This study presents new geochronological and geochemical data for early Paleozoic igneous rocks of the Zhangguangcai Range, northeastern China, and uses these data to further define the early Paleozoic tectonic evolution of the Songnen–Zhangguangcai Range block (SZB) and the Jiamusi block, and provide insights into crustal growth and reworking processes in these blocks of the eastern Central Asian Orogenic Belt. New zircon U-Pb data indicate widespread early Paleozoic magmatism (ca. 516, 496–482, 475–461, and 426 Ma) in the southeastern SZB. The ca. 516 Ma magmatism formed Na-rich tonalites that contain low concentrations of rare earth elements, are enriched in Eu and Sr, have high CaO/Al2O3 and Ba/La ratios, low Rb/Sr and Rb/Ba ratios, and negative Ce anomalies and zircon εHf(t) values. These features suggest an origin from magmas generated by partial melting of ancient accumulated gabbroic rocks with the addition of subducted-sediment–derived fluids. The ca. 496–482 Ma magmatism formed calc-alkaline I-type biotite granodiorites and monzogranites, whereas the ca. 475–461 Ma igneous rocks include biotite monzogranites with K-rich adakitic signatures and alkali-feldspar granites. These ca. 496–461 Ma granitoids have zircon εHf(t) values from –0.82 to +5.44 and two-stage depleted-mantle model (TDM2) ages of 1490–1103 Ma, suggesting they formed from magmas generated by partial melting of heterogeneous Mesoproterozoic lower crustal material. The ca. 426 Ma Na-rich tonalites are geochemically similar to the ca. 516 Ma tonalites and also originated by partial melting of ancient gabbroic or amphibolitic rocks with the involvement of subducted-sediment–derived fluids. The geochemistry of these early Paleozoic igneous rock assemblages is indicative of formation in an active continental margin setting associated with northwestward subduction beneath the southeastern SZB. The zircon Hf isotopic compositions of these early Paleozoic igneous rocks indicate that Paleoproterozoic–Mesoproterozoic crustal material was reworked during the early Paleozoic subduction- and collision-related tectonism.

The Central Asian Orogenic Belt (CAOB, Fig. 1A) is located between the Siberian, North China, and Tarim cratons, and is dominated by a series of microcontinents and accretionary belts. The Central Asian Orogenic Belt has a long-lived orogenic history (>800 m.y.) that involved multiple phases of subduction and continuous accretion (Şengör et al., 1993; Şengör and Natal’in, 1996; Khain et al., 2002; Li, 2006; Windley et al., 2007; Demoux et al., 2009; Li et al., 2009; Wu et al., 2011; Kröner et al., 2014; Xiao and Santosh, 2014; Xiao et al., 2004, 2015; Zhang et al., 2016). Northeastern China is tectonically located in the eastern Central Asian Orogenic Belt (Fig. 1A) and has a Paleozoic tectonic history that is dominated by the amalgamation of microcontinents, including the Erguna block, Xing’an block, Songnen–Zhangguangcai Range block (SZB), Jiamusi block (JB), and Khanka block (Fig. 1B; Li and Ouyang, 1998; Meng et al., 2010, 2011a; Wu et al., 2011; Wang et al., 2012a; Xu et al., 2009, 2012, 2013; Cao et al., 2013; Li et al., 2014; Xu et al., 2014; Wang et al., 2015; Wilde, 2015; Liu et al., 2017). However, the early Paleozoic evolution of these blocks remains controversial, especially the history of the SZB and the JB. This includes the timing of amalgamation between these two blocks; previous research suggested that this occurred during the Archean–Mesoproterozoic (Cao et al., 1992; Heilongjiang Bureau of Geology and Mineral Resources, 1993), the Neoproterozoic or early Paleozoic (Zhang and Sklyarov, 1992; Li et al., 1999; J.F. Liu et al., 2008; Xie et al., 2008a, 2008b; Meng et al., 2010, 2011a; Xu et al., 1994, 2012; Wang et al., 2012a), or the Early–Middle Jurassic (Wu et al., 2007b, 2011; Zhou et al., 2009). More recent research suggests that the amalgamation of the northern SZB and northern JB started by at least ca. 540 Ma and lasted until ca. 496 Ma, and postcollisional extension occurred during the late Cambrian (ca. 490 Ma; J.F. Liu et al., 2008; Wang et al., 2016a). However, the Ordovician tectonic models suggest that the eastern SZB represented an active continental margin setting associated with the northwest-directed subduction of an oceanic plate located between the southern SZB and the JB, as evidenced by the presence of widespread early Paleozoic calc-alkaline igneous rocks in the eastern SZB (Li et al., 1999; Xu et al., 2012; Wang et al., 2012a; Wang et al., 2016a). The Cambrian to Silurian tectonic evolution of the southeastern SZB is somewhat poorly defined relative to the northern segment of the eastern SZB, owing to a lack of systematic research into early Paleozoic igneous rocks. Consequently, the timing of subduction of the oceanic plate beneath the SZB and the genetic linkages between early Paleozoic magmatism and the tectonic evolution of the eastern SZB are still unclear.

