Late Paleozoic-Early Mesozoic tectonic evolution of the Paleo-Asian Ocean: geochronological and geochemical evidence from granitoids in the northern margin of Alxa, Western China

Abstract

The Paleo-Asian Ocean (Southern Mongolian Ocean) ophiolitic belts and massive granitoids are exposed in the Alxa block, in response to oceanic subduction processes. In this work, we report petrographic, geochemical, and zircon U-Pb age data of some granitoid intrusions from the northern Alxa. Zircon U-Pb dating for the quartz diorite, tonalite, monzogranite, and biotite granite yielded weighted mean ²⁰⁶Pb/²³⁸U ages of 302±9.2 Ma, 246.5±4.6 Ma, 235±4.4 Ma, and 229.5±5.6 Ma, respectively. The quartz diorites (~302 Ma) exhibit geochemical similarities to adakites, likely derived from partial melting of the initially subducted Chaganchulu back-arc oceanic slab. The tonalites (~246.5 Ma) display geochemical affinities of I-type granites. They were probably derived by fractional crystallization of the modified lithospheric mantle-derived basaltic magmas in a volcanic arc setting. The monzogranites (~235 Ma) are characterized by low Al₂O₃, but high Y and Yb with notably negative Eu anomalies. In contrast, the biotite granites (~229.5 Ma) show high Al₂O₃ but low Y and Yb with steep HREE patterns and the absence of negative Eu anomalies. Elemental data suggested that the biotite granites were likely derived from a thickened lower crust, but the monzogranites originated from a thin crust. Our data suggested that the initial subduction of the Chaganchulu oceanic slab towards the Alxa block occurred at ~ 302 Ma. This subduction process continued to the Early Triassic (~246 Ma) and the basin was finally closed before the Middle Triassic (~235 Ma). Subsequently, the break-off of the subducted slab triggered asthenosphere upwelling (240–230 Ma).

Keywords

Paleo-Asian Ocean / Alxa / granite / geochemistry

Introduction

The Alxa block occupies a key tectonic position at the junction between the Central Asian Orogenic Belt (CAOB), the Tarim block, the North China Craton and the North Qilian orogenic belt (Wu and He, 1993; Wang et al., 1994; Zhang et al., 1997; Zhai and Bian, 2000; Ge et al., 2009; Geng and Zhou, 2012; Gong et al., 2012, 2013; Song et al., 2013; Zhang et al., 2013). The northern margin of the Alxa block is attached to the southern CAOB which is one of the most important areas for studying Phanerozoic continental growth in the world (Fig. 1). The Alxa block is a critical zone to investigate the tectonic evolution of the Paleo-Asian Ocean. There was a general consensus that successive lateral accretions from the Paleo-Asian Ocean produced the CAOB, with the formation of abundant accretionary complexes. However, the time of the closure of the Paleo-Asian Ocean remained controversial: 1) before late Carboniferous (Gao et al., 1998; Chen et al., 1999; Gao and Klemd, 2003; Xia et al., 2004; Charvet et al., 2007, 2011; Wang et al., 2007a, 2011; Gao et al., 2009; Yang and Zhou, 2009; Han et al., 2010a, b, c, 2011; Hegner et al., 2010); 2) in the late Permian-early Triassic (Li et al., 2002, 2005; Li et al., 2009; Xiao et al., 2008, 2009, 2010a, b, 2013, 2015; Tian et al., 2013, 2015) and 3) during the Triassic (Zhang et al., 2005; Zhang et al., 2007). Most studies have focused on three regions: the northern part of Xinjiang, Mongolia, and the eastern border of Mongolia (Fig. 1) including the Altai orogenic belt (Vladimirov et al., 2001, 2005; Annikova et al., 2006; Wang et al., 2007c, 2008a), the eastern Tianshan orogenic belt (Wang et al., 2008b; Zhou et al., 2010), the Hegenshan orogen and the Suolunshan-Xilamulun regions at the eastern border of Mongolia (Tao et al., 2003; Shi et al., 2004; Bao et al., 2007; Li et al., 2007; Zhao et al., 2007; Miao et al., 2008; Zhang et al., 2008; Chen et al., 2009; Liu et al., 2009; Zhang, 2009; Tong et al., 2010a, b), and the Northern Mongolia-West Baikal orogenic belt (Yarmolyuk et al., 2002; Jahn et al., 2009). However, only a few studies have been carried out in the northern Alxa block (Wang et al., 1998a, b; Xu et al., 2001; Zhang et al., 2002b; Wang et al., 2004; Li et al., 2010; Zhang et al., 2013; Zheng et al., 2014) and western Inner Mongolia (Li et al., 2006a; Wang et al., 2010; Zhang et al., 2011; Li et al., 2012).

