Geochronology and geochemistry of late Paleozoic magmatic rocks in the Yinwaxia area, Beishan: Implications for rift magmatism in the southern Central Asian Orogenic Belt

Accepted Manuscript Geochronology and geochemistry of late Paleozoic magmatic rocks in the Yinwaxia area, Beishan: Implications for rift magmatism in the southern Central Asian Orogenic Belt Rongguo Zheng, Tairan Wu, Wen Zhang, Qingpeng Meng, Zhaoyu Zhang PII: DOI: Reference:

S1367-9120(14)00189-8 http://dx.doi.org/10.1016/j.jseaes.2014.04.022 JAES 1938

To appear in:

Journal of Asian Earth Sciences

Received Date: Revised Date: Accepted Date:

2 December 2013 21 April 2014 23 April 2014

Please cite this article as: Zheng, R., Wu, T., Zhang, W., Meng, Q., Zhang, Z., Geochronology and geochemistry of late Paleozoic magmatic rocks in the Yinwaxia area, Beishan: Implications for rift magmatism in the southern Central Asian Orogenic Belt, Journal of Asian Earth Sciences (2014), doi: http://dx.doi.org/10.1016/j.jseaes. 2014.04.022

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Geochronology and geochemistry of late Paleozoic magmatic rocks in the Yinwaxia area, Beishan: Implications for rift magmatism in the southern Central Asian Orogenic Belt

Rongguo Zheng a, Tairan Wu a, Wen Zhang b, Qingpeng Meng a, Zhaoyu Zhang a a

MOE Key Laboratory of Orogenic Belts and Crustal Evolution, School of Earth and Space Sciences, Peking University, Beijing 100871, China b

Institute of Geology, Chinese Academy of Geological Sciences, Beijing 100037, China

Corresponding author: Tairan Wu MOE Key Laboratory of Orogenic Belts and Crustal Evolution School of Earth and Space Sciences Peking University Beijing 100871 China Tel. +86 10 62765534 Email: [email protected]

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Abstract: Mafic-ultramafic rocks are distributed widely in the Beishan rift, which is located in the southern Beishan, central southern Central Asian Orogenic Belt. The Yinwaxia study area is located in eastern Beishan rift, where mafic-ultramafic rocks occur along major faults. The zircon SHRIMP U-Pb age obtained of a gabbro is 281±11 Ma, and the age of the basalt is constrained by the youngest xenocrystal with an age of 265 Ma, which substantiate that these mafic rocks formed in Permian. Basalts and gabbros exhibit similar geochemical characteristics including: high SiO2, total Fe2O3 and TiO2 contents; low MgO contents and Mg# values; and tholeiitic characteristics. Yinwaxia mafic rocks have relatively high total rare earth element contents, enrichment in light rare earth elements, enrichments in the high field strength elements, and obvious negative Nb-Ta-Ti anomalies. Basalts exhibit low (87Sr/86Sr)i and high εNd(t) values, while gabbros exhibit relatively high (87Sr/86Sr)i and low εNd(t) values. Isotopic compositions of these mafic rocks display a mixed trend between depleted and enriched mantles. Meanwhile, differing εNd(t) values show that basalts were intensively contaminated by juvenile crustal materials, but gabbros were contaminated by older continental crust. We conclude that Yinwaxia mafic rocks were derived from lithospheric mantle metasomatized by fluids and/or melts from subducted slab; parental magmas underwent AFC processes, then emplaced along faults in a continental rift. We collected geochemical and geochronological data in the study area, and collated geochronological data from previous workers in the Beishan orogenic belt to develop a geochronological frequency diagram. From these data and analyses we deduced a model of tectonic evolution for the Beishan orogenic belt. Considering the geochemistry, sedimentological evidence for rifting, and the geochronological frequency diagram, we 2

propose that the Beishan rift had entered a post-collision stage since Early Devonian, and then changed into a continental rift stage around late Carboniferous-early Permian.

Key words: Beishan, mafic rocks, SHRIMP, Sr-Nd isotope, continental rift, Central Asian Orogenic Belt

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1. Introduction The Central Asian Orogenic Belt (CAOB) is the largest area of Phanerozoic continental accretion and crustal growth in the world (Sengör et al., 1993; Jahn et al., 2000; Kovalenko et al., 2004; Windley et al., 2007). The CAOB extends from Kazakhstan in the west to eastern Siberia in the east (Fig. 1), and separates the Siberian craton in the north from the Tarim and north China (or Sino-Korean) cratons in the south (Zonenshain et al., 1990; Mossakovsky et al., 1994; Jahn et al., 2000; Badarch et al., 2002). There is a general consensus that the CAOB grew by successive lateral accretions of arcs, accretionary complexes and a few continental blocks southward from Siberia and southern Mongolia during the evolution of the Paleo-Asian Ocean (Windley et al., 2007; Kröner et al., 2008; Xiao et al., 2009, 2013; Safonova and Santosh, 2014.). Although the timing of the final closure of the Paleo-Asian ocean is still a matter of debate, accumulating evidence suggests a diachronous closing process which resulted in many and diverse tectonic regimes in the late Paleozoic within this immense accretionary orogenic belt. After the amalgamation, the CAOB was affected by continental magmatic activities and modified by intracontinental orogenic reactivations (Windley et al., 1990; Wartes et al., 2002; Khain et al., 2003; Kröner et al., 2007, 2010). The late Paleozoic was an important period of tectonic transition and crustal growth of the CAOB; the crustal growth is represented by large volumes of juvenile granitoids with positive εNd and low initial (87Sr/86Sr)i values (Han et al., 1997; Jahn et al., 2000; Chen and Jahn, 2004). However, this period was also characterized by emplacements of many coeval mafic-ultramafic complexes having controversial origins (Hong et al., 2003; Han et al., 2004; Zhou et al., 2004; Zhao et al., 2006; Windley et al., 2007; Mao et al., 2008; Pan et al., 4

2008; Pirajno et al., 2008). One group of these late Paleozoic mafic-ultramafic complexes occurs in the the Beishan orogenic belt, located in the middle region of the southernmost CAOB (Fig.1a). The Beishan orogenic belt has many Cu-Ni-bearing mafic-ultramafic complexes, especially in the southern portion of the belt (Mao et al., 2008; Pirajno et al., 2008; Zhang et al., 2008). Most of these mafic-ultramafic rocks crop out along regional large-scale faults or sutures, such as mafic rocks in the Pobei and Liuyuan areas. Chemical and isotopic compositions of continental mafic rocks provide the best proxy record for the chemical and physical evolution of the deep continental lithosphere and underlying mantle (Farmer, 2003). Mafic–ultramafic rocks can also provide valuable information for unraveling the geological history of orogenic belts. Although many researches have focused on mafic-ultramafic rocks in the southern Beishan as noted above, there are still controversies regarding their tectonic implications of these rocks. Mafic-ultramafic complexes in the southern Beishan are usually described as the products of within-plate magmatic activity (Jiang et al., 2006; Mao et al., 2008; Pirajno et al., 2008; Zhang et al., 2008), formed as a result of post-orogenic extension or plume related magmatic process (Qin et al., 2011; Su et al., 2011a, 2011b, 2012b). Conversely, it has also been suggested that these mafic-ultramafic complexes are Alaskan-type intrusions, generated in the early Permian subduction-related environment (Xiao et al., 2004a; Mao et al., 2006; Ao et al., 2010). Alternatively, some workers have argued that mafic complexes in the southern Beishan were associated with a mantle plume that resulted in Permian flood basalts in the western part of the Tarim block (Qin et al., 2011; Su et al., 2011a, 2011b, 2012b). This work certainly will provide crucial insights into the mechanism of orogenesis and tectonic history 5

of the southern Beishan. In light of these conflicting theories, we present new geochronological, and major and trace element data, as well as whole rock Sr-Nd isotopic compositions, for the mafic rocks in the Yinwaxia area, in the eastern sector of southern Beishan. To provide further insight into the mechanism of orogenesis and tectonic history of the southern Beishan orogenic belt, we also complied the geochronological data previously reported in the literature to describe magmatic sequences of the Beishan rift. From this large data set, we endeavored to define tectonic settings of mafic rocks in the Yinwaxia area, discuss the formation mechanism, and develop a reasonable tectonic evolution model. 2. Geological background The Beishan orogenic belt, situated in the southernmost CAOB, is a conjunction region of the CAOB and North China and Tarim cratons (Fig. 1a). It is separated from the Tianshan orogenic belt to the west by the Ruoqiang-Xingxingxia fault, and from the Mongolia-Xing’anling orogen to the east by the Altyn Tagh-Alxa fault. The Dunhuang block, part of the Tarim craton (Zhang et al., 2013), is located to the south of the Beishan orogenic belt (Zhou and Graham, 1996; Wu et al., 1998; Zhang et al., 2011a). Tectonically, the Beishan orogenic belt is often regarded as the eastern extension of the Chinese Tian Shan (Li, 1980; Liu and Wang, 1995; Xiao et al., 2010), and it comprises an assemblage of blocks, magmatic arcs and ophiolitic mélanges formed by subduction-accretion processes of the Paleo Asian Ocean. The Beishan area exhibits well-preserved Neoproterozoic to late Paleozoic sequences with intervening ophiolitic zones, and most workers usually divided the orogenic belt into three sub-belts (southern, middle, and northern) by Xiaohuangshan 6

