2. Key Laboratory of Mineral Resources Evaluation in Northeast China, Ministry of Land and Resources, Changchun 130061, China;
3. Department of Earth Sciences, ETH Zurich, Zurich 8092, Switzerland;
4. Physics of Geological Processes, University of Oslo, Oslo 0316, Norway
The central Asian orogenic belt (CAOB) is a giant accretionary orogenic belt among the Siberian Craton, the North China Craton and the Tarim Craton, which is composed of a series of island arcs, forearc or back-arc basins, ophiolitic belts and microcontinents of Neoproterozoic to Mesozoic (Windley et al., 2007; (Sengör and Natal'in, 1996; Sengör et al., 1993). This orogenic belt, as one of the most complex and the biggest accretionary collages, has caused juvenile crustal growths in the Phanerozoic (Safonova and Santosh, 2014; Xiao and Santosh, 2014; Rytsk et al., 2011; Safonova et al., 2011, 2009; Wu F Y et al., 2011, 2005, 2003a, b, 2002, 2000; Kröner et al., 2010, 2007; Turkina et al., 2007; Jahn, 2004; Jahn et al., 2001, 2000). The long-lasting crustal growth processes documented the breakup of the Rodinian supercontinent and the formation of the Eurasian supercontinent (Li et al., 2008; Torsvik and Cocks, 2004; Dobretsov et al., 2003). Northeast China, tectonically termed the Xing'an-Mongolian orogenic belt (XMOB), is in the eastern segment of the CAOB (Wu et al., 2011; Windley et al., 2007; Li, 2006; Jahn, 2004; Sengör and Natal'in, 1996; Sengör et al., 1993). This belt has been traditionally subdivided into a series of massifs, including the Erguna massif in the northwest, the Xing'an and Songliao massifs in the central area, and the Jiamusi and Khanka massifs in the east, separated by major faults (Fig. 1a). The tectonic evolution of NE China had close relationship with the Paleo-Asian oceanic regime and Paleo-Pacific oceanic regime during the Paleozoic-Early Mesozoic (Liu et al., 2017; Han et al., 2012; Windley et al., 2007; Li, 2006), as well as the western circum-Pacific oceanic regime and Mongol-Okhotsk oceanic regime during and following the Mesozoic (Tang et al., 2016; Xu et al., 2013; Wang et al., 2012; Meng et al., 2011, 2010; Wu et al., 2011). Although there is generally no disagreement to the geological evolution from the Paleozoic to the Mesozoic in the Northeast China, the Precambrian history in the area is still controversial due to limited research on the Precambrian units in the NE China. A large number of Precambrian units in NE China have been identified by previous studies and demonstrated by field relationships and zircon data (Zhao et al., 2016a, b, c; Ge et al., 2015; Xu et al., 2015; Tang et al., 2013; Wu G et al., 2012; Miao et al., 2007; HBGMR, 1993; IMBGMR, 1993). However, the limited distribution of Precambrian units and only a small number of studies have been carried out on these rocks limits better knowledge of the Precambrian tectonic evolution in Northeast China.