In this work, LA-ICP-MS zircon U-Pb dating and major and trace element geochemical data have been determined for some felsic intrusions in the northern Alxa block. Our main objectives are to reconstruct the tectonic framework of the Alxa from late Paleozoic to early Mesozoic and to further constraint the closure of the Paleo-Asian Ocean.

Geological setting

The Alxa block includes three important boundary faults and two ophiolitic belts. The three faults from north to south are the Yagan fault belt, the Wutaohai-Enger Us fault belt, and the Chaganchulu fault belt (the Badanjilin fault belt, Zheng et al., 2014) (Fig. 2). The northern margin of the Alxa can be further divided into four tectonic units by the three faults (Fig. 2; Wu and He, 1993): the Yagan tectonic belt (immature island arc), the Zhusileng-Hangwula Paleozoic tectonic belt (early Paleozoic passive continental margin that converted to an active continental margin in the late Paleozoic), the Zongnaishan-Shalazhashan tectonic belt (mature island arc), and the Bayinnuoergong-Langshan tectonic belt (stable block).

The Zhusileng-Hangwula tectonic belt is located between the Yagan fault belt and the Enger Us fault belt. The Precambrian strata include metarhyolite, slate, quartzite, and granitic gneiss. Dalmanitina fossils are discovered in the Ordovician strata. Wu and He (1993) suggested that the tectonic belt represented an early Paleozoic passive continental margin. In addition, the lower Paleozoic strata were mainly carbonate-flysch formations and submarine volcanics. These observations indicated that the Zhusileng-Hangwula tectonic belt transformed from a passive continental margin to an active continental margin in the late Paleozoic (Wu and He, 1993).

The Zongnaishan-Shalazhashan tectonic belt is bounded by the Enger Us fault belt to the south and the Chaganchulu fault belt to the north. The ancient strata are Mesoarchean-Paleoproterozoic strata comprising the Alxa Group metamorphic rocks (Wang et al., 1994). In addition, Paleozoic to Mesozoic granitoids widely occurred in the belt, such as the Zongnaishan granite (single zircon U-Pb ages of 236.6±0.95 Ma, 249.7±2.6 Ma, and 268.4±0.69 Ma) which was interpreted to be a continental margin arc granite (Xie et al., 2014). However, as the principal part of the Zongnaishan-Shalazhashan arc zone, the Wuliji granite (~250.8 Ma) was interpreted to be a post-collisional granite and was produced by partial melting of mantle with assimilation of volcanic arc crust (Zhang et al., 2013). Based on these published data, the Zongnaishan-Shalazhashan tectonic belt was considered as a volcanic arc (Zhang et al., 2013; Xie et al., 2014; Zheng et al., 2014).

In the Bayinnuoergong-Langshan tectonic belt the Precambrian basement are mainly composed of metasedimentary and metavolcanic rocks, and tonalitic-granodioritic gneisses (Geng et al., 2007; Geng and Zhou, 2010, 2011). Abundant Permian granites (289–269 Ma) occurred in this belt, including dioritic gneiss, garnet-bearing tonalitic gneisses, and gneissic granites. They exhibit crust-mantle mixed geochemical characteristics (Geng and Zhou, 2012). Additionally, the Bayinnuoergong granite (zircon U-Pb age of 252.3±0.96 Ma) has been interpreted to be a syn-collisional granite (Xie et al., 2014).

Two ophiolitic belts (Enger Us and Chaganchulu) have been discussed in previous studies (e.g., Zheng et al., 2014). The Enger Us ophiolitic belt extended in NEE direction. The Enger Us ophiolitic suite consisted of ultrabasic rocks, gabbros, basalts, and cherts (Wang et al., 1994). Most of the mafic-ultramafic rocks were highly deformed and had undergone carbonatation and silicification. The Chaganchulu ophiolitic belt consisted of lenticular and striped ultrabasic rocks, gabbros, cherts, and rare basalts (Wang et al., 1994; Zheng et al., 2014). Based on regional geology, paleobiogeography, paleomagnetic and geochemical evidence, Zheng et al. (2014) suggested that the Enger Us ophiolitic belt represented the Paleo-Asian major oceanic basin and the Chaganchulu ophiolitic belt represented a back-arc basin. Recently, Zhang et al. (2015) have carried out statistical analysis of zircon xenocrysts within Permian magmatic rocks from the Zongnaishan-Shalazhashan (ZS) tectonic belt and Bayinnuoergong-Langshan (BL) tectonic belt. The ages and Hf isotopic data of the zircon xenocrysts imply that the basement beneath the ZS tectonic belt is relatively young, resembling the southern Central Asian Orogenic Belt (CAOB), in contrast to the BL tectonic belt. Thus, the boundary of the CAOB with the Alxa Block might be represented by the Chaganchulu ophiolitic belt.