and Niujuanzi-Yueyashan ophiolitic belts, which are southern, middle and northern Beishan (Zuo and He, 1990). The southern Beishan belt is composed of Precambrian strata, Paleozoic volcanic-sedimentary formations, and magmatic intrusions. The upper Paleozoic strata in the southern Beishan are all terrestrial (GSBGMR, 1966), in contrast to the other two belts. The southern Beishan belt is separated from the middle Beishan belt by the Niujuanzi-Yueyashan ophiolite zone. The middle Beishan belt is characterized by early Paleozoic volcanic-sedimentary formations and magmatic intrusions, usually regarded as having resulted from early Paleozoic subduction events (Zuo and He, 1990; Dai et al., 2003; Ao et al., 2010). The middle Beishan belt is separated from the northern Beishan belt by the Xiaohuangshan ophiolitic zone. The northern Beishan belt is relatively complex, and is further subdivided into two zones by Hongshishan ophiolitic belt. The northern zone is mainly composed of lower Paleozoic and magmatic intrusions, however, the southern zone of the northern belt is characterized by Carboniferous strata and late Paleozoic granitoids. The Beishan rift is located in the southern Beishan belt, bordering on the Dunhuang block to the south. The Huaniushan arc is located to the north of the Beishan rift, and bounded by the Gubaoquan-Hongliuyuan shear zone (Fig. 1b; BGMRXUAR, 1993; Xu et al., 2009). The Beishan rift is mainly composed of Carboniferous-Permian strata and magmatic intrusions. Fault-related uplifts and sags are well developed in the rift, and the contact between each pair of strata from Precambrian to Permian is separated by faults (Xu et al., 2009; Su et al., 2012b). The Beishan rift is characterized by exposures of numerous mafic-ultramafic complexes, most of which host Ni-Cu sulfide ore deposits. Mafic-ultramafic complexes such as Poshi, Hongshishan, Bijiashan and Liuyuan 7

complexes, are mainly distributed in the western and central part of the Beishan rift, and intrude the Proterozoic and Carboniferous strata, (Jiang et al., 2006; Qin et al., 2011; Su et al., 2011 a, 2011b; Zhang et al., 2011a). The Yinwaxia area, in which we collected new data, is located in the eastern part of the Beishan rift (Fig. 1b). Strata exposed in the Yinwaxia area are mainly Paleozoic, including upper Silurian, Devonian, Carboniferous and Permian, and the Permian strata are distributed most widely (Fig. 2). The upper Silurian strata in the Yinwaxia area are composed of massive volcanic rocks: basic volcanic rocks dominate the lowest sequence, though the topmost sequence comprises bimodal volcanic rocks, including amygdaloidal basalts and rhyolites (GSBGMR, 1966; Liu et al., 1999). These Silurian basic and acidic volcanic rocks display interbeds, similar to rift-related rock associations. The Permian strata distributed across the southern Beishan include the Shuangbaotang group, the Jinta group (the lower Permian) and the Fangshankou group (the upper Permian). The Shuangbaotang group mainly consists of clastic rocks, including sandstones, pebbly sandstones, limestones, bioclastic limestones. The Jinta group mainly consists of basalts, andesitic basalts, and andesites with siliceous slate and phyllite interlayers. Permian strata in the Yinwaxia area are domanited by the Fangshankou group, which consists predominantly of felsic volcanic and pyroclastic rocks with eruption-explosion facies, including rhyolites, dacites, rhyolite breccia lavas, rhyodacites and homogeneous volcanic tuff. Mafic rocks occur rarely in the Fangshankou group. Additionally, there exist terrestrial plant fossils and other terrestrial materials in the Fangshankou group. Geochemical studies of volcanic rocks in the Fangshankou group demonstrate that they are bimodal volcanic 8

rocks, and formed in a continental rift setting (Liu et al., 1999), and features of Permian sedimentary strata also support the existence of a Permian continental rift. Late Paleozoic granitoids are distributed widely in the Yinwaxia area, and typical examples are biotite granites in the southern Yinwaxia pluton and monzonitic granites in the Xijianquanzi pluton. Previous studies prove that these late Paleozoic granitoids exhibit positive εNd(t) and εHf(t) values, implying additions of depleted mantle (or a juvenile component) in their evolution. These granitoids were mixed products of crustal and mantle derived magmas, and formed in an extensional tectonic setting (Zhang et al., 2010, 2011b, 2012). 3. Field occurrence and petrography Mafic-ultramafic rocks exposed in the Yinwaxia area include basalts, gabbros, and ultramafic rocks (Fig. 2; GSBGMR, 1966). Basalts comprise the upper part of the Permian strata, and the remaining mafic-ultramafic rocks were emplaced into Paleo-Proterozoic strata. The Yinwaxia area ultramafic rocks intruded into the Paleo-Proterozoic strata as an apophysis, including serpentinites and pyroxenites. The mineral assemblages of the serpentinites are mainly serpentine (75–80%), calcite (10–15%), spinel (5–10%) and magnetite (1–5%); they display microscopic crystalloblastic texture and scales crystalloblastic texture microscopically. Mafic rocks in the Yinwaxia area include basalts and gabbros. Yinwaxia area gabbros also intruded into the Paleo-Proterozoic strata. They are dark–light green, and exhibit medium-grained gabbroic texture. The main mineral assemblages of the gabbros are clinopyroxenes and plagioclase, and they also contain minor amphibole, chlorite, magnetite and ilmenite. Many clinopyroxenes underwent 9

intensive amphibolitizations, and plagioclases were usually altered to chlorites. Basalts are a main component of Permian strata in the Yinwaxia area, and display massive structures. These basalts also show porphyritic textures with phenocrysts of plagioclase and pyroxene (0.1–0.2 mm in size) set in a groundmass of plagioclase microlites, granules of pyroxene, and glass. 4. Analytical methods 4.1 SHRIMP zircon analyses Separations of zircon crystals were accomplished by conventional heavy liquid and magnetic techniques. The individual crystals were mounted in epoxy together with the TEMORA standard zircons, and then polished to approximately half thickness. Then the zircons were photographed in reflected and transmitted light as well as SEM cathodoluminescence (CL) images which were all taken at Peking University to study the internal structures in order to identify the suitable target for spot analysis. U-Pb isotopic ratios of zircon crystals were measured using the SHRIMP II in the Beijing SHRIMP Centre, Institute of Geology, Chinese Academy of Geological Sciences, Beijing, China. Instrumental conditions and measurement procedures are the same as those described by Compston et al. (1992). Spots of approximately 20μm-diameter were analyzed. Data for each spot were collected in sets of five scans. The

206

Pb/238U ratios of

the samples were corrected using reference zircon of TEMORA (206Pb/238U = 0.06683; 417 Ma). The data were corrected for common Pb on the basis of the measured decay constants and present-day

238

204

Pb. The

U/235U value given by Steiger and Jager (1977) were

used. Uncertainties given for individual analyses (ratios and ages) are at 1σ level whereas 10

the uncertainties in calculated weighted mean ages are reported as the 95% confidence level. Concordia plots and weighted mean age calculations were carried out using ISOPLOT/Ex 3.23 (Ludwig, 2003). The SHRIMP U-Pb data are reported in Table 1. 4.2 Whole-rock geochemical analyses The major, trace and rare earth elements (REEs) were analyzed at the Laboratory of Orogenic Belts and Crustal Evolution, Peking University. Rock samples for whole rock analyses were crushed and then pulverized in an agate mill. Whole rock major elements were analyzed by X-ray fluorescence (XRF) on fused glass beads, following the analytical procedures of Li et al. (2006) and the analytical precision is within 0.1%. Trace elements were analyzed using Inductively Coupled Plasma Mass Spectrometer (ICP-MS), following the technique of Li (1997). About 50 mg of powder from each sample was dissolved in high-pressure Teflon bombs using a HF+HNO3 mixture. The analytical precision for the common trace elements was superior to 5%, while that of Nb and Ta was superior to 10%. Analytical results are listed in Table 2. Analyses for Sr and Nd isotopes and Sm, Nd, Rb, and Sr concentrations were performed at the Institute of Geology, Chinese Academy of Geological Sciences (IGCAGS), using the Solid Isotope Mass Spectrometer MAT-262 from German Finnigan Corporation. Analytical procedures are described in detail by Yang et al. (2010). During analysis, the NBS-987 standard yielded an average value of

87

Sr/86Sr=0.710274±11 (2σ) and the JMC

standard yielded an average value of 143Nd/144Nd=0.512096±12 (2σ). Mass fractionation of Sr and Nd isotopes were corrected by 86Sr/88Sr =0.1194 and

146

Nd/144Nd = 0.7219. During

analyses, the backgrounds of Rb-Sr and Sm-Nd were 100-300 pg and 50-100 pg, 11

respectively. Analytical results are listed in Table 3. 4.3 Mineral composition analyses Clinopyroxenes were analyzed using a JEOL JXA-8100 wavelength dispersive electron microprobe at Peking University. The operating conditions were 15 kV acceleration voltages with 10 nA beam current and a beam diameter of 1 μm. Analytical results are listed in Table 4. 5. Analytical results 5.1 Geochronology Zircons from the basalt sample (Y-14) have wide ranges of U (79-500 ppm), Th (24-198 ppm) contents and high Th/U ratios (0.23-1.04), which are characteristics of magmatic zircons (Belousova et al., 2002；Wu and Zheng, 2004). Two older 206Pb/238U age (1693.7Ma and 2181.9Ma) were obtained, and the other eight analyses yield