Our focus in this paper is on the Xinghuadukou complex, a unit that has been traditionally considered to be the oldest stratum in the Erguna massif. We investigate meta-sedimentary rocks from the complex and present several new LA-ICP-MS zircon U-Pb, petrological and mineralogical chemical data, in addition with phase equilibria modeling, to constrain the timing of metamorphism, P-T conditions and the P-T path of the Xinghuadukou complex. The results provide important insights into the formation, metamorphism and tectonic evolution of the Xinghuadukou complex.1 GEOLOGICAL SETTING
The Erguna massif, as a significant tectonic unit in NE China, is bordered to the southeast by the Xinlin-Xiguitu suture and to the northwest by the Mongolia-Okhotsk suture (Fig. 1b). This massif was traditionally considered to comprise Precambrian metamorphic basement and Proterozoic to Paleozoic volcanic-sedimentary formations and granitoids (HBGMR, 1993; IMBGMR, 1993). The Precambrian metamorphic basement is called the Xinghuadukou complex, primarily outcropping in the northeast and northwest part of the Erguna massif. Recent geochronological data found in the Erguna massif have shown that this area experienced multi-stage granitic magmatism. Most of granitic intrusions were emplaced during the Cretaceous (~125 Ma), Jurassic to Triassic (220-182 Ma) and Early Paleozoic (517-416 Ma), and lesser granitic magmatism were emplaced during the Late Paleozoic (360-251 Ma), Neoproterozoic (927-737 Ma), Paleoproterozoic (1 854-1 837 Ma) and Neoarchean (2 600 Ma) (Zhao et al., 2016b, c; Ge et al., 2015; Shao et al., 2015; Sun L et al., 2013; Tang et al., 2013; Wu et al., 2011). The Erguna massif comprise the Neoproterozoic (608-949 Ma) meta-sedimentary rocks (the Xinghuadukou complex, the Erguna Group, the Jiageda Group) and a Late Precambrian sequence (the Mohe complex) (Wu G et al., 2012; Zhou et al., 2011a, b; Miao et al., 2007), where khondalites documented a metamorphic event of pan-African (~500 Ma) (Zhou et al., 2011b).
As the metamorphic basement of the Erguna massif, the Xinghuadukou complex is primarily distributed in areas around Mohe, Tahe, Huma, Huzhong, Xinlin and Qiqian (Ge et al., 2015; Biao et al., 2012, 1999; Wu G et al., 2012; Zhou et al., 2011a; Miao et al., 2007) (Fig. 1b). It is comprised of amphibolites, leptynites, leucoleptynites, schists, marbles, gneisses and migmatites. The original Xinghuadukou complex is composed of meta-supracrustal rocks, migmatites and granitic gneisses, however, recent studies suggest that the Xinghuadukou complex should only refer to the meta-supracrustal rocks (Liu et al., 2017; Ge et al., 2015; Biao et al., 2012, 1999). These meta-supracrustal rocks series are a typical volcanic-sedimentary formation including meta-basic and meta-acidic volcanic rocks and meta-sedimentary rocks. These meta-supracrustal rocks are scattered as xenoliths in the meta-granitic plutons of different shapes and sizes because of the effect of multiple granitic magmatic events and structural deformation.2 ANALYTICAL METHODS
Four representative samples were selected for detailed study. The samples were used for whole rock analyses as well as for determining the chemical composition of minerals by electron microprobe. Fresh whole-rock samples were selected to grind in an agate mill to less than 200 mesh in size. The major elements of the samples were analyzed on fused glass disks using a PHILIPS-2404 X-ray fluorescence spectrometer at the Tianjian Institute of Geology and Mineral Resources, China. The analytical results of the major elements in the Xinghuadukou complex are listed in Table 1.
The compositions of minerals in thin section were analyzed with an electron microprobe analyzer (EMPA; JXA-8200, JEOL) with five wavelength-dispersive spectrometers using natural and synthetic minerals, glass, and pure oxides as standards at the Key Laboratory of Mineral Resources Evaluation in Northeast China, Jilin University, China. The concentrations of Si, Ti, Al, Na, Ca, K, Fe, Mg, Mn and Cr in the minerals were determined. Counting times were 20 s at the peak and on the background. The applied acceleration voltage was 15 kV and beam current was 10 nA. The beam diameter for all analyses was typically 2-5 μm.