Geology of granitoids and sampling

Quartz diorite

The quartz diorite is exposed in the Shazaoquan area (Fig. 3; samples AYQ-25, AYQ-26, AYQ-27, AYQ-28, and AYQ-29). It is intruded into the Lower Proterozoic Longshoushan Group. The quartz diorite was intruded by late biotite granite. The mineral assemblage consists of quartz (10%), plagioclases (55%), and amphiboles (35%). Accessory minerals include sphene and minor magnetite (Fig. 4(a)).

Tonalite

The tonalite is exposed in the Dashankou-Yaoquan area. It intruded into the quartz diorite and its middle section was later invaded by a Middle Triassic monzogranite (Fig. 3; samples AYQ-3, YQ-26, YQ-27, YQ-28, and YQ-29). It consists mainly of quartz (25%), plagioclase (60%), biotite (8%), amphibole (5%), and secondary apatite and magnetite (Fig. 4(b)).

Monzogranite

The monzogranite is distributed in the Yaoquan-Hongshanliang area. Its irregular shape has its long axis extending in the EW direction, and it invaded the Paleoproterozoic Longshoushan Group and the Carboniferous quartz diorite (Fig. 3; samples AYQ-9, AYQ-10, AYQ-11, AYQ-12, and AYQ-13). It consists mainly of plagioclases (30%), other feldspars (33%), quartz (30%), and biotite (5%), with a small amount of opaque minerals (Fig. 4(c)).

Biotite granite

The biotite granite is located in western of the quartz diorite, and it invaded in the Paleoproterozoic Longshoushan Group (Fig. 3; samples AYQ-17, AYQ-18, AYQ-19, AYQ-19R, AYQ-20, AYQ-21). It consists mainly of quartz (30%), plagioclases (25%), other feldspars (42%), and biotite (3%) (Fig. 4(d)).

Analytical methods

Whole rock major elements were analyzed at the State Key Laboratory of Continental Dynamics, Northwest University, and were completed by the Rigaku RIX2100 X-ray fluorescence spectrometer (XRF), following the analytical procedures of Liu et al. (2007). The analytical precision is within 0.1%. Trace elements were analyzed at the State Key Laboratory of Continental Dynamics, Northwest University, and the Institute of Geochemistry in Guangzhou. The analyses were completed using the Perkin-Elmer Sciex ELAN 6000 inductively coupled plasma mass spectrometer (Guangzhou) and the 820-MS plasma mass spectrometer (Northwest University), following the analytical procedures of Govindaraju (1994), Li (1997), and Li et al. (2006b). The analytical precision is within 5%. The results of these analyses are shown in Table 1.

Fresh rock samples were crushed to pass a 120-mesh sieve, and separation of zircon crystals was accomplished by conventional artificial panning, heavy liquid separation, magnetic techniques, and binocular microscope observation. Samples were numbered as quartz diorite (TW-4), tonalite (TW-6), monzogranite (TW-2), and biotite granite (TW-3). Zircon cathodoluminescence (CL) images and LA-ICP-MS zircon U-Pb dating were completed by the State Key Laboratory of Continental Dynamics, Northwest University. LA-ICP-MS was performed with the 820-MS plasma mass spectrometer, which has the collision response system and was the newest generation machine of Varian, Inc. (USA). The laser ablation system was the GeoLas 2005-type, ArF 193 nm UV excimer laser, which was produced by the Lambda Physik AG Company (Germany). The analytical procedures followed those of Liu et al. (2007). The analytical results are shown in Table 2.

Results

LA-ICP-MS zircon U-Pb dating

Zircons of the samples are mostly euhedral with short columnar shapes. The zircons display oscillatory zoning consistent with magmatic zircons (Fig. 5; Belousova et al., 2002). They exhibit wide ranges of U (31.89‒4187.52 ppm) and Th (17.69‒1440 ppm) contents with high Th/U ratios (0.16‒1.43 ppm) typical of magmatic zircons (Belousova et al., 2002). Based on the fact that the ordinary lead correction can cause a greater effect on the ²⁰⁷Pb/²³⁵U ratio, we adopted the ²⁰⁶Pb/²³⁸U age-weighted average to represent the formation age of these granitoids. Zircon U-Pb dating results for the quartz diorite, tonalite, monzonitic granite, and biotite granites yielded weighted mean ²⁰⁶Pb/²³⁸U ages of 302±9.2 Ma, 246.5±4.6 Ma, 235±4.4 Ma, and 229±5.6 Ma, respectively (Fig. 6; Table 2).

