466 0.84633 2.07963 17.82325 15.45932 37.66178 33.00 274.39 0.347900 0.705280 0.000019 250 0.704043 7.72 34.51 0.135300 0.512901 0.00001 5.13 7.09 0.512680

0.05399 0.83853 2.06847 18.06274 15.50653 37.76950 48.35 529.09 0.264300 0.704950 0.000019 250 0.704010 2.08 9.62 0.130900 0.512919 0.000011 5.48 7.59 0.512705

0.05426 0.84325 2.07399 17.94683 15.51586 37.75651 51.36 670.54 0.221500 0.704901 0.000020 250 0.704113 2.34 11.75 0.120200 0.512880 0.000011 4.72 7.17 0.512683

0.05460 0.85136 2.08312 17.56567 15.55476 37.92035 54.24 634.70 0.247200 0.704450 0.000018 250 0.703571 4.10 17.64 0.140600 0.512961 0.000008 6.30 8.10 0.512731

0.05513 0.85899 2.09627 18.09149 15.57877 37.98514 18.30 655.10 0.080800 0.704005 0.000018 250 0.703718 2.92 13.39 0.132100 0.512923 0.000009 5.56 7.63 0.512707

0.05419 0.84037 2.07207 17.78009 15.47374 37.67751 33.33 546.87 0.176300 0.704510 0.000019 250 0.703883 3.44 15.40 0.135300 0.512949 0.000011 6.07 8.03 0.512728

Sample

YM-49b

YM-69b

YM-77b

YM-82b

YM-107b

YM-134b

YM-133b

QJ2-5b

BLK3-2b

Location

46°35′06″N

46°58′36″N

46°59′36″N

46°59′36″N

46°22′07″N

42°25′36″N

42°27′25″N

43°29′09″N

43°57′01″N

85°39′09″E

89°22′52″E

89°33′56″E

89°33′56″E

89°09′07″E

88°31′33″E

88°32′22″E

91°07′28″E

92°56′28″E

18.70 558.42 0.096800 0.704491 0.000018 262 0.704130 6.62 35.08 0.114100 0.512799 0.000011 3.14

0.05474 0.84748 2.07271 18.26880 15.48250 37.58400 15.22 286.47 0.153700 0.705212 0.000018 270 0.704622 2.81 8.96 0.189800 0.513006 0.000012 7.18

58.64 372.50 0.455400 0.705570 0.000018 217 0.704165 10.49 50.10 0.126600 0.512848 0.000012 4.10

0.05284 0.82016 2.04700 18.17475 15.48415 38.02845 26.07 512.11 0.147300 0.704428 0.000018 217 0.703974 10.27 48.45 0.128200 0.512845 0.000008 4.04

0.05463 0.84955 2.08540 17.94037 15.53276 37.92628 8.15 806.72 0.029200 0.705120 0.000020 229 0.705025 11.80 52.28 0.136500 0.512892 0.000008 4.95

0.05412 0.84379 2.07850 18.05888 15.56921 38.26953 48.50 625.94 0.224100 0.705563 0.000019 205 0.704910 6.97 34.56 0.122100 0.512742 0.000014 2.03

0.05437 0.85039 2.09166 18.21954 15.63173 38.40337 23.28 838.35 0.080300 0.706491 0.000018 202 0.706260 8.34 42.65 0.118300 0.512669 0.000012 0.60

0.05368 0.83364 2.05998 17.52861 15.47508 38.00607 24.22 953.85 0.073500 0.705689 0.000020 204 0.705476 4.18 16.68 0.151500 0.512868 0.00001 4.49

0.05471 0.84907 2.08101 17.89724 15.49984 37.85664 52.14 713.48 0.211400 0.705750 0.000020 210 0.705119 5.06 26.63 0.114900 0.512692 0.000012 1.05

204 207

204

Pb/206Pb 207 Pb/206Pb 208 Pb/206Pb 206 Pb/204Pb 207 Pb/204Pb 208 Pb/204Pb Rb (μg/g) Sr (μg/g) (87Rb/86Sr)ma (87Sr/86Sr)ma 2σ t (Ma) (87Sr/86Sr)t Sm (μg/g) Nd (μg/g) (147Sm/144Nd)ma (143Nd/144Nd)ma 2σ εNd(0)

(continued on next page)

190


Table 2 (continued) Sample

YM-49b

YM-69b

YM-77b

YM-82b

YM-107b

YM-134b

YM-133b

QJ2-5b

BLK3-2b

Location

46°35′06″N

46°58′36″N

46°59′36″N

46°59′36″N

46°22′07″N

42°25′36″N

42°27′25″N

43°29′09″N

43°57′01″N

85°39′09″E

89°22′52″E

89°33′56″E

89°33′56″E

89°09′07″E

88°31′33″E

88°32′22″E

91°07′28″E

92°56′28″E

5.91 0.512603

7.42 0.512671

6.04 0.512668

εNd(t) (143Nd/144Nd)t

5.94 0.512663

6.72 0.512687

3.98 0.512578

2.63 0.512513

5.67 0.512666

3.25 0.512534

a Isotope ratios for samples at the present day. εNd(t) = [(143Nd/144Nd)t/(143Nd/144Nd)CHUR, t-1] × 104; (143Nd/144Nd)t = (143Nd/144Nd)m-(147Sm/144Nd)m × (eλt-1); (143Nd/144Nd)CHUR, t = (143Nd/144Nd)CHUR-(147Sm/144Nd) CHUR × (eλt-1); (143Nd/144Nd)CHUR = 0.512638; (147Sm/144Nd) CHUR = 0.1967; λ = 6.54 × 10−12 yr−1. b Data from Xu et al. (2008).

by Devonian–Carboniferous mafic volcanic rocks, clastic rocks, and cherts with slight deformation (Samygin et al., 1997; Zhao and He, 2013). In summary, nearly all of the ophiolites in Northern Xinjiang were generated in the early Paleozoic (Fig. 1B; Xiao et al., 1992; Zhang and Huang, 1992; He et al., 2001; Jian et al., 2003, 2005; Xiao et al., 2006b; Xu et al., 2006; Zhu and Xu, 2006; He et al., 2007; Xu et al., 2012). In addition, ophiolites with similar formation times (compared with western Junggar) are also widespread in eastern Kazakhstan (Zhao and He, 2013).

