Early Paleozoic arc magmatism in the Kalamaili orogenic belt, Northern Xinjiang, NW China: Implications for the tectonic evolution of the East Junggar terrane

Journal Pre-proofs Early Paleozoic arc magmatism in the Kalamaili orogenic belt, Northern Xinjiang, NW China: Implications for the tectonic evolution of the East Junggar terrane Qinqin Xu, Lei Zhao, Baogui Niu, Rongguo Zheng, Yaqi Yang, Jianhua Liu PII: DOI: Reference:

S1367-9120(19)30424-9 https://doi.org/10.1016/j.jseaes.2019.104072 JAES 104072

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Journal of Asian Earth Sciences

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30 April 2019 9 October 2019 10 October 2019

Please cite this article as: Xu, Q., Zhao, L., Niu, B., Zheng, R., Yang, Y., Liu, J., Early Paleozoic arc magmatism in the Kalamaili orogenic belt, Northern Xinjiang, NW China: Implications for the tectonic evolution of the East Junggar terrane, Journal of Asian Earth Sciences (2019), doi: https://doi.org/10.1016/j.jseaes.2019.104072

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Early Paleozoic arc magmatism in the Kalamaili orogenic belt, Northern Xinjiang, NW China: Implications for the tectonic evolution of the East Junggar terrane Qinqin Xu a, b, Lei Zhao a, *, Baogui Niu a, Rongguo Zheng a, Yaqi Yang a, Jianhua Liu a a

MNR Key Laboratory of Deep-Earth Dynamics, Institute of Geology, Chinese Academy of

Geological Sciences, Beijing 100037, China b

Shandong Provincial Key Laboratory of Depositional Mineralization and Sedimentary

Mineral, Shandong University of Science and Technology, Qingdao 266590, China * Corresponding author. E-mail address: [email protected]. (L. Zhao)

Abstract: The Kalamaili orogenic belt in northwestern China preserves a record of the magmatic and tectonic evolution of the Paleo-Asian Ocean and is thus an important region for reconstructing the Paleozoic evolution of the East Junggar terrane and the Central Asian Orogenic Belt. This study presents detailed field investigations, new SHRIMP U-Pb ages, whole-rock geochemistry, and in situ zircon Hf isotopic data for the early Paleozoic igneous rocks exposed in the Kalamaili orogenic belt. These igneous rocks are predominantly felsic plutons and mafic–intermediate volcanic rocks but include minor mafic–intermediate dikes. The igneous rocks have zircon U-Pb ages ranging from 463 Ma to 433 Ma and unconformably overlain by Late Silurian and Early Devonian sedimentary rocks. These early Paleozoic igneous rocks exhibit signatures similar to those of arc-related magmatic rocks such as enrichment in large-ion lithophile elements and depletion of high field strength elements. They have high positive εHf(t) values of +10.7 to +19.0, suggesting a juvenile source and considerable continental growth in the East Junggar terrane during the early Paleozoic. The geochemical signatures of the mafic–intermediate rocks suggest that they originated from the partial melting 1

of mantle wedge material that had been metasomatized by slab-derived fluids; the felsic plutons were derived from the partial melting of juvenile crust. This early Paleozoic arc magmatic event, together with other evidence, indicates that the Kalamaili Ocean formed in the early Paleozoic, was subducted during Middle Ordovician–Early Silurian, and closed before Middle Silurian. Stratigraphic correlations reveal that the East Junggar terrane might not have experienced full extension to form a broad ocean during the late Paleozoic.

Keywords: Arc magmatism; Unconformity; Early Paleozoic; East Junggar; Paleo-Asian Ocean

1. Introduction Arc magmatism related to subduction zones is one of the major mechanisms by which continents grow (Kusky and Polat, 1999; McCulloch and Gamble, 1991; Rudnick, 1995; Taylor and McLennan, 1995). Accretionary orogens, characterized by subduction-accretion complexes and arc magmatism (Maruyama, 1997; Sengör et al., 1993; Xiao et al., 2008), were primarily responsible for the growth of the continental crust (Cawood et al., 2009; Jahn, 2004; Kröner et al., 2007; Safonova et al., 2011; Santosh et al., 2009, 2010; Sengör and Natal'in, 1996; Xiao and Santosh, 2014). The Central Asian Orogenic Belt (Fig. 1a) (CAOB; Jahn, 2004; Windley et al., 2007), also known as the Altaids (Sengör and Natal'in, 1996; Sengör et al., 1993), is widely believed to have been formed by accretionary tectonics and host massive amounts of juvenile crust formed during the Phanerozoic (Jahn, 2004; Kröner et al., 2007; Safonova, 2017; Şengör et al., 1993; Windley et al., 2007; Xiao et al., 2009, 2010). The CAOB resulted from the long evolution of the Paleo-Asian Ocean from the late Proterozoic to the late Paleozoic (e.g., Dobretsov et al., 1995; Khain et al., 2002; Windley et al., 2007) or the Mesozoic (e.g., Xiao et al., 2008, 2015). As one of the largest accretionary orogenic collages in the world, the CAOB preserves valuable 2

detailed information on subduction-accretion and arc magmatism, and thus offers a window to explore the formation of accretionary orogens and continental growth (Cai et al., 2018a; Xiao and Santosh, 2014). The East Junggar terrane (Fig. 1b), located in Northern Xinjiang, northwestern China, is an important tectonic unit in the CAOB.

It is an ideal terrane for investigating the continental

growth of Central Asia and the evolution of the Paleo-Asian Ocean (e.g., Coleman, 1989; Han and Zhao, 2018; Han et al., 2006; Hu et al., 2000; Windley et al., 2007; Xiao et al., 2008, 2009). It is also a critical area for exploring the paleotectonic framework of Northern Xinjiang and Northern Xinjiang tectonic evolution models (Chen and Jahn, 2004; He et al., 1994; Li, 2004; Li et al., 2015; Liang et al., 2016; Liu et al., 2017; Xiao et al., 1992b, 2009; Xu et al., 2013a, 2013b, 2015; Zhang et al., 2013). However, the Paleozoic tectonic evolution of the East Junggar terrane remains controversial and, in particular, the question about when the Kalamaili Ocean was formed and closed remains hotly debated (e.g., Cai et al., 2018b; Fang et al., 2015; Han and Zhao, 2018; He et al., 2001; Huang et al., 2012; Li, 1991; Li et al., 1990, 2015; Liu et al., 2017; Shu and Wang, 2003; Su et al., 2012; Wang et al., 2009; Xiao et al., 2009; Zhang et al., 2013). The occurrence of Tuvaella fauna on both the southern and northern sides of the Kalamaili ophiolite belt (KOB) and the evidence of Early Devonian microfossils in cherts in the ophiolite led Li et al. (1990) and Li (1995) to propose that the Kalamaili Ocean opened in the Early Devonian, which evolved from the Late Silurian epicontinental sea on the Siberian plate. According to one interpretation, this ocean probably closed before the Early Carboniferous, a conclusion supported by: a) a depositional age of 345–343 Ma for the sedimentary successions that unconformably overlie the Kalamaili ophiolite (Huang et al., 2012; Zhang et al., 2013); b) the radiometric age of 348 Ma for the syn-collisional quartz diorites that intruded the ophiolite (Xu et al., 2015); and c) the 343 Ma date on the granitic ultramylonite that crosscuts the

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ophiolite belt (Wu et al., 2012). The Early Carboniferous closure time is also supported by 343– 342 Ma and younger dates on post-collisional volcanic rocks in the East Junggar terrane (Du et al., 2018; Zhang et al., 2015). However, a different interpretation suggests that the Early Carboniferous volcanism occurred in an island-arc setting (Li et al., 2013; Su et al., 2012), suggesting that the subduction of the Kalamaili Ocean lasted until Late Carboniferous. A similar proposal was put forward by Long et al. (2012) and Huang et al. (2018) on the basis of their analysis of the provenance of late Paleozoic sedimentary rocks. In their evolutionary model for Northern Xinjiang, Xiao et al. (2008, 2009) argued that the subduction of the Kalamaili Ocean lasted until the Permian. These investigators all maintain that

the Kalamaili Ocean was more

or less a late Paleozoic ocean. However, the existence of Mid–Late Ordovician conodonts in the Kalamaili ophiolite and the regional angular unconformity between the Ordovician and post-Ordovician strata in the Kalamaili orogenic belt led He and Li (2001) and He et al. (2001) to suggest that the Kalamaili Ocean was formed between the late Neoproterozoic and the Cambrian and that it was consumed and closed between the Late Ordovician and Early Silurian. Early Paleozoic magmatism, reported recently in the Kalamaili orogenic belt and adjacent areas (Du et al., 2010; Guo et al., 2009, 2013; Huang et al., 2016a, 2016b; Shi et al., 2015; Xu et al., 2013b), adds a new wrinkle to the debate about the evolution of the Kalamaili Ocean. In this paper, we present observations from field investigations, geochemical data, SHRIMP zircon U-Pb ages, and zircon Hf isotopic data for the early Paleozoic igneous rocks exposed in the Kalamaili orogenic belt (Figs. 1c and 2). We discuss the origin and geodynamic setting of these igneous rocks based on our new data, previously published works, and the regional geology. This analysis can help to better constrain the evolution of the Kalamaili Ocean and thus, clarify the early Paleozoic tectonic evolution of the East Junggar terrane.