Whole-rock geochemistry

Quartz diorite (~302 Ma)

The quartz diorites have SiO₂ contents ranging from 60.66 wt% to 61.37 wt% (Table 1). They exhibit relatively high Al₂O₃ (17.98‒18.44 wt%), CaO (6.08‒6.59 wt%), and Na₂O (4.20%‒4.46 wt%), but low K₂O (0.67%‒0.92 wt%). The MgO contents range from 2.13‒2.51 wt% with variable Mg-number (Mg^#) values of 48 to 51. The sample points plot in the low-K (tholeiitic) series of the SiO₂-K₂O diagram (Fig. 7(a)) and in the diorite field of the SiO₂-(Na₂O+ K₂O) discrimination diagram (Fig. 7(b)). The aluminum saturation index (ACNK) values range from 0.91 to 0.96, and the samples plot in the metaluminous field of the aluminum saturation index diagram (Fig. 8).

In primitive mantle-normalized trace element spider diagrams (Fig. 9(a)), the samples display enrichment in Rb, Ba, Th, U, and Sr, but depletion in Nb, Ta, Ti, and HREE with positive Sr anomalies. In the chondrite-normalized rare earth element (REE) diagrams (Fig. 9(b)), they exhibit moderate LREE enrichment, (La/Yb)_N = 5.22–10.86, and slightly positive Eu anomalies (dEu= 0.97–1.18). The HREEs show flat patterns. In addition, the samples exhibit relatively high Sr (582–620 ppm (parts per million)) and Sr/Y (35.11–53.18), Y/Yb (10.98–11.21), and La/Yb (7.28–15.14) ratios, but low Y (11.41–16.72 ppm), Yb (1.02–1.51 ppm), Ni (10.33–11.55 ppm), Cr (13.35–16.84 ppm), Rb/Sr (0.03–0.05) and La/Ce (0.42–0.50).

Tonalite (~246.5 Ma)

The tonalites have SiO₂ contents ranging from 62.14 wt% to 68.94 wt% (Table 1) and an average Al₂O₃ content of 16.50 wt%. They exhibit relatively high Na₂O (Na₂O/K₂O= 1.72), CaO (average 3.39 wt%), and ^TFe₂O₃ (average 3.76 wt%). The MgO contents of the samples have an average of 1.49%, and the Mg-number (Mg^#) values vary from 41 to 47. The sample points plot in the calc-alkaline series of the SiO₂-K₂O diagram (Fig. 7(a)) and in the granodiorite field of the SiO₂-(Na₂O+ K₂O) discrimination diagram (Fig. 7(b)). The ACNK values range from 0.95 to 1.13, and the samples plot in the metaluminous and slightly peraluminous I-type fields on the aluminum saturation index diagram (Fig. 8).

In the primitive mantle-normalized trace element spider diagrams (Fig. 9(c)), the samples display enrichment in Rb, Ba, Th, and K and depletion in Nb, Ta, P, and Ti. In the chondrite-normalized rare earth element diagrams (Fig. 9(d)), there are extreme differentiations between the LREEs and HREEs and moderately negative Eu anomalies (dEu= 0.75–0.80). In addition, the samples exhibit relatively high abundances of Sr (average of 371.28 ppm), Y (average of 29.68 ppm), and Yb (average of 2.47 ppm), and high Rb/Sr ratios (0.24–0.34), but low La/Ce (0.47–0.52) ratios.

Monzogranite (~235 Ma)

The monzogranites have SiO₂ contents ranging from 72.72 to 78.13 wt% (Table 1) and Al₂O₃ contents ranging from 11.56 wt% to 14.07 wt%. They exhibit relatively equivalent Na₂O (3.15–3.80 wt%) and K₂O (3.61–4.31 wt%, K₂O/Na₂O= 0.97–1.28) values, but low CaO (0.93–2.32 wt%), MgO (0.20–0.56 wt%, Mg^# = 27–37), and TiO₂ (0.12–0.26 wt%). The sample points plot in the high-K calc-alkaline series of the SiO₂-K₂O diagram (Fig. 7(a)) and in the granite field of the SiO₂-(Na₂O+ K₂O) discrimination diagram (Fig. 7(b)). The ACNK values range from 0.93 to 1.04, and the samples plot in the quasi-aluminous field of the aluminum saturation index diagram (Fig. 8).

In the primitive-mantle-normalized trace element spider diagrams (Fig. 9(e)), the samples display enrichment in Th and K, but strong depletion in Nb, Ta, Sr, P, and Ti. In the chondrite-normalized rare earth element diagrams (Fig. 9(f)), they exhibit LREE enrichment [(La/Yb)_N = 1.75‒7.59], extreme differentiation between LREEs and HREEs, and strong negative Eu anomalies (dEu= 0.39‒0.61). In addition, they exhibit relatively high Rb/Sr (0.81‒1.40) but low La/Ce (0.46‒0.48) ratios.