206

Pb/238U

apparent ages of 267.4~314.3Ma. In the CL images (Fig. 3), zircon grains from the basaltic sample display high variable characteristics. Some of them are similar to those in gabbro (i.e. 8.1), and some resemble the magmatic zircon in felsic rocks (i.e. 5.1, 6.1 and 7.1). Different zircon morphologies indicate that most of them (if not none of them) were not formed during basaltic magmatism, but are xenocrystals captured by basaltic magma during its eruption. Thus, the age of the basalt should be constrained by the youngest xenocrystal (~265 Ma, i.e. late Permian), and this is in agreement with the field occurrence. Zircons from the gabbro sample (Y-5) are mostly euhedral and short columnar. Their CL images (Fig. 3) show that they display light color, and parallel banded patterns, composed of light and dark bands, which are typical characteristics of zircons from gabbros 12

(Jian et al., 2003). They have wide ranges of U (63-511 ppm), Th (43-213ppm) contents and high Th/U ratios (0.14-0.86), which are characteristics of magmatic zircons (Belousova et al., 2002；Wu and Zheng, 2004). Two older

206

Pb/238U age (1729.7Ma and 1902.6Ma)

were obtained, and the other seven analyses yield

206

Pb/238U apparent ages of

264.9-297.3Ma, with a weighted mean age of 281±11 Ma (MSWD=3.4, n=7) (Fig.3). 5.2 Whole-rock geochemistry 5.2.1 Basalts The Yinwaxia basalt samples have high SiO2 contents ranging from 46.38 to 54.15 wt. %, and plot in the andesitic basalts and andesite fields on the Nb/Y-Zr/TiO2 diagram (Fig. 4a). The MgO contents and magnesium number (Mg#) values are low (3.14—4.29 wt. % and 36.13—47.59, respectively), suggesting evolving magma. These samples also have relatively low total alkali contents (Na2O+K2O=2.66—6.71 wt. %), and higher Na2O contents (mostly K2O/Na2O=0.30—0.95). They also have relatively low Al2O3 contents (13.03—14.74 wt. %), similar to that of tholeiitic basalts. In addition, they have relatively high total Fe2O3 (10.40—14.11 wt. %, mostly>12 wt. %), TiO2 (2.01—2.88 wt. %) and P2O5 (0.68—1.10 wt. %) abundances, which accord with characteristics of high Fe-Ti basic rocks around the world (FeOt >12 wt. % and TiO2> 2 wt. %; Clague and Bunch, 1976; Clague et al., 1981; Perfit and Fornari, 1983) around the world. In the AFM ternary diagram (Fig. 4b), these basalts exhibit tholeiitic characteristics. These basalts exhibit consistent rare earth elements (REEs) characteristics. They have high total REEs contents (177.89–195.03ppm), and display enrichments in the light REEs (LREEs) relative to middle REEs (MREEs) [(La/Sm)N=1.92–2.38] and heavy REEs 13

(HREEs) [(La/Yb)N=3.63–3.97]. There are essentially no Eu anomalies (δEu=0.93–0.99), indicating

rare

affections

of

plagioclase

fractional

crystallizations.

In

the

Chondrite-normalized REE diagram (Fig. 5a), these basalt samples have similar LREEs contents to that of the ocean island basalt (OIB), but higher MREEs and HREEs contents. Overall, the Yinwaxia basalts have relatively flat REE patterns compare with OIB. In the Primitive mantle-normalized multi-element diagram (Fig. 6a), basalt samples display obvious depletions in the Nb-Ta-Sr-Ti-La-Ce suite. In addition, they have high field strength element (HFSE) contents which are similar to those of OIB, but higher than those of mafic rocks from the Tarim Large Igneous Province (TLIP), the Columbia River basalts and average continental arc rocks (Farmer, 2003; Kelemen et al., 2004; Tian et al., 2010). The Yinwaxia basalts also have higher HREE contents than those of the TLIP, Columbia River, and average continental arc rocks. Relative to HFSEs, there are no obvious enrichments in the large ion lithophile elements (LILEs, e.g., Cs, Rb, Ba, Pb, and Sr). 5.2.2 Gabbros Compared with the basalts, Yinwaxia gabbro samples have higher SiO2 contents (50.37–55.62 wt. %), MgO contents (5.13–6.32wt. %) and Mg# values (50.41–61.49). They also have low total alkali (Na2O+K2O=1.94–5.73 wt. %), CaO (5.60–9.36 wt. %) and Al2O3 contents (13.63–16.06 wt. %). However, the gabbros have relatively low TFe2O3 (8.72–11.08 wt. %), TiO2 (1.32–2.16 wt. %) and P2O5 (0.21–0.36 wt. %) contents. In the Nb/Y-Zr/TiO2 diagram (Fig. 4a), Yinwaxia gabbro samples plot in the andesite field, and display tholeiitic characteristics (Fig. 4b) in the AFM diagram. The Yinwaxia gabbro samples have lower total REE contents (94.69–122.76 ppm) 14

than the basalts. Similarly, they also display enrichments in the LREEs relative to MREEs [(La/Sm)N=2.03–2.72] and HREEs [(La/Yb)N=2.99–3.75]. These gabbro samples display slight negative Eu anomalies, δEu=0.77–0.91, indicating a slight effect of plagioclase fractional crystallization. Compared with OIB, these gabbros have higher HREE contents, and display flatter REE patterns in the Chondrite-normalized REE diagram (Fig. 5b). In the Primitive mantle-normalized multi-element diagram (Fig. 6b), Yinwaxia gabbros display depletions in Nb, Ta, Ti, La and Ce and enrichments in Cs, Rb, U, Pb and Th. They have lower HFSE content than that of basalts, which is a similar pattern to that of mafic rocks from the Tarim Large Igneous province (TLIP) and Columbia River basalts. Relative to HFSEs, there are also no obvious enrichments in LILEs. 5.3 Whole-rock Sr-Nd isotopic composition Whole-rock Rb-Sr and Sm-Nd isotopic compositions for mafic rocks in the Yinwaxia area are listed in Table 3, and plotted in Fig. 7. Initial 87Sr/86Sr and 143Nd/144Nd ratios were calculated using 280 Ma for gabbros, and 260 Ma for basalts. The gabbros have initial 87

Sr/86Sr ratios ranging from 0.706739 to 0.707990, and εNd(t) ranging from −3.42 to −0.68.

The basalts have lower initial Sr isotopic [(87Sr/86Sr)i=0.703774–0.705322], but higher εNd(t) values (4.69–7.28). In Fig.7, samples plot in the ocean island basalts (OIB) field, indicating a mixing trend between depleted mantle and EM2 (enriched mantle type II). 5.4 Clinopyroxenes mineral chemistry Major element analyses of clinopyroxenes from the YInwaxia gabbros are reported in Table 4. Plots of Alz (percentage of tetrahedral sites occupied by Al) vs. TiO2 in augite from gabbroic and ultramafic cumulates show that augite data arrays for arc-related cumulates 15

have a trend of substantially higher Al/Ti in clinopyroxene than that of rift-related tholeiitic rocks (Loucks, 1990). All clinopyroxenes in gabbros from the Yinwaxia study area plot along the rift-related trend (Fig. 8a). In the discrimination diagram (Fig. 8b), all the analyzed clinopyroxenes fall in the compositional fields for the plume-influenced basalts from Iceland and within oceanic plate basalts, distinct from subduction-related settings. 6. Discussion All the rocks in this study underwent strong alteration, consistent with high loss on ignition (LOI) values (1.60-4.92 wt. %). Therefore, the major and trace element geochemistry described here is based on immobile elements during low-temperature alteration and metamorphism up to the greenschist facies (Beccaluva et al., 1979; Pearce and Norry, 1979; Shervais, 1982). Generally, those immobile elements include Al, Ca, Mg, high field strength elements (e.g. Th, Zr, Hf, Nb, Ta, Ti,Y) and REE including Sm-Nd isotopic system. 6.1 Fractional crystallization and crustal contamination No systematic variation trend within the different rock types can be observed in covariations of some selected major elements and their ratios against MgO. In contrast, the good correlation of most of the elements with MgO suggests that common processes controlled the compositions of the different rock types. In particular, positive correlations between CaO, CaO/Al2O3 and MgO support the fractionation of clinopyroxene. TFe2O3 and TiO2 are generally negatively correlated with MgO in all the rocks, suggesting that Fe-Ti oxides played an important role in their evolutions. The presence of phenocrystic plagioclase in most samples also supports significant early crystallization of this mineral, 16

which would account for the coherent negative Sr anomalies observed (Fig. 6), although this seems nominally inconsistent with the absence of obvious negative Eu anomaly in the REE patterns (Fig. 5). Frey et al. (1993) and Xu et al. (2001) have interpreted this phenomenon to result from high Eu3+/Eu2+ ratios in magmas because Eu2+ is compatible with plagioclase, whereas Eu3+ is not. It follows that relatively little Eu may be lost during fractional crystallization of plagioclase in a system with a high oxygen fugacity. Element covariations suggest that clinopyroxene and plagioclase should be the principal fractionating phases in these mafic rocks. Some fractionations of Fe-Ti oxides also occurred. Crustal contaminations could potentially increase SiO2, K2O, Zr, Hf, Th, Cs, Rb and Ba abundances and La/Nb and Zr/Nb ratios, but decrease Ti/Yb and Ce/Pb ratios in mafic magmas (Campbell and Griffiths, 1993; Barker et al., 1997; MacDonald et al., 2001). It is possible to use ratios of highly incompatible elements in mafic rocks to determine these ratios in their mantle source regions, given that elements having similar Kds produce incompatible element ratios that are independent of the degree of partial melting of the mantle source or the amount of subsequent magmatic differentiation (Hanson, 1989; Hofmann, 1997). In addition, Nb is one of the high field strength elements (HFSEs), and generally has low concentrations in the crust. Though La is typically enriched in the crust, Th is commonly enriched in sediments. Therefore, the high (Th/Nb)N ratio (＞1) and low Nb/La ratio (<1) are two reliable indicators for crustal contamination (Saunders et al., 1992; Xia et al., 2007). (Th/Nb)N ratios of Yinwaxia mafic rocks range from 1.51 to 3.90, larger than 1, and their Nb/La ratios range from 0.0029 to 0.084, far away less than 1. In addition, 17