4

2. Geological setting 2.1. Regional tectonics Northern Xinjiang lies in the southern part of the CAOB (Fig. 1a) and comprises three faultbounded tectonic domains. From north to south these belts are the Altay orogenic belt, the Junggar terrane, and the Tianshan orogenic belt (Fig. 1b). The Junggar terrane has been divided into the East and West Junggar terranes and the Junggar basin is transected by structural lines and suture zones (Chen and Jahn, 2004; Xiao et al., 2008) (Fig. 1b). The East Junggar terrane, northeast of the Junggar basin, is separated from the Altay orogenic belt to the north by the Erqis fault and from the Tianshan orogenic belt to the south by the Kalamaili fault. The Erqis fault encompasses a Devonian–Carboniferous ophiolite belt (Deng et al., 2015; Liu, 2011; Ni et al., 2013; Shen et al., 2017; Wang et al., 2011; Wu et al., 2006a; Zhang et al., 2003) that is regarded as the final suture of the Paleo-Asian Ocean between the Altay and Junggar terranes (Briggs et al., 2007; Xiao et al., 2015). Tectonically, the East Junggar terrane is composed of a series of arcs, accretionary complexes, and ophiolite belts generated by subduction–accretion during the Paleozoic due to the consumption of the Paleo-Asian Ocean (Feng et al., 1989; Xiao et al., 2008, 2009). The terrane consists, from the north to the south, of the Dulate-Baytag arc, the Armantai ophiolite belt, the Yemaquan arc, the KOB, and the Jiangjunmiao accretionary complex (Fig. 1b). To the north, the Dulate-Baytag arc consists mainly of Devonian–Carboniferous volcanic and sedimentary rocks (Li et al., 1990; Zhang et al., 2009). The Armantai ophiolite belt was emplaced against Devonian–Carboniferous arc-type volcanic-sedimentary rocks (Xiao et al., 2009) and has an age between 503 Ma and 420 Ma (Jian et al., 2005; Liu et al., 2016; Luo et al., 2017; Xiao et al., 2009; Zeng et al., 2015). The Yemaquan arc mainly comprises Ordovician–Carboniferous volcanic and volcaniclastic rocks (BGMRXUAR, 1993; Du et al., 2018; Li et al., 2013; Long et al., 2012) and Carboniferous–Permian granitoids with ages of 5

342–268 Ma (Chen and Jahn, 2004; Han et al., 2006; Hu et al., 2000; Nie et al., 2016; Su et al., 2006, 2008; Yang et al., 2010, 2011) (Fig. 1b, c). The Kalamaili ophiolite predominantly crops out along the Kalamaili fault. The ophiolite is largely dismembered and is now represented by a variety of tectonic blocks incorporated in a highly deformed volcanic-sedimentary matrix. Mid–Late Ordovician conodonts were found in the siliceous rocks of the KOB by early investigators

(Huang et al., 1990) and this age is consistent with early SHRIMP zircon U-Pb

dates for plagiogranite from this belt (497 Ma, Jian et al., 2005). The siliceous rocks also contain Late Devonian–Early Carboniferous radiolaria (Shu and Wang, 2003) and the age of these strata has been verified by recent radiometric dating of plagiogranites and mafic rocks from the ophiolite; the radiometric ages range from 417 to 330 Ma (Fang et al., 2015; Huang et al., 2012; Liu et al., 2017; Qin, 2012; Tang et al., 2007; Wang et al., 2009; Xu et al., 2015). The Jiangjunmiao accretionary complex is south of the Kalamaili ophiolite. It was formed by the progressive accretion of Devonian, Carboniferous and Permian island arc rocks with evolved chemical signatures (Xiao et al., 2004, 2009). 2.2. Geology of the study area The area of study is located in the southern part of the Yemaquan arc and near the KOB (Fig. 1b, c). No Precambrian strata have been clearly identified in the study area and the exposed strata are dominated by Paleozoic sedimentary, pyroclastic, and volcanic rocks (Fig. 2). This section of the manuscript describes the stratigraphy on the northern and southern sides of the KOB. On the northern side of the KOB, the oldest stratigraphic units are the Middle–Upper Ordovician volcanic rocks (BGMRXUAR, 1993; RGSGXGB, 1977), exposed mainly in the Zhifang area (Figs. 1c and 2a, b). These volcanic rocks are mainly basalt, basaltic trachyandesite, and trachyandesite with minor dacite, rhyolite, breccia, and tuff that underwent greenschistfacies to epidote-amphibolite facies metamorphism (He and Li, 2001; Xu et al., 2013b). In 6

addition, there are many granitic, rhyolitic and dioritic porphyries as well as diabase dikes embedded in the volcanic rocks. The Silurian strata are Late Silurian littoral–neritic facies sediments and pyroclastic rocks with corals and brachiopods. The lithologies present include conglomerate, sandstone, tuffaceous sandstone, limestone, and tuff; the sandstones are crossbedded (Supplementary Fig. S1a). The Late Silurian strata unconformably overlie the Middle–LLate Ordovician units (Figs. 2a, 3a, b and 4a) and the early Paleozoic monzogranite and monzonite (Fig. 2a, b), and have a conformable contact with the Early Devonian rocks (Fig. 2a, b and Supplementary Fig. S1b). The Lower Devonian sequence is a suite of clastic sedimentary and pyroclastic rocks with some limestone that unconformably overlies the Middle–Upper Ordovician (Figs. 2a, b and 3c, d) and the early Paleozoic monzogranite and monzonite (Figs. 2a, b and 3e, f). This unit consists mainly of lithic sandstones, calcareous sandstones, tuffaceous sandstones, and tuffs interbedded with limestones. The Middle Devonian (Supplementary Fig. S1c) has a conformable contact with the Lower Devonian and is mainly composed of lithic sandstones, tuffaceous sandstones, tuffaceous conglomerates, and tuffs interbedded with biohermal limestones. Planar horizontal and wavy bedding is developed in the tuff and sandstone, respectively. The Upper Devonian, conformably overlying the Middle Devonian, is a succession of continental pyroclastic rocks interbedded with terrigenous clastic sedimentary rocks (RGSGXGB, 1977). The lower part of the Upper Devonian unit consists mainly of tuffs interbedded with tuffaceous and lithic sandstones; the sedimentary structures are mainly horizontal bedding. Radiolarians were identified in the tuff from this sequence (XIGS, 2000). The upper part of the Upper Devonian consists mainly of tuffs interbedded with tuffaceous conglomerates (Supplementary Fig. S1d) and coarse sandstones. Large-scale cross-beds are developed in this unit (Supplementary Fig. S1e). The Lower–Middle Devonian units contain abundant brachiopods, corals, crinoids (Fig. 5a, b) and the Upper unit contains plant fossils (Fig. 5c). The two Lower Carboniferous units,

7

the Tournaisian and the Visean, are separated by an angular unconformity (Figs. 2a, c and 3g, h). These units are composed of clastic sedimentary and pyroclastic rocks, mainly siltstones, tuffaceous sandstones, conglomerates, silty mudstones, and thin-bedded tuffs (Supplementary Fig. S1f). The Upper Carboniferous unconformably overlies the Lower Carboniferous (Visean) (Fig. 2a) and consists of basaltic and andesitic porphyries, tuffaceous breccias, breccias, and tuffs. On the southern side of the KOB, the oldest stratigraphic unit contains the Mid-Silurian (Fig. 2c) terrigenous clastic and fine pyroclastic rocks such as lithic sandstones and tuffs. Som of the beds in this unit host Tuvaella fauna (Cai et al., 2015a; He et al., 2001; Li, 1995; Su, 1981; Zhang et al., 1983) (Fig. 5d). The Late Silurian–Early Devonian units conformably overlie the Middle Silurian and consist mainly of medium -bedded tuffs (Supplementary Fig. S2a), sandstones, siltstones, and calcareous sandstones interbedded with limestones. The radiolarians were identified in tuff (XIGS, 2003). The undivided Devonian sequences are littoral–neritic facies assemblages (BGMRXUAR, 1999) that consist mainly of lithic sandstones, tuffs (Supplementary Fig. S2b), and some conglomerates interbedded with sandstone lenses (Supplementary Fig. S2c). Oblique bedding can be seen in the sandstone lenses (RGSGXGB, 1966). Abundant brachiopods, corals, crinoids, and plant fossils are also present in the Late Silurian–Devonian units (Cai et al., 2015a). The Lower Carboniferous (Tournaisian) consists of a suit of continental clastic rocks and has a conformable contact with the Devonian sequences. The Lower Carboniferous unit is composed mainly of conglomerates and sandstones (Supplementary Fig. S2d) with some volcanic rocks and the unit contains Lepidodendralean fossils (XIGS, 2003). The Lower Carboniferous (Visean) unconformably overlies the Tournaisian unit (Fig. 2c) and the Visean consists of a suite of ~350 Ma continental volcanic rocks (Tan et al., 2009; Zhang et al., 2013). The plutonic rocks exposed in the study area can be divided into early and late Paleozoic

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intrusives. The former are mainly monzogranites and quartz monzonites, the latter mainly granitoids and gabbros (Fig. 2). A biotite monzogranite stock (Fig. 4b) and a quartz monzonite stock (Fig. 4c) crop out about 13 km northeast and 42 km southeast of Zhifang, respectively (Figs. 1b and 2a, b). They intrude into Middle–Upper Ordovician volcanic rocks (Figs. 2a, b and 4b) and are unconformably covered by Upper Silurian and Lower Devonian (Figs. 2a, b and 3e, f). At the biotite monzogranite sampling location, the Early Devonian granitic conglomerate (Fig. 4d), composed mainly of in situ rock fragments, directly and unconformably overlies the granite pluton (Figs. 2a and 3e, f). The gravels are poorly sorted with clast diameters ranging from 2 cm to 2 m; and the cement is Lower Devonian tuffaceous sandstone.