Biotite granite (~229.5 Ma)

The biotite granites have SiO₂ contents ranging from 68.89 to 71.04 wt% (Table 1) and Al₂O₃ contents ranging from 13.95 wt% to 15.76 wt%. They exhibit high CaO (1.75–2.54 wt%), relatively equivalent Na₂O (3.63–4.81 wt%) and K₂O (3.52–5.45 wt%, Na₂O/K₂O= 0.67–1.37), and low MgO (0.63–0.81 wt%) and TiO₂ (0.35–0.44 wt%),with low Mg^# (34–36). The sample points plot in the high-K calc-alkaline series of the SiO₂-K₂O diagram (Fig. 7(a)) and in the granite field of the SiO₂-(Na₂O+ K₂O) discrimination diagrams (Fig. 7(b)). The ACNK values range from 0.85 to 1.05, and the samples plot in the metaluminous field of the aluminum saturation index diagram (Fig. 8).

In the primitive-mantle-normalized trace element spider diagrams (Fig. 9(g)), the samples display enrichment in Ba, Rb, and Th, but strong depletion in Nb, Ta, Sr, P, and Ti. In the chondrite-normalized rare earth element diagrams (Fig. 9(h)), they exhibit LREE enrichment [(La/Yb)_N = 41.33‒60.95], extreme differentiation between LREEs and HREEs, and a lack of notably negative Eu anomalies (dEu= 0.74‒1.17).

Discussion

Petrogenesis

Quartz diorite (~302 Ma)

All the quartz diorites share the geochemical affinities of adakites such as high Al₂O₃, Sr, Sr/Y ratio, and depletion in low Y and Yb contents (Defant et al., 1991; Drummond et al., 1996; Martin, 1999; Zhang et al., 2010). The REE data defined listric-shaped REE profiles on chondrite-normalized diagrams, implying the fractionation of amphibole (Richards and Kerrich, 2007). As we all know, the removal of amphibole would produce a decrease in Dy/Yb ratio. However, the negative correlation between Dy/Yb and SiO₂ is not observed in a Harker diagram (not shown).

In the SiO₂-MgO diagram (Fig. 10), all the quartz diorites plot within the adakite field. There are several genetic models proposed to interpret the origin of adakitic rocks: 1) partial melting of a young, hot subducted slab (e.g., Drummond and Defant, 1990); 2) crustal assimilation and fractional crystallization (AFC) of basaltic magmas at high pressure conditions (e.g., Castillo et al., 1999; Macpherson et al., 2006); and 3) partial melting of a thickened lower crust (e.g., Muir et al., 1995).

In general, those adakitic rocks derived from AFC of basaltic magmas are a component of a suite of igneous rocks with basaltic-andesitic-rhyolitic compositions (Castillo et al., 1999). But basaltic and rhyolitic rocks are not observed near these quartz diorites. Additionally, the absence of inherited zircons and relatively high Mg# (48‒51) implied that they did not likely originate from fractional crystallization of primary basaltic magmas with old crustal contaminant (Castillo et al., 1999).

Thickened lower crust-derived adakitic rocks are enriched in K₂O content but depleted in Na₂O content in contrast with our samples. Recently, Hou et al. (2004) suggested that the Gangdese adakitic intrusions originating from the lower crust exhibit high Rb/Sr ratios (>0.05). In contrast, our quartz diorites have relatively low Rb/Sr ratios (0.032‒0.046). More importantly, those adakitic rocks are generally exposed in some specific regions which have undergone crustal thickening, e.g. orogenic belts (Kay et al., 1993; Wang et al., 2006). The quartz diorites (adakites) with high Al and Na₂O contents argue against the model that they were not partial melts of lower continental crust which are characterized by high K₂O and (Na₂O+ K₂O) contents and low Al₂O₃ contents (Atherton and Petford, 1993; Kay et al., 1993; Drummond et al., 1996; Kay and Mpodozis, 2001; Zhang et al., 2001a, b, c, 2002a; Mao et al., 2012). As we will discuss below, there was no crustal thickening process before the emplacement of these quartz diorites. Consequently, it was difficult to envisage that they were produced by partial melting of a thickened lower crust.