4. Field occurrences and petrology of basic rocks in Northern Xinjiang

The lower part of the ophiolite in Tangbale consists of serpentinite peridotite covered by a thick layer of basalts (Fig. 7E). The ophiolite in Sujiquan is similar to that of Tangbale; it was also affected by serpentinization and is overlain by basalts (Fig. 7F). The serpentinite peridotites in A'lashankou and Zhaheba have been extensively weathered, leaving only weather-resistant diabases and basalts (Fig. 7G and H). Gabbros and diabases, including certain diorites and dioritic porphyrites, dominate the dikes that intrude into the Paleozoic strata and granite plutons. The minerals in the diabase are mainly plagioclase and clinopyroxene, hornblende, biotite, chlorite, and opaque minerals. Most crystals are euhedral and semi-euhedral (Fig. 8A–E). Gabbro, diabase, and porphyritic textures are common in these basic rocks

The basic rock samples can be divided into two types: (a) basic rocks of the ophiolites, including gabbros, diabases, and basalts; and (b) basic dikes that intrude the Paleozoic strata and granite plutons. We investigated 23 samples of ophiolites and 56 samples of intermediate–basic dikes (Fig. 3B) and selected results have been published in Chinese (see Xu et al., 2008; Zhou et al., 2008a). The dikes in Karamay on the west side of the Junggar Basin are 2–3 m wide, striking N40E (Fig. 7A). They intrude into the Karamay granodiorite with a zircon U–Pb age of 314–301 Ma (Han et al., 2006). The dikes in Qiakuerte on the northeast side of the Junggar Basin also intrude into the granite pluton (Fig. 7B). Fig. 7C shows the gabbro dikes emplaced into Late Carboniferous strata near Fuyun County on the southern margin of Altay. N330W striking dikes are found in the Early Carboniferous strata in Toksun (Fig. 7D).

Fig. 9. εNd vs. initial 87Sr/86Sr diagram of basic rocks in Northern Xinjiang (reference fields taken from White, 2013). Data sources of the Variscan magmatic rocks (see Table A.1): Kwon et al., 1989; Zhao et al., 1993; Zhou and Wang, 1993; He et al., 1994a; Ni et al., 1994; Wang et al., 1994; Chen et al., 1995; Zhou et al., 1995, 1996a, 1996b; Han et al., 1997; Li et al., 1998; Wang et al., 1998; Gu et al., 2001; Chen et al., 2000; Chen and Jahn, 2004; Zhou et al., 2004; Chen and Arakawa, 2005; Liu et al., 2005; Wang et al., 2005; Chai, 2006; Su et al., 2006a; Tong, 2006; Zhang et al., 2006; Zhu et al., 2006; Wang et al., 2006b, 2007b; Tang et al., 2007; Yuan et al., 2007; Geng et al., 2009; Jiang et al., 2009; Wang et al., 2009; Yuan et al., 2010; Zhang et al., 2011a; Deng et al., 2011; Shen et al., 2011; Xiong et al., 2011.

Fig. 10. 208Pb/204Pb & 207Pb/204Pb vs. 206Pb/204Pb isotope correlation diagrams of basic rocks in Northern Xinjiang (reference fields taken from White, 2013). Data sources of the Variscan magmatic rocks (see Table A.2): Kwon et al., 1989; Jin and Zhang, 1993; Zhao et al., 1993; Bi et al., 1994; Lu and Liu, 1994; Gao et al., 1995; Liu and Yuan, 1996; Zhao et al., 1996; Zhou et al., 1997; Li et al., 1998; Chai, 2006; Tong et al., 2006; Wang et al., 2006b, 2007b; Jiang et al., 2009; Xiong et al., 2011.


(Fig. 8A–E). The phenocrysts are mainly plagioclases (Fig. 8E), and the matrix is composed of fine hornblende, biotite, and plagioclase. The minerals and textures of the basic rocks in the ophiolites are similar to those in the above-mentioned basic rocks intruding the Paleozoic strata and granite plutons (Fig. 8F–H). Indeed, it is difficult to distinguish between the two types of basic rocks except by their field occurrences. 5. Whole-rock Sr, Nd, and Pb isotope geochemistry 5.1. Analytical method Strontium, neodymium, and lead isotope analyses of the basic rocks of the ophiolites were conducted using a multi-collector VG Axiom mass spectrometer. Approximately 100–150 mg of whole-rock powder (200 mesh) was decomposed for Sm–Nd isotopic analyses using a mixture of HF and HClO4 in sealed Teflon beakers set on a hot plate (100–120 °C) for one week. The sample solution was evaporated to dryness to eliminate SiF4 and excessive HF and HClO4. Four milliliters of purified 6 N HCl was subsequently added into the beakers and the solution was dried. Finally, the sample was dissolved using 1 ml of purified 2.5 N HCl and prepared for chemical separation and purification by being left undigested overnight. The Rb, Sr, and light REE were separated using a cationexchange column (packed with Bio-Rad AG50Wx12 resin, 200–400 mesh). The Sm and Nd were further purified using a second cationexchange column (P507), and conditioned and eluted with dilute HCl. The sample dissolution for lead isotope analysis was carried out using acid digestion (HNO3+ HF) in a sealed Teflon beaker on a hot plate for two days, and the sample solution was evaporated to dryness after adding purified HClO4. Next, the sample was dissolved using 1 ml of purified 0.5 N HBr and prepared for chemical separation and purification by being left undigested overnight. The separation of Pb was performed using a cation-exchange column (AG1x8 resin, 200–400 mesh) with 6 N HCl and 0.5 N HBr. The 87Sr/86Sr ratios were normalized to 86Sr/88Sr = 0.1194, and the 143 Nd/144Nd ratios were normalized to 146Nd/144Nd = 0.7219. For the lead isotope analysis, thallium was added to the solution to correct for mass bias, and the correct ratio of 205Tl/203Tl was 2.3871. Repeated analyses of the NBS-987 standard gave 87Sr/86Sr ratios of 0.710259 ± 15