3. Petrography Eight samples from the early Paleozoic felsic intrusive rocks, eight samples from the Middle–Upper Ordovician volcanic rocks, four embedded diorite dikes, and one diabase dike were selected for petrographic analysis (Table 1). The biotite monzogranite (sample ZF-2–5, 7) is grayish in color, medium- to fine-grained, and massive with a granitic texture. It is composed of euhedral to subhedral plagioclase (35– 40%), subhedral to anhedral alkali feldspar (25–35%), anhedral interstitial quartz (30–35%), biotite (3–5%), hornblende (<3%) and accessory minerals (e.g., titanite, Ti–Fe oxide, apatite, and zircon, 2–3%) (Fig. 6a). The plagioclase is platy, polysynthetically twinned, and some grains are zoned. Plagioclase and alkali feldspar were partially altered to clay minerals. The quartz monzonite (ZF14-44–46) is light-colored, medium-grained, and massive with a hypidiomorphic-granular texture and is composed of euhedral to subhedral plagioclase (35– 40%), subhedral alkali feldspar (30–40%), euhedral hornblende (10–15%), anhedral quartz (5– 10%), biotite (3%–5%), clinopyroxene (3–5%), and accessory minerals (<3%) (Fig. 6b, c). The alkali feldspar mainly consists of perthite. The plagioclase is platy and commonly 9

polysynthetically twinned. The gabbroic diorite (ZF-8) is medium- to fine-grained and displays a granular texture. It is composed of euhedral to subhedral plagioclase (60–65%), euhedral to subhedral hornblende (25–30%), clinopyroxene (~5%), alkali feldspar (2–3%), quartz (2–3%), biotite (<3%), and accessory minerals (2–3%) (Fig. 6d). Plagioclase grains are lath-shaped with polysynthetic twinning. The hornblende has been chloritization. The monzodiorite (ZF14-11–13) is grayish-green in color, medium- to fine-grained, and massive with a granular texture. The rock mainly contains euhedral to subhedral plagioclase (55–60%), euhedral to subhedral hornblende (25–30%), alkali feldspar (5–10%), clinopyroxene (<5%), biotite (<5%), quartz (2–3%), and accessory minerals (2–3%) (Fig. 6e). Plagioclase was subject to siallitization, and hornblende suffered chloritization or epidotization. The diabase (ZF14-7) is black-green and shows a typical diabasic texture (Fig. 6f). It is composed of plagioclase (~60%), clinopyroxene (~30%), hornblende (~5%), biotite (<3%), and accessory minerals (<3%). The plagioclase grains are euhedral elongate laths and anhedral clinopyroxene grains are interstitial to plagioclase. The plagioclase has been sericitized and the clinopyroxenechloritized. The andesite (ZF-9, 11, 12) is dark gray, shows a porphyritic texture, and contains a few amygdules. The phenocrysts (10–15%) are mainly platy plagioclase (Fig. 6g) with some hornblende, whereas the pilotaxitic matrix consists of microcrystalline plagioclase (50–60%), hornblende (20%–25%), and quartz (3–5%). The plagioclase phenocrysts have been strongly sericitized. The basaltic andesite (ZF14-2–6) is dark gray in color and porphyritic with plagioclase, clinopyroxene, and minor olivine phenocrysts (up to ~25%, Fig. 6h) set in a matrix of microcrystalline plagioclase (45–55%), clinopyroxene (20%), and minor quartz (1–2%). Platy plagioclase is the most abundant phenocryst. The clinopyroxene phenocrysts are subhedral to

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anhedral and have been chloritized.

4. Analytical methods After removal of the weathered surfaces, hand specimens were crushed in a jaw crusher, sieved, ultrasonically cleaned several times in deionized water and then ground in an agate mill to the size of 200 mesh for whole-rock geochemirsty analyses. Zircons were separated from the crushed samples by conventional magnetic and heavy liquid methods and hand-picked under a binocular microscope. The sample processing and mineral separation were conducted at the Institution of Geological Survey and Mapping of Hebei Province, Langfang city, China. The zircon grains were mounted on epoxy resin and then polished down to expose the grain interiors for optical microscope observations and cathodoluminescence (CL) imaging. CL images were taken using a scanning electron microscope HITACHI S-3000N at the Beijing SHRIMP center, Chinese Academy of Geological Sciences in order to identify the internal textures and select potential targets for U-Pb dating and Hf isotope analysis. 4.1. Whole-rock major and trace element analyses A total of 21 samples were analyzed for whole-rock major and trace elements at the Analytical Laboratory, Beijing Institute of Uranium Geology, China. Major elements were analyzed using a Phillips PW 2404 X-ray fluorescence spectrometer on fused glass discs, which were made from whole-rock powder mixed with Li2B4O7. Loss on ignition (LOI) was calculated after heating the sample powder to 1000°C for 1 h. Ferrous iron was determined by the wet chemical titration method. The analytical uncertainties for major elements range from 1% to 3%. Trace elements (including REE) were determined by inductively coupled plasma-mass spectrometry (ICP-MS) with an ELEMENT XR system after sample dissolution using a mixture of HF and HNO3, following the procedures of Liang and Grégoire (2000). The analytical uncertainties for most trace elements were better than 5% RSD. 11

4.2. Zircon U-Pb analyses Five samples (ZF-2, ZF-3, ZF-7, ZF14-7, and ZF14-44) were selected for zircon U-Pb isotope analyses using the sensitive high resolution ion microprobe (SHRIMP II) at the Beijing SHRIMP center following standard analytical procedures similar to those described by Williams (1998). The current of O2− primary ion beam was 3–5 nA and spot size was about 30 μm in diameter. Five scans through the mass stations were made for each age determination. Inter-element fractionation in the ion emission of zircon was corrected using reference standard TEMORA 2. The average value obtained for the 206Pb/238U ages of TEMORA 2 standard was 416.6 ± 2.6 Ma (2σ), in good agreement with the certified value of 416.8 ± 0.3 Ma (Black et al., 2004). Corrections of U, Th and Pb concentrations were made by normalization to the zircon standard M257 (U = 840 ppm, 206Pb/238U age = 561.3 Ma, Nasdala et al., 2008). The common lead correction was applied using the measured 204Pb abundances. Data processing and assessment was carried out using the SQUID and ISOPLOT programs (Ludwig, 2001, 2003). Uncertainties for the isotopic ratios of individual analyses in Supplementary Table S1 and on the concordia diagrams are given at 1σ, whereas uncertainties for weighted mean ages are quoted at the 95% confidence level (2σ). 4.3. Zircon Hf isotope analyses In-situ zircon Hf isotopic analyses were conducted using a 193 nm Newwave UP193-FX ArF excimer laser ablation microprobe attached to a Neptune multi-collector inductively coupled plasma mass spectrometer (LA-MC-ICPMS) at the Tianjin Institute of Geology and Mineral Resources. Hf isotope analyses were made on the same zircon grains that were previously analyzed for U-Pb isotopes, with spot size of 35 μm in diameter, pulse rate of 10 Hz, and laser beam energy density of with 15 J/cm2. Details of the technique were as described by Geng et al. (2011). The isobaric interference corrections

176Lu

and

176Yb

on

176Hf

were

processed according to the corrections protocols proposed by Iizuka and Hirata (2005) and Wu 12

et al. (2006b). For instrumental mass bias correction, Yb isotope ratios were normalized to 172Yb/173Yb

= 1.35274 (Chu et al., 2002) and Hf isotope ratios to 179Hf/177Hf=0.7325 (Patchett

and Tatsumoto, 1980) using an exponential law. During analyses, the

176Hf/177Hf

ratio of

standard zircon (GJ-1) was 0.282029± 10 (2σ, n=30), in agreement with the recommended values within 2σ error (0.282015± 19, Elhlou et al, 2006; 0.282006± 24, Geng et al., 2011).

5. Analytical results 5.1. Whole-rock major and trace elements Major and trace element data for the samples analyzed are listed in Table 1. In the felsic intrusive rocks (including the quartz monzonite and biotite monzogranite), the SiO2 content ranges from 64.81 to 76.46 wt.% (normalized to 100% anhydrous, similarly hereafter), the K2O from 0.35 to 2.40 wt.%, and the Na2O

from 4.81 to 6.35 wt.%. This

places these rocks in the subalkaline field on a total alkali versus silica (TAS) diagram and in the low-K or medium-K fields on a SiO2 vs. K2O diagram (Fig. 7a, b). These rocks show heterogeneous contents in Fe2O3 (1.09–2.32 wt.%), FeO (0.56–2.41 wt.%) and MgO (0.62– 2.09 wt.%), and all samples plot within the magnesian field on an FeOT/(FeOT+MgO) vs. SiO2 diagram (Fig. 7c). The aluminum saturation index (ASI) values range between 0.93 and 1.07, indicating a weak metaluminous to weak peraluminous character (Fig. 7d). The felsic intrusives have a restricted range of rare earth element (REE) abundance (67.15–92.92 ppm) and show moderate enrichment in light rare earth elements (LREEs) with (La/Yb)N of 4.88–6.67 and weak negative Eu anomalies (Eu/Eu* =0.76–0.95) (Fig. 8a). On a primitive mantle-normalized spider gram (Fig. 8b), all the samples analyzed exhibit similar Cs, Ba, K, and Sr enrichments and similar Nb, Ta, Zr, and Ti depletions. The volcanic rocks of Middle–Upper Ordovician (including basaltic andesite and andesite) have SiO2 contents ranging from 52.20 wt.% to 61.20 wt.%, and varied Al2O3 (15.02–19.84 wt.%), MgO (3.51–5.06 wt.%), Fe2O3 (3.20–4.96 wt.%), FeO (3.32–6.96 wt.%), Na2O (2.94– 13