Ti/Y ratios vary from 168 to 307, slightly higher than those of bulk continental crust and Archean bulk crust (160 and 187, respectively; Taylor and McLennan, 1985). Yinwaxia mafic rocks also exhibit low Nb/U (7.43–15.42) ratios, similar to that of continental crust (8.45, Sun and McDonough, 1989). All these ratios indicate that Yinwaxia magmas might have experienced some degree of crustal material contamination. Trends of crustal contamination also appear on the basis of the correlation between εNd(t), Nb/La and Mg number. It is concluded that Yinwaxia mafic rocks experienced some degree of crustal assimilations. 6.2 Characteristics of magma sources Sr and Nd isotopic ratios are usually used for discrimination of mantle sources. Mixing between Depleted Mantle (DM) and Enriched Mantle I (EM1) or Enriched Mantle II (EM2) will produce a negative trend and OIB-like components in a plot of (87Sr/86Sr)i versus εNd(t) (Zinder and Hart, 1986). Sr-Nd isotopic compositions of all the Yinwaxia rocks have typical OIB-like signatures, and clearly show a mixing trend between DM and EM2 components in the genesis of the magmas (Fig. 7). In addition, several previous studies have attributed the mantle heterogeneity, especially the formation of an EM2 reservoir, to subduction-related modification (e.g., Zindler and Hart, 1986; Rollinson, 1993; Turner et al., 1997; Zhou et al., 2004; Su et al., 2012b). Because the EM2 mantle component is considered to be a mantle contaminated by subduction-related modifications (Weaver, 1991; Greeough at al., 2005), it appears that mantle sources of mafic rocks in the Yinwaxia area were intensively depleted and variably enriched by subduction slab-derived components. In fact, previous studies have proved the existence of Paleozoic subduction 18

events. Previous studies (Mao et al., 2012; Su et al., 2012a; Zheng et al., 2012) discovered numerous Paleozoic subduction-related indicators, such as late Ordovician Nb-enriched basalts (451 Ma) in the Liuyuan area (Mao et al., 2012). Therefore, there might be an early Paleozoic arc (Fig. 1b; Xiao et al., 2010) and the early Paleozoic subduction process could have variably modified mantle sources beneath the Beishan rift. Though both basalts and gabbros were significantly contaminated by crustal materials, they exhibit very different εNd(t) values (Fig 7), which could be ascribed to the different crustal materials incorporated into the primitive magmas of mafic rocks (Zhang and Zou, 2012). Contamination by old crust could quickly decrease εNd values of the mafic rocks, while contamination by juvenile crustal materials would have a weak influence on the εNd values (Zhang and Zou, 2012). Therefore, we suggest that Yinwaxia basalts were intensively contaminated by juvenile crustal materials, while Yinwaxia gabbros were contaminated by the old continental crust further, supported by the fact that gabbros intrude into the Paleo-Proterozoic, which comprises of migmatite, metavolcanic rock, and gneiss. Previous studies (Zhang et al., 2010, 2011b, 2012) show that post-collisional and A-type granitoids with ages of 280–250Ma are widely exposed in this region (Fig. 1b). Sr-Nd isotopic compositions of Yinwaxia basalts display a mixing trend of magma derived from depleted mantle with magma forming coeval A-type granites (Fig. 7). Therefore, we speculate that these basalts were contaminated by parental magmas of coeval A-type granites. We conclude that Yinwaxia mafic rocks were derived from lithospheric mantle metasomatized by fluids and/or melts derived from subducted slab materials. Parental 19

magmas underwent clinopyroxene, plagioclase and Fe-Ti oxide fractionations. In addition, basalts were contaminated by parental magmas of coeval A-type granites, while gabbros were contaminated by the older continental crust, and these different contaminants led to different Sr-Nd isotopic characteristics.

6.3 Tectonic setting In the Primitive mantle-normalized multi-element diagram (Fig. 6), Yinwaxia mafic rocks display obvious negative Nb-Ta-Ti anomalies, similar to those of magmatic rocks formed in a subduction zone. Compared to compositions of average continental arc rocks, it is clear that LILEs of the Yinwaxia mafic rocks are lower, while their HFSEs are obviously higher (Fig. 6), indicating that their mantle sources are totally different from those of arc-derived magmatic rocks. However, the Yinwaxia mafic rocks exhibit trace element characteristics similar to those of continental flood basalts having crustal contamination. Because of crustal contamination, continental flood basalts generally exhibit negative Nb-Ta-Ti anomalies, and positive Pb-Rb-Ba-Th-U anomalies, similar to basalts derived from subduction settings (Kelemen et al., 2004). As discussed above, Yinwaxia mafic rocks experienced crustal assimilation, which caused negative Nb-Ta anomalies. In addition, because oxidation-reduction and aqueous solutions of melting sources in the subduction-related setting are different from those of other tectonic settings, we can use some trace elements ratios and covariant relationship diagrams between related elements to distinguish within plate basalts and subduction-related basalts (Pearce and Norry, 1973; Shervais, 1982; Rollinson, 1993). In the tectonic setting discrimination diagrams (Fig. 9), 20

Yinwaxia mafic rocks exhibit different characteristics than basalts in the arc setting, and are instead similar to that of basalts generated within-plate. Granitoids are distributed widely in the southern Beishan, especially in the Yinwaxia area, most notably including the Yinwaxia and Xijianquanzi granites. The Yinwaxia pluton has a LA-ICP-MS U-Pb age of 281.7±2.9Ma, similar to that of the Yinwaxia mafic rocks. The pluton mainly consists of biotite granites which belong to the middle-K, calc-alkaline series with metaluminous-peraluminous characteristics and high SiO2 and (Na2O + K2O) contents; samples invariably exhibit relatively small Chondrite-normalized light rare earth element (LREE) enrichments with flat heavy rare earth elements (HREE), weak negative Eu anomalies, with depletion of Nb, Ba, P, and Ti and enrichment of Rb, Pb, and K in their primitive mantle-normalized trace elements patterns (Zhang et al., 2010, 2011b). Yinwaxia pluton rocks exhibit positive εHf(t) (4.4–7.8) and εNd(t) (0–1.3), and their isotopic data emphasizes the importance of the depleted mantle (or juvenile component) in its genesis. The Yinwaxia granites are the mixed products of crustal and mantle derived magmas and formed under an extensional tectonic setting in the Early Permian. The LA-ICP-MS zircon U-Pb age of the Xijianquanzi granite is 266.1±2.2 Ma, and εHf(t) values are positive, 1.3–4.7. The Xijianquanzi granitoid mainly consists of monzonitic, alkali-rich, high potassium granites with some “A-type like” granite characteristics, e.g., high 10000×Ga/Al values and weakly V-shaped Chondrite-normalized REE patterns. Chemical characteristics of the Xijianquanzi granitoids indicate that they were mixed products of crustal and mantle derived magmas, and were formed in a rift setting in an extensional period. Studies of these granitiods in the Yinwaxia area suggest an extensional setting, and that these granitiods 21

with positive εHf(t) or εNd(t) values formed with strong crust-mantle interactions, similar to other voluminous CAOB granites with positive εHf(t) or εNd(t) values (Han et al., 1997; Wu et al., 2000, 2002; Chen and Jahn, 2004; Hong et al., 2004; Jahn et al., 2004). The Permian strata distributed extensively in the southern Beishan developed numerous faults with NE strike orientation. As noted in section 2, volcanic rocks in the Fangshankou group are bimodal and formed in a continental rift setting (Liu et al., 1999). Sedimentary formations of the Permian strata also support the existence of a Permian continental rift in the Yinwaxia area. In addition, cherts occur along the regional faults (Fig. 1b), suggesting a relatively extensive Permian continental rift. Considering the geochemical data of the Yinwaxia mafic rocks and the coeval granitoids, and the characteristics of the Permian sedimentary formations, we conclude that the Yinwaxia mafic rocks were formed in a continental rift, rather than a subduction setting.