191

(2σ), comparable to those obtained by high-quality TIMS techniques (e.g., Thirlwall, 1991, 87Sr/86Sr = 0.710241 ± 12) and were in good agreement with the commonly accepted high-precision value of 0.710263 ± 16 (Stein et al., 1997). The average value obtained for the 143 Nd/144Nd ratios of the JNdi-1 standard was 0.512117 ± 8 (2σ), in good agreement with the certified value of 0.512115 ± 7 (Tanaka et al., 2000). During the course of this study, the NBS-981 Pb standard yielded 206Pb/204Pb ratios of 16.9374 ± 13 (2σ), 207Pb/204Pb ratios of 15.4926 ± 14 (2σ) and 208Pb/204Pb ratios of 36.7002 ± 15 (2σ). Our ratios agree with the reference values within the error (Todt et al., 1996, 206Pb/204Pb = 16.9356 ± 7, 207Pb/204Pb = 15.4891 ± 9, 208 Pb/204Pb = 36.7006 ± 34). We also measured the BCR-2 USGS reference material, which yielded 87Sr/86Sr = 0.705132, 143Nd/144Nd = 0.512621, and 207Pb/206Pb = 0.8328 values that are in good agreement with the reference value of 0.704958 within the test errors (Raczek et al., 2003), 0.512633 (Raczek et al., 2003), 0.8333 (Jochum et al., 2005). 5.2. Results Eleven basic rock samples of ophiolites and sixteen basic dike samples were selected for the Sr, Nd, and Pb isotope analyses, and the results are shown in Table 2. According to the ophiolite formation time reported in previous studies (Fig. 1B) and the emplacement age of the basic dikes (Xu et al., 2008; Zhou et al., 2008a) in Northern Xinjiang, the εNd(t) values are in the range of 2.63–8.10 (Table 2). The positive εNd(t) values show that these basic rocks are derived from depleted mantle. In the initial εNd vs. 87Sr/86Sr diagram, nearly all 27 samples are concentrated in the MORB and OIB areas (Fig. 9). Additionally, in the 208Pb/204Pb and 207Pb/204Pb vs. 206Pb/204Pb correlation diagrams (Fig. 10), the majority of the samples fall in the area of MORB and OIB with little crustal composition effects. 6. Laser 40Ar/39Ar dating 6.1. Analytical method The 40Ar/39Ar whole-rock laser fusion method was applied to examine the thermal history characteristics of the two types of basic rocks.

Table 3 Whole-rock laser 40Ar/39Ar dating results of basic rocks associated with the ophiolites in Northern Xinjiang. Sample

SJQ-3 SJQ-7 QIA-2-1 QIA-2-9 QIA-3-3 XZ-3 XZ-5 BJT-1-3 MYL-1 MYL-8 AL-C AL-E AL-8 0820SE-1 0820SE-4 TBL-3-1 TBL-3-3 TBL-4 TBL-7 BKQ-6 BKQ-7 BJT-6 BJT-9

Location Lat (°N)

Long (°E)

44°52′32″ 44°56′57″ 46°31′22″ 46°31′05″ 46°29′11″ 45°46′11″ 45°46′11″ 45°43′43″ 45°37′46″ 45°34′09″ 45°25′29″ 45°25′29″ 45°09′11″ 45°56′57″ 45°56′57″ 45°06′39″ 45°06′39″ 45°06′32″ 45°07′49″ 45°59′39″ 45°59′42″ 45°43′40″ 45°43′41″

90°30′07″ 90°27′58″ 89°03′03″ 89°03′24″ 89°07′02″ 84°′26′32″ 84°26′32″ 85°07′13″ 83°16′41″ 83°13′22″ 82°44′07″ 90°44′07″ 83°10′27″ 90°11′45″ 90°11′45″ 83°21′32″ 83°21′32″ 83°23′10″ 83°24′08″ 85°18′11″ 85°17′37″ 85°07′08″ 85°07′07″

Lithology

Apparent age (Ma)

MSWD

Isochron age (Ma)

(40Ar/36Ar)0

MSWD

Basalt Basalt Basalt Basalt Gabbro Basalt Basalt Gabbro Basalt Gabbro Basalt Basalt Diabase Diabase Diabase Basalt Basalt Gabbro Gabbro Gabbro Gabbro Basalt Basalt

294 ± 3 278 ± 3 295 ± 3 306 ± 3 318 ± 4 275 ± 3 272 ± 3 358 ± 5 251 ± 3 336 ± 3 319 ± 3 302 ± 3 354 ± 3 365 ± 2 603 ± 4 279 ± 4 259 ± 5 220 ± 4 312 ± 4 291 ± 2 288 ± 3 306 ± 3 301 ± 4

11.28 15.37 62.58 24.52 74.59 56.12 6.44 1.45 88.89 186.56 43.26 809.12 36.8 31.11 254.71 23.61 30.09 146.05 13.23 7.41 4.96 3.46 3.17

296 ± 9 284 ± 8 301 ± 7 302 ± 7 328 ± 9 287 ± 12 270 ± 7 348 ± 10 240 ± 6 325 ± 7 323 ± 7 252 ± 7 389 ± 11 361 ± 9 300 ± 40 323 ± 16 258 ± 10 226 ± 9 230 ± 17 293 ± 5 292 ± 7 303 ± 8 303 ± 8

293 ± 5 288 ± 7 252 ± 12 303 ± 3 140 ± 60 287 ± 5 302 ± 6 279.7 ± 0.9 380 ± 20 370 ± 20 240 ± 40 530 ± 20 210 ± 20 320 ± 40 560 ± 30 254 ± 11 303 ± 9 250 ± 40 311 ± 3 290 ± 4 292.5 ± 1.5 299 ± 4 293 ± 3

8.7 9.5 9.3 10 7.7 32 3.6 0.73 6.3 15 4.9 14 9.2 8.8 26 14 7.1 22 4.8 1.5 2.3 1.5 0.38