Figure 1.

Simplified tectonic maps. (A) Eastern Asia (A and M represent the Altaids and Manchurides, respectively, Şengör and Natal’in, 1996). CAOB—Central Asian Orogenic Belt; MOS—Mongol–Okhotsk suture (modified after Li, 2006). (B) Northeastern China (modified after Wu et al., 2007a).

Figure 1.

Simplified tectonic maps. (A) Eastern Asia (A and M represent the Altaids and Manchurides, respectively, Şengör and Natal’in, 1996). CAOB—Central Asian Orogenic Belt; MOS—Mongol–Okhotsk suture (modified after Li, 2006). (B) Northeastern China (modified after Wu et al., 2007a).

It is generally accepted that the Central Asian Orogenic Belt represents the largest known Phanerozoic accretionary orogenic belt on Earth. In addition, Phanerozoic crustal accretion is known to have largely occurred in the Paleozoic orogenic belts (Şengör et al., 1993; Şengör and Natal’in, 1996; Wu et al., 2000; Jahn et al., 2000a, 2000b; Jahn, 2004; Hong et al., 2004; Windley et al., 2007; Wang et al., 2009; Kröner et al., 2014), and lateral subduction played an important role in the generation of continental crustal material (Şengör et al., 1993; Kelemen, 1995; Şengör and Natal’in, 1996; Xiao et al., 2004; Kemp et al., 2009; Xiao and Santosh, 2014; Gazel et al., 2015; Kelemen and Behn, 2016). However, the Paleozoic orogenic belts in the eastern Central Asian Orogenic Belt are also associated with multiple microcontinents. The nature of the deep crust, the presence (or absence) of Precambrian crust, and the timing and processes involved in crustal growth and reworking in these blocks all remain unclear.

Here we report new zircon U-Pb ages, Hf isotopic data, and major and trace element compositions of early Cambrian to middle Silurian igneous rocks from the Zhangguangcai Range in the eastern Central Asian Orogenic Belt. The data are used to constrain the history of amalgamation of the SZB and JB, and to determine the nature of the crust in this region, as well as the processes involved in the growth and reworking of crustal material in these blocks of the eastern Central Asian Orogenic Belt.

The eastern SZB is divided into the Zhangguangcai and Lesser Xing’an Ranges (Fig. 2A). The Zhangguangcai Range consists of widespread Phanerozoic intrusive rocks, Neoproterozoic–early Paleozoic and late Paleozoic volcano-sedimentary rocks, and voluminous Mesozoic–Cenozoic volcanic units (Heilongjiang Bureau of Geology and Mineral Resources, 1993; Meng et al., 2011a, 2011b; Wu et al., 2011; Wang et al., 2012a, 2012b, 2014; Xu et al., 2013; Wang et al., 2015). The Neoproterozoic–early Paleozoic volcano-sedimentary units, including the Neoproterozoic Tadong Group and the Ordovician Xiaojingou and Waibizi Formations, recorded regional greenschist facies metamorphism that locally reached the amphibolite facies (Heilongjiang Bureau of Geology and Mineral Resources, 1993; Jilin Bureau of Geology and Mineral Resources, 1997; Wang et al., 2012a, 2014). The range also contains Phanerozoic intrusive rocks including early to late Paleozoic granitoids that were emplaced ca. 516, 496, ca. 445, ca. 425, and 256–252 Ma, as well as numerous Mesozoic granitoids (Wu et al., 2011; Yu et al., 2013; Wang et al., 2012a, 2014; Qin et al., 2016). The Lesser Xing’an Range is dominated by voluminous granitic intrusions and minor amounts of Neoproterozoic metamorphosed volcano-sedimentary rocks and granitic gneisses, early Cambrian metamorphosed marine sedimentary rocks, Ordovician and late Paleozoic volcano-sedimentary units, and widespread Mesozoic volcanic and sedimentary rocks (Heilongjiang Bureau of Geology and Mineral Resources, 1993; Meng et al., 2011a; Xu et al., 2013; Wang and Liu, 2014; Wang et al., 2014; Wang et al., 2016a). Phanerozoic intrusive rocks in this range include minor amounts of early Paleozoic (ca. 508–496, ca. 490, ca. 470–450, and 432 Ma; Fig. 2A) intermediate to felsic rocks, voluminous late Paleozoic to early Mesozoic (ca. 260, ca. 240, ca. 210, and 199–175 Ma) granitoids, and minor amounts of Mesozoic mafic and ultramafic intrusions (ca. 186–182 Ma) (Wilde et al., 2003; J.F. Liu et al., 2008; Wu et al., 2011; Yu et al., 2012; Wei et al., 2013; Guo et al., 2016; Wang et al., 2016a).