An alternative genetic model was that these quartz diorites were likely derived by partial melting of a subducted oceanic slab. This viewpoint was supported by their high Na₂O, Na₂O/K₂O, and low K₂O varying from 0.67‒0.92 wt% (Martin, 1999). Previous studies indicated that slab-derived adakitic melts show relatively low Rb/Sr ratios with a range from 0.01 to 0.04 (Drummond et al., 1996). The low Rb/Sr ratios in our samples further supported a subducted slab as their source. The relatively high Mg-number (Mg^# = 48‒51) is attributed to gradual assimilation of slab melts by asthenospheric mantle during ascent. All these features suggested that the quartz diorites were derived from a subducted oceanic slab (Sen and Dunn, 1994; Martin, 1999; Rapp et al., 1999; Xu and Ma, 2003; Wang et al., 2007b; Mao et al., 2012). The (La/Yb)_N-Yb_N and Sr/Y-Y diagrams (Fig. 11) showed that the protolith is roughly a garnet amphibolite, which indicates that the source region might not have residual plagioclases, but rather amphiboles, garnets, and Fe-Ti oxides (ilmenites, rutiles, etc.).

Tonalite (~246.5 Ma)

All the tonalites are also calc-alkaline series and share geochemical affinities of I-type granites. They are characterized by high SiO₂, Al₂O₃, and Na₂O, low K₂O, moderate negative Eu anomaly and negative Nb, Ta, Ti anomalies. The occurrence of amphibole in Fig. 4(b) further suggested that they are I-type granites.

The relative high Mg^# (41‒47) suggested that these felsic rocks were likely produced by fractional crystallization of mantle-derived basaltic magmas. The speculation was supported by negative correlations between MgO, ^TFe₂O₃, Mg^#, compatible elements (e.g., Cr and Ni) and SiO₂. These correlations might be attributed to the removal of mafic minerals including biotite and amphibole, inferred by the occurrence of these minerals in Fig. 4(c). The negative Nb-Ta-Ti anomalies in these tonalites suggested that the lithospheric mantle source had been modified by subducted slab-released components. Consequently, we considered that the tonalites were likely derived by fractional crystallization of a modified lithospheric mantle-derived basaltic magma.

Monzogranite (~235 Ma)

All the monzogranites are characterized by high SiO₂ and K₂O, low Al₂O₃, CaO, ^TFe₂O₃, MgO, and TiO₂, and display metaluminous and high-K calc-alkaline signatures. They exhibit geochemical characteristics of mafic rocksderived from partial melts in continental crust (Li et al., 2007). In the Ga/Al diagram (Fig. 12), the samples are plotted within the I-S field. Additionally, standard CIPW calculations show that the rocks contain diopside, but<1 wt% corundum (not shown in table), indicating that the monzogranites are I-type granites.

According to the trace element geochemistry, all the monzogranites display enrichment in Rb, Th, and K, but strong depletion in Ba, Nb, Ta, Sr, P, and Ti. Low P and Ti may be associated with the fractional crystallization of ilmenite, sphene, and apatite. The depletion of Nb and Ta may be associated with the depletion of crustal magma, and the high Rb/Sr ratio indicates that the magma source is a crustal source (Rubatto and Hermann, 2003). The coeval basaltic magmas were required for triggering partial melting of the crustal source.

Biotite granite (~229.5 Ma)

The biotite granites are characterized by high SiO₂ and Al₂O₃ contents, and low Mg and TiO₂ contents with ASI<1.1. A CIPW standard mineralogy calculation shows corundum<1 wt% (no table). All the evidence indicate that the biotite granites belong to I-type granites (Fig. 12).

All the samples display enrichment of Ba, Rb, Th, and LREE, but strong depletion of Nb-Ta-Ti and P, and high Rb/Sr ratio, indicating a crustal source (Rubatto and Hermann, 2003). All the evidence implies that these biotite granites were probably derived by partial melting of metaigneous rocks. Considering that the rocks have low Mg-number (Mg^# = 34‒36) and compatible element contents (e.g., Ni and Cr), we inferred that significant mantle materials were not involved in forming the biotite granites.

The absence of significantly negative Eu anomalies indicates that the source was plagioclase-free due to the high partition coefficient (D) of Eu (D_Eu = 5.417) between felsic melts and plagioclase (Nash and Crecraft, 1985). All the samples are strongly enriched in LREE with high (La/Yb)_N of 41‒61 and exhibit steep HREE patterns with (Gd/Yb)_N of 2.0‒2.9, indicating the presence of garnet in the source. Previous studies indicate that during the dehydration-melting of meta-igneous rocks (biotite gneiss and quartz amphibolite), garnet would occur as one of residual phases at pressures≥12.5 kbar and plagioclase would be unstable at pressures≥15 kbar (Douce et al., 1995). The occurrence of garnet without plagioclase indicates that the source of the biotite granites is relatively deep (>50 km). In summary, the biotite granites were likely produced by partial melting of meta-igneous rocks within the thickened lower crust (>50 km).