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First, samples for laser 40Ar/39Ar dating were petrographically examined. Fresh and suitable samples were initially crushed into small pieces, and samples without weathering edges, inclusions, or fracture fillings were selected for further crushing. The selected pieces were repeatedly crushed using a steel crusher. Next, the broken samples were sieved to a particle size of 140–250 μm. Subsequently, the rock samples, along with Bern4-Ms (18.62 ± 0.06 Ma, Baksi et al., 1996) and ZBH-25 Biotite (132.9 ± 1.3 Ma, Sang et al., 2006) used for neutron monitoring, and K2SO4, CaF2, and KCl used for K, Ca, and Cl isotopic revising, were packaged with aluminum foil and sealed in a vacuum quartz bottle and subsequently irradiated in the H8 canal of the 49-2 nuclear reactor at the Chinese Institute of Atomic Energy (CIAE), Beijing. Next, the samples

were analyzed using the automatic high-precision and high-resolution laser microprobe 40Ar/39Ar dating system. The samples and monitors were degassed and fused using a New Wave CO2 laser (ESI, Portland, OR, USA) operated in continuous mode. The argon isotopes were measured using a VG5400 mass spectrometer (MicroMass Ltd., Manchester, UK). The 40Ar/39Ar ages were calculated based on the argon isotopic ratios measured after corrections for mass discrimination, interfering nuclear reactions, procedural blanks, and atmospheric Ar contamination. We used the values suggested by Steiger and Jäger (1977) for the decay constant (λ). The mass discrimination factor (D) was 1.0045 ± 0.0013, and other correction parameters and values

Fig. 11. Whole-rock laser 40Ar/39Ar age-probability spectrum for basic rocks of ophiolites (A) and for basic dikes that intrude Paleozoic strata and granite plutons (B) in Northern Xinjiang.


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information implied in Fig. 11B shows three peaks of magmatism with ages of 260–250 Ma, 220 Ma and 200–190 Ma, suggesting impulsive intrusions of these basic dikes. During the late Paleozoic, an extensive pan-Asian thermal event of magnitude comparable to the pan-African event produced granite intrusions that account for 5% of the surface exposures in central Asia (Coleman, 1989). These granites and the accompanied intermediate– basic dikes are accordingly widespread in Northern Xinjiang, with ages mainly from the Middle–Late Carboniferous to the Permian (Kwon et al., 1989; Han et al., 1997, 1998, 1999; Hu et al., 2000; Chen and Jahn, 2004; Li et al., 2004; Chen and Arakawa, 2005; Han et al., 2006; Geng et al., 2009; Chen et al., 2010; Tong et al., 2010; Yin et al., 2010). The intrusion of certain basic dikes continued up to the Jurassic (Xu et al., 2008; Zhou et al., 2008a). The age peak of the geothermal event presented in Fig. 11 coincides with the timing of this igneous activity in Northern Xinjiang, which suggests that the early Paleozoic ophiolites may have been reworked by the pan-Asian thermal event (thus causing the 40Ar/39Ar system to be replaced) whereas the basic dikes are the products of this geothermal event. Therefore, the laser 40Ar/39Ar dating results provide the indirect evidence that the basic rocks of the ophiolites were formed in the early Paleozoic and the basic dikes were emplaced in the late Paleozoic. From the perspective of thermochronology, the basic rocks in Northern Xinjiang can be divided into two types although they have similar petrologic characteristics. 7. Discussion 7.1. Sr, Nd, and Pb isotopic data for Variscan magmatic rocks in Northern Xinjiang—a summary

Fig. 12. Isotope diagrams for Variscan magmatic rocks from Northern Xinjiang. (A) εNd(t) vs. intrusive ages. The early–middle Proterozoic and Archean crusts were taken from Jahn et al., 2000b; and (B) εNd(t) vs. TDM. The fields of the Hercynian granites and Himalaya leucogranites were taken from Jahn et al., 2000b. Data sources are the same as in Fig. 9.

were [36Ar/37Ar]Ca = 2.775 × 10−4 ± 2.53 × 10−5, [39Ar/37Ar]Ca = 6.633 × 10−4 ± 3.535 × 10−4, and [40Ar/39Ar]K = 3.9448 × 10−3 ± 1.5421 × 10−3. The procedures for isotope analyses and age calculations were controlled by MassSpec software (Deino, 2010), and the testing procedures can be found in Zhou et al. (2008a). 6.2. Results Twenty three basic rock samples of ophiolites were selected for laser Ar/39Ar dating analysis, and the results are shown in Table 3. Regardless of argon excess or loss, the laser 40Ar/39Ar dating results present ages of 389–226 Ma (Table 3), showing a large difference compared with the previously reported ophiolite ages (Fig. 1B), a result which was totally unexpected. The statistical age-probability spectrum for all the analyses (23 samples) (Fig. 11A) shows that the basic rocks of the ophiolites were affected by a long-term geothermal event which initiated at ca. 400 Ma and lasted for 210 Ma, with the peak of this geothermal event occurring 310–290 Ma. Nine basic dike samples were selected for laser 40Ar/39Ar dating analysis with the ages ranging from 332 Ma to 174 Ma (Zhou et al., 2008a), values that are comparable with the K–Ar ages of basic dikes in Northern Xinjiang (Li et al., 2004; Xu et al., 2008). The statistical age-probability spectrum for these nine samples (Fig. 11B) presents notably different thermal history information than that of the basic ophiolite rock samples (Fig. 11A). These basic dikes intrude into the Paleozoic strata and granite plutons, and clear baking margins and chilled borders are recognized in their contact zones. Thus, the thermal history 40

A large amount of Middle–Late Carboniferous to Permian magmatic rocks is distributed in Northern Xinjiang. These rocks were previously envisaged as a product of post-collisional plutonism (Kwon et al., 1989; Han et al., 1998, 1999; Hu et al., 2000; Jahn et al., 2000b, 2000c; Chen and Jahn, 2004; Chen and Arakawa, 2005; Han et al., 2006, 2010a, 2010b; Chen et al., 2010) or mid-ocean subduction-related magmatism (Liu et al., 2007, 2009; Geng et al., 2009; Yin et al., 2010). These Variscan magmatic rocks, mostly with positive εNd(t) values (Han et al., 1998, 1999; Jahn et al., 2000b, 2000c), not only provide critical information on the Phanerozoic crustal growth but also place important constraints on the tectonic evolution of Northern Xinjiang. The Sr–Nd isotope data, including the derivative parameters for these Variscan magmatic rocks, are summarized in Fig. 9, Fig. 12, and Table A.1. For the depleted-mantle-based model age (TDM), we assume a linear Nd isotope evolution for the depleted mantle from εNd = 0 at 4.56 Ga to +10 at present. The two-stage model (DePaolo et al., 1991) is adopted for TDM calculation with respect to the samples with fSm/NdN −0.2 and the single-stage is used for the samples with an fSm/Nd value limited to the range −0.2 to −0.6. Fig. 12A shows the εNd(t) values as a function of intrusive age. The majority of the analyzed samples show positive εNd(t) values, indicating their relatively juvenile character and their depleted-mantle source. Fig. 12B shows that the Variscan magmatic rocks of Northern Xinjiang have very young model ages concentrated in the range of 400–1000 Ma. This is clearly distinguished from most of the European Hercynian granites, and even more from the leucogranites of the Himalayas (Fig. 12B). In addition, in the εNd(t) vs. initial 87Sr/86Sr ratio diagram, most of these rocks plot in the MORB and OIB areas and its vicinity (Fig. 9). The documented Pb isotope data are presented in Fig. 10 and Table A.2. In the Pb isotope correlation diagrams (Fig. 10) most of the samples cluster in or near the MORB and OIB areas with little crustal composition effect. Despite their high differentiation and often strong hydrothermal alteration (Jahn et al., 2000b), the contamination of the mantle-derived

