Permo-Triassic evolution of the southern margin of the Central Asian Orogenic Belt revisited: Insights from Late Permian igneous suite in the Daheishan Horst, NE China

Accepted Manuscript Permo-Triassic evolution of the southern margin of the Central Asian Orogenic Belt revisited: insights from Late Permian igneous suite in the Daheishan Horst, NE China

Zhi-Gang Song, Zuo-Zhen Han, Li-Hua Gao, Hong-Yan Geng, Xu-Ping Li, Fan-Xue Meng, Mei Han, Wen-Jian Zhong, Jing-Jing Li, Qing-Xiang Du, Jun-Lei Yan, Hui Liu PII: DOI: Reference:

S1342-937X(17)30373-8 https://doi.org/10.1016/j.gr.2017.12.005 GR 1894

To appear in: Received date: Revised date: Accepted date:

31 May 2017 29 November 2017 1 December 2017

Please cite this article as: Zhi-Gang Song, Zuo-Zhen Han, Li-Hua Gao, Hong-Yan Geng, Xu-Ping Li, Fan-Xue Meng, Mei Han, Wen-Jian Zhong, Jing-Jing Li, Qing-Xiang Du, Jun-Lei Yan, Hui Liu , Permo-Triassic evolution of the southern margin of the Central Asian Orogenic Belt revisited: insights from Late Permian igneous suite in the Daheishan Horst, NE China. The address for the corresponding author was captured as affiliation for all authors. Please check if appropriate. Gr(2017), https://doi.org/10.1016/j.gr.2017.12.005

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ACCEPTED MANUSCRIPT Permo–Triassic evolution of the southern margin of the Central Asian Orogenic Belt revisited: insights from Late Permian igneous suite in the Daheishan Horst, NE China

Zhi-Gang Songa,b, Zuo-Zhen Hana,b,c*, Li-Hua Gaoa,b, Hong-Yan Gengd, Xu-Ping Lia,b, Fan-Xue Menga,b, Mei Hana,b, Wen-Jian Zhonga,b, Jing-Jing Lia,b, Qing-Xiang Dua,b, Jun-Lei Yana,b, Hui Liua,b College of Earth Science and Engineering, Shandong University of Science and Technology, Qingdao 266590,

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a

Key Laboratory of Depositional Mineralization & Sedimentary Mineral of Shandong Province, Shandong

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b

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China

University of Science and Technology, Qingdao 266590, China

Laboratory for Marine Mineral Resources, Qingdao National Laboratory for Marine Science and Technology,

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c

d

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Qingdao 266071, China

University Research Facility in Chemical and Environmental Analysis, The Hong Kong Polytechnic

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University, Hung Hom, Kowloon, Hong Kong, China

ABSTRACT

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The Daheishan Horst in Jilin Province, NE China, is the key geological unit that links the Solonker-Xar Moron-Changchun and Changchun-Yanji sutures, which are generally interpreted

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to mark the zone of closure of the Paleo-Asian Ocean along the southeastern margin of the Central Asian Orogenic Belt (CAOB). Here we investigate a suite of volcanic rocks from Daheishan and the Doushantouzi syenogranite intrusion to gain insights into the Permian–Triassic tectonic evolution of the eastern segment of the Paleo-Asian Ocean. Zircon U-Pb data and zircon rare earth element (REE) patterns indicate that the igneous suite formed during the Late Permian (ca. 253–256 Ma) and underwent late-stage alteration during Early Triassic tectono-thermal events. The presence of older magmatic zircon grains in this igneous

ACCEPTED MANUSCRIPT suite suggest multiple pulses of magmatism during the Permian in this region. The Daheishan volcanic rocks predominantly consist of intermediate-felsic rocks, including andesite, rhyodacite and rhyolite with minor basaltic lava. Geochemically, the basaltic rocks belong to calc-alkaline series, whereas the intermediate-felsic rocks classify as calc-alkaline to high-K calc-alkaline

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series. The Doushantouzi syenogranites belong to peraluminous I-type granites, with the A/CNK ratio between 1.06 and 1.15. All these rocks have an arc-like affinity with enriched light rare

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earth elements (LREE) and large ion lithophile elements (LILE; e.g. Rb, Ba and U) and depleted

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high field strength elements (HFSE; e.g. Nb, Ta and Ti). The Daheishan intermediate-felsic rocks

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and Doushantouzi syenogranites have higher SiO2 contents but lower MgO contents and Mg# values and plot within the field of experimentally derived partial melts from metabasaltic rocks

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in MgO vs. SiO2 diagram. These geochemical features, together with the positive εNd(t) (+1.6 to +4.6) and εHf(t) (+1.53 to +7.41) values of zircon grains, indicate that the primary magma of

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these intermediate-felsic rocks likely originated from the partial melting of a juvenile metabasaltic lower crust. In contrast, the basaltic lavas were probably derived from the partial

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melting of a depleted mantle wedge that was metasomatized by fluids from a subducted slab, as suggested by their low initial 87Sr/86Sr ratios and depleted εNd(t) (+3.6 to +4.4) values. These data,

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in conjunction with regional geological investigations, suggest that formation of this Late Permian igneous suite was related to the northward subduction of the Paleo-Asian oceanic plate beneath the Songliao-Xilinhot block and that the eastern segment of the Paleo-Asian Ocean did not close before the Late Permian. Keywords: Zircon geochronology; Geochemistry; Permian igneous suite; Tectonic setting; Paleo-Asian Ocean

1. Introduction

ACCEPTED MANUSCRIPT The Central Asian Orogenic Belt (CAOB), which occupies area wide region between the East European, Siberian, and Tarim-Sino Korean cratons (Fig. 1), is one of the largest and most complex orogenic collages in the world, which preserves crucial evidence for Phanerozoic juvenile crustal growth (Sengör et al., 1993, 1996; Wilde et al., 2000, 2003; Xiao et al., 2003,

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2015; Jahn et al., 2004; Kröner et al., 2007; Windley et al., 2007; Safonova et al., 2011; Zheng et al., 2013; Safonova and Santosh, 2014; Xiao and Santosh, 2014; Wilde, 2015; Wilde and Zhou,

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2015; Cai et al., 2016; Liu et al., 2017a; Safonova, 2017). Despite the recent debate on the

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proportion of recycled older crust versus juvenile material involved in the construction of the

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CAOB (Kröner et al., 2014, 2017; Safonova, 2017), this huge Phanerozoic belt is regarded as a natural laboratory for exploring the geodynamic processes during accretionary orogenesis and

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continental growth (Xiao and Santosh, 2014). The CAOB contains a variety of tectonic components that include arc/back arc systems, oceanic islands/plateaus, ophiolites,

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micro-continental fragments, and collisional and post-collisional complexes (Sengör et al., 1993; Jahn et al., 2000, 2004; Windley et al., 2007; Eizenhöfer et al., 2014; Kröner et al., 2014; Xiao et

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al., 2015; Liu et al., 2017a; Safonova et al., 2017; Zhou et al., 2017). Northeast (NE) China, which has been tectonically referred to as the Xing’an-Mongolian (or Xing-Meng) Orogenic Belt

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(XMOB) in the past, is thought to represent the southeastern segment of the CAOB (Fig. 1, Li, 2006; Wu et al., 2007; Zhou et al., 2014; Song et al., 2015). During its long tectonic evolutionary history, this region witnessed the evolution and final closure of the Paleo-Asian Ocean and the amalgamation of micro-continental blocks, which from southeast to northwest are Jiamusi-Khanka, Songliao-Xilinhot, Xing’an, and Erguna blocks (Fig. 2a, Wu et al., 2007, 2011; Zhou et al., 2009, 2010a,b,c, 2013, 2014, 2017; Cao et al., 2013; Wilde, 2015; Liu et al., 2017a). This region also witnessed overprinting by the circum-Pacific tectonic domain in the east and the

ACCEPTED MANUSCRIPT Mongol-Okhotsk tectonic domain in the northwest during the Mesozoic (Li, 2006; Windley et al., 2007; Wu et al., 2007; Xu et al., 2009; Zhou and Wilde, 2013; Guo et al., 2015; Wilde, 2015; Wilde and Zhou, 2015). The magmatic records left by these complex tectonic processes make NE China a significant region to study the tectonic architecture of NE Asia in general, and the

It

was

traditionally

believed

by

many

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evolution of the CAOB. geologists

that

the

Solonker-Xar

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Moron-Changchun-Yanji suture zone (Fig. 2a) marks the site of the final closure of the

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Paleo-Asian Ocean and this suture has become the focus of many recent studies (Jia et al., 2004;

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Sun et al., 2004; Li, 2006; Wu et al., 2007, 2011; Cao et al., 2013; Zhou and Wilde, 2013; Eizenhöfer et al., 2014; Safonova and Santosh, 2014; Yu et al., 2014; Han et al., 2016, 2017a;

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Wang et al., 2016; Du et al., 2017). However, poor exposure and extensive offset by several Mesozoic faults (Fig. 2a) complicates the precise location of the suture zone with the eastern

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Songliao basin remains controversial (Liu et al., 2017a). This is particularly important in the Daheishan Horst (Fig. 2b), which is the key geological unit that links the Solonker-Xar

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Moron-Changchun suture and the Changchun-Yanji suture. Despite extensive geological, petrological, geochronological, and geochemical investigations in this suture over the last decade

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(Table 1, Jia et al., 2004; Wu et al., 2007; Cao et al., 2013; Yu et al., 2014; Wang et al., 2015a, b, 2016; Du et al., 2017; Liu et al., 2017b; Zhou et al., 2017b), several questions remain unanswered. A number of previous studies suggest that the Changchun-Yanji suture represents the eastern extension of the Solonker-Xar Moron-Changchun suture (Cao et al., 2013; Wang et al., 2015a, b; Wilde et al., 2015; Du et al., 2017; Liu et al., 2017a, b). In contrast, some researchers have recently proposed that the Changchun-Yanji suture is part of the Jilin-Heilongjiang high-pressure metamorphic belt rather than the eastern extension of the

ACCEPTED MANUSCRIPT Solonker-Xar Moron-Changchun suture (Zhou and Wilde, 2013; Zhou et al., 2014; Wang et al., 2016; Zhou and Li, 2017). In addition, controversy has long surrounded the timing of final closure of the Paleo-Asian Ocean, with different perspectives arguing for distinct suturing times ranging from the Middle

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Devonian to the Late Triassic (Xu et al., 2013, 2015; Xiao et al., 2003, 2015; Jia et al., 2004; Jian et al., 2010; Cao et al., 2013; Zhao et al., 2013; Eizenhöfer et al., 2014; Wang et al., 2015a, b,

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2017; Chen et al., 2016; Du et al., 2017; Li et al., 2017a, b; Zhou et al., 2017b). Furthermore, the

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closure style of the eastern segment of the Paleo-Asian Ocean is also an unresolved issue. First,

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it is not clear whether the closure is a scissor-style or a single event (Sun et al., 2004; Cao et al., 2013; Zhang et al., 2014; Wang et al., 2015a, b). Second, the polarity of subduction of the eastern

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segment of the Paleo-Asian Ocean remains open to debate. Some authors have proposed a bidirectional subduction model (Xiao et al., 2003, 2015; Li, 2006; Chen et al., 2009; Eizenhöfer

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et al., 2014; Liu et al., 2017a), whilst others advocate a single, southward subduction (Cao et al., 2013; Wang et al., 2015a, b; Liu et al., 2017b).

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Thus, there has been no consensus so far on the tectonic evolution of the eastern segment of the Paleo-Asian Ocean during the Late Paleozoic–Early Mesozoic. In order to gain more insights

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on these issues, we investigate a suite of igneous rocks, including basaltic lava, andesite, rhyodacite, rhyolite and mylonitized syenogranite, which are exposed in the Daheishan Horst, NE China. We present results from systematic field investigations, zircon U-Pb and Lu-Hf isotope analyses, whole-rock geochemistry and Sr-Nd isotope analyses. These new data not only constrain the ages and petrogenesis of the igneous rocks, but also provide further insights into the Permian–Triassic tectonic evolution of the eastern segment of the Paleo-Asian Ocean.

2. Geological background

ACCEPTED MANUSCRIPT The Daheishan volcanic rocks and Doushantouzi pluton are located at the confluence of the Daheishan Horst and eastern segment of the Solonker-Xar Moron-Changchun-Yanji suture zone, which is located between the North China Craton (NCC) and the combined NE China blocks. To the north of the Changchun-Yanji suture are the Jiamusi block, Khanka block and

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Songliao-Xilinhot block (Fig. 2a). Studies of the various blocks, including their cover sequences, basements, and amalgamation histories, have provided valuable new information in the recent

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years (Supplementary Table 1).

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The Jiamusi block in eastern NE China is traditionally considered as an important tectonic unit

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in the CAOB (Wilde et al., 1999, 2000, 2015; Zhou et al., 2009, 2010b, 2014, 2017a; Wilde and Zhou, 2015) and is connected to the Bureya block to the north in Far East Russia and the Khanka

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block to the south, which are separated by the Jiamusi-Yilan fault and Dunhua-Mishan fault, respectively (Fig. 1). Many geologists have studied the basements of these blocks and argued

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that these blocks form one contiguous crustal unit, referred to as the Bureya-Jiamusi-Khanka block (Natal’in and Borukayev, 1991; Sengör et al., 1993; Wilde et al., 2010; Zhou et al., 2010a,

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b; Wu et al., 2011; Yang et al., 2014, 2017a, b). Other workers proposed that the Jiamusi-Khanka blocks were probably derived from the Gondwana supercontinent (Wilde et al., 2000, 2003, 2010;

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Yang et al., 2015a, 2017a). However, Luan et al. (2017a) questioned this viewpoint and argued that the Jiamusi block may be an exotic massif with an affinity to the Tarim craton. The rock associations in the Jiamusi and Khanka blocks mainly include the granulite-facies Mashan complex, the blueschist-facies Heilongjiang complex, and various stages of igneous rocks (Wilde et al., 2000, 2003; Li et al., 2009, 2010a; Zhou et al., 2010c). The Mashan complex, which is referred to as the “khondalite series” in Chinese literature, is dominated by graphitic schists, felsic granulites and marbles. Recent precise age data indicate that the protoliths of these rocks

ACCEPTED MANUSCRIPT were formed during the Meso–Neoproterozoic and underwent metamorphism at ~500 Ma (Wilde et al., 1999, 2000; Zhou and Wilde, 2013; Wilde and Zhou, 2015; Luan et al., 2017a; Yang et al., 2017a). The Heilongjiang complex, which is structurally interleaved with the Mashan complex, occurs only in the western part of the Jiamusi block with the best preserved outcrops in a

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north-south trending direction from the Luobei, Yilan and Mudanjiang areas (Wu et al., 2007; Zhou et al., 2009; Yang et al., 2015a, b, 2017a, b). This complex is considered to represent the

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remains of an ocean (namely the Mudanjiang Ocean) that existed between the Songliao-Xilinhot

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block and Jiamusi block (Wu et al., 2007; Li et al., 2009; Zhou et al., 2009; Ge et al., 2016,

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2017). In recent studies, Zhu et al. (2015, 2017a, b, c) reported geological, geochronological and geochemical data on various rocks from the Heilongjiang complex, and proposed that the

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Mudanjiang Ocean was not closed by ~141 Ma. The igneous rocks in the Jiamusi and Khanka blocks and their margins are mainly divided into three phases: (1) Early Paleozoic (Wilde et al.,

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2000, 2003; Bi et al., 2014; Yang et al., 2014, 2015a), (2) Carboniferous–Permian (Bi et al., 2015, 2016, 2017; Yang et al., 2015a; Dong et al., 2017a, b; Yang et al., 2017b), and (3) Triassic (Wang

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et al., 2015c; Yang et al., 2015a, b; Liu et al., 2017c). To the west of the Jiamusi block is the Songliao-Xilinhot block that occupies an extensive area

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in NE China (Fig. 2a). This block consists of the southern Great Xing’an Range in the west, the Zhangguangcai Range in the east, the Lesser Xing’an Range in the northeast and the Songliao basin in the centre part. Data from several hundred drill holes reveal that the basement beneath the Songliao basin is dominated by Paleozoic volcanic strata and Paleozoic–Mesozoic granitoids (Pei et al., 2007; Zhou et al., 2012), with minor Precambrian components (Wang et al., 2006; Pei et al., 2007). In addition, Precambrian basement rocks from the Dongfengshan Group and the Tadong Group have also been recognized in the northeastern part of the block (Quan et al., 2013;

ACCEPTED MANUSCRIPT Wang et al., 2013; Luan et al., 2017b). Most recently, Mesoproterozoic ages of ca. 1400 Ma were acquired from gneissic granites in the Xilinhot complex, suggesting that the Xilinhot block probably is a microcontinental fragment with Precambrian basement rather than a Paleozoic accretionary complex as previously considered (Han et al., 2017b).