Figure 2.

(A) Distribution of early Paleozoic igneous rocks and the Heilongjiang Complex in the Zhangguangcai Range, Lesser Xing’an Range, and adjacent areas of northeastern China. (B, C) Detailed geological maps of Shuguang and Tadong areas of the Songnen–Zhangguangcai Range block.

Figure 2.

(A) Distribution of early Paleozoic igneous rocks and the Heilongjiang Complex in the Zhangguangcai Range, Lesser Xing’an Range, and adjacent areas of northeastern China. (B, C) Detailed geological maps of Shuguang and Tadong areas of the Songnen–Zhangguangcai Range block.

The JB is separated from the SZB to the west by the Jiayin-Mudanjiang fault and is dominated by minor amounts of Neoproterozoic intrusions and voluminous Paleozoic granitoids, and the Mashan Complex as well as minor amounts of late Paleozoic and Mesozoic volcano-sedimentary rocks (Cao et al., 1992; Heilongjiang Bureau of Geology and Mineral Resources, 1993; Wilde et al., 2000, 2003; Wu et al., 2011; Xu et al., 2013; Bi et al., 2014; Yang et al., 2014, 2017). The Paleozoic intrusive rocks in this area were emplaced between the Cambrian and Early Ordovician (ca. 530–484 Ma), and in the Permian (ca. 299–254 Ma) (Wilde et al., 2003; Wu et al., 2001, 2011; Bi et al., 2014; Yang et al., 2014, 2015). The Mashan Complex recorded granulite facies metamorphism at ca. 560 and 500 Ma (Wilde et al., 1997, 2000; Yang et al., 2017).

The SZB and JB are separated by a narrow high-pressure metamorphic belt, defined as the Heilongjiang Complex, which crops out along the Jiayin-Mudanjiang fault from Luobei in the north to the area around Yilan and to Mudanjiang in the south (Fig. 2A). The complex has been thought to represent an originally north-south–oriented suture belt between the two blocks, and consists of serpentinite, blueschist, greenschist, marble, amphibolite, muscovite-albite schist, and quartzite units (Cao et al., 1992; Heilongjiang Bureau of Geology and Mineral Resources, 1993; Wu et al., 2007b; Wang et al., 2012a). Blueschists from the Heilongjiang Complex yielded whole-rock Ar–Ar ages of ca. 645 Ma and glaucophane plateau ages of ca. 600 Ma (Zhang, 1992), and this complex was also overprinted by intense Early Jurassic compressional deformation (ca. 180 Ma; Li et al., 1999; Wu et al., 2007b; Zhou et al., 2013).

This study focuses on seven granitic plutons that crop out in the Tadong and Shuguang areas of the Zhangguangcai Range, in the southeastern SZB (Figs. 1 and 2). Representative photographs and photomicrographs of these plutons are given in Figure 3, and their field and petrographic characteristics are described in the following.

Figure 3.