194



195

Fig. 14. Tectonic map of the Tethyan belt showing distribution of remnant ocean basins and ophiolites (modified from Dilek et al., 2007). SCB — South Caspian Basin. The distribution of remnant ocean basins are from Brunet et al. (2003) (for SCB); Yilmaz et al. (1997) and Gürer and Aldanmaz (2002) (for Taurides–Anatolia region); Platt (2007) and Li et al. (2009) (for Mediterranean Sea and Black Sea).

magmas by crustal materials (Han et al., 1998), and the controversy regarding their petrogenesis (e.g. Kwon et al., 1989; Han et al., 1998, 1999; Hu et al., 2000; Chen and Arakawa, 2005; Han et al., 2006; Geng et al., 2009; Yin et al., 2010), all the Variscan magmatic rocks in Northern Xinjiang have similar isotopic components. This suggests that most of the granitoids and mafic–ultramafic intrusions originated from the same source with MORB and OIB affinity.

also show that these basic rocks have low initial 87Sr/86Sr and high initial 143Nd/144Nd values, indicative of clear oceanic affinity, as shown by Variscan granitoids (Fig. 9; Kwon et al., 1989; Wang et al., 1993; Wu et al., 1997). Overall, the basic rocks in this study and other Variscan magmatic rocks were derived from the same source. Their isotope geochemical characteristics suggest that since the Paleozoic, a large geochemical province with MORB and OIB affinity has existed in Northern Xinjiang.

7.2. Characteristics of the source area The Sr, Nd, and Pb isotope analyses of 11 basic rocks associated with ophiolites in Northern Xinjiang show that they have similar isotope ratios (Table 2), which together with their wide distribution (Fig. 1B) suggests that a relatively uniform and integrated source has existed in the region since the Paleozoic. The Sr, Nd, and Pb isotope ratios (Table 2) of the intermediate–basic dikes are similar to those of the ophiolites with some small changes, implying that these basic rocks were derived from the same source area. The positive εNd(t) (2.63–8.10, Table 2) values indicate that their magma source may be depleted mantle. Furthermore, the young model ages and positive εNd(t) values of the Variscan magmatic rocks (Table A.1; Fig. 12) as well as the Cretaceous–Paleogene magmatic rocks (Ji et al., 2006) in Northern Xinjiang indicate a Paleozoic juvenile depleted mantle. The initial value of the εNd vs. 87Sr/86Sr diagram (Fig. 9) and the 208 Pb/204Pb & 207Pb/204Pb vs. 206Pb/204Pb diagrams (Fig. 10) of all the basic rocks reveals that the relatively uniform and integrated source has a clear affinity with MORB and OIB. The Sr–Nd isotopic analyses

7.3. The basement nature of the Junggar Basin based on geophysical data The basement structure and properties of the Junggar Basin has been a subject of ongoing debate (Jiang, 1984; Yang et al., 1986; Coleman, 1989; Geology Department, Chinese Academy of Sciences and Xinjiang Petroleum Administration Bureau, 1989; Carroll et al., 1990; Xiao et al., 1992; Zhao, 1992b; Peng, 1994; Yuan et al., 1994; Zhou, 1994; Zuo et al., 1999; Zeng et al., 2002; Hu and Wei, 2003; Zhang et al., 2004; Zhao et al., 2008b). Some researchers consider that the basement of Junggar Basin is a Precambrian continental block (Yang et al., 1986; Geology Department, Chinese Academy of Sciences and Xinjiang Petroleum Administration Bureau, 1989; Zhao, 1992b; Peng, 1994; Yuan et al., 1994; Zuo et al., 1999; Zeng et al., 2002; Hu and Wei, 2003; Zhang et al., 2004) while others believe that it is the oceanic crust (Jiang, 1984; Coleman, 1989; Feng et al., 1989; Kwon et al., 1989; Carroll et al., 1990; Xiao et al., 1992; Zhou, 1994; Hu et al., 2000; Chen and Arakawa, 2005).

Fig. 13. 2D P-wave velocity structure of the crust and upper mantle derived from transects in the Junggar Basin. (A) Xayar–Burjing profile (modified from Zhang et al., 2011b) runs from the northern margin of the Tarim Basin through the Junggar Basin and ends at the southern slopes of the Altay range (Fig. 15), with a length of 995 km. (B) Emin–Qitai profile (modified from Xue et al., 2006; Zhao, 2012). This NW-trending seismic profile begins near Emin County, west of the Junggar Basin, runs through the central part of the basin, and terminates near Qitai County, north of the Bogda Shan mountain range (Fig. 15), with a total length of 620 km. (C) Karamay–Karmst profile (modified from Zhao et al., 2008b). Located at the northern part of the Junggar Basin, running from Karamay, west of the Junggar Basin to Karmst, east of the Junggar Basin in a nearly E–W direction (Fig. 15). C1 — interface of the upper crust and middle crust; C2 — interface of the middle crust and lower crust. In the upper part of the figures, the heavy dotted line is the Bouguer anomaly in mGal, and the heavy dash-dotted line is the aeromagnetic anomaly in nT. The Bouguer and aeromagnetic anomalies are from Zhao, 2012.