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The Mesozoic–Cenozoic Daheishan Horst lies between the eastern margin of the Songliao Basin and the Jiamusi-Yilan fault (Fig. 2a). Assuming that the Changchun-Yanji suture is the

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eastern extension of the Solonker-Xar Moron-Changchun suture, preserved relics of the suture

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should exist in the Daheishan Horst. Paleozoic volcano-sedimentary strata in the study area

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mainly include the Jingtai marbles, Fangniugou volcanic (intermediate-felsic) rocks, Taoshan Formation (black siltstone, silty slate) and Daheishan volcanic rocks (Fig. 2b, JBGMR, 1988,

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1997; Jiang et al., 2014). Based on regional stratigraphic correlations or stratigraphic relationships with fossil-bearing strata, this suite of rocks were previously considered to be

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Precambrian, Ordovician, or Silurian in earlier studies (e.g., Liu et al., 1982; JBGMR, 1988, 1997; Jia, 1990). Gao (1985) argued that the Daheishan volcanic rocks (which he named as the

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Daheishan conglomerate) formed at ca. 246 Ma based on a Rb–Sr whole-rock isochron age. However, this age should represent the post-magmatic event, considering the rocks in the study

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area have undergone greenschist facies metamorphism (JBGMR, 1988, 1997). Recent geochronological studies show that the Fangniugou volcanic rocks formed during the Silurian (Jiang et al., 2014; Pei et al., 2016); however, the age of the Daheishan volcanic rocks remains unclear because of a lack of precise geochronological data. The granitoids that are exposed in the study area were previously determined to be Phanerozoic in age (Fig. 2b). In earlier research, these granitoids were considered to have formed during the Middle Paleozoic (JBGMR, 1997) based on their lithostratigraphic relationships, K-feldspar K-Ar dating and other analytical

ACCEPTED MANUSCRIPT approaches. This understanding is also supported by the recently reported zircon U-Pb ages of granodiorites (419 ± 3 Ma and 414 ± 5 Ma) from the Ximangzhang pluton and monzogranite (400 ± 2 Ma) from the Houmiaoling pluton (Pei et al., 2016). Some of the contact relationships of the various plutons in the study area are still ambiguous because of serious coverage and

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fewer outcrops, so further research is still required to better constrain the ages of the plutons.

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3. Sampling and petrography

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Representative field photographs and photomicrographs of the Daheishan volcanic rocks

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(DYS1, DYS2, SD7, XQ1, XDC1 and XDM1) and the Doushantouzi pluton (DST1) are shown in Fig. 3 and Fig. 4, respectively. The sample localities, GPS readings, rock types, field

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observations and mineralogy of the different rock types are described below. Samples DYS1 and DYS2 (foliated rhyodacite) were collected ~0.5 km west of Dayushu

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village (43°34′33.3″N, 125°06′04.6″E). Sample DYS1 is khaki in colour and has a porphyritic texture and rhyolitic structure, while sample DYS2 is grey green in colour and has a porphyritic

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texture and massive structure (Fig. 3a, Fig. 4a, b). The phenocrysts (~30%) of both samples consist of alkali feldspar (~23%), minor hornblende (~3%), quartz (~2%) and plagioclase (~2%).

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The groundmass (~70%) is dominated by felsic minerals (~68%) and minor accessory minerals (~3%). Sericitic alteration is common in both the groundmass and phenocrysts. Sample SD7, which is greyish green massive andesite with volcanic and carbonatite enclaves (Fig. 3b), was collected ~0.7 km northwest of Jiapigou village (43°36′37.1″N, 125°13′22.4″E). This rock shows chloritization on the surface and contains minor phenocrysts (~10%) that have experienced secondary alteration and a groundmass (~90%) that consists of plagioclase (~34%), hornblende (~15%), minor quartz (~1%) and glass (~40%) and has a hyalopilitic texture (Fig.

ACCEPTED MANUSCRIPT 4c). Sample XQ1 (foliated rhyolite) was collected ~0.1 km north of Lvhua village (43°34′49.0″N, 125°14′58.1″E). The rock is greyish green and has a porphyritic texture and massive structure (Fig. 3c, Fig. 4d). The phenocrysts (~35%) are mainly quartz (~15%), orthoclase (~10%) and

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plagioclase (~10%), while the groundmass (~65%) consists of felsic minerals (~40%), sericite (~20%) and glass (~5%).

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Sample XDC1 (basaltic lava) was collected ~0.5 km northwest of Yongjiu village

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(43°37′10.1″N, 125°16′14.6″E). The rock is dark grey in colour and has a cryptocrystalline

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texture and massive structure with microcrystalline plagioclase, pyroxene and glass (Fig. 3d, Fig. 4e). Chloritization is common on the surface of the rocks.

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Sample XDM1 was collected ~0.8 km southeast of Xiadamen village (43°36′32.3″N, 125°15′0.2″E). This sample is a grey foliated andesite with a porphyritic texture and massive

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structure (Fig. 3e, Fig. 4f). The phenocrysts (~20%) are mainly fine-grained plagioclase and minor quartz, and the groundmass (~80%) consists of microcrystalline plagioclase, sericite,

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minor quartz and opaque minerals.

Sample DST1 (mylonitized syenogranite) was collected ~0.2 km east of Doushantouzi village

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(43°30′31.6″N, 125°07′36.9″E). The sample is light red in colour and has a granitic texture and massive structure (Fig. 3f, Fig. 4g, h). This sample is dominated by orthoclase (~30%), plagioclase (~20%), perthite (10%), and quartz (~30%), which are surrounded by microcrystalline biotite (~5%), sericite (~3%), and opaque minerals (~2%). Petrographic observations of chorite and sericite, foliation, and mylonitization suggest that all the igneous rocks in this study have undergone varying degrees of alteration and deformation during post-magmatic tectono-thermal events.

ACCEPTED MANUSCRIPT 4. Analytical methods

4.1. Major and trace element analyses

The least altered and homogeneous portions of the represented whole-rock samples were

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crushed and powdered to 200 mill in an agate mortar for geochemical analyses after petrological studies. Major and trace element analyses were conducted at the Supervision and Inspection

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Center of Mineral Resources, Ministry of Land and Resources of Jinan, Shandong Province,

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China. The concentrations of SiO2 and Al2O3 were analysed using the gelatine coagulation

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gravimetric method and the xylenol orange method, respectively. The other major oxides and some trace elements (Ba, Sr, V, and Cr) were determined by IRIS-Intrepid ICP atomic emission

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spectrometer (AES) using the standard GB/T14506-2010 for the oxides. The other trace element concentrations were determined using an XSeries 2 inductively coupled plasma mass

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spectrometer (ICP-MS). Precisely weighed sample powders (50 mg) were dissolved in Teflon bombs in HF + HNO3. An internal standard solution that contained the single element Rh was

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used to monitor signal drift during counting. Analytical uncertainties were generally less than 5%.

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The major and trace element data are provided in Table 2.

4.2. Zircon U-Pb dating

Zircon grains were separated from the whole-rock samples by conventional magnetic and heavy liquid separation techniques and were manually picked under a binocular microscope at the Langfang Regional Geological Survey, Hebei Province, China. The handpicked zircon grains were embedded in epoxy and polished down to half size and cleaned in an acid bath before analysis. Cathodoluminescence (CL) images were obtained using a JEOL scanning electron

ACCEPTED MANUSCRIPT microscope at the State Key Laboratory of Continental Dynamics, Northwest University, Xi’an, China. Distinct domains within these zircons were selected for analysis based on the CL images. U-Pb dating and trace element analyses of the zircons were conducted synchronously with an LA-ICP-MS at the State Key Laboratory of Continental Dynamics, Northwest University, Xi’an,

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China. The zircon 91500 was used as an external standard for age calibration, and the NIST SRM 610 silicate glass was applied for instrument optimization. The laser spot size was 32 μm

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during the analyses. The programs ICPMSDataCal (Ver. 6.7; Liu et al., 2008, 2010) and Isoplot

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(Ver. 3.0; Ludwig, 2003) were used for data reduction. Common Pb corrections were conducted

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following the method of Andersen (2002).

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4.3. In-situ Lu-Hf isotopic analyses

Zircon Hf isotope analyses were conducted using a Nu Plasma HR MC-ICP-MS that was

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coupled with a Geolas 2500 laser-ablation system at the State Key Laboratory of Continental Dynamics, Northwest University, Xi'an, China. A laser spot size of 44 µm and a repetition rate of

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10 Hz were used. The analytical procedures were similar to those in Xu et al. (2004). Interference from

176

175

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interference-free

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Lu while measuring

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Hf was corrected based on the intensity of the

Lu isotope and a recommended

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Lu/175Lu ratio of 0.02655 to calculate the

Lu/177Hf ratios (Chu et al., 2002). Similarly, interference from

was corrected based on the intensity of the interference-free

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176

Yb while measuring

Yb isotope and a

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176

Hf

Yb/172Yb

ratio of 0.5886 (Chu et al., 2002) to calculate the176Yb/177Hf ratios. Zircon 91500 was used as the reference standard, with a recommended respectively). The decay constant for

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Hf/177Hf ratio of 0.282314 ± 0.000011 (n = 34, 2σ,

Lu of 1.867×10-11year-1 (Albarède et al., 2006) and the

present-day chondritic ratios of 176Hf/177Hf = 0.282785 and 176Lu/177Hf = 0.0336 (Bouvier et al.,

ACCEPTED MANUSCRIPT 2008) were adopted to calculate the εHf(t) values. Single-stage model ages (TDM1) were calculated using a depleted mantle with a present-day 176

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Hf/177Hf ratio of 0.28325 and

Lu/177Hf ratio of 0.0384 (Griffin et al., 2000). Two-stage model ages (TDM2) were calculated

using an assumed 176Lu/177Hf ratio of 0.015 (Rudnick and Gao, 2003) for the average continental

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crust.

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4.4. Whole-rock Sr-Nd isotopic analyses

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Rb-Sr and Sm-Nd isotopic analyses were performed on a Micromass Isoprobe multi-collector

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ICPMS (MC-ICP-MS) in static mode at the Institute of Geochemistry, Chinese Academy of Science (Guangzhou), China. The analytical procedures followed those of Li et al. (2004). Rb

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and Sr isotopes were separated using cation columns, whereas Sm and Nd isotopes were separated using HDEHP-coated Kef columns. The measured

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Sr/86Sr and

143

Nd/144Nd ratios

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were normalized to 87Sr/86Sr = 0.1194 and 143Nd/144Nd = 0.7219, respectively. During the course of this study, the standards NBS 987 and Jndi-1 yielded 87Sr/86Sr ratios of 0.710257 ± 19 (2σ, n =

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5. Results

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5) and 143Nd/144Nd ratios of 0.512096 ± 5 (2σ, n = 5).

5.1. Zircon U-Pb geochronology and REE geochemistry

Representative CL images of zircons from the various lithologies in the igneous suite are shown in Figs. 5 and 6, together with the analytical spots. The U-Pb age data are reported in Supplementary Table 2, which are plotted in concordia diagrams together with histograms and bar charts (Figs. 7 and 8). Zircon trace element data are presented in Supplementary Table 3, and the rare earth element (REE) distribution patterns are shown in Fig. 9. The characteristics and

ACCEPTED MANUSCRIPT results of the zircons in individual samples are discussed below.

5.1.1. Daheishan volcanic rocks

The majority of zircon grains from the six Daheishan volcanic rock samples are euhedral to

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subhedral and display long prismatic to stumpy grain morphology with lengths that vary from 40 to 200 μm and aspect ratios from 1:1 to 4:1. In the CL images, most of the zircon grains exhibit

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complex internal structures (Fig. 5). Some of the grains display dark oscillatory zoned cores that

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are surrounded by wide bright outer rims with distinct oscillatory zonings, which indicate

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multiple growth stages, while some grains display transgressive zones or patches, which indicate recrystallization (Li et al., 2015). Additionally, very few zircon grains show thin and bright rims,

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which suggest weak metamorphism. The zircon trace elements suggest that these zircons have high Th/U ratios between 0.27 and 1.77 (barring one value of 6.64; Supplementary Table 2). All

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these features reveal that these zircons are of magmatic origin but they may have undergone recrystallization in the magma chamber, alongside varying degrees of superposition and

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reconstruction during subsequent thermal events (Koschek, 1993; Corfu et al., 2003; Turkina et al., 2012; Santosh et al., 2017). This conclusion is supported by the normalized REE patterns,

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which show scattered Ce and Eu anomalies (Ce/Ce* = 0.88–278.94, Eu/Eu* = 0.04–0.53) (Fig. 9). Most of the grains display distinct positive Ce anomalies and negative Eu anomalies, which are characteristic of magmatic zircons (Belousova et al., 2002), whereas the zircon grains with young ages (ca. < 252 Ma) are characterized by relatively high concentration of light rare earth elements (LREEs) and show a flat LREE distribution pattern with distinct Eu anomalies and no significant or weak Ce anomalies (Fig. 9), which are considered to be the geochemical imprints of hydrothermal activity (Hoskin and Ireland, 2000; Rubatto, 2002; Hoskin, 2005). Although

ACCEPTED MANUSCRIPT minor zircon grains with older ages also display “hydrothermal” REE patterns, we argue that these grains might represent mixed ages of magmatism and later thermal events and have no effect on our interpretations. Twenty-four spot analyses on 24 zircons from sample DYS1 (rhyodacite) produced

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Pb/238U

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ages from 242 ± 4 to 273 ± 5 Ma. As discussed above, some of the zircons have undergone late-stage superposition and reconstruction, and the young zircons (ca. < 252 Ma) display

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geochemical characteristics of hydrothermal alteration; thus, the weighted mean age (246.6 ± 3

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Ma) of these zircons might represent the timing of the late-stage thermal event, whereas the

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remaining ages reflect that continuous and multiple pulses of magmatism occurred during the Early to Late Permian. Zircons, which are initially derived from country rocks or sources and

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significantly older than the host igneous rocks, are often captured during magma emplacement or ascent (Zhang et al., 2015; Du et al., 2017). The discussion by Miller et al. (2007) and

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experiments by Matzel et al. (2006) suggested that the recycling of zircon antecrysts during successive magmatic injections is the primary cause of age dispersion. Therefore, the youngest

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group age of magmatic zircons in volcanic rocks best represents the eruption time (Su et al., 2010; Koch et al., 2015); hence, the weighted mean age (254.3 ± 2.5 Ma) of the Late Permian

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ages from sample DYS1 is interpreted as the eruption time of the rhyodacite. Seventeen spot analyses on 17 zircons from sample DYS2 (rhyodacite) produced

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Pb/238U

ages from 240 ± 5 to 266 ± 5 Ma (Fig. 7b). Similar with sample DYS1, the ages of the hydrothermal-like zircons yield a weighted mean age of 245.8 ± 3.2 Ma (MSWD = 1.12, n = 7), which correspond to the timing of a late-stage thermal event, while the older ages suggest that continuous and multiple magmatic events occurred from the Middle to Late Permian. The weighted mean age of 255.4 ± 4.2 Ma from the Late Permian zircons is interpreted as the

ACCEPTED MANUSCRIPT eruption time of the rhyodacite. A total of 21 analyses of zircons from sample SD7 (andesite) show

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Pb/238U ages that range

from 241 ± 5 to 448 ± 7 (Fig. 7c). The weighted mean age of 244.8 ± 3.7 Ma is interpreted as the timing of a late-stage thermal event because of the hydrothermal-type REE patterns of these

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zircons. The remaining Permian ages indicate multiple phases of magmatism, whereas the two much older ages represent the ages of inherited zircons. The Late Permian zircons yield a

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weighted mean age of 255.5 ± 3 Ma, which represents the eruption time of the andesite.