Representative photographs and photomicrographs of early Paleozoic igneous rocks from the Zhangguangcai Range, northeastern China. Af—alkali-feldspar, Bi—biotite, Hb—hornblende, Pl—plagioclase, Q—quartz. (A) Tonalite outcrop (sample 11HNA7-1). (B) Tonalite (11HNA7-1). (C) Tonalite (11HNA13-1). (D) Porphyritic biotite monzogranite (15XH30-1) containing alkali-feldspar megacrysts. (E) Porphyritic biotite monzogranite (15XH30-1). (F) Porphyritic monzogranite (15XH32-1). (G) Biotite granodiorite (HDL2-2). (H) Biotite monzogranite (15XH9-1). (I) Alkali-feldspar granite (15XH10-1).

Figure 3.

Representative photographs and photomicrographs of early Paleozoic igneous rocks from the Zhangguangcai Range, northeastern China. Af—alkali-feldspar, Bi—biotite, Hb—hornblende, Pl—plagioclase, Q—quartz. (A) Tonalite outcrop (sample 11HNA7-1). (B) Tonalite (11HNA7-1). (C) Tonalite (11HNA13-1). (D) Porphyritic biotite monzogranite (15XH30-1) containing alkali-feldspar megacrysts. (E) Porphyritic biotite monzogranite (15XH30-1). (F) Porphyritic monzogranite (15XH32-1). (G) Biotite granodiorite (HDL2-2). (H) Biotite monzogranite (15XH9-1). (I) Alkali-feldspar granite (15XH10-1).

Samples from the Tadong area include tonalites (samples 11HNA7, 11HNA17, and 11HNA13), and monzogranites (15XH30 and 15XH32). Samples 11HNA7 and 11HNA17 were taken from a tonalite that was emplaced into amphibolite, biotite-plagioclase gneiss, marble, and schist units of the Zhudundian Formation, part of the Tadong Group (Jilin Bureau of Geology and Mineral Resources, 1997; Wang et al., 2014). This tonalite contains quartz (∼30% by volume), plagioclase (∼65%), minor biotite (∼5%), and accessory magnetite, zircon, and apatite (Figs. 3A, 3B). The other tonalite (samples 11HNA13-1 and 11HNA13-2) was emplaced into the Lalagou Formation of the Tadong Group (Jilin Bureau of Geology and Mineral Resources, 1997; Wang et al., 2014), and samples were taken from the Tadong iron deposit. These tonalite samples are medium grained, have a granitic texture, are massive, and contain hornblende (∼5%), quartz (∼30%), plagioclase (∼65%), and accessory zircon and magnetite (Fig. 3C). The porphyritic monzogranites (samples 15XH30 and 15XH32) are porphyritic and massive. Samples 15XH30 (including 15XH30-1, 15XH30-2, 15XH30-3, and 15XH30-4) contain alkali-feldspar megacrysts with plagioclase and quartz inclusions, and consist of quartz (∼28%), plagioclase (∼30%), alkali feldspar (∼37%), minor biotite (∼5%), and accessory zircon, magnetite, and apatite (Figs. 3D, 3E). In contrast, porphyritic monzogranite samples 15XH32-1 to 15XH32-5 contain ∼45% (by volume) phenocrysts of plagioclase (∼15%), alkali feldspar (∼20%), and quartz (10%) in a fine-grained groundmass of quartz, feldspar, and biotite (Fig. 3F).

The Shuguang area hosts biotite granodiorite (HDL2-2, HDL2-3), biotite monzogranite (15XH9-1, 15XH9-2, and 15XH9-3), and alkali-feldspar granite (15XH10-1, 15XH10-3, and 15XH10-4) intrusions, all of which are medium-grained, have granitic textures, and are massive. The biotite granodiorites contain biotite (∼10%), quartz (∼20%), plagioclase (∼62%), alkali feldspar (∼8%), and accessory minerals (Fig. 3G). Biotite monzogranite samples 15XH9-1 and 15XH9-3 contain biotite (∼8% and ∼15%, respectively), quartz (∼25% and ∼20%, respectively), plagioclase (∼37% and ∼30%, respectively), alkali feldspar (∼30% and ∼35%, respectively), and accessory zircon, apatite, and magnetite (Fig. 3H). The alkali-feldspar granites are dominated by quartz (∼30%), alkali feldspar (∼58%), and plagioclase (∼6%), with minor hornblende (∼4%), biotite (∼2%), and accessory minerals (Fig. 3I).