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Fig. 15. Topographic map of Northern Xinjiang showing the extent of the remnant ocean (shaded area). The heavy white lines show the location of the three geophysical transects (Fig. 13) in the Junggar Basin. NTT — North Tarim Thrust, AISNQF — Atbashy-Inylchek-South Nalati-Qawablak Fault, NL — Nikolaev Line, NTSF — North Tian Shan Fault, MJF — Major Junggar Fault, DF — Darbut Fault, ETST — East Tian Shan Fault, KF — Kalamaili Fault, IF — Irtysh Fault, and FF — Fuyun Fault. The “Yili Block” represents a large continental block, which mainly occurs in Kazakhstan where it is known as the “Ili” block (Sengör, and Natal'in, 1996). It is considered a microcontinent with a Precambrian basement separating the Chinese Tian Shan into northern and southern branches (Allen et al., 1992; Chen et al., 1999; Wang et al., 2006a, 2007a).

A series of wide-angle reflection/refraction seismic profiles and natural converted-wave seismic profiles were carried out in the Junggar Basin and P-wave velocity models were obtained (Zhao et al., 2003; Li et al., 2006b; Xue et al., 2006; Zhao et al., 2008b, 2010a, 2010b; Zhang et al., 2011b, 2013; Teng et al., in press). These seismic profiles, incorporated with gravity and magnetic data, describe the fine structures of the crust and uppermost mantle in the study area, which enables us to study the basement and framework of the Junggar Basin. The seismic data discussed in this study include two wide-angle reflection/refraction profiles (Fig. 13A and B) and one profile of natural converted seismic waves (Fig. 13C). In this study, we do not reinterpret these deep seismic datasets, but attempt to study the basement structure and properties of the Junggar Basin with a combination of gravity and aeromagnetic data. The 2D P-wave velocity structure of the crust and upper mantle of three geophysical transects in the Junggar Basin shows that the average crustal velocity of the Junggar Basin is higher (6.5 km/s) than that of the Tarim Basin (6.3 km/s) (Zhao et al., 2003, 2008a, 2010a; Zhao, 2005). High-velocity layers are present in the lower part of the upper crust (Fig. 13A) and upper part of the middle crust (Fig. 13B and C) at depths of 18–22 km. The velocity reaches 6.5–6.7 km/s, which indicates that the basement is not Precambrian crystallized basement (Zhao et al., 2008b). The velocity increases from 6.5–6.7 km/s at the top of the high-velocity layers to 7.2–7.4 km/s at the bottom of the lower crust (Fig. 13). The velocity of oceanic crust is generally 6.5–6.9 km/s (Salisbury and Christensen, 1978; Teng, 2003), and occasionally up to 7.5 km/s (Teng, 2003). Therefore, the velocity range of the middle and lower crusts in Junggar is similar to that of oceanic crust.

The Bouguer anomaly shows relatively high gravity values for the Junggar Basin (Fig. 13). Moreover, the 2D density structure of the crust along these three profiles indicates that the average density of the Junggar Basin is significantly high (Zhao et al., 2003, 2008b, 2010a). Similarly, the aeromagnetic anomaly (Fig. 13) and average magnetic intensity (Zhao et al., 2003, 2008b, 2010a) are relatively high for the Junggar Basin. In addition, the distribution of high magnetization is consistent with that of high density within the Junggar Basin (Zhao, 2012). Existing geophysical data (Fig. 13; Jiang, 1984; Xiao et al., 1992; Zhou, 1994; Zhao et al., 2003; Xue et al., 2006; Zhao et al., 2008a, 2008b, 2010a; Zhang et al., 2011b; Zhao, 2012) show that the Junggar Basin is characterized by high velocity, high density, and high magnetization. Jiang (1984) and Zhou (1994) pointed out the presence of preVariscan highly magnetic layers in the Junggar Basin. Jiang (1984) attributed this high aeromagnetic anomaly to oceanic basalts at a depth of 20 km and considered the crust of the Junggar Basin to be oceanic. Zhao et al. (2008b) inferred that several sections of the basement along the Karamay–Karmst profile (Fig. 13C) are basic to ultra-basic materials. The inference from the geophysical data, therefore, is that the basement of the Junggar Basin is oceanic crust. 7.4. Previously proposed tectonic models and the remnant ocean model Previous researchers have proposed four models for the tectonic evolution of Northern Xinjiang (Sengör et al., 1993; Mossakovsky et al., 1994; Sengör and Natal'in, 1996; Buslov et al., 2001; Badarch et al., 2002; Windley et al., 2002; Wang et al., 2003; Liu et al., 2007, 2009; Xiao et al., 2008, 2009; Geng et al., 2009; Zhang, 2009; Yin et al.,


2010). Sengör et al. (1993) and Sengör and Natal'in (1996) envisaged a continuous and oceanward migration of a single subduction zone and strike–slip stacking that created the entire Altaid collage. However, this model for the development of the CAOB faces several difficulties in light of recent geologic observations (Badarch et al., 2002; Windley et al., 2002; Bykadorov et al., 2003; Windley et al., 2007; Kröner et al., 2010). Sengör et al. (1993) and Sengör and Natal'in (1996) suggested that the migration of magmatic fronts progressed southward. However, there is no simple north-to-south progression of volcanism, and even south-to-north accretion has been reported in some arc terranes in southern Mongolia and the Altay mountains (Perelló et al., 2001; Badarch et al., 2002; Windley et al., 2002). The margins of the East European craton and southwest Mongolia were passive during the Paleozoic with evidence of carbonate platform development (Bykadorov et al., 2003; Windley et al., 2007; Kröner et al., 2010), which is not compatible with active subduction on the southeastern margin of the East European craton and southern Mongolia as envisaged by Sengör et al. (1993). The various terrane collision models (Didenko et al., 1994; Mossakovsky et al., 1994; Buchan et al., 2001; Buslov et al., 2001; Badarch et al., 2002; Windley et al., 2002) can explain the distribution of many different ophiolites, but fail to explain the fact that the syncollisional granites in Northern Xinjiang are rare (Xiao et al., 2006a). The accretionary wedge model (Xiao et al., 2008, 2009; Zhang, 2009) is able to better evaluate the contributions of many different ophiolite slices, but ignores the deformation sequence. The model does not distinguish between the Paleozoic deformation and the Cenozoic deformation. In fact, many faults in the Paleozoic imbricate fold-and-thrust belt are strike–slip faults developed in the Cenozoic (Xu et al., 2009;