Tectonic implications

In recent years, based on comprehensive studies on ophiolites, magmatism, structure geology, sedimentary rocks, and HV/EHV metamorphic rocks (Zonenshain et al., 1990; Chen and Jahn, 2004; Gao et al., 2007; Zhang et al., 2007; Hegner et al., 2010), knowledge about the tectonic evolution of the CAOB has been tremendously improved. However, granitoids from each tectonic unit have been interpreted to be emplaced in different tectonic settings, and their tectonic implications remain unclear and controversial (Kozakov et al., 1997; Budnikov et al., 1999; Yarmolyuk et al., 2002; Jahn et al., 2004, 2009; Annikova et al., 2006; Orolmaa et al., 2008).

Using 227 Ma as the dividing line, the early Mesozoic granitoids of the CAOB were emplaced in two magmatic episodes (Li et al., 2010) (Fig. 1). In the western Baikal orogen, the first-stage magmatic rocks (251‒227 Ma, alkaline A-type granites) were emplaced in a post-orogenic or intraplate tectonic setting (Yarmolyuk et al., 2002; Jahn et al., 2004, 2009). In the Altai orogen, the first-stage granites consist of post-orogenic I-type and A-type granites, and most scholars believed that the main orogenic event ended during the Late Permian (Fig. 1; Pavlova et al., 2008). In the eastern Tianshan and Beishan orogens, the first-stage granites are composed of high-K calc-alkaline A-type and I-type granites, emplaced in a post-collision tectonic setting (Fig. 1; Li et al., 2006a; Zhang et al., 2007; Wang et al., 2008b; Li et al., 2010; Wang et al., 2010; Zhou et al., 2010; Zhang et al., 2013). However, in the middle of Mongolia, the first-stage granites are high-K calc-alkaline and calc-alkaline S- and I-type granites, formed in a late syn-orogenic setting (Orolmaa et al., 2008) (Fig. 1). In the Inner Mongolia and Jilin orogen, the first-stage granites consist of high-K calc-alkaline I-type and S-type granites with arc affinities (Fig. 1) (Tao et al., 2003; Bao et al., 2007; Li et al., 2007; Miao et al., 2008; Zhang et al., 2008; Chen et al., 2009; Zhang, 2009; Tong et al., 2010a).

The second stage ranged from Late Triassic to Early Jurassic (226‒195 Ma). These magmatic rocks include high-K calc-alkaline A-type granites and I-type granites, formed in a post-orogenic tectonic setting or an extensional environment (Vladimirov et al., 2001; Tao et al., 2003; Shi et al., 2004; Annikova et al., 2006; Ma et al., 2007; Wang et al., 2007c, 2008a; Li et al., 2010) (Fig. 1). However, granites in the Okhotsk belt and adjacent Mongolia were emplaced in a syn-orogenic tectonic setting. Their emplacement might be in response to Mesozoic back-arc basin closure and arc-continent collision (Yarmolyuk et al., 2002; Jahn et al., 2004, 2009). Therefore, the two-stage granitoids from different tectonic units were emplaced in different tectonic environments.

Abundant Early Permian granites also intruded into the Alxa metamorphic basement, indicating that the studied area was strongly modified by the late Paleozoic orogeny (Geng and Zhou, 2012). Based on the spatial and temporal distribution of the Paleozoic granites, ophiolite-complex rocks, and volcanic-sedimentary assemblage, we believe that the late Paleozoic Enger Us ophiolitic belt (~302 Ma for pillow lava; Zheng et al., 2014), Shalazhashan granite (continental margin arc), and Chaganchulu ophiolitic belt (~275 Ma for gabbro; Zheng et al., 2014) made up a trench-arc-basin system as a product of the Paleo-Asian Oceanic southward subduction. Xie et al. (2014) have found some radiolarian fossils as young as the Late Permian, implying that the subduction of the Enger Us ocean might have lasted to at least the Late Permian. Furthermore, the ca. 250 Ma Wuliji post-collisional granite intruded into the upper section of the Late Permian, which is molasse formation composed of sandstone, gravel-bearing sandstone, silty shale, and conglomerate (Zhang et al., 2013). Thus, the Paleo-Asian branch ocean, represented by the Enger Us ophiolitic belt, was probably closed at the end of the Permian.