For the zircon dating, zircons were separated using standard heavy liquids and magnetic methods and were handpicked under a binocular microscope. Purified zircon grains were then mounted in epoxy and polished to about half thickness to expose the cores of the grains. Zircons were photographed in transmitted and reflected light, and imaged using cathodoluminescence (CL). CL images were collected using a CL spectrometer (Garton Mono CL3+) equipped on a Quanta 200F ESEM with a 2-min scanning time at conditions of 15 kV and 120 nA. LA-ICP-MS zircon U-Pb isotopic analyses were performed using an Agilent 7500a inductively coupled plasma mass spectrometer (ICP-MS) equipped with 193-nm laser, housed at the Key Laboratory of Mineral Resources Evaluation in Northeast Asia, Ministry of Land and Resources, China. Zircon 91500 was used as an external standard for age and the standard silicate glass NIST SRM 610 was applied for instrument optimization. The concentrations of U, Th and Pb elements were calibrated using 29Si as an internal calibrant. 207Pb/206Pb, 206Pb/238U and 207Pb/235U ratios and apparent ages were calculated using ICPMSDataCal Ver. 6.7 (Liu et al., 2010, 2008). The age calculations and concordia plots were made using Isoplot (Ver. 3.0) (Ludwig, 2003).3 SAMPLE DESCRIPTIONS AND MINERAL COMPOSITIONS 3.1 Description of the Studied Samples
Along the road from Tahe to the Lulin Forest, where the Xinghuadukou complex outcrops (Fig. 2), rock samples were taken a zone of the supracrustal Xinghuadukou complex. The outcrop shows that the rock has experienced partial melting with evidence of small-scale leucosome bands and migmatite (Fig. 3).
Based on thin-section study using a polarizing microscope the samples can be characterized as follows. Sil-Grt-Bt-Ms paragneisses (samples Z1545-5, Z1545-9, Z1545-10, and Z1545-11) for this study contain garnet porphyroblasts in a matrix comprising biotite, muscovite, sillimanite, quartz and plagioclase with sericitic alteration. The rock shows an obvious gneissic schistosity defined by oriented biotite and sillimanite (Fig. 4). The grain size distribution of the garnet varies. A small number of several-mm-sized garnets occur, but still several 0.2-mm-sized relic garnet grains are resorbed by muscovite and biotite. Garnet porphyroblasts are generally xenomorphic and in contact with biotite, quartz and sillimanite. Certain mm-sized garnet grains are nearly free of inclusions, whereas inclusions of quartz and biotite are common in the smaller garnet grains. The inclusions in the garnet are commonly biotite, quartz and sillimanite. There are three different types of biotite. The first type occurs as 0.2-mm-size inclusions of garnet with a reddish-brown color without obvious cleavage, and the second is a sheet type of reddish-brown biotite with rutile inclusions in a matrix along with tablet and fibrous sillimanite. The third type is the common brown biotite ranging from 0.2 to 1 mm in size with obvious cleavage oriented along the dominant foliation. Sillimanite occurs as a tablet and fibrous shape ranging from 0.05 to 0.5 mm in size. Sillimanite in studied rocks can be subdivided into four different occurrences: (1) distributed in matrix within the host rock, (2) distributed as fibrous shape along the biotite cleavage, (3) as inclusions in late muscovite and quartz, and (4) as inclusions in garnet. Muscovite occurs as a coarse-to-medium grained sheet crystal crosscutting the dominant foliation, distributed around relic garnet porphyroblasts and the core of selvages of biotite and sillimanite. Plagioclase has sericitic alteration, and usually formed around relic garnet. Quartz usually appears as coarse-to-medium grain size, with rounded or elongated shape in the matrix. Accessory phase minerals are rutile, monazite, zircon and zoisite. According to petrographic observation, the peak mineral assemblage may contain sillimanite, garnet, quartz, biotite with or without ilmenite, while muscovite and plagioclase in matrix should be retrograde phases. The main minerals list and their modal proportions (vol.%) of different samples are listed in Table 2.
EPM spot analysis was performed in garnet to show composition changes from core to rim. In general, garnet is rich in almandine and shows weak compositional variations from core to rim (Fig. 5). Garnet contains 72 mol.% to 77 mol.% almandine, 7 mol.% to 10 mol.% pyrope, 11 mol.% to 17 mol.% spessartine and 2 mol.% to 3 mol.% grossular (see Table S1).