197

Zhao et al., 2012). In addition, Xiao et al. (2008, 2009) placed the termination of the Paleo-Asian Ocean during the end of the Permian and the Triassic, but the Northern Xinjiang was in a post-collisional setting during the Permian, and there is no evidence of Late-Permian to Middle-Triassic accretion–collision processes (Han et al., 2010a, 2011). The mid-ocean ridge subduction model (Liu et al., 2007, 2009; Geng et al., 2009; Yin et al., 2010) can account for the geochemical characteristics of the granitoids, high-Mg diorites and mafic rocks but is not compatible with their widespread occurrences. Thus, no single mechanism proposed by these models can adequately explain the tectonics of Northern Xinjiang. Ophiolites are generally accepted as a product of plate collision. However, the distribution of the ophiolites in Northern Xinjiang bears no relationship to the final suture of the closure of the Paleo-Asia Ocean (Coleman, 1989; Feng et al., 1989). Sengör et al. (1993) pointed that the ophiolites in the Altaids occur in a haphazard pattern and carry no significance as structural markers to delineate former paleotectonic entities as they do in Himalayan- or Alpine-type collisional orogens. The Northern Xinjiang ophiolites spread over a large area (Fig. 1B), with a formation age in the Cambrian–Silurian. The large spatial distribution and large age span are similar to those of the ophiolites in the eastern Mediterranean (Fig. 14), which indicates that the tectonic evolution of Northern Xinjiang is similar to that of the Paleo-Tethys–Neo-Tethys in the eastern Mediterranean and in line with the evolution pattern of a remnant ocean (Fig. 14; Yilmaz et al., 1997; Li et al., 2009). The tectonic comparison between West Junggar and East Kazakhstan has proved the existence of a limited late Paleozoic remnant ocean (Fig. 1A), and consequently, the absence of large-scale subduction and

Fig. 16. Lithology section showing the sedimentary sequence of the remnant ocean and the sequence of oceanic crust in Northern Xinjiang (letters refer to photo numbers in Fig. 17).

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Fig. 17. Photographs of sediments in the remnant ocean and the underlying ophiolites. (A) Serpentinized peridotite; (B) Gabbro; (C) Conglomerate; (D) Pillow basalts; (E) Chert; (F) Purple red siltstone; (G) Black siltstone with oil-generating potential; (H) Flysch mudstone with hydrocarbon generation potential.

an arc–basin–trench system (He and Xu, 2003; He et al., 2004; Zhao and He, 2013). Chen and Guo (2010) proposed that the late Paleozoic tectonic evolution in West Junggar is dominated by a passive basin filling in a remnant Paleo-ocean. Additionally, the remnant ocean is well developed

along the convergent plate boundaries following the closure of the Paleo-Asian and Paleo- and Neo-Tethyan oceans (Figs. 1A and 11; Aplonov, 1995; Yilmaz et al., 1997; Gürer and Aldanmaz, 2002; Brunet et al., 2003; Li et al., 2009).


Therefore, the ophiolites in Northern Xinjiang may be regarded as representing a large area of oceanic crust rather than a suture boundary of plates. Evidence is mounting that the Junggar Basin may be entirely underlain by oceanic crust (Fig. 13; Jiang, 1984; Coleman, 1989; Feng et al., 1989; Kwon et al., 1989; Hu et al., 2000; Chen and Arakawa, 2005) and thus comprises a trapped or remnant ocean from the late Paleozoic (Carroll et al., 1990). Its overlying Paleozoic sediments are a continuation of the marine sediments exposed in the east and west of Junggar (Coleman, 1989). A complete lack of metamorphosed lower crustal rocks or zones of high strain-rate within the areas around the Junggar Basin indicates that its basement is similar to that of the adjacent areas (Hsü, 1988). It is possible to infer that a remnant ocean (Fig. 15) existed in the late Paleozoic and was consequently underlain by the early Paleozoic remnant oceanic crust (Fig. 16) in Northern Xinjiang of NW China. According to our detailed field investigation combined with regional geological data, we have reconstructed the basin sedimentary sequence of this late Paleozoic remnant ocean (Figs. 16 and 17). 7.5. Tectonic evolution and continental crust growth of Northern Xinjiang 7.5.1. Tectonic evolution of Northern Xinjiang The two types of basic rocks in the study area are geochemically identical (Table 4), implying that they may represent different stages of Paleozoic tectonic evolution in Northern Xinjiang. The evolution history can be separated into three main stages (Fig. 18). Assuming that the Junggar Basin is underlain by an oceanic crust, the widespread Cambrian–Ordovician ophiolites (Fig. 1B) and overlying Ordovician–Silurian volcanic-sedimentary rocks (Table 1) indicate that an early Paleozoic ocean existed in Northern Xinjiang. The ocean crust formation began in the Early–Middle Cambrian, and the ophiolites were generated in a seafloor-spreading environment accompanied by oceanic island basalts (Fig. 18A). As shown in Section 3, nearly all of the ophiolites in Northern Xinjiang were dated as early Paleozoic, suggesting that the Paleo-ocean in Northern Xinjiang, called the Junggar Ocean (Coleman, 1989; Dobretsov et al., 2003) reached its maximum area during the early Paleozoic. During the late Paleozoic, a remnant ocean was retained, bound by the Altay range in the north and the Yili Block in the south (Figs. 15 and 18A). The Yili block is considered a microcontinent with a Precambrian basement separating Tian Shan into northern and southern branches (Fig. 15; Allen et al., 1992; Chen et al., 1999; Wang et al., 2006a, 2007a). A relic from the previous tectonic stage, the remnant ocean was underlain by the early Paleozoic oceanic crust composed of serpentinite peridotites, gabbros, diabases, basalts, and cherts (Figs. 16 and 18A). This remnant ocean was filled with Devonian–Carboniferous sediments dominated by variegation siltstones, mudstones, black mudstones, and siltstones with oil-generating potential, flysch siltstones and mudstones, and coarse clastic rocks (Figs. 16 and 17). The most spectacular event that accompanied the filling of the late Paleozoic remnant ocean was the widespread and pervasive emplacement of Variscan granites, intermediate–basic dikes and their volcanic equivalents (Fig. 18B and C), known as the pan-Asian thermal event (Coleman, 1989). These Variscan magmatic rocks formed mainly during the Middle–Late Carboniferous to Permian (Kwon et al., 1989; Han et al., 1997, 1998, 1999; Hu et al., 2000; Chen and Jahn, 2004; Li et al., 2004; Chen and Arakawa, 2005; Han et al., 2006; Geng et al., 2009; Chen et al., 2010; Yin et al., 2010), making a thermal imprint on the early