4.22 wt.%), K2O (0.18–1.40 wt.%), CaO (6.02–10.07 wt.%), and TiO2 (0.62–0.80 wt.%) contents. The Mg# [100 × Mg/(Mg + Fe2+)] ranges from 37.68 to 54.44. On a TAS diagram diagram (Fig. 7a), all the volcanic rocks plot as subalkaline and fall in the basaltic andesite and andesite fields. The volcanics also plot in the andesite and basalt fields on a Zr/TiO2 vs. Nb/Y diagram (Fig. 7e). The basaltic andesite and andesite, respectively, show low-K tholeiitic and medium-K calc-alkalic affinities (Fig. 7b, f). The volcanic rocks have a large range in REE abundance (43.86–152.50 ppm) (Table 1) and show LREEs enrichment with a (La/Yb)N ration of 2.74–8.59 and weak Eu anomalies (Eu/Eu* =0.83–1.13) (Fig. 8a). On a primitive mantlenormalized spider gram (Fig. 8b), the REEs exhibit similar enrichments in the large ion lithophile elements (LILEs), including Cs, Ba, K, and Sr, and depletions in several high field strength elements (HFSEs) such as Nb, Ta, Zr, and Ti. The mafic–intermediate intrusive rocks (including diabase, monzodiorite, and gabbroic diorite) display only small variations in bulk rock composition (Table 1; Fig. 7). Their SiO2 contents range from 51.26 to 54.74 wt.%, Al2O3 from 14.72 to 15.44 wt.%, MgO from 4.33 to 5.16 wt.%, Fe2O3 from 4.83 to 5.65 wt.%, FeO from 7.42 to 8.34 wt.%, CaO from 5.47 to 8.20 wt.%, and TiO2 from 1.67 to 1.91 wt.%. The Mg# ranges from 38.63 to 42.40. They have high concentrations of Na2O (3.82–5.40 wt.%) and low concentrations of K2O (0.32–0.89 wt.%). The diabase and gabbroic diorite fall in the subalkaline field and the two monzodiorite samples straddlethe alkaline–subalkaline field boundary on the TAS diagram (Fig. 7a). All the mafic– intermediate rock analyses fall in the subalkaline fields on the Zr/TiO2 vs. Nb/Y diagram (Fig. 7e). These mafic–intermediate intrusions are low-K to medium-K (Fig. 7b). They have a more restricted range in REE abundance (59.86–88.67 ppm) and display relatively flat REE patterns with weak positive Eu anomalies (Eu/Eu* = 1.02–1.24) (Fig. 8a). The primitive mantlenormalized spider gram (Fig. 8b) shows that they are enriched in Cs, Ba, K, and Sr, and depleted in Nb, Ta, Zr, and Ti.

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5.2. Zircon U-Pb ages and Hf isotopes Five samples, three biotite monzogranites, one quartz monzonite, and one diabase were selected for zircon U-Pb dating and Hf isotope analyses. A summary of the results is listed in Table 2 and detailed data for the SHRIMP zircon U-Pb age determinations and LA-ICPMS zircon Hf isotope analyses are presented in Supplementary Tables S1 and S2, respectively. The zircons from the three biotite monzogranite samples show oscillatory zoning and are predominantly euhedral stubby to elongate crystals with grain sizes in the 80–200 μm range (Fig. 9a–c). Zircons from sample ZF-2 have low Th (40–97 ppm) and U (91–155 ppm) contents with high Th/U ratios (0.41–0.65) (Supplementary Table S1). Twelve analyzed spots on 12 grains are concordant or nearly concordant and yield 206Pb/238U ages ranging from 450 Ma to 468 Ma with a weighted mean age of 458 ± 4 Ma (Supplementary Table S1; Fig. 10a). The Hf isotope analyses from sample ZF-2 show initial 176Hf/177Hf ratios of 0.282801–0.283025, εHf(t) values of +11.3 to +19.0 (average +13.8, Table 2, Supplementary Table S2). Zircons from sample ZF-3 have low Th (43–180 ppm) and U (97–269 ppm) contents with high Th/U ratios (0.46–0.87) (Supplementary Table S1). Nine analyses from nine grains are concordant or nearly concordant and yield

206Pb/238U

ages ranging from 448 Ma to 476 Ma with a weighted mean

age of 463 ± 7 Ma (Fig. 10b). The Hf isotope data from sample ZF-3 show initial 176Hf/177Hf ratios of 0.282776–0.282958, εHf(t) values of +10.7 to +17.0 (average +14.3, Table 2, Supplementary Table S2). Zircons from sample ZF-7 have low Th (39–171 ppm) and U (92– 287 ppm) contents with high Th/U ratios (0.44–0.67) (Supplementary Table S1). Eleven analyses from spots on 11 grains are concordant or nearly concordant and yield 206Pb/238U ages ranging from 427 Ma to 452 Ma with a weighted mean age of 436 ± 4 Ma (Fig. 10c). Hafnium isotope analyses conducted on sample ZF-7 give initial 176Hf/177Hf ratios of 0.282851–0.282990, εHf(t) values of +12.4 to +17.4 (average +14.0, Table 2, Supplementary Table S2). Zircons from the ZF14-44 quartz monzonite sample) have no visible oscillatory zoning and

15

are predominantly euhedral stubby to elongate crystals with grains ranging in size from 120– 250 μm (Fig. 9d). These zircons have low Th (17–158 ppm) and U (42–301 ppm) contents with moderate to high Th/U ratios (0.27–0.68) (Supplementary Table S1). Twelve analyses on 12 grains are concordant or nearly concordant and yield 206Pb/238U ages ranging from 418 Ma to 442 Ma with a weighted mean age of 433 ± 3 Ma (Fig. 10d). The Hf isotope analyses yield initial

176Hf/177Hf

ratios of 0.282902–0.282998 and εHf(t) values of +14.3 to +17.4 (average

+15.7, Table 2, Supplementary Table S2). Zircons from diabase sample ZF14-7 show oscillatory zoning and are predominantly euhedral stubby crystals with grain sizes ranging from 70 to 120 μm in size (Fig. 9e). The zircons from this sample have low Th (61–361 ppm) and U (131–361 ppm) contents with moderate to high Th/U ratios (0.48–1.03) (Supplementary Table S1). Ten spot analyses from 10 grains are concordant or nearly concordant and yield 206Pb/238U ages ranging from 443 Ma to 460 Ma with a weighted mean age of 448 ± 3 Ma (Fig. 10e). The Hf isotope initial 176Hf/177Hf ratios range from 0.282843–0.282979with εHf(t) values of +12.3 to +17.1 (average +14.5, Table 2, Supplementary Table S2).

6. Discussion 6.1. Early Paleozoic magmatism in the East Junggar terrane Our new zircon U-Pb radiometric ages for mafic to felsic plutons in the Zhifang area suggest that they were emplaced between 463 and 433 Ma. The Middle Ordovician (463 Ma) biotite monzogranite is apparently the oldest pluton on the northern side of the KOB and the Early Silurian (433 Ma) quartz monzonite (Figs. 1c and 2b) is a newly identified outcrop of an early Paleozoic plutons in the East Junggar terrane. Some of the other granitic rocks and quartz monzonites in this region have zircon U-Pb ages of 455–432 Ma (Xu et al., 2013b) (Fig. 1c). Basalt from the Ordovician volcanic sequence at Hongliuxia, about 30 km south of Zhifang, yielded an isotopic Pb-Pb age of 446 Ma (Ding and Tang, 1999). The Ordovician intermediate16

felsic volcanic rocks northwest of Zhifang have zircon U-Pb ages of 454–442 Ma (Xu et al., 2013b). Our new data combined with previous results demonstrate that early Paleozoic magmatism in the Zhifang area began at about 463 Ma and continued until at least 432 Ma. In the southeastern part of the East Junggar terrane, early Paleozoic magmatism is represented by a monzogranite pluton and a gabbro-diorite suite in the Barkol area with zircon U-Pb ages of 441 Ma (Guo et al., 2013) and 444–440 Ma (Xiao et al., 2019), respectively. In the eastern part of the terrane, zircon cores from a granite porphyry in the Qiongheba area yielded a U-Pb age of 442 Ma (Zhang et al., 2010) and the zircon U-Pb dates show that the Kfeldspar granites in the Hersai area were emplaced between 432 Ma and 429 Ma (Du et al., 2010). On the southern side of the KOB, plagiogranite with a zircon U-Pb age of 468 Ma was identified in the northeastern Laojunmiao area (Shi et al., 2015). Recently, an Early Silurian quartz diorite (443 Ma, Huang et al., 2016a) and a quartz monzodiorite (433 Ma, Huang et al., 2016b) were identified in the western part of the East Junggar terrane. These ages indicate that significant early Paleozoic magmatism occurred in the East Junggar terrane (Fig. 1c) and that this magmatic event probably began in the late Early Ordovicianand peaked in Late Ordovician–Early Silurian (445–435 Ma) (Fig. 11a). The geochronology of detrital zircons from sedimentary rocks in the East Junggar terrane (Fig. 1c) also provides evidence for this early Paleozoic magmatic activity (Cai et al., 2015b; Huang et al., 2013; Long et al., 2012). The U-Pb dates on detrital zircons from the Lower–Middle Devonian lithic sandstones on the southern side of the KOB yielded age peaks of 450 Ma (Fig. 11b, Cai et al., 2015b) and similar age peaks for the Silurian–Early Devonian graywacke zircons from the northern side were ~440 Ma (Fig. 11c, Long et al., 2012). In addition, U-Pb dates for detrital zircons from sedimentary sequences in the Armantai ophiolitic mélange yielded an age peak at 446 Ma (Fig. 11d, Huang et al., 2013). All these data indicate that peak magmatism in the East Junggar terrane occurred in the Late Ordovician to Early Silurian.