Zircon U-Pb Dating

Zircon grains were separated from whole-rock samples using the conventional heavy liquid and magnetic techniques and then by hand-picking under a binocular microscope at the Langfang Regional Geological Survey, Hebei Province, China. Laser ablation–inductively coupled plasma–mass spectrometry (LA-ICP-MS) zircon U-Pb analyses were performed using an Agilent 7500a ICP-MS equipped with a 193 nm laser, housed at the Geological Laboratory Centre, China University of Geosciences, Beijing, and the State Key Laboratory of Geological Processes and Mineral Resources, China University of Geosciences, Wuhan, China. A zircon 91500 was used as an external standard for age calibration, and the NIST SRM610 silicate glass was applied for instrument optimization. The crater diameter was 36 μm in the analyses in the Beijing laboratory, and 32 μm during the analyses in the Wuhan laboratory. The instrument parameter and detail procedures were described by Liu et al. (2010). The GLITTER (version 4.4, Macquarie University, Sydney, Australia [Jackson, 2001]), ICPMSDataCal (Liu et al., 2010), and Isoplot (version 3.0; Ludwig, 2003) programs were used for data reduction. Correction for common Pb was made following Andersen (2002). All LA-ICP-MS zircon isotope ratio and age uncertainties are quoted at the 1σ level.

Major and Trace Element Determinations

For geochemical analysis, whole-rock samples, after the removal of altered surfaces, were crushed in an agate mill to ∼200 mesh. Whole-rock major and trace element compositions of the samples analyzed during this study were determined at the State Key Laboratory of Geological Processes and Mineral Resources, China University of Geosciences, Wuhan. Major element compositions were determined by X-ray fluorescence (XRF; Rigaku RIX 2100 spectrometer) using fused-glass disks. Trace element compositions were analyzed by ICP-MS (Agilent 7500a with a shield torch) after acid digestion of samples in Teflon bombs. The detailed procedures are the same as descriptions by Y.S. Liu et al. (2008). Analytical uncertainties are in the range 1%–3%. The analytical results for the BHVO-1 (basalt), BCR-2 (basalt), and AGV-1 (andesite) standards indicate that the analytical precision for major elements is better than 5%, and for trace elements generally better than 10% (Rudnick et al., 2004).

Hf Isotope Analyses

In situ zircon Hf isotope analyses were undertaken using a Neptune Plus multicollector (MC) ICP-MS in combination with a Geolas 2005 excimer ArF LA system (193 nm) that was hosted at the State Key Laboratory of Geological Processes and Mineral Resources, China University of Geosciences, Wuhan. All data were acquired on zircon grain in single spot ablation mode with a spot size of 44 μm. (For the details of the operating conditions for the LA system and the MC-ICP-MS instrument as well as the analytical method, see Hu et al., 2012.) Present-day chondritic ratios of 176Hf/177Hf = 0.282772 and 176Lu/177Hf = 0.0332 (Blichert-Toft and Albarède, 1997) were used to calculate εHf(t) values, and Hf model ages were calculated using the methods of Griffin et al. (2000, 2002).

Zircon U-Pb Dating

Zircon grains were extracted from seven granitic plutons for LA-ICP-MS U-Pb dating. These zircon grains are euhedral to subhedral, show oscillatory zoning, and have high Th/U ratios (0.17–1.05; Fig. 4; Table 1), suggesting a magmatic origin (Pupin, 1980; Koschek, 1993). The youngest ages reflect the timing of crystallization of the plutons, whereas the older ages are indicative of the timing of crystallization of captured zircon grains. Zircon U-Pb concordia are shown in Figure 5 and LA-ICP-MS analytical data are given in Table 1.

Figure 4.

Cathodoluminescence images of selected zircon grains in early Paleozoic igneous rocks from the Zhangguangcai Range, northeastern China. Solid and dash circles represent U-Pb and Lu-Hf analysis spots, respectively. Numbers in and adjacent to circles are analysis numbers, and values below the images are corresponding zircon U-Pb ages.

Figure 4.

Cathodoluminescence images of selected zircon grains in early Paleozoic igneous rocks from the Zhangguangcai Range, northeastern China. Solid and dash circles represent U-Pb and Lu-Hf analysis spots, respectively. Numbers in and adjacent to circles are analysis numbers, and values below the images are corresponding zircon U-Pb ages.

TABLE 1.

LA–ICP–MS ZIRCON U-Pb DATA FOR THE EARLY PALEOZOIC IGNEOUS ROCKS FROM THE ZHANGGUANGCAI RANGE

Figure 5.