Pb/238U ages from 243 ± 5 to 305 ± 7 Ma (Fig. 7d). Considering the hydrothermal-type REE

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206

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A total of 54 spots were analysed for 54 zircons from sample XQ1 (rhyolite), yielding

patterns of the zircons with ca. < 252 Ma ages, the weighted mean age (245 ± 5.7 Ma) of these

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zircons should represent the timing of a late-stage thermal event. The remaining ages reflect multiple pulses of magmatism during the Permian or even the Late Carboniferous. Among these

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ages, the weighted mean age (254.7 ± 2.6 Ma) of the Late Permian zircons is interpreted as the eruption time of the rhyolite.

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A total of 28 analyses were conducted on 28 zircons from sample XDM1 (andesite), yielding Pb/238U ages from 249 ± 4 to 353 ± 7 Ma (Fig. 7e). The weighted mean age (249.9 ± 3.3 Ma)

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of the young zircons that show “hydrothermal” REE patterns is interpreted to reflect the late-stage thermal event. The remaining Permian ages indicate that multiple phases of magmatism occurred during the Permian, whereas the one much older age is considered to represent the age of an inherited zircon. The weighted mean age (254 ± 3.6 Ma) of the Late Permian zircons is interpreted as the eruption time of the andesite. Twenty-four spot analyses on 24 zircons from sample XDC1 (basaltic lava) produced 206

Pb/238U ages from 243 ± 4 to 289 ± 4 Ma (Fig. 7f). The weighted mean age of 247.7 ± 3.2 Ma

ACCEPTED MANUSCRIPT from the hydrothermal-like zircons represents the timing of a late-stage thermal event, whereas the remaining ages indicate continuous and multiple magmatic events from the Early to Late Permian. The weighted mean age (253 ± 4.5 Ma) of the Late Permian zircons is interpreted as the formation time of the basaltic lava.

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5.1.2. Doushantouzi pluton

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The complex structures of the zircons from sample DST1 indicate multiple tectono-thermal

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activity in the study area, which is also supported by the mylonitization of the pluton (Fig. 3f).

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The zircons are euhedral to subhedral and show long prismatic to stumpy grain morphology with sizes from 80 to 150 μm and aspect ratios from 1:1 to 3:1 (Fig. 6). These grains display typical

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oscillatory zonings in the CL images; however, some grains show core-rim structures, which indicate multiple growth stages, while others show transgressive zones, which indicate

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recrystallization. Trace element analyses indicate that the zircons have high Th/U ratios between 0.48 and 1.03 (Supplementary Table 2), suggesting a magmatic origin (Koschek, 1993; Corfu et

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al., 2003; Turkina et al., 2012; Santosh et al., 2017). However, the zircons have two types of distinct REE patterns with scattered Ce and Eu anomalies (Ce/Ce* = 1.07–80.44, Eu/Eu* =

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0.05–0.34) (Fig. 9f). The majority exhibit significant depletion in LREEs, enrichment in heavy rare earth elements (HREEs), and prominent negative Eu anomalies and positive Ce anomalies, which are typical features of magmatic zircons (Belousova et al., 2002). Meanwhile, the younger zircon grains (ca. < 252 Ma) are characterized by relatively high concentration of LREEs with no Ce anomalies (Fig. 9f), which are considered to be the geochemical imprints of hydrothermalism (Hoskin and Ireland, 2000; Rubatto, 2002; Hoskin, 2005). Similar to the zircons from the Daheishan volcanic rocks, several older zircons from sample DST1 display “hydrothermal” REE

ACCEPTED MANUSCRIPT patterns, which are interpreted as the mixing ages of magmatism and later thermal events and have no effect on our interpretations. The 206Pb/238U ages from 27 analyses of sample DST1 vary from 242 ± 5 to 294 ± 5 (Fig. 7g). Considering the “hydrothermal” REE patterns, the weighted mean age of 245.5 ± 5.4 Ma from

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the hydrothermal-like zircons is interpreted to represent the timing of a late-stage thermal event. The remaining ages are interpreted to reflect continuous and multiple pulses of magmatism

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represents the emplacement age of the syenogranite.

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during the Permian. The weighted mean age (256.5 ± 3.3 Ma) of the Late Permian zircons

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5.2.1. Daheishan volcanic rock samples

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5.2. Whole-rock major and trace element geochemistry

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The complex structures of the zircons, scattered zircon U-Pb ages and REE patterns,

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petrographic observations of chlorite and sericite, and moderate to high loss on ignition (LOI) of most of the samples (2.29–3.93, except for samples DYS3 and DYS4) suggest that the

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Daheishan volcanic rocks have undergone varying degrees of subsequent sub-marine

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hydrothermal alteration together with greenschist metamorphism and accretion-related deformation, to which some elements (e.g., K, Na, Rb, Ba, Sr, and Zn) are sensitive (Safonova et al., 2008, 2012, 2016a; Volkova et al., 2009). LREEs are known to be more mobile than middle and heavy REEs during secondary processes that occur only during carbonatization or at high water/rock ratios (Humphris, 1984), which are not observed in our case. Thus, our interpretation emphasizes REEs, high field strength elements (HFSEs), Cr, Ni, and some major elements that are generally immobile during secondary processes (Polat et al., 2002). The major element data show that the Daheishan volcanic samples possess large SiO2

ACCEPTED MANUSCRIPT variations, with a gap in basic to intermediate compositions (46.94–56.68 wt.%) that divides the samples into two groups: a basaltic group of basaltic lava, and an intermediate-felsic group with andesite, rhyodacite and rhyolite compositions. The Nb/Y-SiO2 diagram (Winchester and Floyd, 1977) and Co-Th diagram (Hastie et al., 2007) are used as proxies for the total alkali-silica (TAS)

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diagram and SiO2-K2O diagram, respectively (Fig. 10). The intermediate-felsic samples are compositionally andesitic/rhyodacitic/rhyolitic and belong

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to calc-alkaline or high-potassic series (Fig. 10b). In the Harker diagram (Fig. 12), the Al2O3,

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CaO, MgO, Fe2O3t, P2O5, and TiO2 concentrations are negatively correlated with SiO2, whereas

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K2O and Na2O have no significant correlation with SiO2, which implies that the alkalis were probably largely affected by weathering processes. The intermediate-felsic samples have uniform

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chondrite-normalized REE patterns (Fig. 11a; Boynton, 1984) and are characterized by enrichment in LREE (LaN/YbN = 3.28–9.36) and depleted in HREE (LREE/HREE = 3.83–7.63).

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In the primitive mantle-normalized multi-element diagram (Fig. 11b; Sun and McDonough, 1989), these samples all show remarkable enrichment in large ion lithophile elements (LILEs)

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and depletion in HFSEs. The rhyolites, rhyodacites, and XDM andesites have remarkably negative Eu and Sr anomalies, which are thought to have been caused by plagioclase

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fractionation, either in the solid residuum from partial melting or via fractional crystallization (Green, 1994). However, the SD andesites show weak negative Eu anomalies and no Sr anomalies, suggesting that plagioclase played a minor role in the partial melting and fractional crystallization. Moreover, the SD andesites exhibit weak Zr-Hf depletion (Fig. 11b), probably because of the involvement of a certain amount of carbonatite and some mantle-derived materials, which will be discussed in Section 6.2.2. The basaltic samples are marked by low SiO2 (44.73–46.94 wt.%) and TiO2 (1.41–1.54 wt.%)

ACCEPTED MANUSCRIPT with relatively high Al2O3 (18.14–18.66 wt.%) concentrations and plot within the calc-alkaline field (Fig. 10b), similar to those of high-Al basalts, which are generally regarded to be related to arcs or mid-ocean ridges (Kuno, 1960; Crawford et al., 1987; Grove et al., 1988; Sisson and Grove, 1993; Ozerov, 2000; Eason and Sinton, 2006). However, these samples have higher alkali

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contents compared modern high-Al basalts (i.e., the Aleutian, Izu, Japan, and Kamchatka Islands), probably because of Na2O and K2O mobility during alteration. A primitive

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mantle-normalized multi-element diagram and the chondrite-normalized REE patterns are shown

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in Fig. 11. The Daheishan basaltic samples are enriched in LREEs and LILEs but depleted in

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HFSEs with no obvious Eu anomaly (Eu/Eu* = 0.89–1.05), resembling typical arc rocks (or the familiar “arc signature” or continental crustal signature) (Elliott, 2003; Rudnick and Gao, 2003;

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Santosh et al., 2017). Sample XDC3 has low SiO2 and plots on the boundary of andesite and dacite (Fig. 10b), probably because of uneven distribution of minerals or significant

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post-eruptional alteration; therefore, sample XDC3 was not used to discuss the petrogenesis and tectonic setting of the basaltic rocks.

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5.2.2. Doushantouzi pluton samples

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The DST granitic samples preserve their primary minerals and have low LOI values (0.88–1.52 wt.%). This observation, alongside the homogeneous major element concentrations and consistent variations in both HREEs and HFSEs, suggests that these elements remained essentially immobile during post-magmatic alteration. However, the LREEs vary widely in the chondrite-normalized REE diagram (Fig. 11c), which indicates that these elements were mobile during post-magmatic alteration and deformation. Therefore, these elements are not used in the geochemical interpretations below.

ACCEPTED MANUSCRIPT The syenogranites are silica-rich, with SiO2 from 71.37 wt.% to 73.38 wt.%. These samples have high alkali contents, with K2O = 3.37–3.80 wt.% and Na2O = 3.44–4.06 wt.%, but low amounts of MgO (0.69–0.73 wt.%), Fe2O3t (2.87–3.04 wt.%), MnO (0.05 wt.%), TiO2 (0.33–0.39 wt.%), CaO (1.40–1.43 wt.%) and P2O5 (0.07 wt.%). Al2O3 ranges from 13.31 wt.%

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to 14.16 wt.%, with the index A/CNK ranging from 1.06 to 1.15, which indicates that these rocks are peraluminous. In the primitive mantle-normalized multi-element diagram, the granitic

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samples exhibit remarkable enrichment in LILEs and depletion in HFSEs (Fig. 11d).

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5.3. In-situ zircon Hf isotopic analyses

In-situ Hf isotope analysis was performed on zircons from five samples of the Daheishan

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volcanic rocks (DYS1, SD7, XDC1, XDM1, and XQ1) and one sample from the Doushantouzi pluton (DST1). The results are listed in Supplementary Table 4 and are shown in Fig. 13a.

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The Permian zircons from all the samples have relatively homogeneous Hf isotopic compositions. The initial 176Hf/177Hf ratios vary from 0.282668 to 0.282769, and the εHf(t) values

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range from +1.53 to +5.13. The single-stage and two-stage Hf model ages (TDM1 and TDM2) range from 835 to 681 Ma and from 1115 to 927 Ma, respectively. The Hf isotopic compositions of two

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inherited zircons (443 and 448 Ma) from sample SD7 are similar to those of the Permian zircons, with initial

176

Hf/177Hf ratios of 0.282715 and 0.282711, respectively. The εHf(t), TDM1 and TDM2

of these two zircons are 7.41 and 7.36, 752 and 758 Ma, and 929 and 937 Ma, respectively. All the analysed zircons have Hf isotopic compositions that are similar to those of Phanerozoic igneous rocks in the CAOB (Fig. 13a; Liu et al., 2010b; Cao et al., 2013; Yu et al., 2014; Bi et al., 2017) but distinct from those of Neoarchean and Paleoproterozoic zircons from the Paleozoic to upper Mesozoic strata in the Yanshan Fold and Thrust Belt (Yang et al., 2006a).

ACCEPTED MANUSCRIPT 5.4. Whole-rock Sr-Nd isotopic analyses

The results of the Sr-Nd isotopic analyses of the igneous rocks in this study are shown in Table 3 and Fig. 13b. The initial 87Sr/86Sr and εNd(t) values were calculated based on their U-Pb zircon ages. All the samples show similar and relatively uniform initial Sr-Nd isotopic compositions.

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For these samples, (87Sr/86Sr)i = 0.703102–0.704742 and εNd(t) = +1.6 to +4.6, and the Nd model

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ages (TDM) range from 997 to 648 Ma. These values are consistent with those of the Phanerozoic

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granitoids in the CAOB (Hong et al., 2000; Jahn et al., 2000; Wu et al., 1999, 2000, 2003a).