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6.2. Origin of the early Paleozoic igneous rocks All the analyzed igneous rocks show high positive εHf(t) values (+10.7 to +19.0, Supplementary Table S2) and plot near the depleted mantle boundary (Fig. 12) indicating their relatively juvenile character and their depleted-mantle source. Note that the εHf(t) values of the igneous rocks analyzed for this study are similar to those of most of the other early Paleozoic igneous rocks in the East Junggar terrane suggesting a dominantly juvenile source region. The compilation of ages and Hf isotopic data for zircon xenocrysts from the Paleozoic felsic igneous rocks in the East Junggar terrane also confirms the predominantly juvenile source signature (Zhang et al., 2017). 6.2.1 Petrogenesis of the mafic–intermediate rocks The mafic–intermediate rocks analyzed for this study show lower Mg#s (37.68–52.89), Cr (5.06–73.5 ppm), and Ni (8.21–34.5 ppm) (Table 2) than mantle-derived primary melts (Mg# = 73–81, Cr>1000 ppm, Ni>400 ppm (Sharma, 1997; Wilson 1989). This indicates that it is not likely that these rocks were crystallized from primary magma derived from a mantle source. These mafic–intermediate rocks have negative Nb-Ta-Ti anomalies and are enriched in LREEs and LILEs and depleted in HFSEs (Fig. 8); this is consistent with the geochemical signature of arc magmas (e.g., Luhr and Haldar, 2006; Tamura et al., 2014). Moreover, they display relatively low Nb contents (0.95‒3.41 ppm) and low Nb/La ratios (0.12‒0.41) similar to typical arc rocks that commonly contain less than 4 ppm Nb and have Nb/La ratios less than 0.9 (e.g., Tamura et al., 2014). On a V–Ti diagram the basaltic andesites and andesites predominantly show Ti/V ratios between 10 and 20 and plot in the island arc basalt (IAB) field (Fig. 13a). This IAB affinity is also demonstrated in the Th/Yb–Nb/Yb diagram (Fig. 13b), where these volcanic rocks plot at higher Th/Yb values indicating a volcanic arc array with clear a subduction zone influence (Pearce, 2008). In addition, low Nb/La (0.12‒0.41) and Nb/Zr (0.02‒0.03) ratios most probably reflect a subduction component dominated by slab-derived fluids (Hofmann, 1997;

18

Kepezhinskas et al., 1997). Therefore, the mafic–intermediate rocks were likely generated by partial melting of a mantle wedge metasomatized by fluids released from the subducted slab. 6.2.2 Petrogenesis of the felsic plutons The East Junggar felsic plutons have low ASI values (<1.1, Fig. 7d) consistent with the geochemical characteristics of I-type granites (Chappell and White 1992; Chappell, 1999). The East Junggar plutons have high SiO2 (64.81–76.46 wt.%) and low MgO (0.62–2.09 wt.%) contents with Mg#s (35.97–48.87) that imply these rocks are not likely generated by direct melting of a mantle source. The rocks are enriched in LREEs and LILEs and depleted in HFSEs (e.g., Nb, Ta and Ti) indicating island arc affinity (Fig. 8). In addition, the felsic plutons have relatively low Rb (3.01–39.20 ppm), low Y (12.70–15.40 ppm), and Nb contents (2.34–3.54 ppm) that plot in the volcanic arc granite (VAG) field (Fig. 13c). Experimental petrogenetic studies have demonstrated that partial melting of mafic rocks can generate melts with SiO2 > 60 wt.% that are low in Mg, TiO2, and Fe2O3T (Koepke et al., 2004; Rapp and Watson, 1995). It is apparent that the siliceous rocks in the KOB thatcontain Mid–Late Ordovician conodonts (Huang et al., 1990) and the plagiogranite with azircon U-Pb age of 497±12 Ma (Jian et al., 2005) preserve a record of formation of the early Paleozoic oceanic crust. Thus, we speculate that the felsic plutons were generated by partial melting of this juvenile crust, crust likely derived from Late Neoproterozoic to earliest Paleozoic juvenile depleted mantle. In addition, the predominantly juvenile signature of the early Paleozoic igneous rocks further implies that arc magmatism has played an important role in the growth of the continental crust in the East Junggar terrane and the CAOB. 6.3. Implications for tectonic evolution Rconstruction of the Paleozoic tectonic evolution of the East Junggar terrane inevitably hinges on the evolution of the Kalamaili Ocean, but the ocean’s evolution remains ambiguous because of the controversy surrounding the timing of its opening and closing. Some authors 19

have proposed that the Kalamaili Ocean was an early Paleozoic ocean by emphasizing the angular unconformity between the Ordovician and post-Ordovician strata and the occurrence of Tuvaella fauna in the Kalamaili region (He and Li, 2001; He et al., 2001). Alternatively, on the basis of the Early Devonian–Early Carboniferous ages for the ophiolitic blocks in the KOB, other authors have suggested a late Paleozoic ocean (Fang et al., 2015; Huang et al., 2012; Li, 1995; Liu et al., 2017; Qin, 2012; Shu and Wang, 2003; Tang et al., 2007; Wang et al., 2009; Xu et al., 2015). Although there is no consensus on the formation of the Kalamaili Ocean, the geological fact that the Tuvaella fauna exists in the Kalamaili region rocks (Cai et al., 2015a; He et al., 2001; Li, 1995; Su, 1981; Zhang et al., 1983) suggests that this ocean did not last from the early Paleozoic to the Devonian (Xia et al., 2007). An important issue related to the above controversy concerns what sedimentary environment one ascribed to the Middle Silurian–Early Carboniferous strata in the Kalamaili region. On the northern side of the KOB, the Middle–Upper Ordovician is unconformably overlain by Upper Silurian units (Figs. 2a, 2b, 3a, and 4a) and the Lower–Middle Silurian is absent. The Upper Silurian, Devonian, and Lower Carboniferous (Tournaisian) form a continuous sedimentary sequence composed mainly of pyroclastic and clastic sedimentary rocks with abundant brachiopod, coral, crinoid, and plant fossils (Figs. 5, 14 and Supplementary Fig. S1). Moreover, from the Upper Silurian to the Lower Carboniferous, the sediment grain sizes display coarse-fine-coarse variations that record transgressive-regressive sedimentary sequences. On the southern side of the KOB, the Middle Silurian rocks are the oldest stratigraphic unit and they are terrigenous clastic sedimentary and fine pyroclastic rocks and host the Tuvaella fauna (Figs. 5d and 14). Tuvaella brachiopods most commonly occur in continental margin and at a specific depth (not more than 20–30 m) in marine environments (Boucot, 1975; Su, 1981). This setting indicates a littoral environment for the southern side of the KOB during the Middle Silurian. The Middle Silurian is overlain by a continuous sequence of Late Silurian through

20

Early Carboniferous (Tournaisian) units that are also dominated by clastic sedimentary and pyroclastic rocks (Supplementary Fig. S2) and yield fossils similar to those in the northern side (Fig. 14). Additionally, the grain size of these sediments also displays a coarse-fine-coarse variation, reflecting a deepening-shallowing-deepening change of the sea. Field investigations and stratigraphic correlation show that the angular unconformities are widely developed on the northern side of the KOB between the Upper Silurian, the Lower Devonian, the underlying Middle–Upper Ordovician, and the early Paleozoic felsic plutons (Figs. 2a–b, 3a–b and 4a). The Middle–Late Ordovician rocks experienced greenschist facies to epidote-amphibolite facies metamorphism (He and Li, 2001; Xu et al., 2013b). Notably, the sedimentary and structural characteristics of the Middle–Upper Silurian, Devonian, and Lower Carboniferous (Tournaisian) on both sides of the KOB are similar, including conformable contact, flat occurrence, and weak deformation and metamorphism (Supplementary Figs. S1, S2). Moreover, the lithologies on each side are predominantly by sedimentary clastic and pyroclastic rocks with abundant brachiopod, coral, crinoid, and plant fossils (Figs. 5 and 14). This suggests a littoral–shallow marine environment in the Kalamaili region during the Middle Silurian–Early Carboniferous. A biostratigraphic study of the Devonian sedimentary successions in Northern Xinjiang revealed that the Emsian brachiopods of this region were similar to contemporaneous faunas in the Armantai-Beitashan and Bakol regions (Gong et al., 1996; Xiao et al., 1992a). This uniformity of sedimentary environments and biota implies that a vast ocean may not have existing in the region around East Junggar during the late Paleozoic. Therefore, the early Paleozoic and Early Devonian–Early Carboniferous ophiolites along the KOB may have different tectonic environments. The zircon U-Pb age of 497 Ma for the Kalamaili ophiolite (Jian et al., 2005) and the occurrence of Mid–Late Ordovician conodonts in the siliceous rocks (Huang et al., 1990) indicate that a paleo-oceanic domain, that is, the Kalamaili Ocean, formed in the early Paleozoic.