6.4 Tectonic implications The Beishan orogenic belt is considered to be a typical and key area for understanding the formation and tectonic evolution of the central southern Altaids (the CAOB) between the Tianshan and Inner Mongolia (Ao et al., 2010; Xiao et al., 2010; Guo et al., 2012). The CAOB underwent a complex Paleozoic evolution including amalgamation of many disparate tectonic terranes, and some fundamental tectonic problems have remained unsolved, especially for the Beishan rift (Ao et al., 2010; Liu et al., 2011; Su et al., 2011a, b, 2012b; Zhang et al., 2011). To better understand the tectonic evolution of the Beishan rift, we made a geochronological frequency diagram (Fig. 10) using published zircon ages, and 22

from this developed a tectonic evolution model of the Beishan rift (Fig. 11). The age data in the frequency diagram (Fig. 10) were collected from mean and concordant ages of magmatic (mostly) and metamorphic rocks in the Beishan rift and its adjacent region. Only zircon ages measured by more precise analytical methods were compiled for Fig. 10 (Table 5), such as LA-ICP-MS, SHRIMP and SIMS zircon U-Pb ages. Rock types included in the Fig. 10 are ultramafic rocks, mafic-ultramafic rocks, granitoids, dacites, rhyolites and eclogites. In constructing the plots of Fig. 10, we quote the

206

Pb/238U age for zircons,

instead of their mean or concordant ages, and filters were applied to screen out results with unacceptably large analytical errors and unacceptable discordance. Based on these age data quoted, we developed the relative probability plot of magmatic zircons from Paleozoic rocks of the Beishan rift (Fig. 10), and from this plot, we obtained four peak ages, including 534Ma, 450Ma, 415Ma and 282Ma. The peak age of 534 Ma mainly reflects ages of ophiolitic mélanges. Dating results of cumulate gabbros in the Hongliuhe ophiolite suggest it formed 516 ± 7 Ma (Zhang and Guo, 2008). In the eastern part of the same ophiolitic belt, the Yueyashan ophiolite was formed in the early Cambrian, with a SHRIMP U-Pb age of 533±1.7 Ma (Ao et al., 2012). This ophiolitic belt is the oldest one in the Beishan orogenic belt, and their geochemical characteristics of these ophiolites suggest affinities with SSZ-type ophiolites (Ao et al., 2012; Zheng et al., 2012). From these data we conclude that the major ocean, as a branch of the Paleoasian Ocean represented in the Beishan orogenic belt, was formed in the Precambrian (Fig. 11a), and that some micro-continental blocks might have been distributed in this ocean, such as the Mazongshan and Hanshan blocks. Subduction of 23

intra-oceanic lithosphere generated the Hongliuhe-Yueyashan ophiolite in the period 533–516Ma. Subduction between the Mazongshan and Dunhuang blocks continued and generated various magmatic and metamorphic rocks (Fig.11b) with peak ages of 450Ma (Fig. 10). Previous studies reported that U-Pb isotope analyses of zircons from the Liuyuan Nb-enriched basalts and dacites yielded concordant ages of 450.5±3.9 Ma and 441.8±3.1 Ma, respectively. Petrogenetic studies show that the Nb-enriched basalts resulted from partial melting of mantle wedge peridotites, which were previously metasomatized by adakites (Mao et al., 2012). A narrow eclogite zone was reported along the boundary between granitic gneisses and paragneisses in the Gubaoquan-Liuyuan area (Liu et al., 2002), and the U-Pb dating of the Gubaoquan eclogites indicates an Ordovician age of c. 465 Ma for the eclogite facies metamorphism. Petrologic studies proved that eclogites started as oceanic crust in the Palaeoasian Ocean, which was subducted to eclogite depths in the Ordovician (Qu et al., 2011). Occurrences of early Paleozoic eclogites, Nb-enriched basalts and arc-related dacites prove progressive subduction in the southern Beishan area during the early Paleozoic (Liu et al., 2011; Zhang et al., 2011; Mao et al., 2012). In the Beishan rift, there developed a series of Silurian-Devonian post-collisional granitoids whose ages form a peak age of 415Ma in the geochronological frequency diagram (Fig. 10), including the Shuangfengshan A-type granite (415±3Ma, Li et al., 2009), the Huitongshan A-type granite (397±3Ma, Li et al., 2011) and high potassium calc-alkaline granitoids (396-436 Ma, Zhao et al., 2007) in the Liuyuan area. The Shuangfengshan A-type granite is the oldest A-type granite in the Beishan rift yet 24

discovered (Li et al., 2009). Previous studies suggest that these granitoids were formed in a post-collisional setting, and resulted from partial melting of continental crust owing to the under-plating of the mantle-derived magma related to slab break-off (Zhao et al., 2007; Li et al., 2009, 2011). In addition, a molasse sedimentary sequence was found in the upper Devonian strata around the Dundunshan area, also indicating a post-collisional setting in the late Devonian (He et al., 2004). Thrust-nappe structures were reported in the Precambrian to Ordovician strata, which displayed N-S compressions, W-E strike-slips and were formed in the Late Silurian to Early Devonian (Liu et al., 2002). Characteristics of the granitoids and strata in the Beishan rift all confirm a large-scale collisional event in the late stage of early Paleozoic, then Beishan rift came into post-collisional stage (Fig. 11c). The statistical results of the geochronological analysis (Fig. 10) show that magmatic activities were rare in the Beishan rift, and no magmatic rocks have yet been reported with an age of 340–380Ma (Table 5). Rock types of Upper Devonian and Lower Carboniferous units in the area also correspond to the statistical results of the geochronological analysis. The Upper Devonian is composed of keratophyre, breccia, tuff and sandstone, in unconformable contact with the Middle Devonian. The Upper and Middle Devonian units are continental sediments, and constitute the molasse formation in the Beishan rift (He et al., 2004). There are also rare magmatic rocks in the Lower Carboniferous section, which is mainly composed of phyllite, limestone, sandstone and conglomerate and characterized by shallow-marine to nonmarine sedimentary rocks (GSBGMR, 1966). We speculate that this period (340–380Ma) included continued orogenesis (the late Devonian) and marine regression (the early Carboniferous) processes, with rare magmatic activities. 25

In recent studies, Zhang et al. (2008) collected age data of 18 mafic and 21 granitic rocks from Tarim and its marginal areas, and obtained an age span of 260–320 Ma with a peak age of 275 Ma. In addition, Qin et al. (2011) collected zircon U-Pb age data from basalts and mafic dykes in the Beishan rift, eastern Tianshan and Tarim basin, and obtained a peak age of 280 Ma. In the current study, we collected zircon U-Pb ages of magmatic rocks including mafic-ultramafic rocks, diorites, granitoids and rhyolites in the Beishan rift, and obtained a similar peak age of 282 Ma (Fig. 10). It is obvious that an important magmatic event took place near 280Ma in the Beishan rift, even in the NW China. Most studies suggest that magmatic rocks of this period formed in the within-plate setting (Mao et al., 2006, 2008; Qin et al., 2011; Su et al., 2011a, 2011b, 2012b; Zhang et al., 2011a), though there are still controversies concerning the geodynamics models, including plume-related magmatisms (Qin et al., 2011; Su et al., 2012b) and lithospheric delamination (Zhang et al., 2011a). As discussed above, Yinwaxia mafic rocks exhibit obvious geochemical differences from mafic rocks in the Tarim Large Igneous Province (TLIP). Specifically, Yinwaxia mafic rocks have obvious negative Nb-Ta-Ti anomalies, and have higher HFSE and HREE abundances, which indicate that they derived from partial melting of spinel lherzolites with relatively low melting degrees, unlikely typical mafic rocks associated with plume. In addition, Yinwaxia mafic rocks also display different Sr-Nd isotopic compositions (Fig. 7). It is notable that distributions of mafic-ultramafic rocks in the Beishan area are linear along faults, rather than having the planar distribution expected of a mantle plume (Fig. 1b). For example, the Poyi, Poshi and Luodong intrusions occur along the Baidiwa fault (Mao et al., 2008; Pirajno et al., 2008), and the Liuyuan and 26

Yinwaxia mafic rocks occur along the Hongliuyuan fault (Zhang et al., 2011a; Cai et al., 2012). It is also totally different from magmatic activities associated with a mantle plume, which are always planar distributions (Campbell and Griffiths, 1990; Renne and Basu, 1991; Coffin and Eldholm, 1994). According to discussions above, it is hard to attribute formations of Yinwaxia mafic rocks to magmatic activities associated with a mantle plume. On the basis of the geochemical characteristics of Yinwaxia mafic rocks and a regional tectonic evolutionary model using regional geochronological data, we propose that the Yinwaxia mafic-ultramafic rocks resulted from magmatism associated with lithospheric delamination and asthenosphere upwelling in a continental rift setting. In this scenario, the Mazongshan and Dunhuang blocks collided in the late stage of the early Paleozoic, thickening the lithosphere of the Beishan area, leading to a gravitational instability, which provides the driving force for the lithosphere delamination. Afterwards the asthenosphere upwelled and provided heat leading to partial melting of the lithopheric mantle, which had been metasomatized by fluids or melts derived from subducted slab. In the course of rifting in the late Carboniferous-Permian, the partial melting continued regionally, and parental magmas of mafic and ultramafic rocks were produced owing to decompression. Consequently, undergoing fractional crystallization (AFC) processes, parental magmas were emplaced along faults and formed mafic and ultramafic rocks in the Beishan rift (Fig. 11d).

7. Conclusion The Yinwaxia mafic rocks formed in the Permian (265 Ma and 281 Ma). The overall 27

isotopic compositions of these mafic rocks imply an OIB-like mantle source, display DM-EM2 mixed trends, and indicate derivation from lithospheric mantle metasomatized by fluids and/or melts derived from subducted slab. Parental magmas underwent fractionations of clinopyroxene, plagioclase and Fe-Ti oxides. Yinwaxia basalts were contaminated by parental magmas of coeval A-type granites, while gabbros were contaminated by the older continental crust, causing different Sr-Nd isotopic characteristics for these rock types. Taking into account the geochemistry of the Yinwaxia mafic rocks and coeval granitoids, and characteristics of the Permian sedimentary formations, we conclude that Yinwaxia mafic rocks were formed in a continental rift, and derived from magmatism associated with lithospheric delamination and asthenosphere upwelling in a continental rift setting.

Acknowledgement This study was financially supported by the National Natural Science Foundation of China (Grant no. 41372225). We sincerely thank Wenjiao Xiao, Keda Cai and one anonymous reviewer for their constructive comments that greatly helped to improve the manuscript. We also appreciate the assistances of Libing Gu, Fang Ma, Hangqiang Xie, Jinrong Li and Qiannan Li for geochronological and geochemical analyses.