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6. Discussion

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6.1. Ages of the Daheishan volcanic rocks and Doushantouzi pluton

The volcanic rocks in the Fangniugou-Leshan area in the Daheishan Horst (Fig. 2b) were

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previously thought to have formed during Early Paleozoic based on regional stratigraphic correlations or information from fossil-bearing strata, whereas the Doushantouzi pluton was

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considered to have formed during Middle Paleozoic based on lithostratigraphic relationships, K-feldspar K-Ar dating and other analytical approaches (JBGMR, 1988, 1993). Recent

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geochronological studies showed that the Fangniugou volcanic rocks formed during the Silurian (Jiang et al., 2014; Pei et al., 2016); however, the eruption time of the Daheishan volcanic rocks has not been well constrained because of the lack of precise geochronological evidence. In this study, zircon grains from the igneous suite exhibit complex internal structures (Figs. 5 and 6), such as core-rim structures and transgressive zones, indicating multiple growth stages, recrystallization or varying degrees of superposition and reconstitution during post-magmatic thermal events (Li et al., 2015). This inference is also supported by petrographic observations of

ACCEPTED MANUSCRIPT sericite and chorite and the moderate to high LOI of most of the samples. As indicated by the chondrite-normalized REE patterns, the zircons with younger ages (ca. < 252 Ma) are characterized by high concentration of Y, Nb, Ta, Hf, and LREEs and flat LREE distribution patterns with no significant or weak Ce anomalies and distinct Eu anomalies (Supplementary

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Table 3, Fig. 9), which are considered to be the geochemical imprints of hydrothermal activity (Hoskin and Irland, 2000; Rubatto, 2002; Hoskin 2005). Therefore, the Early Triassic ages of the

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zircons represent the timing of post-magmatic thermal events, which are identical to not only the

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Rb–Sr whole-rock isochron age (246 Ma) of the Daheishan volcanic rocks (Gao, 1985), but also

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many metamorphic ages reported from the adjacent areas (Hu et al., 2003; Zhang et al., 2005; Wu et al., 2007; Liu et al., 2017b). In contrast, almost all the remaining zircons display distinct

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positive Ce anomalies and negative Eu anomalies (Fig. 9), which are characteristic of magmatic zircons (Belousova et al., 2002). However, the

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Pb/238U ages of these zircons widely range

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from the Late to Early Permian or even earlier, which indicates multiple pulses of magmas and inheritance. As suggested by the experiments from Matzel et al. (2006) and discussion by Miller

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et al. (2007), the recycling of zircon antecrysts during successive magmatic injections is the primary cause of age dispersion, so only the youngest age group in the igneous rocks best

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represent the crystallization time (Su et al., 2010; Koch et al., 2015). Consequently, the youngest magmatic group of the Daheishan volcanic rocks (253 ± 4.5~255.5 ± 3 Ma) and the Doushantouzi pluton (256.5 ± 3.3 Ma) suggest that these rocks formed during the Late Permian, which is markedly different from the Silurian volcanic rocks in the Fangniugou area (Jiang et al., 2014; Pei et al., 2016). Thus, we consider these rocks to represent different tectonic slices with different ages, and the Late Permian igneous rocks are likely the residual geologic units of the Solonker-Xar Moron-Changchun-Yanji suture zone in the Daheishan Horst.

ACCEPTED MANUSCRIPT A compilation of the age data from this study is shown in Fig. 8, which suggests that continuous and multiple pulses of magmatism occurred throughout the Early to Late Permian, with Triassic ages corresponding to metamorphic or thermal post-magmatism events. Previous studies also reported Permian magmatic ages in central Jilin Province (such as the Daheshen

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Formation, ca. 302–279 Ma; Hulan Group, ca. 287 Ma; Xiaolihe pluton, ca. 260 Ma; Seluohe complex, ca. 252 Ma; etc.) (Li et al., 2007a; Wu et al., 2007; Cao et al., 2012, 2013; Yu et al.,

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2014; Wang et al., 2015a). Permian–Triassic magmatic events have also been widely reported

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from the adjacent Yanbian and Inner Mongolia areas (see Table 1 and Fig. 14 for details). Our

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new data, in conjunction with the above data from the literature, suggest that important

the Permian–Triassic.

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6.2. Nature of the magma sources

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magmatic events occurred along the Solonker-Xar Moron-Changchun-Yanji suture zone during

6.2.1. Origin of the Daheishan basaltic rocks

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The basaltic lavas have low SiO2 and high Fe2O3t and TiO2 contents, along with high Mg#

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values, which is distinct from what is generally observed in crustally derived melts (Patiño Douce and Beard, 1995; Patiño Douce, 1997) or crustal materials (Rudnick and Gao, 2003) and implies a mantle source for the magma (Frey and Prinz, 1978). Understanding the effects of crustal contamination on basaltic rocks is important, because mantle-derived magmas could be contaminated by upper crustal material during emplacement or ascent (Guo et al., 2015). The basaltic lavas in this study are characterized by depletion in Nb and Ta, which can be related to crustal contamination or magma mixing during the ascent of magmas (Wilson, 1989) or reflect magma generation in a subduction-related environment (Pearce, 1983). Nevertheless, magma

ACCEPTED MANUSCRIPT mixing or crustal contamination en route can be ignored for the following reasons: (1) the presence of negative Zr-Hf anomalies (Fig. 11b) suggests that the magmas that formed the basaltic rocks assimilated little or no crustal material because minor crustal contamination could produce positive Zr-Hf anomalies, as these elements are enriched in crustal materials (Puchtel et

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al., 1998; Zhao and Zhou, 2007; Yu et al., 2014; Bi et al., 2017); (2) the abundances of crustal-affinity elements (e.g. Th = 2.84 ppm, U = 0.61 ppm, on average) in the basaltic lavas are

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much lower than those of average upper crust (Rudnick and Gao, 2003), which suggests a

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limited addition of crustal materials (Zhu et al., 2008); (3) the low Th/Ce (0.05–0.07) and Th/La

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(0.11–0.17) ratios of the basaltic lavas indicate that crustal contamination played an insignificant role during the magma evolution because mantle-derived magmas have low Th/Ce (0.02–0.05)

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and Th/La (~0.12) ratios (Sun and McDonough, 1989), while continental crust has relatively high Th/Ce (~0.15) and Th/La (~0.3) ratios (Taylor and McLennan, 1995; Plank, 2005); (4) the

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basaltic lavas have higher La/Nb ratios (1.7–3.9) than those of continental in eastern China (average value = 1.7), ruling out simple crustal contamination as a possible scenario (Yu et al.,

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2014); and (5) the narrow ranges of the initial 87Sr/86Sr ratios (0.7036–0.7039) and εNd(t) values (+3.6 to +4.4) for the basaltic lavas further confirm the absence of crustal contamination (Table

source.

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3). Thus, the arc-like trace element features of the XDC basaltic lavas were inherited from their

Geochemically, the basaltic lavas plot within the calc-alkaline field (Fig. 10b) and are marked by high Al2O3 concentrations (18.14–18.66%), similar to those of high-Al basalts, which are generally regarded to be related to arcs or mid-ocean ridges (Kuno, 1960; Crawford et al., 1987; Grove et al., 1988; Sisson and Grove, 1993; Ozerov, 2000; Eason and Sinton, 2006). These lavas display enrichment in LREEs and LILEs and depletion in HFSEs (Fig. 11a, b), which are

ACCEPTED MANUSCRIPT consistent with the source being a mantle wedge that was modified by components from the down-going slab (McCulloch and Gamble, 1991; Pearce and Peate, 1995; Elliott, 2003; Wei et al., 2017). The slab components include (1) the fluxing of fluids from the dehydration of hydrous oceanic crust and/or overlying sediments (Hawkesworth et al., 1993, 1997; Turner et al., 1997;

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Class et al., 2000) and (2) the melting of oceanic crust and/or sediments (Peacock et al., 1994; Stern and Kilian, 1996; Elliott et al., 1997; Hawkesworth et al., 1997; Münker, 2000). However,

a

lithologic

association

of

high-Mg

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forms

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the reaction between slab-derived silicate melts and mantle peridotite in an arc setting generally andesite–adakitic

TTG/D

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(tonalite-trondhjemite-granodiorite/dacite) (Drummond et al., 1996; Polat and Kerrich, 2001; Wang et al., 2007); therefore, the absence of this rock association in the study area precludes the

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involvement of slab melts. Accordingly, we suggest that the basaltic magma was generated by the partial melting of a mantle wedge that was previously metasomatized by slab-derived fluids.

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Moreover, the basaltic lavas have low initial 87Sr/86Sr ratios (0.7036–0.7039) and positive εNd(t) values (+3.6 to +4.4) and plot in the field that surrounds the mantle array (Fig. 13b), which

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reveals that their mantle source must have been depleted. Thus, the primary magma of the basaltic lavas in the study area was likely generated by the partial melting of a depleted mantle

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wedge that was metasomatized by fluids from a subducted slab in the ocean basin between the NCC and Songliao-Xilinhot block. This interpretation is further supported by the whole-rock geochemical and Sr-Nd isotopic studies of the Permian basalts along the Changchun-Yanji suture (Cao et al., 2011, 2012; Yu et al., 2014; Du et al., 2017).

6.2.2. Origin of the Daheishan intermediate-felsic rocks

The basaltic lavas and the intermediate-felsic rocks from the Daheishan volcanic rocks yielded

ACCEPTED MANUSCRIPT similar Late Permian formation ages in our study, suggesting that these rocks formed nearly coevally. Two different pathways have been proposed to explain the generation of intermediate-felsic magmas in an arc-related setting: (1) extensive fractional crystallization from associated arc-basaltic magmas (Pin and Paquette, 1997; Bonin, 2004), where intermediate-felsic

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and mafic rocks generally have similar geochemical and isotopic compositions and with mafic rocks being volumetrically significant (Garland et al., 1995); and (2) the partial melting of

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mid/lower crust that is underplated by mantle-derived melt, with intermediate-felsic rocks being

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volumetrically significant (Doe et al., 1982; Hildreth et al., 1991). However, the

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intermediate-felsic rocks in the study area were more likely generated through partial melting rather than the fractional crystallization of the associated basaltic magmas because (1) a distinct

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compositional gap in the SiO2 contents exists between the basaltic and intermediate-felsic rocks; (2) the scatter is greater than can be explained by fractional crystallization in the La versus

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La/Sm diagram (Fig. 15a); and (3) the exposure of the Daheishan intermediate-felsic rocks is much larger than that of the basaltic lavas, which cannot be explained by the crystallization of

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basaltic magmas.

The Daheishan volcanic rocks are mainly composed of basaltic lavas, andesites, rhyodacites

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and rhyolites, which closely resemble the lithological assemblages of volcanic rocks from subduction-related settings (Xu and Qiu, 2010). The intermediate-felsic rocks are classified as calc-alkaline volcanic rocks and are characterized by enrichment in LREEs and LILEs and depletion in HFSEs (Fig. 11a, b), alongside relatively higher SiO2 (56.68–76.18%) contents but lower MgO (0.25–3.3%) and Mg# values (18.1–46.6) than those of the lower crust (e.g., SiO2 =53.4, MgO = 7.24, Mg# = 60.1; Rudnick and Gao, 2003). In addition, these rocks contain high concentrations of Y (18.5–38.35 ppm) and Yb (1.98–4.59 ppm), yielding low (La/Yb)N

ACCEPTED MANUSCRIPT (3.28–9.36) values, which resemble those of subduction-related rocks rather than adakitic rocks (Fig. 15b). Thus, the lithological assemblages and geochemical imprints of the Daheishan volcanic rocks provide a scenario for a subduction setting that is similar to the modern Andes (Li, 2006; Yu et al., 2014) and suggest that the primary magma of the intermediate-felsic rocks was

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likely derived from the partial melting of lower crust material (Zen, 1986; Barbarin, 1999; Nabelek et al., 2001; Cao et al., 2013; Yu et al., 2014; Yang et al., 2015; Bi et al., 2017). This

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scenario similar to what was found for the adjacent Andean-type volcanic rocks in the Daheshen

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Formation as reported by Yu et al. (2014).

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Experimental studies have demonstrated that partial melts that are derived from different protoliths under variable melting conditions can be identified based on major element oxides

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(Patiño Douce and Beard, 1995; Patiño Douce, 1995, 1999; Altherr et al., 2000). The Daheishan intermediate-felsic rocks have higher SiO2 but lower MgO and Mg# and plot within the field of

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experimentally derived partial melts from metabasaltic rocks (Fig. 15c). In addition, the negative Nb-Ta-Ti anomalies in the primitive mantle-normalized multi-element diagram and the

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concave-upward chondrite-normalized REE patterns suggest residues of hornblende and garnet in the source (Rollinson, 1993). Accordingly, the Daheishan intermediate-felsic rocks were most

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likely derived from the partial melting of metabasaltic rocks at the root of an active continental margin, with hornblende and garnet as the major residue mineral phases in the source region (Wang et al., 2015d). Previous studies suggested that a juvenile basaltic lower crust, which is a mixed lithology of Precambrian lower crust that was underplated by mantle-derived basaltic magma during the Meso-Neoproterozoic, played an important role in the source region for the Phanerozoic granitoids and some felsic volcanic rocks with positive εNd(t) values in NE China (Wu et al.,

ACCEPTED MANUSCRIPT 2000, 2002, 2003a; Liu et al., 2010; Cao et al., 2013; Yu et al., 2014; Du et al., 2017). A simple mixing model with Sr-Nd isotopic data was employed to calculate the proportions of the mantle and crustal components (Fig. 16; Wu et al., 2000). The results show that the upper crustal component had no role in the generation of the Daheishan intermediate-felsic rocks, whereas the

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lower crust and mantle-derived basaltic magmas were the two major components. Fig. 16 shows that the mantle component represents approximately 75–90%. However, this observation does

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not indicate that the Daheishan intermediate-felsic rocks were formed by mixing lower crustal

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melts and basaltic materials in such proportions. Rather, the intermediate-felsic magmas were

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produced by the melting of a mixed lithology that involved pre-existing lower crustal material that was underplated or intruded by mantle-derived magma in such proportions (Wu et al., 2003a;

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Liu et al., 2010). Thus, the Nd model ages are mixed ages, which represent the average resident ages of pre-existing crust and juvenile mantle materials (Wu et al., 1999), which explains why

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the Nd model ages are younger than the Hf model ages. In addition, the zircons from the Daheishan intermediate-felsic rocks are characterized by positive εHf(t) values from 1.53 to 7.41,

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with young model TDM2 ages from 1155 to 929 Ma (Supplementary Table 4), which are consistent with previously obtained values for Phanerozoic felsic rocks in NE China and

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suggests that significant growth of the continental crust in this region occurred during the Meso-Neoproterozoic, as previously reported (Wu et al., 2000, 2002, 2003a; Liu et al., 2010; Cao et al., 2013; Yu et al., 2014). The primitive mantle-normalized multi-element diagram (Fig. 11b) shows that the SD andesites display negative Zr-Hf anomalies, which can be caused by liquid immiscibility in a system with carbonatite, the presence of ore accessories, mixing between silicate and carbonatite melts, or the melting of magmas from metasomatized mantle rocks with broad Zr-Hf anomalies

ACCEPTED MANUSCRIPT (Ivanov et al., 1998). As shown in Fig. 2b and observed in the field, carbonatite bodies accompany the Daheishan volcanic rocks, and carbonatite enclaves are present in the SD andesites (Fig. 3b). In addition, the SD andesites have significantly higher CaO contents than the other Daheishan intermediate-felsic rocks (except sample XDM3, Table 2). These lines of

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evidence imply the addition of a certain amount of carbonatite materials during the evolution and eruption of the magma. However, calculations indicate that a significant decrease in Zr and Hf

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contents requires the addition of a substantial amount of carbonatite material, so the appearance

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of negative Zr-Hf anomalies from only the involvement of carbonatite material does not seem

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probable (Ivanov et al., 1998). As noted above, the geochemical features indicate that the Daheishan volcanic rocks have an affinity to active continental margin volcanic rocks, which

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implies that the involvement of metasomatized mantle melts was another possible reason for the

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the samples (Table 2).

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negative Zr-Hf anomalies. This inference is supported by the relatively higher MgO contents of

In particular, the andesites (SD and XDM samples) have lower εNd(t) values than the

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rhyodacites and rhyolites (Table 3). This feature, together with the presence of older inherited zircons in Samples XDM1 (353 Ma) and SD7 (443 and 448 Ma), suggest the involvement of

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older crustal materials during the eruption and ascent of the magmas. In summary, the primary magma of the Daheishan intermediate-felsic rocks likely originated from the partial melting of a Meso-Neoproterozoic juvenile metabasaltic lower crust, with hornblende and garnet as the major residue mineral phases in the source region. In addition, the andesitic magmas were contaminated by some old crustal materials, and a certain amount of carbonatite materials and metasomatized mantle melts were probably involved in the SD andesitic magmas.