(A–G) Laser ablation–inductively coupled plasma–mass spectrometry zircon U-Pb concordia for early Paleozoic igneous rocks from the Zhangguangcai Range, northeastern China (this paper; Wu et al., 2011; Wang et al., 2012a, 2014). MSWD—mean square of weighted deviates. (H) Probability density plot.

Figure 5.

(A–G) Laser ablation–inductively coupled plasma–mass spectrometry zircon U-Pb concordia for early Paleozoic igneous rocks from the Zhangguangcai Range, northeastern China (this paper; Wu et al., 2011; Wang et al., 2012a, 2014). MSWD—mean square of weighted deviates. (H) Probability density plot.

Tadong Area

A total of 17 analyses of zircon grains from tonalite sample 11HNA7-1 yielded 2 groups of concordant ages at 516 ± 4 (mean square weighted deviates, MSWD = 0.01, n = 7) and 534 ± 3 (MSWD = 0.04, n = 8; Fig. 5A), with 2 further single zircon grain ages at 591 and 1018 Ma, indicating that the tonalite was emplaced ca. 516 Ma (Wang et al., 2014).

A total of 17 analyses of zircon grains from porphyritic monzogranite sample 15XH30-1 yielded 4 concordant age groups at 496 ± 11 (MSWD = 0.01, n = 2), 529 ± 11 (MSWD = 0.07, n = 2), 560 ± 5 (MSWD = 0.48, n = 12), and 585 ± 12 Ma (MSWD = 0.01, n = 2), in addition to a single zircon grain age of 512 Ma (Fig. 5B). The youngest age group (ca. 496 Ma) represents the timing of emplacement of this intrusion and is consistent with the age reported by Wu et al. (2011).

A total of 18 analyses of zircon grains from porphyritic monzogranite sample 15XH32-1 yielded 3 groups of weighted mean 206Pb/238U ages at 475 ± 7 (MSWD = 0.04, n = 4), 509 ± 5 (MSWD = 0.34, n = 8), and 531 ± 6 Ma (MSWD = 0.06, n = 6; Fig. 5C), indicating that this monzogranite was emplaced ca. 475 Ma.

The 19 analyses of zircon grains from tonalite sample 11HNA13-1 yielded 5 groups of weighted mean 206Pb/238U ages at 426 ± 6 (MSWD = 0.01, n = 4), 450 ± 7 (MSWD = 0.02, n = 4), 485 ± 7 (MSWD = 0.05, n = 2), 514 ± 6 (MSWD = 0.36, n = 4), and 531 ± 6 Ma (MSWD = 0.09, n = 5; Fig. 5D); the youngest group (ca. 426 Ma) of ages indicates the timing of emplacement of the tonalite (Wang et al., 2014).

Shuguang Area

A total of 24 analyses of zircon grains from biotite granodiorite sample HDL2-2 yielded two groups of weighted mean 206Pb/238U ages at 482 ± 8 (MSWD = 0.29, n = 6) and 525 ± 5 Ma (MSWD = 0.03, n = 17; Fig. 5E); the former indicates the timing of crystallization of the granodiorite.

The 20 analyses of zircon grains from biotite monzogranite sample 15XH9-1 yielded 4 groups of concordant 206Pb/238U ages at 461 ± 5 (MSWD = 0.13, n = 9), 501 ± 7 (MSWD = 0.08, n = 5), 535 ± 9 (MSWD = 0.07, n = 4), and 554 ± 12 Ma (MSWD = 0.06, n = 2; Fig. 5F). The youngest age group (ca. 461 Ma) indicates that the biotite monzogranite was emplaced during the Middle Ordovician.

A total of 16 analyses of zircon grains from alkali-feldspar granite sample 15XH10-1 yielded 3 groups of concordant 206Pb/238U ages at 462 ± 6 (MSWD = 0.10, n = 7), 495 ± 7 Ma (MSWD = 0.01, n = 5), and 533 ± 8 Ma (MSWD = 0.07, n = 4; Fig. 5G). The youngest of these groups (ca. 462 Ma) indicates that this alkali-feldspar granite crystallized during the Middle Ordovician.

Combining these new zircon U-Pb ages with previously published geochronological data indicates that the Zhangguangcai Range of the southeastern SZB records widespread early Cambrian to middle Silurian magmatism with age peaks ca. 516, 497, 478–462, 451, and 425 Ma (Fig. 5H; Wu et al., 2011; Wang et al., 2012a, 2014).