21

The occurrence of 463–433 Ma arc-related magmatic rocks in our study area indicates that oceanic subduction took place during the Middle Ordovician–Early Silurian (Fig. 14). The regional angular unconformity between the Upper Silurian and the underlying Middle–Upper Ordovician and Middle Ordovician–Early Silurian felsic plutons (Figs. 2a–b, 3a–b and 4a) together with the littoral–shallow marine environment in the Kalamaili region during the Middle Silurian–Early Carboniferous implies that the Kalamaili Ocean might have closed before the Middle Silurian (Fig. 14). The East Junggar terrane may be part of an epicontinental sea on the Siberian plate during this period. This scenario is consistent with the occurrence of Tuvaella fauna—an endemic fauna confined to the southern margin of the Siberian craton (Su, 1981)—in the Kalamaili region. After the pre-Middle Silurian closing of the Kalamaili Ocean, the East Junggar terrane most likely experienced extension and rifting in the Early Devonian (Fig. 14) as indicated by the formation of a ~417 Ma ophiolite (Huang et al., 2012) and a ~413 Ma granite (Li et al., 2009). As mentioned above, the uniformity of the sedimentary environments and the biota indicates that the East Junggar terrane had not experienced full extension during the late Paleozoic and that a broad ocean between the southern and northern sides of the KOB was not formed. As to the origin of the Late Devonian–Carboniferous ophiolitic fragments in the Kalamaili region, some authors have proposed that these rocks are non-ophiolitic blocks mixed into the mélange by later tectonism (He et al., 2001). Alternatively, others have suggested that these ophiolitic blocks represent new oceanic crust formed in the late Paleozoic (Qin, 2012; Shu and Wang, 2003; Tang et al., 2007; Wang et al., 2009; Xu et al., 2015; Han and Zhao, 2018). Therefore, the tectonic significance of Late Devonian–Carboniferous ophiolites remains controversial and their origin is an important issue that needs to be investigated further. However, that investigation is beyond the scope of this paper. Owingto the above controversy, whether the tectonic event in the Early Carboniferous (~340 Ma) indicates intracontinental orogenesis (Ren

22

et al., 2016) or was the final amalgamation in the East Junggar terrane (Han and Zhao, 2018; Huang et al., 2012; Li et al., 1989; Wu et al., 2012; Xu et al., 2015; Zhang et al., 2013, 2015) that also needs further investigation. Regionally, the early Paleozoic ocean in northern West Junggar (Du and Chen, 2017; Du et al., 2019; Yang et al., 2018, 2019; Zhao and He, 2014; Zheng et al., 2019) is also proposed to have closed before Middle–Late Silurian (Chen et al., 2015; Du et al., 2019; Yang et al., 2019). After the Middle–Late Silurian initial amalgamation, northern West Junggar probably underwent extension and re-rifting in the Devonian (Chen et al., 2015; Du et al., 2019; Yang et al., 2019). Therefore, it is possible that the mid-Paleozoic initial amalgamation and subsequent extension and re-rifting are geodynamic events that occurred inof East and West Junggar at the same time. If this is the case, the realization that these events were coeval could lead to a better understanding of the southwestern CAOB’s tectonic history.

7. Conclusion The igneous rocks unconformably overlain by Upper Silurian and Lower Devonian in the Zhifang area, East Junggar terrane, were generated during the Middle Ordovician–Early Silurian (463–433 Ma). This confirms the occurrence of early Paleozoic magmatism in the Kalamaili orogenic belt. These early Paleozoic igneous rocks consist predominantly of felsic plutons and mafic–intermediate volcanic rocks but include minor mafic–intermediate dikes. These igneous rocks have geochemical signature of arc magmas and were derived from a juvenile source, indicating that considerable continental growth took place in the East Junggar terrane during the early Paleozoic. The mafic–intermediate rocks were generated by partial melting of a mantle wedge metasomatized by slab-derived fluids, whereas the felsic plutons were produced by the partial melting of juvenile crust. The Kalamaili Ocean might have formed in the early Paleozoic and been subducted during the Middle Ordovician–Early Silurian; this timing is indicated by the formation of the 463–433 23

Ma arc-related igneous rocks. This ocean probably closed before Middle Silurian, the age of the regional angular unconformity between the Upper Silurian and the underlying Middle– Upper Ordovician and Middle Ordovician–Early Silurian felsic plutons. After the early Paleozoic closing of the Kalamaili Ocean, the East Junggar terrane most likely experienced extension and rifting in the Early Devonian. The development of apaleo-oceanic domain in the Kalamaili region during the late Paleozoic is a subject that needs further study.

Acknowledgements We appreciate helpful discussions with Acad. Jishun Ren during the preparation of this manuscript. We are grateful to the Editor and anonymous reviewers for their detailed comments and constructive suggestions that have improved the manuscript. Thanks are also due to Wei Zhang for her assistance with SHRIMP zircon U-Pb dating. This work was supported by the National Natural Science Foundation of China (Grant No. 41572206) and the China Geological Survey (Grant No. DD20190358).

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Early Carboniferous volcanic rocks in East Junggar, North Xinjiang: implications for postcollisional magmatism and geodynamic process. Gondwana Research 28 (4), 1466– 1481. Zhang, Z.C., Zhou, G., Kusky, T.M., Yan, S.G., Chen, B.L., Zhao, L., 2009. Late Paleozoic volcanic record of the Eastern Junggar terrane, Xinjiang, Northwestern China:major and trace element characteristics, Sr–Nd isotopic systematics and implications for tectonic evolution. Gondwana Research 16 (2), 201–215. Zhang, Z.X., Rong, J.Y., Di, Q.L., 1983. Silurian Tuvaella gigantean faunule (brachiopoda) of the Barkol area, northeastern Xinjiang. Acta Palaeontologica Sinica 22 (3), 278–294 (in Chinese with English abstract). Zhao, L., He, G.Q., 2014. Geochronology and geochemistry of the Cambrian (~518 Ma) Chagantaolegai ophiolite in northern West Junggar (NW China): constraints on spatiotemporal characteristics of the Chingiz–Tarbagatai megazone. International Geology Review 56, 1181–1196. Zheng, R.G., Zhao, L., Yang, Y.Q., 2019. Geochronology, geochemistry and tectonic implications of a new ophiolitic mélange in the northern West Junggar, NW China. Gondwana Research, doi.org/10.1016/j.gr.2019.01.008 (in press).

Fig. 1. (a) Tectonic outline of the Central Asian Orogenic Belt (modified from Jahn et al., 2000; Sengör and Natal'in, 1996). (b) Simplified tectonic map of Northern Xinjiang (modified from Chen and Jahn, 2004; Xiao et al., 2008): AOB, Armantai ophiolite belt; KOB, Kalamaili ophiolite belt; JAC, Jiangjunmiao accretionary complex EF, Erqis fault; KF, Kalamaili fault. (c) Simplified geological map of the East Junggar terrane (modified from BGMRXUAR, 1993). The labeled boxes indicate the locations of the areas shown on the geologic maps in the Fig. 2a, b, and c. 42

Fig. 2. Geologic maps of the study area showing sample locations and the unconformities discussed in the text. (a) Northern Zhifang area (modified from RGSGXGB, 1977); (b) Southern Zhifang area (modified from RGSGXGB, 1981; XIGS, 2000); (c) Southern Kalamaili ophiolite belt (modified from XIGS, 2003). Fig. 3. Angular unconformities between Upper Silurian and Middle–Upper Ordovician (a–b), Lower Devonian and Middle–Upper Ordovician (c–d), Lower Devonian and the early Paleozoic granite (e–f) and Lower Carboniferous (Tournaisian and Visean) (g–h). The locations of the unconformities are shown in Fig. 2a. Fig. 4. Photographs of outcrops showing the contacts between strata and plutons in the Zhifang area, East Junggar terrane. (a) The angular unconformity between Late Silurian and Middle– Late Ordovician rocks with the basal Silurian conglomerate above the unconformity; (b) Early Paleozoic monzogranite intruded into Middle–Upper Ordovician; (c) Early Paleozoic quartz monzonite stock; (d) Early Devonian granitic conglomerate unconformably overlying an early Paleozoic monzogranite pluton. Fig. 5. Fossils in Silurian–Devonian strata of the Kalamaili region Fig. 6. Photomicrographs of representative igneous rocks from the Zhifang area, East Junggar terrane. (a) Biotite monzogranite (ZF-3); (b) Quartz monzonite (ZF14-44); (c) Quartz monzonite (ZF14-46); (d) Gabbroic diorite (ZF-8); (e) Monzodiorite (ZF14-12); (f) Diabase (ZF14-7); (g) Andesite (ZF-9); (h) Basaltic andesite (ZF14-2). Qt, Quartz; Kf, K-feldspar;Pl, Plagioclase; Bt, Biotite; Hb, Hornblende; Cpx, Clinopyroxene. Fig. 7. Selected major element diagrams for the igneous rocks from the Zhifang area, East Junggar terrane. (a) TAS diagram (after Le Maitre, 2002); (b) K2O vs. SiO2 diagram (after Peccerillo and Taylor, 1976); (c) FeOT/(FeOT+MgO) vs. SiO2 diagram (after Frost and Frost, 2008); (d) ASI vs. SiO2 diagram (after Frost and Frost, 2008); (e) Zr/TiO2 vs. Nb/Y diagram