ACCEPTED MANUSCRIPT 6.2.3. Origin of the Doushantouzi syenogranites

The mylonitized syenogranites in this study are peraluminous, with A/CNK between 1.06 and 1.15. They have relatively high content of Na2O (3.44–4.06%) but low initial

87

Sr/86Sr ratios

(0.7039–0.7045) and positive εNd(t) values (3.7–4.1), which suggest that these rocks may be

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I-type and/or A-type, but not S-type granites according to the criteria of Chappell and White

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(1974, 1992, 2001) and Chappell (1999). Additionally, these rocks are characterized by low

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contents of Nb (6.50–11.58 ppm), Ta (0.49–0.58 ppm), Y (19.44–26.40 ppm), and Zn (29.70–47.99 ppm) alongside low 10000*Ga/Al (2.26–2.61) and Fe2O3t/MgO (3.97–4.22) values,

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similar to the highly fractionated I-type granites in NE China (Wu et al., 2003a) but distinct from

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those of A-type granites (Whalen et al., 1987). This result is also supported by the absence of mafic alkaline minerals such as riebeckite or arfvedsonite.

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The significant enrichment in LILEs and depletion in HFSEs in the primitive

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mantle-normalized multi-element diagram (Fig. 11d) and the high SiO2 concentrations imply that the primary magma of the Doushantouzi syenogranites was derived from the partial melting of

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crustal material (Xu et al., 2009; Cao et al., 2013; Bi et al., 2017). Similar to the Daheishan

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intermediate-felsic rocks, the Doushantouzi syenogranites have higher SiO 2 but lower MgO and Mg# and plotting within the field of experimentally derived partial melts from metabasaltic rocks (Fig. 15c). This result and the negative Nb-Ta-Ti anomalies (Fig. 11d) suggest residues of hornblende and garnet in the source (Rollinson, 1993). In addition, the samples have low initial 87

Sr/86Sr ratios (0.7039–0.7045), positive εNd(t) (3.7–4.1) and zircon εHf(t) (2.14–5.03) values,

and young second-stage Hf model ages from 1118 to 935 Ma, which are consistent with the isotopic features of the Daheishan intermediate-felsic rocks and indicate crustal growth during the Meso-Neoproterozoic. Therefore, the primary magma of the Doushantouzi syenogranites

ACCEPTED MANUSCRIPT likely originated from the partial melting of a Meso-Neoproterozoic juvenile metabasaltic lower crust, with hornblende and garnet as the major residue mineral phases in the source region.

6.3. Tectonic implications

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6.3.1. Eastward termination of the Solonker-Xar Moron-Changchun suture

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The Solonker-Xar Moron-Changchun suture represents the middle domain of the Paleo-Asian

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Ocean closure during the Permian. However, its eastward continuation in NE China has remained controversial. Traditionally, the Changchun-Yanji suture was considered as the eastern

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extension of the Solonker-Xar Moron-Changchun suture by most researchers (e.g., Cao et al.,

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2012, 2013; Yu et al., 2014; Wilde, 2015; Liu et al., 2017a). Nevertheless, some workers suggested that the Changchun-Yanji suture is part of the Jilin-Heilongjiang high-pressure

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metamorphic belt rather than the eastern extension of the Solonker-Xar Moron-Changchun

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suture (Zhou and Wilde, 2013; Zhou et al., 2014; Wang et al., 2016; Zhou and Li, 2017), based on new Mesozoic age data from granitoids in western Yanji which were previously considered as

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Paleozoic (Zhang et al., 2004), Permian ages of the Qinglongcun, Seluohe and Wanbao

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complexes, the confirmation of the Permian Kaishantun accretionary complex in southeastern Yanji (Wu et al., 2003c; Tang et al., 2004), metamorphic age data (ca. 230 Ma) for the Hulan, Seluohe, Qinglongcun and Kaishantun meta-complexes (Zhou et al., 2013), and detrital zircon U-Pb data for the Triassic Dajianggang Formation in central Jilin Province (Wang et al., 2016). The controversy stems from several issues: (1) the Solonker-Xar Moron suture is hidden in the Songliao basin, so how the suture extends through the basin remains unclear; (2) evidence along the Daheishan Horst, a key geological unit that links the eastern Inner Mongolia and Jilin regions, is lacking; and (3) the outcrops are poor and strong offset is observed along several Mesozoic

ACCEPTED MANUSCRIPT NNE strike-slip faults in central-eastern Jilin Province (Fig. 14). Within the Songliao basin, more than 200 drill holes reached the pre-Mesozoic basement rocks, some of these core samples have been dated. Zircon U-Pb ages indicated that the basement rocks can be divided into two groups: a southern group with old ages (1000–1800 Ma and > 2000 Ma)

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and a northern group with relatively young ages (161–181 Ma & 236–368 Ma) (Pei et al., 2007; Han et al., 2009; Liang et al., 2009). The old age group is markedly similar to that in the NCC, so

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the southern section of the Songliao basin and the NCC should share the same basement (Pei et

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al., 2007; Han et al., 2009). Based on these studies, Liu et al. (2017a) located the Solonker-Xar

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Moron-Changchun suture along the northern side of the old basement drilling locations within the Songliao basin (Fig. 2a).

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In terms of location, the present study area is much closer to the Songliao-Xilinhot block than the Jiamusi block (Fig. 2a). In addition, the Late Permian Daheishan volcanic rocks and

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Doushantouzi pluton show features of subduction related igneous rocks as discussed in Section 6.2. These rocks, together with many recently recognized subduction related igneous rocks (Cao

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et al., 2011, 2012, 2013; Yu et al., 2014; Wang et al., 2015a; Liu et al., 2016), form a Permian magmatic belt most probably produced by the subduction of the Paleo-Asian Ocean (Fig. 14).

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Moreover, our zircon U-Pb dating results indicate that the Late Permian igneous rocks in this study underwent overprinting of Early Triassic post-magmatic thermal events, the ages of which are consistent with those of metamorphic events (ca. 250 Ma) reported from the adjacent areas (Hu et al., 2003; Zhang et al., 2005; Wu et al., 2007; Liu et al., 2017b), older than the age (ca. 230 Ma) of the Jilin-Heilongjiang high-pressure metamorphic belt (Zhou et al., 2013). Considering the above features, we suggest that the Solonker-Xar Moron-Changchun suture passes from west to east through the Songliao basin and Daheishan Horst and extend to central

ACCEPTED MANUSCRIPT Jilin and the Yanbian area. This conclusion is also supported by more recent studies of the Dayushan syn-collisional granite (Sun et al., 2004) and Sandaohe post-collision extensional syenogranite (Sun et al., 2005), sedimentary and paleogeography data (Yang et al., 2006b; Qu et al., 2013), paleontological information (Li et al., 2011), detrital zircon ages of Permian sediments

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(Sun et al., 2013; Wang et al., 2015b; Han et al., 2016; Zhou et al., 2017b) and the recognition of

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Shitoukoumen (east of Changchun) (Zhou et al., 2013).

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two type of high-pressure metamorphic rocks in Yantongshan (north of Panshi) and

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6.3.2. Closure of the eastern segment of Paleo-Asian Ocean

The timing of the final closure of the Paleo-Asian Ocean along the Solonker-Xar

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Moron-Changchun-Yanji suture has been a controversial issue and the various proposals include Middle Devonian to Late Carboniferous or earlier (Tang, 1990; Xu and Chen, 1997; Zhao et al.,

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2013; Xu et al., 2015; Chen et al., 2016), late Early Permian (Feng et al., 2010), or Middle Mesozoic (Nozaka and Liu, 2002). However, several other studies have suggested that the final

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closure of the Paleo-Asian Ocean probably occurred during the Late Permian–Late Triassic (Sengör et al., 1993; Xiao et al., 2003, 2015; Sun et al., 2004; Li, 2006; Wu et al., 2007, 2011;

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Cao et al., 2012, 2013; Eizenhöfer et al., 2014; Han et al., 2016, 2017; Wang et al., 2015a; Du et al., 2017; Li et al., 2017a, b; Wang et al., 2017; Zhou et al., 2017b). Abundant new geological, geochronological, geochemical and paleomagnetic data that have been obtained over the past decade provide convincing evidence concerning the tectonic evolution of the middle segment of the Paleo-Asian Ocean along the Solonker suture during the Permian–Triassic, and provide comparisons and references for the tectonic evolution of the eastern segment of the Paleo-Asian Ocean.

ACCEPTED MANUSCRIPT Based on information from sedimentary sequences, spatial and temporal distribution of ophiolites, evolution of paleobiogeographical realms and magmatic zonal patterns (Li, 2006) and some syn-collisional granites (e.g., Shuangjingzi, Halatu and Dayushan plutons) along the Solonker-Xar Moron-Changchun-Yanji suture, Li et al. (2007b) proposed that the Paleo-Asian

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Ocean finally closed during the Late Permian–Middle Triassic, leading to the formation of the Solonker-Xar Moron-Changchun-Yanji suture zone.

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The detrital zircon U-Pb dating of a series of Permian–Triassic strata (e.g., Huanggangling,

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Zhesi, Linxi Formations, etc.) in southern Inner Mongolia and provenance studies indicated that

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the middle section of the Paleo-Asian Ocean did not close until the Early Triassic (Han et al., 2015; Eizenhöfer et al., 2014; Li et al., 2014a; Li et al., 2017b; Wang et al., 2017). This opinion

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is also supported by Early Permian paleomagnetic data from the Siberian and Sino-Korean paleoplates, as summarized by Li (2006), and rock magnetic and paleomagnetic studies of

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Permian sandstone from the Taohaiyingzi area in Inner Mongolia (Qin et al., 2013). The identification of Late Permian radiolarians in cherts from the Xingshuwa ophiolite belt on

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the northern side of the Xar Moron River (Wang and Fan, 1997) and the presence of a pillow basalt with an age of 277 Ma in the Wudaoshimen ophiolite (Wang et al., 2014) suggest that an

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ocean that extended along the Linxi ophiolite belt did not close until the Late Permian. In addition, the mixing of Cathaysian and Angaran flora appeared in the upper Permian strata in the Ondor Sum subduction-accretion complex (Sun et al., 2013), which is not evident in the lower Permian strata (Wilde, 2015) and suggests that the ocean in this area probably closed during the Late Permian or later. Recently, a Late Carboniferous–Middle Permian E-W-trending calk-alkaline magmatic belt was identified along the northern margin of the NCC in Inner Mongolia, and the plutons within

ACCEPTED MANUSCRIPT the belt were proposed to be the products of an Andean-type continental margin arc that was related to the subduction of the Paleo-Asian oceanic plate (Wang et al., 2009; Zhang et al., 2010; Liu et al., 2012, 2013; Ma et al., 2013; Li et al., 2017a). Furthermore, Chen et al. (2009) proposed that the final suturing of the Solonker zone occurred from 296 to 234 Ma based on

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geochronological and geochemical studies on the Baolidao arc rocks. Moreover, Li et al. (2014b) documented the Xilingol Complex along the southern margin of the Xilinhot block and

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suggested that the timing of the suturing along the Solonker suture belt can be constrained to be

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between 269 and 231 Ma. All the above-mentioned studies strongly suggest that the final

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collision along the Solonker suture zone occurred during the Late Permian–Early Triassic (Xiao et al., 2003, 2015; Li et al., 2017a; Liu et al., 2017a; Wang et al., 2017).

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As discussed earlier, the Daheishan volcanic rocks predominantly consist of rhyolitic, rhyodacitic, and andesitic rocks with minor basaltic rocks. The lithological assemblage, high-Al

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basalt affinity of the basaltic rocks, arc-like trace element features of both the basaltic and intermediate–felsic rocks, and their low initial

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Sr/86Sr ratios, positive εNd(t) and zircon εHf(t)

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values indicate that the Late Permian igneous rocks in the study area may have formed in an active continental margin setting (Gill, 1981; Pearce, 1982; Grove and Kinzler, 1986; Grove et

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al., 2003; Yu et al., 2014).

The zircon U-Pb data in this study indicate continuous and multiple magmatic events during the Permian in the Daheishan Horst, which connects the Solonker-Xar Moron-Changchun suture and Changchun-Yanji suture. Similar Permian arc magmatism has been recorded along the Changchun-Yanji suture belt (Fig. 14), thus confirming extensive Permian subduction-related magmatic activity that was associated with the evolution of the Paleo-Asian Ocean. In this study, immobile elements were used to identify the tectonic setting of the Daheishan basaltic rocks and

ACCEPTED MANUSCRIPT associated rocks along the Changchun-Yanji suture from the literature by utilizing a series of discriminant diagrams. All the Permian basaltic rocks plot within the volcanic arc basalt/active continental margin basalt field in the Th-Hf-Ta, La/Nb-La, and Th/Yb-Nb/Yb discrimination diagrams (Fig. 17). Moreover, almost all the Permian intermediate–felsic rocks plot within the

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typical arc-rock field in the (La/Yb)N-(Yb)N diagram, distinct from Triassic syn-collisional granites with adakitic affinities (Fig. 15b). An active continental margin setting that was related

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to the subduction of the Paleo-Asian oceanic plate likely created the above-mentioned Permian

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igneous rocks. Therefore, the eastern segment of the Paleo-Asian Ocean was undergoing

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subduction during the Late Permian, and the collision between the combined NE China blocks and NCC occurred after this time. This conclusion is further supported by the following lines of

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evidence.

First, recently reported geochronological data for detrital zircons from Permian–Triassic

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sedimentary rocks along the Changchun-Yanji suture zone provide important constraints on the Permian–Triassic tectonic evolution of this region (Wang et al., 2015b; Han et al., 2016; Zhou et

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al., 2017b). Wang et al. (2015b) and Zhou et al. (2017b) conducted provenance studies of the Permian–Triassic sedimentary units (such as Shoushangou, Fanjiatun, Yangjiagou, Dasuangou,

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Kedao Formations, etc.) with detrital zircon U-Pb dating and Hf isotope methods and proposed that the initiation of the final closure of the Paleo-Asian Ocean along the Changchun-Yanji suture occurred during the end of the Permian at the latest and that the subsequent collisional orogeny terminated at ~233 Ma. Han et al. (2016) concluded that the Paleo-Asian Ocean probably closed between 250 and 222 Ma in the Huadian area in central Jilin Province based on detrital zircon U-Pb geochronological and whole-rock geochemical studies of metamorphic sandstones from the Dongnancha Formation.