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List of Figures Fig. 1 (a) Geological sketch map of the Central Asian Orogenic Belt (modified after Şengör et al., 1993; Jahn et al., 2000). (b) The geological map of the Beishan rift and its adjacent region (modified after Wang et al., 2007; age data are from Zhao et al., 2007; Li et al., 2009, 2011, 2013; Su et al., 2011 and references therein; Zhang et al., 2011a; Mao et al., 2012; Jiang et al., 2013).

Fig. 2 The geological map of the Yinwaxia area (modified after GSBGMR, 1966).

Fig.3 Cathodoluminescence (CL) images and SHRIMP dating results of zircons from the basalt and the gabbro in the Yinwaxia area.

Fig. 4 (a) The Zr/TiO2 versus Nb/Y chemical classification diagram for mafic rocks from the Yinwaxia (after Winchester and Floyd, 1977); (b) AFM diagram (A=Na2O+K2O, F=FeO, M=MgO) for mafic rocks from the Yinwaxia area. The boundary line between tholeiitic and calc-alkaline rock types is from Miyashiro (1974).

Fig. 5 The Chondrite-normalized REE pattern of basalts (a) and gabbros (b) from the Yinwaxia area (normalized data and N-MORB values are from Sun and McDonough, 1989). Mafic rocks in the Tarim large igneous province are basalts from boreholes at depths between 5,166 and 6,333 m in the northern Tarim uplift (Tian et al., 2010). 47

Fig.6 The N-MORB-normalized trace elements spiderdiagram of basalts (a) and gabbros(b) from the Yinwaxia area (normalized data and E-MORB values are from Sun and McDonough, 1989).Mafic rocks in the Tarim large igneous province are basalts from boreholes at depths between 5,166 and 6,333 m in the northern Tarim uplift (Tian et al., 2010). Data of basalts in the East African Rift, average continental arc and OIB are collected from Shinjo et al. (2011), Kelemen et al. (2004) and Sun and McDonough (1989).

Fig. 7 εNd(t) vs. (87Sr/86Sr)i diagram for mafic rocks from the Yinwaxia area (Zindler and Hart,1986). Data for mafic rocks in the western part of Beishan rift are from Su et al., (2012). Data for mafic rocks in the Liuyuan area are from Zhang et al. (2011a). Data for granites in the Yinwaxia area are from Zhang et al. (2011b) and Feng et al. (2012). DMM, depleted mantle-derived melt; EM1, enriched mantle1; EM2, enriched mantle2; εNd(280Ma)=10.96 and (87Sr/86Sr)i=0.7022 are values of DMM.

Fig. 8 (a) Plots of Alz (percentage of tetrahedral sites occupied by Al) vs. TiO2 in clinopyroxenes from gabbros of the Yinwaxia area (after Loucks, 1990). (b) The TiO2-Na2O-SiO2/100 diagram, using compositions of clinopyroxenes for discriminating different tectonic settings (after Beccaluva et al., 1989). Abbreviations, E-MORB: enriched mid-ocean ridge basalt; N-MORB: normal mid-ocean ridge basalt; WOPB: within oceanic plate basalts; ICB: Iceland basalts; SSZ: supra-subduction 48

zone basalts.

Fig. 9 Tectonic setting discrimination diagrams for Yinwaxia mafic rocks. (a) Zr/Y-Zr diagram (after Pearce and Norry, 1979); (b) Ti–Zr diagram (after Pearce, 1982).

Fig. 10 Relative probability plots of zircon

206

Pb/238U ages from magmatic and

metamorphic rocks in the Beishan rift and its adjacent area.

Fig. 11 The tectonic evolution model for the Beishan rift. See text for detailed discussions.

List of tables Table 1 SHRIMP U-Pb data of the basalt (Y-14) and the gabbro (Y-5) in the Yinwaxia area. Notes: Errors are 1-sigma; Pbc and Pb* indicate the common and radiogenic lead, respectively. Common Pb corrected using measured 204Pb. Table 2 Major and trace elements compositions of basalts and gabbros in the Yinwaxia area. Table 3 Whole rock isotopic compositions of basalts and gabbros in the Yinwaxia area. Table 4 Major element analyses of clinopyroxenes from gabbros of the Yinwaxia area. Table 5 Compiled ages of magmatic and metamorphic rocks in the Beishan rift and its adjacent area.

49

Table 1

Spots

U

Th

(ppm)

(ppm)

206

Th/U

Pb*

Age(Ma)

206

Pbc

207

*

206

*

Pb

Pb

±%

(ppm)

(%)

8.9

1.25

0.0557

9.9

Pb

207

*

±%

235

206

Pb

*

±%

238

U

206

Pb

238

U

U

207

1σ

Pb

206

Pb

1σ

The gabbro sample ( Y-5) 1.1

216

174

0.84

0.36

10.2

0.0472

2.5

297.3

7.3

441

220

2.1

63

52

0.86

2.6

5.69

0.0576

49.7

0.35

50.0

0.0445

6.0

280.9

16.4

516

1091

3.1

303

42

0.14

11.5

1.57

0.0471

8.8

0.28

9.1

0.0437

2.1

275.8

5.7

54

210

4.1

263

55

0.22

77.9

0.24

0.1139

1.7

5.39

2.6

0.3433

2.0

1902.6

32.2

1862

30

5.1

398

213

0.55

14.4

0.56

0.0530

6.1

0.31

6.5

0.0420

2.0

264.9

5.3

328

139

6.1

511

196

0.40

19.8

1.54

0.0542

5.2

0.33

5.8

0.0445

2.6

280.5

7.0

380

118

7.1

133

58

0.45

5.2

1.02

0.0470

3.7

0.29

4.3

0.0454

2.2

286.2

6.1

50

87

8.1

117

64

0.56

4.7

0.79

0.0470

7.5

0.30

7.8

0.0468

2.2

295.0

6.4

49

179

9.1

162

81

0.52

43.0

0.06

0.1032

0.8

4.38

2.2

0.3078

2.0

1729.7

30.5

1683

15

0.31

27.6

0.67

.1437

2.1

7.98

3.0

.4028

2.2

2181.9

41.3

2273

36

The basalt sample ( Y-14) 1.1

79

24

2.1

201

183

0.94

7.2

2.44

.0510

16.0

0.29

16.2

.0406

2.4

256.7

6.0

242

369

3.1

500

112

0.23

21.6

0.66

.0516

2.9

0.36

3.5

.0500

1.9

314.3

5.8

268

66

4.1

233

139

0.62

9.3

1.37

.0487

9.5

0.31

9.7

.0456

2.1

287.3

5.9

136

222

5.1

241

152

0.65

8.9

1.88

.0479

6.9

0.28

7.2

.0423

2.1

267.4

5.5

92

163

6.1

133

58

0.45

5.2

1.02

.0470

3.7

0.29

4.3

.0454

2.2

286.2

6.1

50

87

7.1

117

64

0.56

4.7

0.79

.0470

7.5

0.30

7.8

.0468

2.2

295.0

6.4

49

179

8.1

197

198

1.04

7.8

1.81

.0466

24.1

0.29

24.2

.0454

2.5

286.2

7.1

27

577

9.1

200

117

0.60

8.2

4.55

.0318

32.9

0.20

33.0

.0454

2.5

286.1

6.9

-983

972

10.1

142

63

0.46

36.6

0.00

.1133

1.2

4.69

2.4

.3005

2.1

1693.7

31.1

1852

22

Notes: Common Pb corrected using the measured

204

Pb. Pb* indicate the radiogenic lead, Pbc indicate the common lead.

50

Table 2 No.