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Paleozoic ophiolites (Fig. 11A). It is possible that these Variscan magmatic rocks originated from the partial melting of the remnant oceanic crust, but the heat source has not yet been determined (Coleman, 1989). The termination of marine sedimentation at the end of the Early Permian (Carroll et al., 1990) signals the end of the remnant ocean (Fig. 18C). The remnant ocean was not closed by large-scale subduction and extensive arc–basin–trench systems (He and Xu, 2003; He et al., 2004). The early Paleozoic oceanic crust still exists despite the termination of the remnant ocean. Subsequent Middle Permian through Cenozoic sedimentation in the Junggar Basin was entirely non-marine. The tectonic evolution of Northern Xinjiang has been in a state of intracontinental deformation since the Mesozoic (Xiao et al., 1992; Li and Xiao, 1999; Li et al., 2006a). In particular, during the late Cenozoic, the building and growth of the modern Tian Shan and Altay orogens were initiated by the far-field effects of the India–Eurasia collision (Avouac et al., 1993; Hendrix et al., 1994; Abdrakhmatov et al., 1996; Yin et al., 1998; Bullen et al., 2001, 2003; Charreau et al., 2005, 2006; Huang et al., 2006; De Grave et al., 2007; Zhang et al., 2007; Sun et al., 2009; Wang et al., 2010; Xu, 2011; Liu et al., 2012). Denudation of the orogens generated clastic sediments that were deposited in the Tarim and Junggar basins. The denudation and uplift have been ongoing, bringing the rocks to their present outcrop positions. The Northern Xinjiang at present is under a compressional regime (Ma, 1987; Wu, 1987; Xiao et al., 1992). 7.5.2. Continental crust growth of Northern Xinjiang The exact mechanism of the Phanerozoic continental crust growth in Northern Xinjiang remains debatable. Sengör et al. (1993) attributed the formation of the juvenile continental crust to subduction zone magmatism alone. Alternatively, an increasing number of authors consider the extensive post-collisional granitoids in Northern Xinjiang to be important to the Phanerozoic continental crust growth (Han et al., 1997, 1998, 1999; Hu et al., 2000; Jahn et al., 2000a, 2000b, 2000c; Chen and Jahn, 2004; Chen and Arakawa, 2005; Han et al., 2006; Chen et al., 2010). However, the post-collision concept has been challenged by recent studies that assigned the intrusions of granitoids and intermediate–basic dikes to ridge-subduction related magmatism (Geng et al., 2009; Yin et al., 2010). Whether or not the magmatism associated with Variscan granitoids is post-collisional is beyond the scope of this paper, but it is discussed in several related works (Han et al., 1997; Geng et al., 2009; Yin et al., 2010). Importantly, there is no doubt that the Variscan magmatic rocks are widespread in Northern Xinjiang. More importantly, these Variscan magmatic rocks were derived from the same source. It has been well established that the average composition of continental crust is granodioritic (Taylor and McLennan, 1985), whereas the crust of Northern Xinjiang was oceanic. How was the oceanic/basic crust converted into a felsic one? In Northern Xinjiang, very large volumes of granitoids and their volcanic equivalents were generated during the late Paleozoic. The isotopic data suggest that these granitoids were derived from a source with MORB and OIB affinity (Figs. 9 and 10). This source is the remnant oceanic crust discussed in this study. The parental magma of the granitoids originated from the partial melting of the remnant oceanic crust formed in the early Paleozoic (Fig. 18). The continental crust consists mainly of granitoid rocks, accompanied by smaller amounts of mafic and ultramafic rocks (Jahn et al., 2000a). We propose that the Variscan granitoids and basic dikes

Table 4 Comparison of the geological, geochemical, and geochronological features of the two types of basic rocks in Northern Xinjiang. Type

Occurrence

Formation age

Laser 40Ar/39Ar age

Age-probability

εNd

Source area

Component of ophiolite Dikes

Overlying peridotite Intruding Paleozoic strata and granite plutons

Early Paleozoic 332–174 Ma

389–226 Ma 332–174 Ma (Zhou et al., 2008a)

310–290 Ma age peak 260–250 Ma, 220 Ma and 200–190 Ma age peaks

4.94–7.87 2.63–8.10

MORB and OIB MORB and OIB

200


Fig. 18. Illustration showing the Paleozoic tectonic evolution of Northern Xinjiang. (A) Oceanic crust formation in the Cambrian–Silurian; (B) A remnant ocean was retained during the late Paleozoic and was filled with Devonian–Carboniferous sediments; (C) Widespread and pervasive emplacement of granites, intermediate–basic dikes and their volcanic equivalents between the Late Carboniferous and the Early Permian. The marine sedimentation ceased at the end of the Early Permian.

represent the formation of continental crust transferred from the basic crust in the Paleozoic. In other words, the Phanerozoic continental crust growth of Northern Xinjiang was completed by mass transfer from the early Paleozoic remnant oceanic crust; this approach may considerably change our views of continental growth. 8. Conclusions (1) The widespread basic rocks in Northern Xinjiang can be divided into two categories: (i) gabbros, diabases, and basalts in the early Paleozoic ophiolites; and (ii) late Paleozoic basic dikes that intrude into Paleozoic strata and granite plutons. (2) The basic rocks and other Variscan magmatic rocks in this study were derived from the same source. Their isotope geochemical

characteristics and widespread distribution suggest that since the Paleozoic, a large geochemical province with MORB and OIB affinity has existed in Northern Xinjiang; this region is related to a longlived remnant ocean and the underlying early Paleozoic oceanic crust. (3) Geophysical data confirm the existence of remnant oceanic crust in Northern Xinjiang at present. (4) The growth of the continental crust in Northern Xinjiang was completed by mass transition from the early Paleozoic remnant oceanic crust. (5) The remnant ocean model can better explain the tectonic evolution of Northern Xinjiang. Supplementary data to this article can be found online at http://dx. doi.org/10.1016/j.earscirev.2013.08.005.


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