ACCEPTED MANUSCRIPT Second, many Triassic syn-collisional granites (such as the Jianpingzhen, Dayushan, Liushugou plutons, etc.) have been recognized in central-eastern Jilin Province (Sun et al., 2004, 2005; Zhang et al., 2004; Cao et al., 2013; Wang et al., 2015a). These granites constitute an Early–Middle Triassic syn-orogenic granitic belt along the Changchun-Yanji suture (Fig. 14),

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which implies that the eastern segment of the Paleo-Asian Ocean should have been closed during the Early–Middle Triassic. The Early Triassic hydrothermal/metamorphic ages from the igneous

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rocks in this study and the ca. 250 Ma (whole-rock Rb-Sr age) metamorphic age from the Hulan

some

collision-related

metamorphism

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Third,

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Group (Wu et al., 2007) probably correspond to the above-mentioned collisional events. of

metamorphic

Changchun-Yanji suture has been elucidated from recently reported

40

rocks

along

the

Ar/39Ar ages, which are

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commonly used to constrain the cessation of events (Reichow et al., 2009). The Ar-Ar dating of the Hulan complex produced a phengite age of 229 ± 5 Ma for a two-mica schist and a biotite

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age of 224 ± 0.8 Ma for a garnet-biotite gneiss in the Hongqiling area (Xi et al., 2003, 2006). Li et al. (2010b) reported hornblende age of 228 ± 3 Ma from an amphibolite at Hongqiling. In

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addition, the youngest zircon U-Pb age of 239 ± 11 Ma (Zhang et al., 2009) in the silicolite of the Yantongshan Complex provides a Middle Triassic upper age limit for the metamorphism of this

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complex. Based on the above data, Zhou and Wilde (2013) concluded that the phengite Ar-Ar age of 229 ± 5 Ma from Xi et al. (2003) probably records the termination of metamorphism in this area and that the final collisional events in this region occurred during the Middle Triassic at ~230 Ma. Fourth, paleontological data indicate that the boundary between Angara and Cathaysia flora along Wangqing-Hunchun in the Yanbian area should mark the closure position of the easternmost Paleo-Asian Ocean (Sun et al., 2013). Fossil data suggested that some floral mixing

ACCEPTED MANUSCRIPT had begun during the Late Permian (Lopingian), which implies that the Jiamusi-Khanka blocks and the NCC probably collided during the Late Permian or later (Sun et al., 2013; Du et al., 2017). Thus, we suggest that the final closure of the Paleo-Asian Ocean along the Changchun-Yanji

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suture occurred during the Late Permian–Middle Triassic. However, recent studies have strongly indicated that the Paleo-Asian Ocean completely closed during the Late Carboniferous to Early

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Permian along the northern margin of the Tarim Craton (Han et al., 2011; Huang et al., 2012; Liu

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et al., 2014; Alexeiev et al., 2015; Ge et al., 2015; Safonova et al., 2016b; Cheng et al., 2017;

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Han and Zhao, 2017; Hu et al., 2017; Xiao et al., 2017; Zhong et al., 2017), at the latest Early Permian to Middle Permian in the Alxa Terrane (Zhang et al., 2013; Liu et al., 2017d, e), and

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during the Late Permian to Early Triassic along the Solonker suture zone (Xiao et al., 2003, 2015; Li, 2006, 2017a; Liu et al., 2017a). We conclude therefore the final closure of the Paleo-Asian

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Ocean might have taken place along the northern margin of the Tarim-Sino Korean cratons in a scissors style, and the closure time of the western-middle segment is earlier than that of the

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eastern segment.

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6.3.3. Permian subduction polarity of the Paleo-Asian oceanic plate

Similar to the Solonker-Xar Moron-Changchun suture, a Late Paleozoic subduction-related magmatic belt developed along the Changchun-Yanji suture, as stated above. Recently, Pei et al. (2016) studied geochemical, zircon U-Pb and Hf isotopic data for some Early–Middle Paleozoic igneous rocks in the central Jilin Province and divided the Early–Middle Paleozoic magmatism in this region into four episodes, namely, Late Cambrian, Middle Ordovician, Late Ordovician–Early Silurian and Late Silurian–Middle Devonian, and proposed a model of Late

ACCEPTED MANUSCRIPT Cambrian–Early Silurian subduction followed by an arc-continental collision during the Late Silurian–Early Devonian, which can be compared with the tectonic evolution along the Solonker suture proposed by Jian et al. (2008). This observation and the widely distributed Late Paleozoic subduction-related magmatism along the Changchun-Yanji suture suggest that the wide region

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from Inner Mongolia to eastern Jilin along the Sonlonker-Xar Moron-Changchun-Yanji suture records a common geodynamic setting.

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Several early papers summarized available data from geological surveys of Inner Mongolia

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and argued that a bidirectional subduction of the Paleo-Asian oceanic plate caused the two

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opposing active continental margins to collide, which created the Solonker suture during the Late Permian–Middle Triassic (Xiao et al., 2003, 2015; Li, 2006; Chen et al., 2009; Liu et al., 2017a).

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Based on a comprehensive analysis of geological, paleobiogeographical, paleomagnetic and geochemical data, Li (2006) proposed that the bidirectional subduction model was also

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reasonable for the tectonic evolution of the Changchun-Yanji suture. However, whether the igneous rocks in the study area were related to the southward or northward subduction of the

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Paleo-Asian oceanic plate remains unclear. We argue that the northward subduction is more suitable based on the following discussion.

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In the paleogeographic and palinspastic reconstruction by Wen et al. (1996), Jilin Province was treated as four separate terranes (from west to east, the Baicheng, Songliao, Jizhong and Yanbian terranes) (not shown) based on biostratigraphic, sedimentological, paleomagnetic and geochemical studies. During the Permian, the Songliao terrane was characterized by Angaran flora and a temperate-frigid zone biota, whereas the Jizhong terrane was characterized by the Tethys bio-province. The distributions of the fauna and flora indicate paleogeographic affinities among the geological terranes and suggest that the Songliao terrane, which contains the studied

ACCEPTED MANUSCRIPT igneous rocks, was close to the northern Songliao-Xilinhot block during the Permian (Wen et al., 1996). Moreover, Liu et al. (2017a) proposed that the Changchun-Yanji suture, which has been cut and strongly deformed by Mesozoic thrusting and strike slipping, extends from the western Changchun through Panshi, Huadian, and Dunhua, to eastern Yanji and that our study area lies to

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the north of the suture (Fig. 14), which suggests a close affiliation with the Songliao-Xilinhot block to the north.

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The above discussion suggests that the Late Permian igneous rocks in the study area might

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have formed in the active continental margin of the Songliao-Xilinhot block, which was

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associated with the northward subduction of the Paleo-Asian oceanic plate. Recent studies of some Permian igneous rocks in the central Jilin and Yanbian areas also support this conclusion

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(Wen et al., 1996; Cao et al., 2011, 2012; Yu et al., 2014; Du et al., 2017). Wen et al. (1996) suggested that the lower Permian Daheshen Formation was an exotic fragment that was related

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to the Jiamusi-Khanka blocks according to the angular unconformable contact between the Daheshen Formation and the underlying tectonic mélange, an analysis of the sedimentary facies

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and the presence of Angara flora. Additionally, the volcanic rocks from the Daheshen Formation were suggested to have formed in an active continental margin setting (Cao et al., 2012; Yu et al.,

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2014), so these volcanic rocks were likely formed by the northward subduction of the Paleo-Asian oceanic plate. In addition, the Qianshan gabbro, Wudaogou diorite, Guanmenzuizi basalts (Cao et al., 2011, 2012), and Shangsidaogou basalts (Du et al., 2017) in the Yanbian area, which have Permian ages, were regarded as products in an active continental margin setting that was related to the northward subduction of the Paleo-Asian oceanic plate beneath the Jiamusi-Khanka blocks. The above-mentioned igneous rocks and the upper Permian igneous rocks in this study form

ACCEPTED MANUSCRIPT an approximately E-W-trending Permian epicontinental arc-magmatic belt to the north of the Changchun-Yanji suture (Fig. 14), suggesting the northward subduction of the Paleo-Asian oceanic plate beneath the combined NE China block, as illustrated in Fig. 18.

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7. Conclusions

Zircon U-Pb geochronological and Hf isotopic analyses and whole-rock major-trace element

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and Sr-Nd isotopic analyses on igneous rocks from the Daheishan Horst, together with a review

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of the recent data and models allow us the following salient conclusions.

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1. The zircon U-Pb data and zircon REE patterns indicate that continuous and multiple magmatism occurred in the study area during the Permian and that the Daheishan volcanic rocks

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and Doushantouzi pluton formed during the Late Permian (ca. 253–256 Ma) and underwent late-stage alteration during Early Triassic tectono-thermal events. These rocks should be the

Daheishan Horst.

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residual geologic units of the Solonker-Xar Moron-Changchun-Yanji suture zone in the

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2. The primary magma of the Daheishan basaltic lavas was derived from the partial melting of a depleted mantle wedge that had been metasomatized by fluids from a subducted slab.

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3. The Daheishan intermediate-felsic rocks and Doushantouzi syenogranites were the products of the partial melting of a Meso-Neoproterozoic juvenile basaltic lower crust. However, the andesitic magmas were probably contaminated by some old crustal materials, and a certain amount of carbonatite materials and metasomatized mantle melts were probably involved in the SD andesitic magmas. 4. The Late Permian igneous rocks in this study, which formed in an active continental margin setting, and the previously reported Permian subduction-related volcanic rocks, form an

ACCEPTED MANUSCRIPT approximately E-W-trending Permian epicontinental arc-magmatic belt to the north of the Changchun-Yanji suture, suggesting that the northward subduction of the Paleo-Asian oceanic plate beneath the combined NE China block and indicates that the eastern segment of the Paleo-Asian Ocean was not closed before the Late Permian.

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Acknowledgments

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We are grateful to Gondwana Research Associate Editor Inna Safonova, and the anonymous

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referees for their constructive comments which greatly improved this paper. We thank Prof M.

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Santosh for proof reading the English. We thank the staff of the State Key Laboratory of Continental Dynamics, Northwest University, Xi’an, China, for their advice and assistance

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during zircon U-Pb and Hf-isotopic dating. We appreciate the Supervision and Inspection Center of Mineral Resources, the Ministry of Land and Resources of Jinan, China, for their assistance in

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the major and trace element analysis. We are also grateful to the Institute of Geochemistry, Chinese Academy of Science (Guangzhou), China, for help with collecting the whole rock Sr-Nd

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isotopic data. This work was financially supported by the National Natural Science Foundation of China (grants: 41372108 and 41702131), SDUST Research Fund (grant: 2015TDJH101),

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Major Scientific and Technological Innovation Projects of Shandong province (grants: 2017CXGC1602 and 2017CXGC1603), and Natural Science Foundation of Shandong province (grant: ZR201702150228).

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Figure and table captions

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Fig.1. Schematic tectonic map that shows the main subdivisions of central and eastern Asia and

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the location of the study area (modified from Zhou et al., 2014).

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Fig.2. (a) Tectonic sketch map of NE China, modified after Liu et al. (2017a). (b) Tectonic

area in central Jilin Province.

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sketch map of the Daheishan Horst. (c) Detailed geological map of the Fangniugou-Jingtai

Fig.3. Representative field photographs of rocks from the Daheishan volcanic rocks and

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Doushantouzi pluton. (a) Foliated rhyodacites DYS1 and DYS2. (b) Andesite SD7 carrying carbonatite enclaves. (c) Foliated rhyolite XQ1. (d) Basaltic lava XDC1. (e) Foliated andesite

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XDM1. (f) Mylonitized syenogranite DST1. Fig.4. Photomicrographs of selected samples from the Daheishan volcanic rocks (a-f) and

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Doushantouzi pluton (g-h) in the Daheishan Horst, Jilin Province (all the photos were taken under cross-polarized light). (a) DYS1, rhyodacite; (b) DYS2, rhyodacite; (c) SD7, andesite; (d) XQ1, rhyolite; (e) XDC1, basaltic lava; (f) XDM1, andesite; (g) and (h) DST1, syenogranite. Af-alkali feldspar. Hb-hornblende. Or-orthoclase. Pl-plagioclase. Pt-perthite. Q-quartz. Fig.5. Representative cathodoluminescence (CL) images of zircons from the Daheishan volcanic rocks. The solid-line and dash-line circles represent the locations for U-Pb dating and in-situ

ACCEPTED MANUSCRIPT Lu-Hf isotope analysis, respectively. The numbers below the images show the ages of the zircons and the εHf(t) values. C-core. R-rim. TZR-transgressive zone of recrystallization. Fig.6. Representative cathodoluminescence (CL) images of zircons from the sample DST1 from the Doushantouzi pluton. The solid-line and dash-line circles represent the locations for U-Pb

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dating and in-situ Lu-Hf isotope analysis, respectively. The numbers below the images show the ages of the zircons and the εHf(t) values. C-core. R-rim. TZR-transgressive zone of

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recrystallization.

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Fig.7. LA-ICP-MS U-Pb concordia diagrams of zircons from the igneous rocks in this study. The

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weighted mean age and MSWD are shown in each figure. Fig.8. Histogram of the age spectra of all the analysed zircons from the igneous rocks in this

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study.

Fig.9. Chondrite-normalized REE patterns of zircons from the igneous rocks in this study.

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Fig.10. Rock classification diagrams of (a) Nb/Y versus SiO2 (after Winchester and Floyd, 1977) and (b) Co versus Th (after Hastie et al., 2007). The literature data of Permian volcanic rocks

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along the Changchun-Yanji suture are from Li et al. (2007a), Cao et al. (2012) and Yu et al. (2014). B, basalt; BA/A, basaltic andesite/andesite; D/R, dacite/rhyolite; IAT, island arc

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tholeiitic; CA, calc-alkaline; H-K/SHO, high potassic/shoshonitic. Fig.11. Chondrite-normalized REE patterns and primitive mantle (PM)-normalized trace element patterns of the Daheishan volcanic rocks (a, b) and the Doushantouzi syenogranites (c, d). The data for the Andes-type Daheshen volcanic rocks are from Cao et al. (2012) and Yu et al. (2014). The chondrite and primitive mantle values are from Boynton (1984) and Sun and McDonough (1989), respectively. Fig.12. Harker variation diagrams for major oxides vs. the SiO2 content for the Daheishan

ACCEPTED MANUSCRIPT intermediate-felsic rocks. Fig.13. εHf(t) versus t diagram (a) and εNd(t) versus (87Sr/86Sr)i diagram (b) of the igneous rocks in this study. CAOB: Central Asian Orogenic Belt; YFTB: Yanshan Fold and Thrust Belt (Yang et al., 2006a).

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Fig.14. Simplified geological map of central-eastern Jilin, which shows the localities of the Permian–Triassic igneous rocks (geochronological data see Table 1 for details) along the

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Changchun-Yanji suture (modified after Wang et al., 2015a).

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Fig.15. (a) Chemical variation diagram for the Daheishan volcanic rocks. (b) (La/Yb) N versus

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YbN diagram (Martin, 1986). The literature data are from Sun et al. (2004), Cao et al. (2012, 2013), Yu et al. (2014), Wang et al. (2015a) and Liu et al. (2016). (c) MgO versus SiO2

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diagram (Martin et al., 2005) for the Daheishan intermediate-felsic rocks and Doushantouzi syenogranites. PMB, experimental partial melts derived from basalts or amphibolites; LSA,

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low silica adakite; HAS, high silica adakite. Fig.16. εNd(t) versus (87Sr/86Sr)i diagram for the intermediate-felsic samples in this study. The

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numbers indicate the percentages of participation of the crustal materials. The data for the Phanerozoic intermediate-felsic igneous rocks in NE China are from Wu et al. (2000, 2002,

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2003a, 2003b, 2007), Liu et al. (2010b), Yu et al. (2014), and Du et al. (2017). The calculated parameters for Sr (ppm), (87Sr/86Sr)i, Nd (ppm) and εNd(t) are 20, 0.703, 1.2 , and +8 from asthenospheric mantle (DM); 200, 0.704, 15, and +8 for basalt; 250, 0.740, 30, and -12 for upper continental crust (UCC); and 230, 0.708, 20, and -15 for lower continental crust (LCC). All the data are from Wu et al. (2000). Fig.17. Trace element discrimination diagrams for the tectonic setting of the Permian basaltic rocks along the Changchun-Yanji suture. (a) Th-Hf-Ta diagram (Wood, 1980). IAT, island arc

ACCEPTED MANUSCRIPT tholeiite; CAB, calc-alkaline basalt; MORB, middle ocean ridge basalt; WPT, within-plate tholeiite; WPAB, within-plate alkali basalt. (b) La/Nb versus La diagram (Li, 1993). IAB, island arc basalt; OIB, ocean island basalt. (c) Th/Yb versus Nb/Yb diagram (Pearce and Peate, 1995). Literature data are from Cao et al. (2012), Yu et al. (2014) and Du et al. (2017).