Y-1

Y-2

Y-3

Y-4

Y-5

Y-6

Y-7

Y-8

Y-9

Rocks

Gabbro

Gabbro

Gabbro

Gabbro

Gabbro

Gabbro

Gabbro

Basalt

Basalt

SiO2

55.62

52.81

52.96

53.33

51.70

53.27

50.37

53.29

54.15

1.32

1.335

2.56

2.56

15.27

13.63

14.74

14.53

TiO2 Al2O3

1.59 15.48

1.38 16.06

2.06 15.31

2.16

1.36

15.25

15.27

TFe2O3

9.54

9.28

11.04

11.08

9.20

8.96

8.72

12.89

11.98

MnO

0.16

0.15

0.20

0.16

0.15

0.12

0.164

0.24

0.22

6.16

6.32

3.72

3.14

MgO

5.46

6.22

5.36

5.13

6.26

CaO

5.60

9.36

6.70

6.45

8.96

7.49

8.46

6.35

6.59

Na2O

2.79

1.00

1.80

1.87

2.33

2.22

4.97

1.64

1.51

K2O

0.52

0.93

0.88

0.90

0.91

1.19

0.76

1.35

1.69

P2O5

0.27

0.21

0.34

0.36

0.23

0.23

0.237

0.78

0.80

LOI

2.83

2.47

3.21

3.20

3.52

3.64

4.92

2.29

2.72

99.89

Total Mg#

99.85 55.75

99.88 59.59

99.88 51.66

99.89 50.49

99.88 59.99

99.87 60.21

61.49

99.86 38.81

99.89 36.59

Sc

31.0

30.9

36.3

34.0

33.8

26.5

36.1

26.3

25.8

152

220

174

177

V

183

165

230

225

178

Co

33.8

35.8

30.6

31.4

35.3

32.8

34.2

21.8

18.1

Ga

17.9

17.2

18.0

18.7

17.4

17.1

19.0

22.9

21.4

14.7

10.7

8.05

7.37

Li

13.9

14.5

15.7

16.5

14.5

Be

1.29

1.03

1.05

1.33

1.03

1.32

1.129

1.66

1.41

Cs

0.38

0.41

0.58

0.68

0.90

0.73

0.8182

0.89

0.80

43.4

20.7

14.7

17.5

Rb

3.71

72.3

15.6

13.8

11.6

Sr

232

341

197

187

406

347

588

220

331

Ba

222

107

164

118

102

125

175

226

254

3.97

2.48

2.77

2.52

1.07

1.13

0.923

0.864

Th

3.54

2.55

2.49

3.00

2.65

U

1.30

0.825

0.741

0.934

0.813

1.15

0.855

Ta

0.593

0.462

0.581

0.591

0.434

0.573

0.417

Nb

8.84

7.02

8.77

9.30

7.19

8.55

7.38

14.6

13.5

Hf

6.34

4.84

5.48

6.31

5.02

6.39

6.74

8.53

7.96

303

248

406

377

Zr

299

228

262

299

229

Y

40.1

33.1

40.8

42.0

34.3

37.4

40.9

63.3

61.7

Pb

11.8

11.7

15.9

4.86

7.77

9.55

10.7

7.84

6.46

15.3

14.4

24.7

23.5

La

15.5

12.4

12.2

15.8

12.5

Ce

39.5

31.5

33.4

40.5

32.3

38.4

37.2

61.2

58.8

Pr

5.09

4.02

4.52

5.25

4.16

4.86

5.10

8.23

7.97

22.1

22.9

39.8

38.9

Nd

23.5

18.6

22.1

24.6

19.5

Sm

5.96

4.80

5.99

6.37

5.08

5.61

5.85

10.2

10.1

Eu

1.74

1.52

1.95

1.89

1.62

1.51

1.79

3.47

3.37

6.49

6.84

11.9

11.7

1.92

1.89

12.0

11.7

Gd

7.04

5.62

7.11

7.43

6.02

Tb

1.16

0.943

1.18

1.24

1.01

1.08

1.14

Dy

7.42

5.99

7.51

7.79

6.39

6.99

7.21

51

Ho

1.52

1.23

1.54

1.59

1.31

1.43

1.47

2.43

2.37

Er

4.50

3.65

4.45

4.64

3.87

4.26

4.30

7.02

6.83

Tm

0.642

0.513

0.622

0.662

0.551

0.622

0.612

0.994

0.953

6.51

6.19 0.914 185.19

Yb

4.25

3.41

4.08

4.34

3.64

4.08

3.97

Lu

0.621 118.45

0.494 94.69

0.591 107.25

0.632 122.76

0.533 98.49

0.592 113.33

0.584 113.33

0.972 191.35

3.65

3.64

2.99

3.64

3.43

3.75

3.62

3.79

3.79

0.82

0.89

0.91

0.84

0.86

0.77

0.87

0.97

0.95

∑REE （La/Nb）N δEu

Table 2 continued No. Y-10

Y-11

Y-12

Y-13

Y-14

Y-15

Y-16

Y-17

Rocks

Basalt

Basalt

Basalt

Basalt

Basalt

Basalt

Basalt

Basalt

SiO2

51.46

51.73

53.90

52.82

51.77

53.18

53.33

47.38

TiO2

2.77

2.69

2.57

2.88

1.99

2.10

2.05

2.62

Al2O3

14.00

14.16

14.66

14.65

14.55

13.74

13.03

13.73

TFe2O3

13.61

13.72

12.94

13.23

11.00

11.61

10.40

14.11

MnO

0.24

0.28

0.25

0.23

0.21

0.21

0.21

0.20

MgO

4.03

4.08

3.33

3.55

3.90

3.93

4.29

3.99

CaO

6.56

6.43

5.84

5.26

6.43

6.10

6.84

6.59

Na2O

2.50

1.59

1.61

2.05

4.40

4.45

3.81

4.14

K2O

1.04

1.07

1.54

1.72

1.30

2.26

1.96

1.36

P2O5

0.82

0.86

0.80

0.88

0.75

0.69

0.68

1.10

LOI

2.87

3.28

2.43

2.41

3.59

1.60

3.30

4.69

Total Mg#

99.89 39.44

99.88 39.53

99.87 36.13

99.68 37.13

99.90 43.84

99.89 42.70

99.9. 47.59

99.91 38.39

Sc

27.5

26.0

26.2

27.4

28.0

28.6

27.7

35.1

V

209

176

170

179

174

197

188

336

Co

24.5

23.4

19.3

22.9

18.0

20.8

17.8

37.5

Ga

21.2

21.2

22.7

22.8

24.7

25.5

23.6

23.1

Li

8.89

9.64

8.72

10.1

9.11

5.27

6.64

22.2

Be

1.56

1.41

1.54

1.45

1.40

1.64

1.46

1.07

Cs

0.462

0.361

1.00

0.794

1.31

0.711

0.911

3.43

Rb

11.0

10.8

16.8

16.9

13.0

21.7

23.1

31.4

Sr

239

290

220

299

317

332

266

346

Ba

150

131

265

250

442

379

247

179

Th

2.32

2.47

2.82

2.63

2.82

2.99

2.70

1.56

U

0.872

0.903

0.921

1.13

1.31

1.17

0.878

0.561

Ta

0.813

0.851

0.872

0.864

0.749

0.813

0.427

0.554

Nb

12.8

13.6

14.2

13.9

13.8

14.2

10.9

8.64

Hf

7.44

7.59

8.71

8.24

10.4

11.2

10.0

7.27

Zr

353

363

418

388

388

421

377

238

Y

59.8

59.9

64.6

63.9

70.9

72.1

66.1

55.7

Pb

6.18

8.29

11.3

11.2

10.7

9.45

10.7

16.9

La

22.0

22.6

24.5

24.5

25.0

25.9

24.5

19.1

52

Ce

55.8

56.8

61.9

62.1

65.6

65.9

61.7

50.2

Pr

7.63

7.72

8.32

8.39

9.06

9.01

8.35

7.18

Nd

37.6

37.6

40.2

41.0

45.4

43.9

41.3

37.1

Sm

9.84

9.76

10.3

10.7

11.4

11.2

10.5

9.91

Eu

3.25

3.25

3.45

3.58

3.69

3.74

3.59

3.45

Gd

11.6

11.5

12.1

12.4

12.9

12.9

12.0

11.4

Tb

1.86

1.85

1.96

1.99

2.12

2.13

1.95

1.79

Dy

11.6

11.5

12.2

12.3

13.0

13.1

12.1

10.7

Ho

2.32

2.32

2.48

2.49

2.61

2.65

2.44

2.09

Er

6.61

6.69

7.25

7.14

7.56

7.69

7.05

5.78

Tm

0.924

0.933

1.02

0.992

1.05

1.09

0.992

0.774

Yb

5.99

6.09

6.67

6.47

6.90

7.06

6.47

4.81

Lu

0.872 177.89

0.891 179.50

0.983 193.33

0.952 195.03

1.03 207.43

1.05 207.24

0.961 193.86

0.692 164.90

3.67

3.71

3.67

3.78

3.63

3.67

3.79

3.97

0.93

0.94

0.95

0.95

0.93

0.95

0.98

0.99

∑REE （La/Nb）N δEu

Note: Major elements are analyzed using XRF (in wt %), trace elements using ICP-MS (in ppm)

53

Table 3 Sample No.

Y-8 Y-9 Y-10 Y-11 Y-12 Y-6 Y-7 Y-5 Y-2 Y-1

87

Rb/86Sr

Rb [ppm] 13.4

Sr [ppm] 204.3

0.1895

17.2

265.6

0.1874

9.75

189.8

0.1487

87

Sr/86Sr

(87Sr/86Sr)i

0.704475

Error (2σ) 0.000012

0.706015

0.000013

0.704405

0.000015

147

Sm/144Nd

0.703774

Sm [ppm] 10.4

Nd [ppm] 39.6

0.1598

0.705322

10.5

39.2

0.1617

0.703854

9.20

34.7

0.1608

(143Nd/144Nd)i

εNd(t)