Biotite is mainly ferribiotite with FeOT and MgO contents ranging from 19.25 wt.% to 22.89 wt.% and 5.64 wt.% to 9.39 wt.%, respectively. The TiO2 contents vary from 0.28 wt.% to 3.99 wt.%. Biotite in matrix shows lower XMg of 0.31-0.33 than that as inclusions in garnet (XMg=0.37-0.46). FeOT and MgO content in biotite show a significant negative correlation, which is the result of Fe-Mg exchange during metamorphism.
Muscovite around relic garnet, has Si ranging from 3.06 to 3.07 p.f.u. (Table S1) with a relative high value of XMg (XMg= 0.42-0.44).4 METAMORPHIC P-T CONDITIONS 4.1 Methods
Metamorphic P-T conditions for the metapelites from the Xinghuadukou complex were obtained from pseudosection and conventional geothermobarometry method. On the basis of the internally consistent thermodynamic database, the Ti-in-biotite (Wu and Chen, 2015) and the garnet-biotite (GB) (Holdaway, 2000) geothermometers and garnet-biotite-muscovite-aluminosilicate-quartz (GBMAQ) geobarometer (Wu and Zhao, 2007) were applied to calculate P-T conditions.
The P-T section of the Sil-Grt-Bt-Ms paragneisses was calculated using Gibbs energy minimization method by Perple_X software (Connolly, 2005) with the thermodynamic dataset of Holland and Powell (1998, as updated in 2003). Melt- and garnet-solution models from White et al. (2007) and biotite-solution model from Tajčmanová et al. (2009) were applied. Plagioclase-solution model was used from Newton et al. (1980) and Kfeldspar-soulution was applied from Thompson and Hovis (1979). Mn solution in staurolite and garnet was explained by the Mn end-members proposed by Tinkham et al. (2003). Based on the dominant minerals and their chemical compositions, the phase equilibria of the Xinghuadukou complex metapelites from the Lulin Forest can be modeled in the system MnNCKFMASH (MnO-Na2O-CaO-K2O-FeO-MgO-Al2O3-SiO2-H2O) for the P-T ranges from 3 to 8 kb and 500 to 800 ℃. The Mn content in these rocks is low. However, MnO was included in the phase equilibria modeling due to its effect on the garnet stability at low temperatures (Tinkham et al., 2003; Spear, 1995). The influence of Fe3+ is ignored in this study since no magnetite was found in the metapelites forming Xinghuadukou complex, and the amount of Fe3+ in biotite and garnet was too low to be significant. The effective bulk composition for modeling in this study was calculated by integrating mineral compositions and their modal abundances (Table 2). The upper left of the P-T phase diagram shows the resulting effective bulk compositions (in mol.%) (Figs. 6-7).4.2 Results
On the basis of the mineral assemblage of the metapelite of the Xinghuadukou complex, the peak mineral assemblage should be Bt+Grt+Sil+Qtz, and the retrograde mineral assemblage should be Bt+Mus+Grt+Qtz±Pl. From evidence of surrounding migmatites with leucosome, K-feldspar from partial melting (Fig. 4f) and late, crosscutting muscovite in the cores of biotite and sillimanite, this rock should have experienced fluid-saturated solidus melting, muscovite dehydration melting, and partial biotite dehydration melting reactions. This finding means that the muscovite in the metapelite was consumed during prograde metamorphism. The stable field should be in the area of Bt+Grt+Kfs+Sil+liq (Fig. 6). The metamorphic reaction within this area along the isobaric heating path is Bt+Sil+Qtz= Gt+Kfs+liq. Garnet was once considered to be one of the most resistant minerals to diffusion in most metamorphism. However, it has been found that diffusion of divalent cations such as Fe2+, Mg2+ can relax the compositional zones in garnet at temperatures during upper amphibolite facies (Florence and Spear, 1991); therefore, Fe-Mg minerals like biotite and garnet, may not document their mineral compositions at peak metamorphic condition. On the other hand, the biotite inclusions of garnet will likely preserve the prograde metamorphism or near-peak metamorphic conditions. The XMg and Ti in biotite increase with an increase in temperature. We obtained two different high XMg values (0.37 and 0.46) for biotite inclusions in garnet. However, the latter one with the higher XMg value has notably little Ti, showing the latter biotite is a prograde product during the growth of the garnet. The former one with an XMg value of 0.37 and 0.21 p.f.u. of Ti can be regarded as the near-peak T condition of metamorphism. Based on the isopleth of the XMg of biotite, XGrs of garnet and Ti-in-biotite geothermometer, the near-peak T conditions are 710- 740 ℃ with a pressure ranging of 5-6 kb. The latter biotite should be at the P-T conditions of 6.1 kb and 645 ℃ as a prograde metamorphic record.