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Fig.18. Simplified cartoon model that shows the Permian volcanism at the southern margin of the combined NE China blocks. PAO, Paleo-Asian Ocean; CNECB, combined NE China blocks;

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CC, continental crust; SCLM, subcontinental lithospheric mantle.

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Table 1

Moron-Changchun-Yanji suture zone.

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Table 2

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Geochronological data for the Permian-Triassic igneous rocks along the Solonker-Xar

Major (wt.%) and trace (ppm) elements of the igneous rocks in this study.

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Note: LOI: Loss on ignition; A/CNK = mole [Al2O3/(Cao + Na2O + K2O)]; A/NK = mole [Al2O3/(Na2O + K2O)]; δEu = 2*EuN/(GdN + SmN); LREE = La + Ce + Pr + Nd + Sm + Eu;

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HREE = Gd + Tb + Dy + Ho + Er + Tm + Yb + Lu; ∑REE = ∑LREE + ∑HREE; (La/Yb)N =

Table 3

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(La/0.310)/(Yb/0.209).

Sr-Nd isotopic data of the igneous rocks in this study. Note: The 87Rb/86Sr and 147Sm/144Nd ratios were calculated using the Rb, Sr, Sm and Nd contents. The εNd(t) values were calculated using the present-day (147Sm/144Nd)CHUR = 0.1967 and (143Nd/144Nd)CHUR = 0.512638 values (Lugmair and Marti, 1978). The TDM values were calculated using the present-day (147Sm/144Nd)DM = 0.2136 and (143Nd/144Nd)DM = 0.513151 (Miller and O’Nions, 1985).

ACCEPTED MANUSCRIPT Table 1 GPS

Location

RI

SC

NU

MA

43°00′54.6″, 127°09′44.8″ Seluohe

13JH1-1

CE

AC

252 ± 5

10

43°05′57″, 125°50′52″

Dahongshilazi

LA-ICP-MS 260 ± 3

11

42°54′26″, 125°50′28″ 43°02′31″, 125°06′22″

Qingyang Anyi

LA-ICP-MS LA-ICP-MS SHRIMP TIMS SHRIMP SHRIMP SHRIMP SHRIMP

259 ± 3 252 ± 2 264 ± 5 267 ± 1 243 ± 5 242 ± 7 243 ± 5 254 ± 8

11 11

SHRIMP

261 ± 20

11

LA-ICP-MS LA-ICP-MS LA-ICP-MS LA-ICP-MS

279 ± 3 293 ± 2 286 ± 2 249 ± 1

12 12 12 13

LA-ICP-MS 270 ± 1

13

LA-ICP-MS SIMS LA-ICP-MS SIMS LA-ICP-MS LA-ICP-MS LA-ICP-MS LA-ICP-MS LA-ICP-MS

255 ± 2 257 ± 2 259 ± 2 258 ± 2 299 ± 2 299 ± 2 300 ± 2 300 ± 2 260 ± 1

13 13 13 13 14 14 14 14 15

LA-ICP-MS 252 ± 1

15

D

11

Granodiorite 42°22′19″, 124°55′44″ Songshuzui Monzogranite 42°32′52″, 125°11′14″ Xiaosiping Granodiorite 42°25′11″, 124°40′46″ Jianshanzi Monzogranite 42°17′19″, 124°45′34″ Jianshanzi Gt-Mus 42°16′10″, 124°46′52″ Fangniugou monzogranite Rhyolite Daheshen Metadacite Daheshen Rhyolite Daheshen Biotite Jianpingzhen monzogranite Garnet-bearing Fangniugou monzogranite Granodorite Sanmenmojia Monzogabbro Fangshendingzi Biotite syenogranite Youyi Olivine gabbro Suangfengshan Rhyolite 43°22′16.3″, 126°43′51.1″ Daheshen Basalt 43°22′09.8″, 126°44′05.3″ Daheshen Rhyolitic tuff 43°21′20.7″, 126°44′54.1″ Daheshen Rhyolite 43°21′16.3″, 126°44′58.0″ Daheshen Biotite 43°23′09″, 126°01′17″ Xiaolihe monzogranite Quartz monzonite 43°26′42″, 125°59′52.7″ Shanhe

LK24-1 LK25-9 LK36-1 LK37-1 DH008 DH012 DH037 DH041 12JL1-1

SHRIMP

1 2 3 3 3 4 4 4 4 4 4 4 4 5 5 6 7 7 7 7 8 9

LA-ICP-MS 251 ± 2

MG-28 MG-64 MG-103 MG-119

LK12-1

SHRIMP 276 ± 2 SHRIMP 277 ± 4 SHRIMP 279 ± 10 SHRIMP 260 ± 12 SHRIMP 256 ± 3 SHRIMP 252.2 ± 1.7 SHRIMP 251.8 ± 1.1 SHRIMP 291.8 ± 2.3 SHRIMP 250.2 ± 2.4 SHRIMP 252.5 ± 2.3 SHRIMP 274.4 ± 2.5 SHRIMP 284 ± 4 SHRIMP 288 ± 6 SHRIMP 278.5 ± 3 SHRIMP 273.7 ± 1 LA-ICP-MS 287.4 ± 1.7 LA-ICP-MS 263–237 LA-ICP-MS 306–253 LA-ICP-MS 299–245 LA-ICP-MS 246 ± 3 LA-ICP-MS 277 ± 3 LA-ICP-MS 280 ± 3

Dahongshilazi

Quartz syenite

JH4-1 JK11-1 JK11-4 LK04-2

Age (Ma) Ref.

43°06′40″, 125°57′29″

FW00-121

MG-143

Method

PT

43°52.4′, 116°11.2′ Xilinhot 43°06′11.7″, 117°12′46.1″ Kedanshan Solon Obo Ondor Sum Banlashan Solon Obo Solon Obo Mandula Mandula Mandula Mandula Chaganhadamiao Chaganhadamiao Hujierte Chaganhadamiao 44°15′40″, 116°56′42″ Maodeng Haerhada Haerhada Baiyinnuoer Baiyinnuoer 43°35′27″, 117°15′58″ Wudaoshimen Xingshuwa

PT E

Sample no. Lithology Solonker-Linxi area A-type granite Plagiogranite 2002MWZ04 Cumulate gabbro 2002MW-03 Pillow lava 2002ML06-4 Cumulate gabbro MSL13-1 Anorthosite MSL13-2 Diorite SLS01 Gabbro-diorite SLS03 Andesite SLS04 Diabase SLS06 Diabase CGHDM01 Gabbro CGHDM02 Plagiogranite NM08-188 Gabbro NM08-73 Basalt GSY47 Keratophyre WD03-1 Meta-pillow basalt WD04-2 Meta-basalt WD05-1 Meta-basalt WD05-3 Meta-basalt WD1019-01 Pillow basalt 13LX34 Rodingite Central Jilin Province JHS3-1 High-Mg andesite Alkali feldspar DY123-4 granite Alkali feldspar DYS124-2 granite DY126-2 Syenogranite DY143-2 Monzogranite

42°58′11″, 125°51′48″

Dakangshan

11 11 11 11 11

ACCEPTED MANUSCRIPT Hornblende gabbro 43°59′42″, 127°20′48″ Tuding LA-ICP-MS 252 ± 1 15 Mylonitized JK13-1 42°58′59.2″, 127°41′31.1″ Huangniling LA-ICP-MS 252 ± 1 15 monzogranite KY12-33-4 Andesite Kaiyuan SIMS ca. 249 16 KY13-12-4 Basaltic andesite Kaiyuan SIMS 250 ± 4 16 The Yanbian area YZ02-2 Monzogranite 42°52′15″, 128°30′32″ Dakai LA-ICP-MS 249 ± 4 17 YZ02-12-3 Tonalite 42°28′46″, 128°56′12″ Bailiping LA-ICP-MS 285 ± 9 17 YZ02-22-2 Monzogranite 42°12′14″, 128°49′21″ Bailiping LA-ICP-MS 245 ± 6 17 YZ02-25-2 Monzogranite 42°10′57″, 128°44′58″ Bailiping LA-ICP-MS 245 ± 3 17 YZ02-27-2 Monzogranite 42°03′10″, 128°49′32″ Bailiping LA-ICP-MS 248 ± 2 17 JXNC-1-2 Gabbro Wudaogou SHRIMP 270 ± 10 18 D0832-6 High-Mg diorite 43°15′27″, 130°59′51″ Taipinggou LA-ICP-MS 241 ± 1 19 D0849-1 High-Mg diorite 42°12′52″, 130°53′17″ Xiaoxinancha LA-ICP-MS 240 ± 1 19 SMZG-1 Granodiorite 43°05′24″, 127°52′42″ Liukesong LA-ICP-MS 263 ± 1 20 SMZM-1 Gabbro 43°05′24″, 127°52′42″ Liukesong LA-ICP-MS 262 ± 1 20 YH7-1 Diorite Qianshan LA-ICP-MS 255 ± 3 21 YH8-1 Gabbro Wudaogou LA-ICP-MS 282 ± 2 21 FW00-34 Diorite 43°08′45″, 127°53′47″ Liukesong SHRIMP 247 ± 1 11 FW00-45 Quartz diorite 42°54′19″, 128°23′19″ Xiaobutun LA-ICP-MS 260 ± 2 11 Alkali feldspar FW00-104 43°01′30″, 129°28′24″ Shiren LA-ICP-MS 253 ± 2 11 granite FW00-110 Monzogranite 43°07′33″, 128°51′02″ Liangbing TIMS 247 ± 1 11 FW00-122 Monzogranite 43°18′00″, 128°56′48″ Xibeicha LA-ICP-MS 240 ± 2 11 Alkali feldspar FW02-191 43°06′43″, 128°00′53″ Erhedian LA-ICP-MS 253 ± 2 11 granite FW02-191 Quartz diorite 43°26′46″, 128°27′16″ Liushuhe LA-ICP-MS 253 ± 2 11 YZ02-40 Granodiorite 42°11′05″, 129°09′22″ Liudong LA-ICP-MS 246 ± 3 11 YZ02-49 Quartz diorite 42°39′39″, 129°30′00″ Longyandong LA-ICP-MS 251 ± 2 11 YH2-1 Basalt Tumen LA-ICP-MS 250 ± 5 12 YH15-1 Basaltic andesite Guanmenzuizi LA-ICP-MS 275 ± 5 12 13YB3-2 Quartz monzonite 42°46′13.2″, 128°52′20″ Mengshanbeidong LA-ICP-MS 249 ± 1 15 13JH28-1 Biotite 42°45′59.2″, 128°57′31.1″ Liushugou LA-ICP-MS 245 ± 1 15 monzogranite SSD-1 Basalt Shangsidaogou LA-ICP-MS 261.8 ± 3.9 22 HD-1 Basalt Houdidong LA-ICP-MS 252 ± 3 22 References are as follows: 1. Shi et al. 2004; 2. Jian et al. 2007; 3. Miao et al. 2007; 4. Jian et al. 2010; 5. Chen et al., 2012; 6. Cheng et al. 2013; 7. Chu et al. 2013; 8. Wang et al. 2014; 9. Song et al. 2015; 10. Li et al. 2007a; 11. Wu et al. 2011; 12. Cao et al. 2012; 13. Cao et al. 2013; 14. Yu et al. 2014; 15. Wang et al. 2015a; 16. Yuan et al. 2016; 17. Zhang et al. 2004; 18. Zhao et al. 2008; 19. Fu et al. 2010; 20. Liu et al. 2010b. 21. Cao et al. 2011; 22. Du et al. 2017.

AC

CE

PT E

D

MA

NU

SC

RI

PT

13JH3-1

ACCEPTED MANUSCRIPT XD SD7 SD8 SD9 XQ1 XQ2 XQ3 DY DY DY DY DST DST DST M3 S1 S2 S3 S4 1 1 1 quartz

CE

AC

70. 08 0.3 7 15. 19 2.9 4 0.0 6 0.3 6 1.6 3 4.1 3 3.2 0 0.0 8 2.2 9 100 .33

76. 18 0.2 6 10. 67 2.2 4 0.1 3 0.2 5 2.3 6 3.3 4 2.4 4 0.0 7 2.5 4 100 .48

66. 64 0.6 3 15. 36 3.8 0 0.0 9 0.8 3 1.5 5 5.2 0 3.1 5 0.1 5 2.3 9 99. 79

68. 66 0.6 5 14. 12 3.6 2 0.0 8 0.7 4 1.9 8 4.0 0 3.1 8 0.1 8 2.6 3 99. 84

PT

73. 60 0.3 0 11. 36 2.5 0 0.1 5 0.2 8 2.7 5 3.4 0 2.8 2 0.0 8 2.9 1 100 .15

RI

57. 40 0.9 9 15. 66 7.3 9 0.1 3 3.1 7 7.8 2 3.0 1 1.1 8 0.3 9 3.9 3 101 .07

NU

56. 78 1.1 5 16. 89 7.5 0 0.0 9 3.3 0 5.2 8 4.2 8 0.9 1 0.3 7 2.6 9 99. 23

MA

59. 82 1.2 2 15. 67 7.1 3 0.0 9 2.8 3 3.7 6 4.2 1 1.8 1 0.3 8 2.6 7 99. 57

trachyte

syenogranite 66. 14 0.6 5 15. 31 3.8 4 0.1 0 1.1 1 2.8 1 4.9 7 2.8 2 0.1 7 1.6 4 99. 56

66. 86 0.7 3 15. 42 4.0 1 0.0 8 1.2 4 2.5 4 4.9 9 3.0 1 0.1 9 1.5 4 100 .62

30. 3 0.7 6 9.4 0 75. 51 75. 24 11. 30 12. 70 15. 70 57. 90 10. 50 52. 20 708 .10