0.512938

Error (2σ) 0.000011

0.512644

7.09

0.512945

0.000010

0.512648

7.16

0.512950

0.000012

0.512654

7.28

143

Nd/144Nd

10.4

198.4

0.1516

0.704556

0.000014

0.703996

9.52

35.9

0.1605

0.512949

0.000011

0.512654

7.27

16.4

196.5

0.2422

0.704718

0.000013

0.703822

10.3

38.8

0.1605

0.512817

0.000013

0.512521

4.69

41.4

345.8

0.3612

0.709434

0.000011

0.707990

5.61

22.10

0.1359

0.512351

0.000015

0.512101

-3.41

20.5

568.0

0.1018

0.707188

0.000015

0.706781

5.71

22.4

0.1276

0.512410

0.000014

0.512175

-1.97

11.4

407.3

0.0824

0.708110

0.000015

0.707781

5.18

19.71

0.1405

0.512500

0.000010

0.512241

-0.68

71.3

336.1

0.6137

0.709193

0.000012

0.706739

4.60

18.53

0.1283

0.512463

0.000011

0.512227

-0.96

3.51

231.2

0.0462

0.707304

0.000011

0.707119

5.76

23.13

0.1385

0.512497

0.000012

0.512242

-0.66

54

Table 4 Sample

1.1

2.1

2.2

3.1

3.2

4.1

4.2

9.2

SiO2

51.13

51.58

52.16

50.83

50.9

51.17

51.37

51.39

TiO2

0.93

0.88

0.87

0.79

0.8

0.94

0.99

0.86

Al2O3

3.15

2.28

2.42

3.17

2.3

1.93

2.17

2.46

Cr2O3

0.62

0.03

0.12

0.76

0.45

0.02

0

0.27

Fe2O3

2.31

1.66

1.87

2.05

2.57

2.05

1.78

2.42

FeO

4.5

5.85

5.46

4.35

4.22

6.42

6.3

4.45

MnO

0.13

0.23

0.21

0.19

0.2

0.24

0.23

0.14

MgO

16.25

15.93

16.38

16.33

16.17

15.55

15.8

15.72

CaO

20.67

20.73

20.84

20.35

20.81

20.26

20.21

21.7

Na2O

0.41

0.31

0.33

0.37

0.35

0.36

0.36

0.39

K2O

0

0

0.02

0.02

0.02

0.01

0.02

0

Totals

100.1

99.48

100.68

99.21

98.79

98.94

99.23

99.8

Cations normalized to 6 oxygens Si

1.88

1.913

1.909

1.883

1.897

1.915

1.913

1.899

Ti

0.026

0.025

0.024

0.022

0.022

0.026

0.028

0.024

Al

0.137

0.1

0.104

0.138

0.101

0.085

0.095

0.107

Cr

0.018

0.001

0.003

0.022

0.013

0.001

0

0.008

3+

0.064

0.046

0.051

0.057

0.072

0.058

0.05

0.067

2+

Fe

0.138

0.181

0.167

0.135

0.131

0.201

0.196

0.138

Mn

0.004

0.007

0.007

0.006

0.006

0.008

0.007

0.004

Mg

0.89

0.881

0.893

0.901

0.898

0.867

0.877

0.866

Ca

0.814

0.824

0.817

0.808

0.831

0.813

0.807

0.859

Na

0.029

0.022

0.023

0.027

0.025

0.026

0.026

0.028

K

0

0

0.001

0.001

0.001

0

0.001

0

Sum

4

4

4

4

4

4

4

4

Wo

42.69

42.63

42.35

42.48

42.99

41.91

41.79

44.52

En

46.7

45.58

46.32

47.43

46.48

44.75

45.46

44.87

Fs

10.61

11.79

11.33

10.09

10.53

13.34

12.75

10.61

Fe

Continued Sample

4.3

5.1

5.2

6.1

6.2

6.3

7.1

7.2

SiO2

52.03

52.15

52.24

52.41

51.63

51.07

50.92

51.67

TiO2

0.85

0.71

0.72

0.73

0.84

1.05

0.85

0.81

Al2O3

2.23

2.45

2.14

2.04

2.01

2.04

2.2

2.51

Cr2O3

0.21

0.41

0.26

0.44

0.01

0

0.05

0.57

Fe2O3

1.53

1.19

0.05

0.94

2.43

1.11

2.69

2.6

FeO

5.44

5.1

6.44

5.46

5.54

7.72

5.27

3.98

55

MnO

0.21

0.17

0.14

0.18

0.25

0.25

0.15

0.17

MgO

16.49

16.4

16.23

16.67

15.96

15.11

15.77

16.42

CaO

20.55

20.92

20.53

20.57

20.53

19.82

20.72

20.97

Na2O

0.34

0.36

0.28

0.35

0.42

0.37

0.34

0.48

K2O

0.01

0.01

0

0

0

0

0

0

Totals

99.89

99.87

99.04

99.78

99.62

98.54

98.96

100.18

Cations normalized to 6 oxygens Si

1.916

1.918

1.939

1.929

1.914

1.923

1.902

1.896

Ti

0.024

0.02

0.02

0.02

0.023

0.03

0.024

0.022

Al

0.097

0.106

0.094

0.089

0.088

0.091

0.097

0.109

Cr

0.006

0.012

0.008

0.013

0

0

0.001

0.017

0.042

0.033

0.001

0.026

0.068

0.031

0.075

0.072

3+

Fe

2+

Fe

0.168

0.157

0.2

0.168

0.172

0.243

0.165

0.122

Mn

0.007

0.005

0.004

0.006

0.008

0.008

0.005

0.005

Mg

0.905

0.899

0.898

0.914

0.882

0.848

0.878

0.898

Ca

0.811

0.824

0.816

0.811

0.815

0.8

0.829

0.825

Na

0.024

0.026

0.02

0.025

0.03

0.027

0.025

0.034

K

0

0

0

0

0

0

0

0

Sum

4

4

4

4

4

4

4

4

Wo

42.1

43.08

42.62

42.25

42.1

41.6

42.58

43.02

En

47

47

46.88

47.64

45.54

44.12

45.09

46.87

Fs

10.9

9.92

10.51

10.11

12.37

14.28

12.33

10.12

Sample

7.3

8.1

8.2

9.1

SiO2

51.46

50.85

51.32

51.73

TiO2

0.74

0.96

0.96

0.91

Al2O3

2.3

3.17

2.34

2.2

Cr2O3

0.56

0.46

0.03

0.07

Fe2O3

2.32

1.41

2.22

1.36

FeO

3.96

5.03

5.6

6.13

MnO

0.21

0.14

0.22

0.18

MgO

16.9

15.94

15.67

15.95

CaO

20.48

20.43

20.82

20.57

Na2O

0.36

0.4

0.39

0.34

K2O

0

0.02

0

0

Totals

99.29

98.81

99.56

99.45

Continued

Cations normalized to 6 oxygens Si

1.902

1.892

1.905

1.919

Ti

0.021

0.027

0.027

0.025 56

Al

0.1

0.139

0.102

0.096

Cr

0.016

0.014

0.001

0.002

3+

0.065

0.04

0.062

0.038

2+

Fe

0.122

0.156

0.174

0.19

Mn

0.007

0.004

0.007

0.006

Mg

0.931

0.884

0.867

0.882

Ca

0.811

0.814

0.828

0.818

Na

0.026

0.029

0.028

0.024

K

0

0.001

0

0

Sum

4

4

4

4

Wo

42.04

42.99

42.88

42.41

En

48.27

46.67

44.91

45.76

Fs

9.69

10.34

12.21

11.83

Fe

57

Table 5 Location

Rock type

Age (Ma)

Analytical method

Data source

Xuanwoling

Gabbro

260.7±2.0

SIMS

Su et al., 2010a

Poshi

Olivine gabbro

275.5±1.2

SIMS

Ao, 2010

Poshi

Olivine gabbro

284±2.2

SIMS

Qin et al., 2011

Poyi

Alkaline granite vein

251.4±1.4

SIMS

Su et al., 2010b

Poyi

Gabbro

271±6.2

SIMS

Ao, 2010

Luodong

Gabbro

284.0±2.3

SIMS

Su et al., 2010b

Luodong

Gabbro

283.8±1.1

LA-ICP-MS

Ao, 2010

Hongshishan

Olivine gabbro

281.8±2.6

LA-ICP-MS

Ao et al., 2010

Hongshishan

Troctolite

286.4±2.8

SIMS

Su et al., 2010c

Hongshishan

Diorite

279.7±4.8

SIMS

Qin et al., 2011

Hongshishan

Dacite

279.1±2.9

SIMS

Qin et al., 2011

Hongshishan

Ryholite

321.7±3.4

SIMS

Su et al., 2011

Bijiashan

Gabbro

279.2±2.3

SIMS

Qin et al., 2011

Liuyuan

Nb-enriched basalts

451.6±4.4

SIMS

Mao et al., 2012

Liuyuan

Dacites

442.23±3.1

SIMS

Mao et al., 2012

Liuyuan

Diorite

272.7±4.4

SHRIMP

Zhang et al., 2011

Liuyuan

Diorite

291.4±4.9

SHRIMP

Zhang et al., 2011

Liuyuan

Ultramafic rocks

250.4±9.0

SHRIMP

Zhang et al., 2011

Liuyuan

Granodiorite

423±8

SHRIMP

Zhao et al., 2007

Liuyuan

Monzonitic granite

396±15

SHRIMP

Zhao et al., 2007

Liuyuan

K-feldspar granite

436±9

SHRIMP

Zhao et al., 2007

Shuangfengshan

A-type granite

415±3

LA-ICP-MS

Li et al., 2009

Huitongshan

A-type granite

397±3

LA-ICP-MS

Li et al., 2011

Gubaoquan

Eclogite

465±10

LA-ICP-MS

Liu et al., 2011

Hongliuhe ophiolite

Cumulate gabbro

516.2±7.1

SHRIMP

Zhang and Guo, 2008

Hongliuhe ophiolite

Biotite granite

404.8±5.2

SHRIMP

Zhang and Guo, 2008

Qiaowan

Granodiorite

303.7±2.4

LA-ICP-MS

Feng et al., 2012

58

Yinwaxia

Biotite granite

281. 7 ± 2. 9

LA-ICP-MS

Zhang et al., 2011

Xijianquanzi

Monzonitic granite

266.1±2.2

LA-ICP-MS

Zhang et al., 2010

Yueyashan ophiolite

Plagiogranite

533±1.7

SIMS

Ao et al., 2011

Yueyashan ophiolite

Plagiogranite

536±7

SHRIMP

Hou et al., 2012

59

60

61

62

63

64

65

66

67

68

69

70

Highlights 

Yinwaxia mafic rocks were formed in Permian (265 Ma and 281 Ma).



They derived from lithospheric mantle metasomatized by fluids and/or melts.



Yinwaxia mafic rocks formed in a continental rift.

71