The observed retrograde assemblage Bt-Grt-Mus-Pl-Sil- Qtz defines a P-T range of 3.5-5.4 kb and 585-660 ℃ from phase equilibria modeling. At an average pressure of 4.5 kb, the measured XPy and XMg of the garnet core isopleth yield a temperature of 645 ℃, under solidus conditions. To constrain the retrograde P-T conditions, the relic garnet contacting the muscovite and biotite are chosen to apply to the GB geothermometer. The result yields a retrograde temperature of approximately 625 ℃. GBMAQ geobarometer results show that the retrograde pressure is approximately 4.4 kb. In this case, a nearly isobaric cooling path is suggested from the results.
For cooling along the path from point B to point C, the melt of the rock begins to crystallize with reactions of Grt+Kfs+liq=Bt+Sil. With the crystallization of melt, it will release H2O to provide an internal source of H2O support for retrograde metamorphism. Garnet will be consumed, and both of sillimanite and biotite will be produced. Since some parts of garnet are likely to be preserved, the fact that exactly the same amounts of garnet grown during prograde metamorphism will be consumed is not likely to happen. The biotite and sillimanite forms selvages and are dispersed in the matrix with quartz (Wei, 2016; Spear et al., 1999). With continued cooling of the rock, it will encounter the retrograde of muscovite dehydration melting reaction, and enter the field of Bt-Mus-Grt-Kfs-Sil-Qtz-liq with the reaction of Grt+Sil+Kfs+liq=Mus+Bt+Qtz. With the decrease in temperature, the rock continuously retrogrades and produces plagioclase with the reaction Grt+Sil+Kfs+liq= Mus+Bt+Pl+Qtz. The rock will continuously consume the garnet until the K-feldspar and melt totally disappear. If there is no melt-loss occurring, the rock will reach equilibrium without garnet present. However, if a degree of melt-loss occurs, the retrograde reactions will not reach equilibrium, and a fraction of relic garnet will be preserved. When melt-loss occurs, domains without sufficient garnet will have an assemblage of Bt-Mus-Sil-Qtz(±)Pl, and domains without sufficient sillimanite will have an assemblage of Bt-Mus-Grt-Qtz(±)Pl.5 U-Pb ZIRCON ISOTOPE
Zircons from Sample Z1545-5 are colorless to light yellow, transparent to translucent with various shapes, including ovals, pyramid and plates (Fig. 8). The zircons range in size from 50 to 120 μm, with length to width ratios of 1 : 3. The CL imaging reveals that a few large grains show complex internal structures, with irregular cores and narrow rims, indicating they are probably caused by overgrowth during the late metamorphic processes or alteration. The grains generally comprise unzoned or weak-zoned thin dark rims from CL images, whereas the majority of the cores of zircons are characterized as well-defined oscillatory zoning (Fig. 8a). Sixty-eight analyses were performed on 67 zircon grains, yielding ages ranging from a 206Pb/238U age of 479 Ma (spot Z1545-5-15) to a 207Pb/206Pb age of 2 605 Ma (spot Z1545-5-10). The results of zircon data are listed in Table S2 and shown in Fig. 9. Most of the 67 zircons from Sample Z1545-5 plot near the Concordia curve barring eight analyses that record Pb loss. Four Neoarchean grains yield ages from 2 509 to 2 606 Ma with peak at 2 518 Ma. Nine analyses of Paleoproterozoic zircon grains yield ages ranging from 1 618 to 2 343 Ma with three peaks at 1 825, 2 179, and 2 351 Ma. Fourteen analyses of Mesoproterozoic zircon grains present ages ranging from 1 011 to 1 567 Ma with three peaks at 1 012, 1 390, and 1 576 Ma. Thirty-four Neoproterozoic zircon grains yield ages from 592 to 999 Ma with the main peaks at 602, 698, 797, 848, and 917 Ma. The remaining seven Cambrian zircons have ages from 479 to 530 Ma with peaks at 518 Ma.