44

1.8 3 15. 80 130 .50 150 .30 19. 50 34. 20 40. 50 92. 60 19. 10 44. 50 550 .00

71. 37 0.3 9 14. 16 2.9 0 0.0 5 0.7 3 1.4 0 3.4 4 3.8 0 0.0 7 1.5 2 99. 83

73. 38 0.3 7 13. 31 3.0 4 0.0 5 0.7 2 1.4 1 3.8 6 3.3 7 0.0 7 0.9 6 100 .53

72. 18 0.3 3 14. 08 2.8 7 0.0 5 0.6 9 1.4 3 4.0 6 3.4 7 0.0 7 0.8 8 100 .10

1.1 1.0 1.0 5 6 8

D

58. 46 0.6 6 18. 11 7.5 3 0.1 1 1.6 5 5.3 5 1.7 9 3.1 7 0.1 5 2.8 9 99. 87

rhyolite

SC

andesite

PT E

Table 2 Samp XD XDC XD XD XD le C1 2 C3 M1 M2 Lit alkali olo andesite basalt gy Si 44.7 46. 45. 56. 59. O2 3 94 46 68 18 Ti 1.4 1.4 1.1 0.8 1.54 O2 1 3 4 8 Al2 18.3 18. 18. 19. 18. O3 4 66 14 69 32 Fe2 10.1 9.6 9.7 7.8 7.0 O3T 1 2 6 0 0 Mn 0.2 0.1 0.0 0.0 0.26 O 2 4 6 8 Mg 10.4 9.4 9.8 2.5 2.7 O 1 5 7 3 9 Ca 4.6 3.8 2.0 2.4 4.53 O 9 0 6 5 Na2 2.1 2.4 2.4 2.8 2.14 O 9 2 6 7 K2 4.1 4.4 4.1 3.0 4.13 O 2 9 1 6 P2 0.4 0.7 0.2 0.2 0.45 O5 7 5 1 0 LO 2.7 3.4 3.1 2.8 3.56 I 5 3 7 7 Tot 100. 100 99. 99. 99. al 20 .51 70 91 70 A/ CN K A/ NK Mg 66. 66. 39. 44. 67.1 # 1 7 1 1 1.6 1.4 1.8 1.6 Be 1.28 2 1 5 7 21. 12. 12. 11. Sc 9.02 90 90 88 70 207. 218 118 133 125 V 62 .60 .00 .93 .3 50.7 75. 34. 31. 98. Cr 9 89 10 11 93 28.4 38. 14. 15. 16. Co 2 10 20 11 60 56.9 49. 3.5 17. 19. Ni 6 00 0 98 80 12.4 20. 120 16. 14. Cu 2 20 .70 33 20 90.1 165 107 71. 82. Zn 2 .10 .20 05 70 15.2 25. 20. 20. 24. Ga 1 80 40 96 80 87.2 78. 65. 106 44. Rb 1 30 40 .58 60 644. 579 349 272 303 Sr 46 .30 .90 .84 .50

46. 6 1.3 3 17. 60 123 .50 101 .10 23. 20 43. 00 9.4 0 97. 70 17. 90 38. 60 692 .30

45. 9 1.4 3 19. 10 120 .50 107 .70 24. 20 45. 60 10. 90 98. 90 19. 90 44. 90 879 .30

18. 2 1.4 1 5.7 0 16. 10 84. 10 2.9 0 4.0 0 28. 80 49. 60 15. 20 74. 60 106 .00

19. 5 1.7 2 7.6 0 28. 11 95. 08 3.6 0 5.3 0 4.8 0 49. 20 20. 20 56. 80 112 .90

18. 1 1.0 1 5.5 0 17. 58 79. 54 2.5 0 3.9 0 4.4 0 35. 90 14. 60 56. 60 143 .40

30. 2 2.4 0 9.3 5 59. 40 14. 58 10. 41 6.8 5 9.8 9 107 .40 19. 29 103 .94 166 .26

28. 8 1.9 2 4.8 0 52. 89 114 .20 6.0 0 6.4 0 6.5 0 78. 60 21. 00 38. 00 79. 38

36. 4 1.7 7 8.3 7 49. 08 12. 22 6.6 2 5.6 3 10. 62 75. 67 18. 44 70. 57 328 .77

38 2.0 4 5.5 0 54. 71 79. 19 5.3 0 5.2 0 7.6 0 76. 60 21. 10 27. 30 280 .50

1.4 5 33. 3 1.1 9 7.3 4 46. 79 5.4 6 7.5 2 2.7 2 17. 48 47. 99 16. 95 112 .72 246 .95

1.3 3 31. 9 1.2 8 6.1 0 41. 52 295 .50 5.3 0 7.4 0 17. 40 31. 50 18. 40 75. 50 190 .10

1.3 5 32. 3 1.0 9 6.1 0 38. 18 152 .10 4.6 0 6.3 0 14. 00 29. 70 19. 30 78. 00 223 .80

ACCEPTED MANUSCRIPT

Tb Dy Ho Er Tm Yb Lu Hf Ta Pb Th U

27. 60 107 .40 5.1 0 0.5 0

27. 60 81. 10 9.5 0 2.8 0

34. 10 86. 00 9.2 0 1.4 0

34. 80 91. 00 8.8 0 1.9 0

18. 50 205 .00 5.4 5 2.9 9

22. 60 231 .80 6.6 0 3.2 0

22. 70 172 .40 4.4 0 2.5 0

38. 35 385 .61 11. 71 14. 62

24. 2 224 .6 10. 00 8.9 0

29. 59 272 .21 11. 92 4.1 0

32. 60 244 .10 11. 50 1.8 0

19. 44 298 .82 11. 58 9.5 7

19. 50 53. 20 6.5 0 3.8 0

26. 40 53. 10 5.7 0 3.0 0

844 745 541 305 291 666 647 545 964 638 742 692 818 662 737 .20 .56 .80 .70 .60 .00 .80 .70 .07 .40 .02 .50 .17 .10 .80 27. 50 54. 60 7.5 2 30. 50 6.0 3 1.6 3 5.5 3 0.9 7 6.2 3 1.1 7 3.2 2 0.4 2 2.2 8 0.2 9 2.4 0 0.7 2 15. 40 5.0 0 1.0 5 0.5 8 147 .89

27. 90 61. 70 7.8 7 31. 00 6.6 8 1.7 2 6.3 2 0.9 6 6.8 4 1.1 1 3.4 8 0.4 8 2.3 6 0.2 9 2.1 7 0.6 8 19. 70 5.3 0 1.1 6 0.5 9 158 .71

δE u ∑R EE ∑L 117. 115 91. 108 82. 82. 127 RE 17 .94 72 .49 33 88 .25 E ∑H 15.3 18. 18. 17. 17. 15. 17. RE 5 31 73 23 13 92 87 E (La 4.94 5.2 5.6 6.7 7.6 4.9 4.8

19. 40 38. 00 4.6 0 17. 30 3.4 6 0.6 5 3.1 6 0.5 3 3.2 9 0.6 7 2.2 5 0.3 8 2.7 1 0.4 7 6.8 5 0.5 1 6.7 0 9.5 0 2.5 4 0.7 6 96. 87

15. 90 39. 90 5.3 4 20. 40 3.9 6 0.6 8 3.0 2 0.5 8 3.7 0 0.8 9 2.5 6 0.4 5 3.0 8 0.5 1 7.1 2 0.6 4 5.4 0 5.5 0 1.2 9 0.8 5 100 .97

18. 60 36. 30 4.3 9 16. 20 3.6 6 0.6 7 3.0 9 0.5 5 3.6 7 0.8 5 2.5 1 0.4 6 2.9 5 0.4 9 6.3 1 0.6 0 5.7 0 5.4 0 1.3 1 0.8 1 94. 39

42. 15 87. 83 11. 65 43. 41 8.0 9 1.9 8 7.6 1 1.2 3 6.9 0 1.4 1 4.4 6 0.6 7 4.5 9 0.7 3 10. 18 0.8 3 22. 91 21. 91 2.0 0 0.7 3 222 .71

17. 10 28. 30 4.9 2 19. 40 4.0 0 1.0 9 3.7 2 0.6 9 4.7 3 0.9 6 3.1 0 0.4 8 2.9 9 0.4 6 5.8 0 0.7 4 11. 7 5.9 0 0.6 9 0.8 6 91. 94

PT

27. 50 55. 40 7.5 0 29. 50 5.7 6 1.5 9 5.3 5 0.8 8 5.3 6 0.9 7 2.7 1 0.3 6 1.9 8 0.2 6 2.3 4 0.6 5 11. 30 4.2 0 1.0 5 0.5 9 145 .12

RI

18. 20 35. 10 4.8 4 19. 60 4.0 6 1.0 8 4.0 1 0.7 3 4.6 9 0.9 4 2.7 0 0.3 7 2.1 9 0.2 9 2.7 8 0.4 1 12. 60 4.2 0 1.1 9 0.9 1 98. 80

SC

16. 30 34. 40 4.9 0 20. 60 4.6 9 1.4 4 4.3 8 0.8 0 5.3 0 1.0 3 2.9 0 0.3 7 2.0 8 0.2 7 3.0 3 0.5 1 10. 00 3.6 0 1.3 6 0.9 7 99. 46

NU

Gd

24. 60 127 .80 5.3 0 5.0 0

MA

Eu

16. 10 35. 90 5.9 1 26. 30 5.8 3 1.6 8 5.2 8 0.9 5 5.6 8 1.0 4 2.9 4 0.3 7 2.1 8 0.2 9 2.8 2 0.5 6 6.0 0 2.4 0 0.5 5 0.8 1 110 .45

AC

Sm

22. 50 49. 30 6.8 5 29. 30 6.1 5.79 2 1.8 1.72 7 5.4 5.21 2 0.9 0.78 8 5.5 4.02 8 1.0 0.74 2 2.7 2.17 3 0.3 0.28 5 1.9 1.87 8 0.2 0.28 5 2.0 4.66 4 0.4 0.37 6 8.8 7.01 0 2.4 3.27 0 0.6 0.62 0 0.9 0.97 6 132. 134 52 .25

23. 43 196 .57 11. 75 11. 86 105 8.5 4 19. 62 45. 44 10. 55 25. 96 5.2 8 1.6 4 4.8 5 0.8 0 4.5 0 0.8 9 2.7 0 0.3 9 2.6 8 0.4 2 5.3 3 0.4 8 7.5 9 9.7 1 0.7 2 0.9 4 125 .72

D

703. 678 367 67 .60 .20

18.7 3 48.7 Ce 3 12.1 Pr 5 30.0 Nd 5 La

26. 70 126 .10 7.6 0 4.6 0

CE

Ba

29. 40 85. 70 5.8 0 12. 70

PT E

18.7 2 176. Zr 12 11.0 Nb 5 13.4 Cs 0 Y

30. 85 68. 14 14. 20 34. 82 6.7 7 1.7 5 6.2 1 1.0 1 5.6 4 1.1 2 3.5 1 0.5 2 3.5 0 0.5 6 7.4 9 0.6 8 14. 31 16. 32 1.7 7 0.8 5 178 .60

17. 60 33. 10 5.2 3 21. 20 5.5 4 1.2 9 5.0 4 0.9 7 6.0 8 1.2 2 3.8 6 0.5 8 3.6 2 0.5 4 6.5 4 0.9 1 10. 40 5.7 0 0.7 0 0.8 0 105 .87

11. 89 31. 97 7.0 6 15. 98 3.3 3 0.9 6 3.1 5 0.5 5 3.4 0 0.7 0 2.2 3 0.3 6 2.5 4 0.4 2 8.5 8 0.5 5 11. 43 24. 05 1.1 1 0.6 9 84. 54

5.9 0 12. 30 1.9 7 10. 21 2.1 1 0.6 4 2.0 7 0.4 6 3.3 9 0.7 4 2.4 1 0.4 1 2.8 5 0.4 7 2.2 0 0.5 8 11. 2 6.9 0 1.0 2 0.8 9 45. 93

31. 90 61. 20 7.0 7 24. 50 4.2 2 0.9 1 4.0 2 0.6 7 4.1 5 0.8 3 2.7 5 0.4 3 2.8 8 0.4 6 2.2 5 0.4 9 11. 00 9.7 0 1.3 8 0.9 3 145 .99

127 136 83. 86. 79. 195 74. 156 83. 71. 33. 129 .78 .87 41 18 82 .11 81 .53 96 19 13 .80 20. 21. 13. 14. 14. 27. 17. 22. 21. 13. 12. 16. 11 84 46 79 57 60 13 07 91 35 80 19 3.4 4.2 6.1 3.8 5.9 3.2 9.3 8.1 7.9 9.1 3.1 1.4

ACCEPTED MANUSCRIPT 8

0

5

6

Rb

Sr

8

3

8

5

9

6

4

8

6

3

7

5

6

0

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/Yb )N

t

87

Rb/86 87Sr/86 ±2 Sr Sr σ

MA

Samp Litholog le y

NU

Table 3

Sri

Sm

Nd

147

Sm/144 143Nd/144 ±2 εNd( TDM1 TDM2 Nd Nd σ t)

AC

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D

(M (pp (pp (pp (pp (M (M a) m) m) m) m) a) a) XDC alkali 254 87.21 644.4 0.3918 0.7050 8 0.7036 0.11648 0.51273 5.79 30.05 5 4.4 658 663 1 basalt 6 40 25 09 6 2 XDC alkali 254 78.30 579.3 0.3913 0.7053 8 0.7039 0.12627 0.51270 6.12 29.30 6 3.6 780 733 2 basalt 0 78 53 39 7 4 XDM andesite 253 106.5 272.8 1.1311 0.7071 9 0.7031 0.12296 0.51263 5.28 25.96 6 2.3 871 837 1 8 4 14 73 02 1 3 XDM andesite 253 44.60 303.5 0.4255 0.7047 9 0.7032 0.13764 0.51266 4.69 20.60 7 2.4 982 831 2 0 15 81 50 0 1 SD8 andesite 257 38.60 692.3 0.1614 0.7049 9 0.7043 0.11952 0.51259 6.03 30.50 4 1.7 898 886 0 48 38 47 4 7 SD9 andesite 257 44.90 879.3 0.1478 0.7049 9 0.7044 0.13027 0.51260 6.68 31.00 5 1.6 997 900 0 59 47 06 3 6 DYS quartz 255 103.9 166.2 1.8102 0.7113 8 0.7047 0.11266 0.51268 8.09 43.41 6 3.7 702 725 1 trachyte 4 6 29 08 42 7 7 DYS quartz 257 70.57 328.7 0.6215 0.7068 9 0.7045 0.11754 0.51268 6.77 34.82 6 3.6 737 737 3 trachyte 7 36 45 73 4 7 XQ1 rhyolite 255 74.60 106.0 2.0378 0.7110 11 0.7036 0.12091 0.51273 3.46 17.30 7 4.4 684 670 0 47 27 35 2 5 XQ2 rhyolite 255 56.80 112.9 1.4567 0.7087 9 0.7034 0.11735 0.51274 3.96 20.40 6 4.6 648 649 0 77 28 44 6 2 DST syenogra 257 112.7 246.9 1.3216 0.7088 10 0.7040 0.12598 0.51271 3.33 15.98 6 3.8 761 718 1 nite 2 5 93 81 49 2 4 DST syenogra 257 75.50 190.1 1.1500 0.7080 8 0.7038 0.12493 0.51272 2.11 10.21 4 4.1 730 695 2 nite 0 15 99 95 8 6 DST syenogra 257 78.00 223.8 1.0091 0.7081 10 0.7044 0.10413 0.51267 4.22 24.50 5 3.7 665 723 3 nite 0 90 58 68 2 4 Note: 87Rb/86Sr and 147Sm/144Nd ratios were calculated using Rb, Sr, Sm and Nd contents. εNd(t) values were calculated using present day (147Sm/144Nd)CHUR = 0.1967 and (143Nd/144Nd)CHUR = 0.512638 (Lugmair and Marti, 1978). TDM values were calculated using present day (147Sm/144Nd)DM = 0.2136 and (143Nd/144Nd)DM =

ACCEPTED MANUSCRIPT

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0.513151 (Miller and O’Nions, 1985).

ACCEPTED MANUSCRIPT Highlights · A suite of Late Permian igneous rocks in the Daheishan Horst were reported in this paper. · A tectonic setting at an active continental margin was proposed. · A link to the northward subduction of the Paleo-Asian oceanic plate was found beneath the

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Songliao-Xilinhot Block.

Graphics Abstract

Figure 1

Figure 2

Figure 3

Figure 4

Figure 5

Figure 6

Figure 7

Figure 8

Figure 9

Figure 10

Figure 11

Figure 12

Figure 13

Figure 14

Figure 15

Figure 16

Figure 17