43

(after Winchester and Floyd, 1977); (f) FeOT/MgO vs. SiO2 diagram (after Miyashiro, 1974). FeOT=FeO+0.9Fe2O3, ASI=molecular Al/(Ca−1.67P+Na+K). Fig. 8. Chondrite-normalized rare earth element patterns (a) and primitive mantle-normalized trace element spidergrams s (b) for the igneous rocks from the Zhifang area, East Junggar terrane. The chondrite and primitive mantle normalization values are from Sun and McDonough (1989). Data for the island arc basalts in the Mariana arc are from Tamura et al. (2014). Fig. 9. Cathodoluminescence images of selected zircon grains from igneous rocks from the Zhifang area, East Junggar terrane. Samples ZF-2, -3, and -7 are monzogranites, ZF14-44 is a quartz monzonite, and ZF14-7 is a diabase. Fig. 10. Zircon U-Pb concordia diagrams for the igneous rocks from the Zhifang area, East Junggar terrane. Samples ZF-2, -3, and -7 are monzogranites, ZF14-44 is a quartz monzonite, and ZF14-7 is a diabase. Fig. 11. Single-grain U-Pb age-probability density plots of magmatic zircons for the early Paleozoic igneous rocks (a) and detrital zircons for the Paleozoic sedimentary rocks (b–d) in the East Junggar terrane. Original data sources for the early Paleozoic igneous rocks are Du et al., 2010, Guo et al., 2009, 2013, Huang et al., 2016a, 2016b, Shi et al., 2015, Xu et al., 2013b, Zhang et al., 2010, and this study. Fig. 12. Graph showing εHf(t) vs. U-Pb age for zircons from early Paleozoic igneous rocks in the Zhifang area, East Junggar terrane. CHUR = chondritic uniform reservoir (DePaolo and Wasserburg, 1976). Data source: a, Guo et al., 2013; b, Huang et al., 2016a; c, Huang et al., 2016b; d, Xu et al., 2013b. Fig. 13. Discrimination diagrams for igneous rocks from the East Junggar terrane showing: (a) V vs. Ti (after Shervais, 1982) and (b) Th/Yb vs. Nb/Yb (after Pearce, 2008) for the Middle– Upper Ordovician volcanic rocks; (c) Rb vs. Y+ Yb (after Pearce et al., 1984) for the Middle 44

Ordovician–Early Silurian felsic plutons. IAB, island arc basalt; MORB, mid-ocean ridge basalt; CFB, Continental flood basalt; OIB, oceanic island basalt; N-MORB, normal MORB; EMORB, enriched MORB; syn-COLG, syn-collision granite; WPG, within plate granite; ORG, ocean ridge granite; VAG, volcanic arc granite. Fig. 14. Three partially illustrative stratigraphic columns for the Kalamaili Ophiolite Belt showing the space-time and tectonic relationships among the tectonostratigraphic units that comprise the orogenic belt.

45

Table 1 Petrographic data, major (wt.%) and trace element (ppm) concentrations for the igneous rocks from the Zhifang area, East Junggar terrane, China Biotite monzogranite