Sample Z1545-10 has a population of colorless to light yellow, transparent to translucent zircon crystals that are mostly prism or oval in shape. The zircon grains have lengths of 50-130 μm and length/width ratios of 1 : 3 (Fig. 8b). A total of 76 points were performed from 75 zircon crystals; the results of zircon data are listed in Table S2 and shown in Fig. 9. The data show ages ranging from a 206Pb/238U age of 496 Ma (spot Z1545-10-06) to a 207Pb/206Pb age of 2 695 Ma (spot Z1545-10-33). Generally, the ages of Sample Z1545-10 can be approximately divided into five groups. One Neoarchean zircon grain shows an age of 2 695 Ma, the oldest recorded from the studied sample. Ten analyses of Paleoproterozoic grains yield a range of ages from 1 653 to 1 965 Ma with peaks at 1 660, 1 841, 1 915, and 1 978 Ma. Eighteen analyses of Mesoproterozoic grains present ages varying from 1 030 to 1 525 Ma with peaks at 1 250, 1 360, and 1 488 Ma. Forty-six Neoproterozoic analyses show a range of ages from 758 to 992 Ma, with age peaks at 808, 901, and 996 Ma. The younger Cambrian zircon grains aged 496 and 509 Ma show an obvious low Th/U ratio compared to the other zircon grains from the studied sample, which indicate a metamorphic origin.6 DISCUSSION 6.1 Ages of the Rocks from the Xinghuadukou Complex
The Xinghuadukou complex was originally regarded as a sequence of Paleoproterozoic-to-Mesoproterozoic volcanic-sedimentary rocks named the Xinghuadukou "Group" (Sun et al., 2002; Biao et al., 1999). However, this division was on the basis of lithostratigraphic relationships and unreliable isotopic ages (Biao et al., 2012, 1999; HBGMR, 1993). In recent years, an increasing number of new geochronological data by LA-ICP-MS, SHRIMP and single-grain U-Pb dating method of zircons in the Erguna massif have been published (Wang Z Y et al., 2018; Wang Y et al., 2017; Zhou et al., 2017, 2011a; Ge et al., 2015, 2005; Wu G et al., 2012; Miao et al., 2007). Miao et al. (2007) dated the Xinghuadukou complex in the Xinlin-Hanjiayuanzi area using the U-Pb SHRIMP method and suggested that it formed in the Cambrian or Neoproterozoic. Zhou et al. (2011a) analyzed detrital zircons from Mohe complex in the Hongqi-Mohe area with the LA-ICP-MS method. The results showed that the deposition of the sedimentary rocks from the Mohe complex (as regarded as Xinghuadukou complex in our study) occurred no earlier than 608 Ma and recorded the pan-African metamorphic age of ~500 Ma. Wu G et al. (2012) analyzed 96 detrital zircons using the U-Pb LA-ICP-MS method from the Xinghuadukou complex in the Lulin Forest yielding ages concentrated from 749 to 2 765 Ma (igneous zircons) and 448-716 Ma (metamorphic zircons), suggesting that the sedimentary rocks from the Xinghuadukou complex were deposited at least after 749 Ma. Ge et al. (2015) dated 12 samples (one is a meta-volcanic protolith) from the so-called "Xinghuadukou Group" with different ages at Mangui (957 Ma), Huma (~850 Ma), Hongqi Forest (844 and 820 Ma), Qiqian (810 and 418 Ma), Pangu (794 and 786 Ma), Mohe (479 and 265 Ma), and Lulin Forest (192 and 209 Ma). Although this previously published data show the presence of the Precambrian basement of the Erguna massif, the formation age of the supracrustal rocks of the Xinghuadukou complex remains controversial.