The quartz diorites (adakites) (~302 Ma) in this study are close to the southern Chaganchulu ophiolitic belt. As indicated by the petrogenesis of the quartz diorites, they likely originated from a subducted slab. Given the temporal and spatial distribution of the Chaganchulu ophiolitic belt and quartz diorites, we suggest that the quartz diorites might be derived by partial melting of the Chaganchulu back-arc oceanic slab. In the Langshan area, eastern Alxa block, some 292‒285 Ma deformed granitic-granodioritic porphyries show typical arc affinities and might have been emplaced before the collision of the Zongnaishan-Shalazhashan arc with the Alxa block (Lin et al., 2014). Similarly, Feng et al. (2013) have investigated some 306‒262 Ma, EW-trending mafic-ultramafic rocks which occurred in the Bijiertai, Honggueryulin, and Qinggele areas along the Bayinnuoergong-Langshan tectonic belt. All the mafic-ultramafic rocks represent arc magmatism as products of the Chaganchulu back-arc oceanic subduction. At ca. 246.5 Ma, the emplacement of the tonalites with arc affinities implied a subduction setting rather than a collisional or post-collisional setting. The Chaganchulu oceanic continued to the Early Triassic.

Considering the location and age of the Enger Us and Hegenshan ophiolitic belt, the Enger Us oceanic basin might be equivalent to the Hegenshan ocean which was also closed not later than the Permian (Fig. 1; Miao et al., 2008). Similarly, another Paleo-Asian branch ocean, represented by the Suolunshan-Xilamulun suture zone, might be equivalent to the Chaganchulu back-arc oceanic basin, which was closed in the late Permian-early Triassic, consistent with the final amalgamation of the Sino-Korean and Siberian cratons (Wang and Fan, 1997; Li et al., 2006b; Li et al., 2007; Tong et al., 2010b).

From 235 Ma to 229.5 Ma, partial melting of the continental lower crust produced the monzogranites and biotite granites. In the tectonic discrimination diagrams of granites (Fig. 13), all the samples are plotted within the post-collision field (Pearce et al., 1984). It is then inferred that the area might have undergone tectonic transformation from a collisional orogenic compressional environment to a post-orogenic extensional environment.

On the basis of the new petrological, geochemical, and geochronologic data, together with studies on regional geology, we propose an integrated model for the Late Paleozoic to Early Mesozoic tectonic evolution of the Alxa block as illustrated in Fig. 14:

1) Southward subduction of the Paleo-Asian Ocean produced a trench-arc-basin system (Southern Mongolian Ocean+ Zongnaishan-Shalazhashan island arc+ Chaganchulu back-arc basin) (Fig. 14(a)).

2) From 302 Ma to Late Permian, the Chaganchulu back-arc oceanic slab subducted south underneath the Alxa block (Feng et al., 2013; Lin et al., 2014; Zheng et al., 2014), and partial melting of the subducted slab produced the quartz diorites with adakitic affinities (Fig. 14(b)).

3) From Late Permian to 240 Ma, the southward subducted slab-released fluids induced partial melting of the overlying enriched mantle wedge. Then, fractional crystallization of the modified mantle-derived basaltic magmas produced the tonalites (Fig. 14(c)). In this period, the Enger Us oceanic basin was closed due to the collision of South Mongolian Block with the Zongnaishan-Shalazhashan arc (Xie, 2014; Zheng et al., 2014; Xiao et al., 2015).

4) From 240 Ma to 230 Ma, the Chaganchulu back-arc basin was finally closed and the Zongnaishan-Shalazhashan island arc was welded to the Alxa block, followed by crustal thickening. Then, asthenosphere mantle upwelling triggered by the slab break-off induced partial melting of the thickened lower crust to produce the biotite granites and partial melting of relatively thin crust to form the monzogranites, respectively (Fig. 14(d)).

Conclusions

1) The quartz diorites (302±9.2 Ma) were derived by partial melting of the initially subducted Chaganchulu back-arc oceanic slab, The tonalites (246.5±4.6 Ma) were produced by fractional crystallization of a modified lithospheric mantle-derived basaltic magma, and were formed in a volcanic arc setting. The monzogranites (235±4.4 Ma) had a crustal source, and coeval basaltic magmas were required for triggering partial melting of the crustal source. The biotite granites (229.5±5.6 Ma) were derived by partial melting of meta-igneous rocks within the thickened lower crust.

2) The Enger Us oceanic basin was likely closed in the Late Permian, and the Chaganchulu back-arc oceanic basin began to subduct underneath the Alxa block at 302 Ma. This subduction process continued to the Early Triassic (246 Ma) and the basin closed fully before the Middle Triassic (235 Ma). The compressional environment led to crustal thickening. Then the asthenosphere mantle upwelling was triggered by slab breakoff and mantle-derived magma underplating near the crust-mantle boundary (240‒230 Ma). The tectonic environment then converted to a post-orogenic extensional environment, and the area might also have undergone a conversion process from crustal thickening to thinning.

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