Sample

Quartz monzonite

Andesite

ZF-2

ZF-3

ZF-4

ZF-5

ZF-7

ZF14-44

ZF14-45

ZF14-46

ZF-9

ZF-11

ZF-12

SiO2

75.49

74.95

76.46

75.99

74.25

65.78

66.23

64.81

58.76

59.31

61.20

TiO2

0.25

0.24

0.22

0.23

0.25

0.46

0.43

0.49

0.63

0.65

0.62

Al2O3

13.56

13.35

13.11

12.76

13.94

16.68

16.61

16.81

15.02

15.61

15.13

Fe2O3

1.22

1.11

1.09

1.39

1.36

1.92

1.45

2.32

3.48

3.96

3.20

FeO

0.91

0.90

0.76

0.56

0.69

2.17

2.41

1.98

3.32

3.99

3.51

MgO

0.63

0.97

0.62

0.65

0.67

1.90

1.91

2.09

4.07

5.06

4.02

CaO

1.62

1.68

1.28

2.12

1.63

2.97

2.71

3.09

8.81

6.41

6.52

MnO

0.04

0.05

0.06

0.02

0.06

0.07

0.07

0.08

0.15

0.15

0.12

Na2O

5.48

6.35

5.87

5.68

4.81

5.71

5.63

5.96

4.18

3.31

4.22

K2O

0.73

0.35

0.47

0.55

2.27

2.18

2.40

2.23

1.40

1.35

1.28

P2O5

0.05

0.05

0.05

0.05

0.06

0.14

0.13

0.14

0.19

0.19

0.18

LOI

1.38

1.59

1.32

2.01

1.20

3.54

3.07

3.64

6.92

4.36

4.80

46

Biotite monzogranite

Sample

Quartz monzonite

Andesite

ZF-2

ZF-3

ZF-4

ZF-5

ZF-7

ZF14-44

ZF14-45

ZF14-46

ZF-9

ZF-11

ZF-12

Total

99.85

99.84

99.90

99.90

99.89

99.69

99.69

99.75

99.59

99.53

99.57

Mg#

35.97

47.62

38.98

39.04

38.37

46.50

47.87

47.84

52.89

54.44

52.86

Li

5.09

7.50

5.30

4.65

4.54

10.40

12.20

14.10

18.80

32.50

22.40

Be

0.72

0.75

0.63

0.77

0.91

0.67

0.52

0.67

0.79

1.01

0.84

Sc

4.46

4.52

4.14

4.38

4.93

9.74

9.10

9.63

21.50

24.20

23.10

V

20.20

31.60

20.60

23.30

22.30

91.20

81.50

88.80

167.00

205.00

194.00

Cr

4.39

1.36

0.86

0.96

0.79

6.08

11.40

6.88

96.10

104.00

100.00

Co

2.75

3.12

2.52

2.58

2.82

10.80

10.90

11.30

20.00

22.80

21.20

Ni

0.60

2.46

0.65

0.65

0.57

5.35

7.99

5.88

20.90

32.80

25.60

Cu

1.36

7.97

1.21

1.01

1.42

31.70

28.50

32.80

51.70

51.60

64.70

Zn

12.10

19.10

17.70

9.66

20.40

39.90

40.30

42.90

56.90

73.80

60.30

Ga

11.40

11.00

10.90

10.10

12.60

15.20

14.80

15.40

14.00

19.20

16.30

Rb

6.83

3.01

4.28

4.54

17.30

36.70

39.20

34.50

12.20

11.80

14.40

Sr

494.00

506.00

778.00

790.00

340.00

600.00

617.00

523.00

293.00

433.00

433.00

47

Biotite monzogranite

Sample

Quartz monzonite

Andesite

ZF-2

ZF-3

ZF-4

ZF-5

ZF-7

ZF14-44

ZF14-45

ZF14-46

ZF-9

ZF-11

ZF-12

Y

13.40

15.40

13.70

12.70

15.30

12.90

13.90

12.80

17.80

19.90

19.20

Zr

98.70

102.00

91.90

86.70

105.00

111.00

113.00

104.00

186.00

198.00

197.00

Nb

3.36

3.31

3.03

3.04

3.54

2.34

2.50

2.64

2.90

3.41

3.17

Cs

0.09

0.08

0.08

0.13

0.11

0.69

0.74

0.54

0.35

0.28

0.22

Ba

538.00

153.00

110.00

113.00

1005.00

1230.00

1019.00

943.00

627.00

533.00

676.00

La

15.70

16.90

16.10

13.90

18.40

12.30

12.10

13.20

22.90

26.30

25.40

Ce

34.10

40.50

37.20

32.30

42.20

25.10

25.20

26.50

58.60

68.30

66.20

Pr

3.46

3.95

3.63

3.18

4.06

3.30

3.33

3.37

6.14

6.99

6.80

Nd

12.30

14.10

13.00

11.40

14.50

13.90

14.40

14.10

24.60

28.10

26.90

Sm

2.39

2.87

2.57

2.31

2.85

2.86

2.99

2.76

5.25

5.94

5.68

Eu

0.68

0.67

0.75

0.63

0.73

0.67

0.71

0.72

1.31

1.46

1.45

Gd

2.17

2.52

2.30

2.05

2.55

2.25

2.63

2.16

4.33

4.87

4.51

Tb

0.38

0.42

0.38

0.35

0.42

0.41

0.42

0.41

0.66

0.76

0.72

Dy

2.09

2.47

2.17

2.05

2.39

2.29

2.59

2.37

3.35

3.75

3.57

48

Biotite monzogranite

Sample

Quartz monzonite

Andesite

ZF-2

ZF-3

ZF-4

ZF-5

ZF-7

ZF14-44

ZF14-45

ZF14-46

ZF-9

ZF-11

ZF-12

Ho

0.48

0.53

0.46

0.45

0.53

0.44

0.46

0.46

0.67

0.76

0.72

Er

1.54

1.68

1.52

1.46

1.68

1.45

1.32

1.28

2.03

2.34

2.24

Tm

0.28

0.31

0.27

0.26

0.31

0.23

0.28

0.22

0.31

0.36

0.34

Yb

1.84

2.02

1.79

1.72

1.98

1.71

1.78

1.57

1.96

2.23

2.12

Lu

0.30

0.33

0.29

0.28

0.33

0.24

0.26

0.24

0.30

0.34

0.33

Hf

3.19

3.10

2.82

2.82

3.28

3.49

3.57

3.08

5.04

5.80

5.53

Ta

0.31

0.29

0.28

0.27

0.33

0.17

0.17

0.19

0.21

0.26

0.23

Tl

0.02

0.01

0.02

0.02

0.05

0.17

0.10

0.08

0.06

0.06

0.06

Pb

2.02

2.58

2.57

1.64

3.26

4.92

3.59

3.64

5.81

7.40

7.58

Bi

0.02

0.02

0.04

0.03

0.01

0.00

0.02

0.03

0.09

0.06

0.08

Th

3.76

3.83

3.42

3.38

3.92

3.58

3.62

3.58

4.03

4.72

4.44

U

1.17

1.05

0.83

0.66

1.07

1.15

1.03

0.96

1.62

1.77

1.58

ΣREE

77.70

89.26

82.44

72.33

92.92

67.15

68.46

69.35

132.40

152.50

146.98

(La/Sm)N

4.24

3.80

4.04

3.88

4.17

2.78

2.61

3.09

2.82

2.86

2.89

49

Biotite monzogranite

Sample

Quartz monzonite

Andesite

ZF-2

ZF-3

ZF-4

ZF-5

ZF-7

ZF14-44

ZF14-45

ZF14-46

ZF-9

ZF-11

ZF-12

(Gd/Yb)N

0.98

1.03

1.06

0.99

1.07

1.09

1.22

1.14

1.83

1.81

1.76

(La/Yb)N

6.12

6.00

6.45

5.80

6.67

5.16

4.88

6.03

8.38

8.46

8.59

Eu/Eu*

0.91

0.76

0.95

0.88

0.83

0.81

0.77

0.91

0.84

0.83

0.88

Table 1 Continued. Basaltic andesite

Sample

Gabbroic diorite

Monzodiorite

Diabase

ZF14-2

ZF14-3

ZF14-4

ZF14-5

ZF14-6

ZF-8

ZF14-11

ZF14-12

ZF14-13

ZF14-7

SiO2

53.81

53.37

53.75

53.14

52.20

54.74

53.70

52.87

53.00

51.26

TiO2

0.73

0.79

0.80

0.77

0.73

1.67

1.78

1.81

1.81

1.91

Al2O3

19.04

19.67

19.84

19.29

19.10

14.72

15.25

15.35

15.35

15.44

Fe2O3

4.65

4.15

3.36

3.86

4.96

5.29

5.13

4.83

5.64

5.65

FeO

5.75

6.31

6.96

6.87

5.80

7.50

7.73

8.34

7.50

7.42

MgO

3.51

4.19

4.28

3.51

3.73

4.33

4.80

4.71

4.63

5.16

CaO

8.12

6.76

6.02

8.24

10.07

6.88

5.52

5.39

5.47

8.20

50

Basaltic andesite

Sample

Gabbroic diorite

Monzodiorite

Diabase

ZF14-2

ZF14-3

ZF14-4

ZF14-5

ZF14-6

ZF-8

ZF14-11

ZF14-12

ZF14-13

ZF14-7

MnO

0.16

0.16

0.15

0.16

0.15

0.23

0.25

0.22

0.23

0.24

Na2O

3.74

3.85

4.16

3.61

2.94

3.82

5.13

5.40

5.25

4.15

K2O

0.36

0.59

0.54

0.41

0.18

0.60

0.49

0.85

0.89

0.32

P2O5

0.13

0.15

0.15

0.15

0.13

0.23

0.23

0.22

0.23

0.25

LOI

3.03

3.90

4.05

3.13

3.72

1.44

2.21

1.88

2.14

2.14

Total

99.35

99.28

99.18

99.20

99.32

99.15

99.15

99.07

99.16

99.14

Mg#

38.64

42.66

43.31

37.68

39.35

38.63

40.93

39.83

39.61

42.40

Li

13.90

27.80

27.70

19.30

18.90

5.74

22.60

21.60

22.80

17.30

Be

0.41

0.35

0.77

0.73

0.28

0.57

0.85

0.49

0.40

0.58

Sc

30.80

31.20

32.50

31.20

32.10

34.30

38.50

38.40

35.90

40.30

V

324.00

315.00

327.00

326.00

328.00

478.00

425.00

433.00

420.00

402.00

Cr

9.01

5.06

8.07

6.73

8.01

13.10

6.83

8.01

6.18

73.50

Co

30.70

29.60

28.80

30.00

31.70

34.00

41.70

35.40

35.60

40.70

Ni

9.96

8.21

10.30

9.98

10.40

14.30

16.70

15.70

14.90

34.50

51

Basaltic andesite

Sample

Gabbroic diorite

Monzodiorite

Diabase

ZF14-2

ZF14-3

ZF14-4

ZF14-5

ZF14-6

ZF-8

ZF14-11

ZF14-12

ZF14-13

ZF14-7

Cu

24.90

257.00

144.00

80.20

160.00

157.00

157.00

117.00

111.00

162.00

Zn

88.90

92.40

96.30

90.30

87.30

104.00

91.20

124.00

98.10

113.00

Ga

18.00

20.10

18.80

18.50

19.40

19.00

20.20

18.90

18.70

19.20

Rb

4.71

7.86

6.85

5.17

2.62

7.05

6.00

14.20

12.70

3.56

Sr

462.00

405.00

475.00

466.00

391.00

407.00

312.00

656.00

562.00

437.00

Y

13.70

15.00

16.20

14.70

12.80

31.10

31.40

30.40

29.90

35.80

Zr

39.70

45.10

47.40

42.80

40.30

88.60

85.70

79.10

77.50

115.00

Nb

1.00

1.09

1.22

1.09

0.95

2.70

2.52

2.45

2.41

2.60

Cs

0.31

0.45

0.38

0.33

0.19

0.11

0.09

0.28

0.27

0.17

Ba

93.10

85.70

104.00

142.00

147.00

429.00

177.00

536.00

443.00

118.00

La

6.50

8.13

8.33

7.08

6.52

6.69

6.51

5.91

5.92

10.80

Ce

14.00

17.10

17.50

15.20

14.10

19.20

15.30

14.80

14.60

25.00

Pr

2.02

2.45

2.49

2.18

2.04

2.73

2.48

2.38

2.38

3.83

Nd

9.79

11.60

12.00

10.40

9.55

13.40

13.10

12.60

12.60

19.00

52

Basaltic andesite

Sample

Gabbroic diorite

Monzodiorite

Diabase

ZF14-2

ZF14-3

ZF14-4

ZF14-5

ZF14-6

ZF-8

ZF14-11

ZF14-12

ZF14-13

ZF14-7

Sm

2.41

2.87

2.90

2.53

2.34

4.15

3.82

3.80

3.78

5.27

Eu

0.79

0.99

1.06

0.91

0.84

1.58

1.45

1.37

1.38

1.84

Gd

2.31

2.53

2.83

2.76

2.19

3.67

4.31

4.38

4.55

5.48

Tb

0.45

0.44

0.52

0.40

0.43

0.80

0.86

0.87

0.82

1.04

Dy

2.52

3.00

3.16

2.66

2.45

5.24

5.83

5.81

6.00

6.71

Ho

0.50

0.57

0.56

0.53

0.46

1.12

1.18

1.05

1.06

1.31

Er

1.39

1.57

1.83

1.40

1.21

3.19

3.38

3.18

3.03

3.75

Tm

0.25

0.25

0.27

0.23

0.21

0.53

0.56

0.50

0.52

0.63

Yb

1.70

1.48

1.69

1.54

1.33

3.16

3.36

3.24

2.74

3.51

Lu

0.24

0.21

0.29

0.23

0.20

0.47

0.51

0.50

0.49

0.50

Hf

1.26

1.71

1.79

1.48

1.45

2.74

2.99

2.54

2.37

3.53

Ta

0.08

0.11

0.11

0.07

0.07

0.22

0.18

0.18

0.17

0.15

Tl

0.02

0.03

0.02

0.02

0.01

0.03

0.01

0.04

0.08

0.02

Pb

4.48

3.34

3.54

4.22

5.03

1.50

2.00

1.26

1.02

4.96

53

Basaltic andesite

Sample

Gabbroic diorite

Monzodiorite

Diabase

ZF14-2

ZF14-3

ZF14-4

ZF14-5

ZF14-6

ZF-8

ZF14-11

ZF14-12

ZF14-13

ZF14-7

Bi

0.13

0.05

0.05

0.04

0.04

0.01

0.09

0.04

0.04

0.03

Th

1.00

1.16

1.20

1.07

0.98

0.62

0.63

0.58

0.55

1.25

U

0.32

0.44

0.45

0.34

0.39

0.37

0.43

0.34

0.37

0.56

ΣREE

44.87

53.19

55.44

48.04

43.86

65.93

62.65

60.39

59.86

88.67

(La/Sm)N

1.74

1.83

1.85

1.81

1.80

1.04

1.10

1.00

1.01

1.32

(Gd/Yb)N

1.12

1.41

1.39

1.48

1.36

0.96

1.06

1.12

1.37

1.29

(La/Yb)N

2.74

3.94

3.54

3.30

3.52

1.52

1.39

1.31

1.55

2.21

Eu/Eu*

1.02

1.13

1.13

1.06

1.13

1.24

1.09

1.03

1.02

1.05

Note: Mg#=100 × Mg/(Mg+Fe2+); Eu/Eu* =

EuN / (SmN × GdN)1/2. The major elements were normalized to 100% anhydrous.

54

Table 2 Summary of zircon U-Pb ages and Hf isotope analyses for the igneous rocks from the Zhifang area, East Junggar terrane Sample

Rock type

U-Pb age (Ma)

(176Hf/177Hf)i

εHf(t)

ZF-2

Biotite monzogranite

458 ± 4 Ma

0.282876

13.8

ZF-3

Biotite monzogranite

463 ± 7 Ma

0.282888

14.3

ZF-7

Biotite monzogranite

436 ± 4 Ma

0.282897

14.0

ZF14-44

Quartz monzonite

433 ± 3 Ma

0.282949

15.7

ZF14-7

Diabase

448 ± 3 Ma

0.282902

14.5

Note: (176Hf/177Hf)i and εHf(t) are the average values of samples.

55

Highlights 

Middle Ordovician–Early Silurian arc magmatism is confirmed.



The early Paleozoic igneous rocks were derived from a juvenile source.



Closure of the Kalamaili Ocean occurred before Middle Silurian.



Early Paleozoic is an important continental growth episode in the East Junggar.

56

Declaration of interests ☒ The authors declare that they have no known competing financial interests or personal relationships that could have appeared to influence the work reported in this paper.

☐The authors declare the following financial interests/personal relationships which may be considered as potential competing interests:

57