As described in Section 3, the dated samples Z1455-5 and Z1545-10 are garnet-sillimanite gneisses, and their protoliths are pelitic-sandy sedimentary rocks. The rock ages indicate the ages of their origin or metamorphism since the zircons from the gneisses should be detrital or metamorphic in origin. The U-Pb dating results from Z1545-5 reveal a range of ages from 479±8 to 2 605±55 Ma with clusters at 2 029-2 606, 1 263-1 851, 909- 1 169, 849-850, 749-832, 640-731, 592-608, and 479-530 Ma. The youngest group, with a range of ages from 479 to 530 Ma and low Th/U ratios (0.03-0.36) with a weighted mean 206Pb/238U age of 500±18 Ma (MSWD=2.8), and the zircon grains with analytical spots of Z1545-5-15, Z1545-5-14 and Z1545-5-12 (Fig. 8), have distinguished core and rim structures, which indicate that they are of metamorphic origin. The zircons of the second youngest group of 592-608 Ma age are euhedral to subhedral and prismatic showing clear oscillatory zoning and a relatively high Th/U ratio (0.33-0.78), which shows typical characteristics of igneous origin. Consequently, we interpret that the 592-608 Ma group with a weighted mean age of 601±12 Ma (MSWD=0.12) should define the maximum depositional age of the protolith of the garnet-sillimanite gneisses within the Xinghuadukou complex. The dating results from Z1545-10 show a range of ages ranging from 496±15 to 2 695±16 Ma with clusters at 2 695, 1 835-1 965, 1 236-1 653, 884-1 030, 758-853, and 496-509 Ma. We interpret that the youngest zircon age group of 496-509 Ma with relative low Th/U ratios (0.13-0.14) and obvious core-rim structure is of metamorphic origin. The youngest detrital zircon of Sample Z1545-10, having an age of 758±18 Ma, limits the maximum depositional age of the protolith of the garnet-sillimanite gneisses. In summary, our new geochronological data on zircons from these two samples show that the Xinghuadukou complex in the Lulin Forest was deposited during the Neoproterozoic, rather than the Paleoproterozoic-to-Mesoproterozoic (Sun et al., 2002; Biao et al., 1999; HBGMR, 1993), and they recorded the time of metamorphism as pan-African (~500 Ma).6.2 Early Paleozoic Metamorphic P-T Path and Tectonic Implications
A few studies have focused on the age of the Xinghuadukou complex (Ge et al., 2015; Biao et al., 2012; Wu G et al., 2012; Miao et al., 2007). Mineral assemblages of various rocks from the Xinghuadukou complex are described by Biao et al. (1999) and HBGMR (1993). Although research on preliminary ages and geochemistry data has been reported, a reliable metamorphic condition and reliable P-T for the Xinghuadukou complex has not been identified.
Based on our studies, the Xinghuadukou complex is characterized by a clockwise P-T path. This clockwise P-T path can be divided into three main stages: a prograde stage, a near-peak upper amphibolite facies stage, and a near-isobaric cooling stage. The three stages of this clockwise P-T path provide important insights into the collision orogenic process and allow one to build a model for the tectonic evolution of the Xinghuadukou complex within the Erguna massif. The prograde metamorphic stage relates to the garnet growth and occurred at P-T conditions of 6.1 kb and 640 ℃ (point A in Fig. 6). The metapelites from the Xinghuadukou complex document the collision process between the Erguna massif and the Xing'an massif, and they underwent amphibolite facies metamorphism. During the collision and an increase in temperature, pressure, and depth, the metapelites experienced fluid-saturated solidus melting, muscovite dehydration melting, and part-of-biotite dehydration melting reactions to reach near-peak upper-amphibolite facies metamorphism with P-T conditions of 5-6 kb and 710-740 ℃. Thereafter, during a near-isobaric cooling, the rock experienced retrograde metamorphism with muscovite crystallization forming in the core of biotite and sillimanite to reach the observed mineral assemblage at P-T conditions of 4.4 kb and 625 ℃.
The timing of collision between the Erguna and the Xing'an massifs has been studied by many researchers for several decades. The most commonly quoted age of collision is from ~647 until ~500 Ma. The Xinlin-Gaxian-Jifeng ophiolitic melanges crop out in the northern Great Xing'an Range, which marks the closure of the Xinlin-Xiguitu Ocean between the Erguna and Xing'an massifs, having U-Pb ages of 539-510 (Xinlin), ~630 (Gaxian), and 647 Ma (Jifeng) (Feng et al., 2016; Feng, 2015). The presence of the Toudaoqiao blueschist of N-normal-mid-ocean ridge basalt (MORB) and ocean island basalt (OIB) protoliths with an age of ~500 Ma along the south margin indicates the final collision between the Erguna and the Xing'an massifs (Zhao et al., 2017; Miao et al., 2015; Zhou et al., 2015).
Widespread Late Paleozoic granites and diorites formed along the Xinlin-Xiguitu suture including the Luguhe granite with an age of ~510 Ma (Wu G N et al., 2005), the Tahe intrusion with an age of ~500 Ma (Ge et al., 2005), the Menduli granite with an age of ~502 Ma (Qin et al., 2007), the Halabaqi granite with an age of ~500 Ma (Sui et al., 2006), the Xikouzi granite with an age of ~500 Ma (Sun D Y et al., 2013), and the Hanjiayuanzi diorite with an age of 501 Ma (Zhao, 2017). The geochemical features of all these rocks indicate they formed in a post-collision or post-orogenic tectonic setting. In that case, the Xinghuadukou complex is generated by the collision of the Erguna and the Xing'an massifs, and documented the subduction, subsequent uplift and post-collision process.7 CONCLUSIONS
A detailed and systematic petrological, zircon-U-Pb-age and metamorphic study of metapelites within the Xinghuadukou complex has led to the following primary conclusions.
(1) The protoliths of the Sil-Grt-Bt-Ms paragneiss within the Xinghuadukou complex from the Lulin Forest of NE China are claystones.
(2) Garnet porphyroblasts in the metapelites show diffusion zoning from core to rim with inclusions of sillimanite, biotite and quartz. From the analysis of different mineral assemblage, the metamorphic evolution can be divided into three stages: an early prograde metamorphic stage, a near-peak upper amphibolite facies stage, and a near-isobaric cooling stage.
(3) Zircon U-Pb dating yielded metamorphic ages between 496 and 509 Ma and an upper limit deposition age of at least ~601 Ma, which shows the Xinghuadukou complex formed during the Neoproterozoic, rather than during the Paleoproterozoic to Mesoproterozoic, and recorded pan-African metamorphic events of the Erguna massif.
(4) The P-T conditions together with mineral assemblages define a clockwise P-T path, revealing important information related to the subduction, subsequent uplift and post collision process of the Erguna and Xing'an massifs.ACKNOWLEDGMENTS
This study was supported by the National Natural Science Foundation of China (No. 41472164) and the MADE-IN-EARTH ERC starting grant of Switzerland (No. 33577). We are grateful to the anonymous reviewers who helped to improve the paper, and the editors for handling and editing. The final publication is available at Springer via https://doi.org/10.1007/s12583-018-0843-z.
Electronic Supplementary Materials: Supplementary materials (Tables S1-S2) are available in the online version of this article at https://doi.org/10.1007/s12583-018-0843-z.
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