Major and Trace Element Geochemistry

The ca. 516 Ma tonalites contain 71.2–75.7 wt% SiO2, 14.8–16.9 wt% Al2O3, 0.32–0.71 wt% total Fe2O3, 0.16–0.47 wt% MgO, 2.84–3.44 wt% CaO, 4.64–5.30 wt% Na2O, and 0.05–0.62 wt% K2O (Table 2), are classified as low-K tholeiitic series in K2O versus SiO2 diagram, and are weakly to strongly peraluminous [A/CNK, i.e., Al2O3/(CaO + Na2O + K2O) = 1.06–1.14; Figs. 6A–6C; Streckeisen and Le Maitre, 1979; Peccerillo and Taylor, 1976; Maniar and Piccoli, 1989]. They contain low total rare earth element (REE) concentrations (4.75–6.53 ppm) and have U-shaped REE patterns with prominently positive Eu anomalies (Eu/Eu* = 13.2–20.3), slightly negative Ce anomalies (Ce/Ce* = 0.75–0.85), and (La/Yb)N values of 3.40–10.5 (Fig. 7A; Table 2; Boynton, 1984). These tonalites are also enriched in large ion lithophile elements (LILE; e.g., K, Rb, Ba, and Sr) and Pb, and are depleted in high field strength elements (HFSE; e.g., Nb and Ta) and Th (Fig. 7B).

TABLE 2.

MAJOR AND TRACE ELEMENT DATA FOR THE EARLY PALEOZOIC IGNEOUS ROCKS FROM THE ZHANGGUANGCAI RANGE

Figure 6.

(A) Q’–ANOR [Q’ = 100*Q/(Q + Ab + Or + An); ANOR = 100*An/(An + Or)] diagram for early Paleozoic igneous rocks from the Zhangguangcai Range, northeastern China. (B) A/NK [Al2O3/(Na2O + K2O)] versus A/CNK [Al2O3/(CaO + Na2O + K2O)] values. (C, D) K2O, FeOT/MgO versus SiO2 for early Paleozoic igneous rocks from the Zhangguangcai Range, northeastern China. Field boundaries in A–C are from Streckeisen and Le Maitre (1979), Peccerillo and Taylor (1976), and Maniar and Piccoli (1989), respectively. Gray shadow areas are from Wang et al. (2012a). Arrows in D represent reducing (exp—experimental) and oxidizing conditions from Berndt et al. (2004).

Figure 6.

(A) Q’–ANOR [Q’ = 100*Q/(Q + Ab + Or + An); ANOR = 100*An/(An + Or)] diagram for early Paleozoic igneous rocks from the Zhangguangcai Range, northeastern China. (B) A/NK [Al2O3/(Na2O + K2O)] versus A/CNK [Al2O3/(CaO + Na2O + K2O)] values. (C, D) K2O, FeOT/MgO versus SiO2 for early Paleozoic igneous rocks from the Zhangguangcai Range, northeastern China. Field boundaries in A–C are from Streckeisen and Le Maitre (1979), Peccerillo and Taylor (1976), and Maniar and Piccoli (1989), respectively. Gray shadow areas are from Wang et al. (2012a). Arrows in D represent reducing (exp—experimental) and oxidizing conditions from Berndt et al. (2004).

Figure 7.

Chondrite-normalized rare earth element and primitive mantle–normalized multielement variation diagrams for early Paleozoic igneous rocks from the Zhangguangcai Range, northeastern China. Chondrite and primitive-mantle values are from Boynton (1984) and Sun and McDonough (1989), respectively. Data for pegmatite in the Catalina Schist are from Sorensen and Grossman (1989), and those for early Paleozoic igneous rocks of the northern Songnen–Zhangguangcai Range block (SZB) are from Wang et al. (2016a); other symbols are as in Figure 6.

Figure 7.

Chondrite-normalized rare earth element and primitive mantle–normalized multielement variation diagrams for early Paleozoic igneous rocks from the Zhangguangcai Range, northeastern China. Chondrite and primitive-mantle values are from Boynton (1984) and Sun and McDonough (1989), respectively. Data for pegmatite in the Catalina Schist are from Sorensen and Grossman (1989), and those for early Paleozoic igneous rocks of the northern Songnen–Zhangguangcai Range block (SZB) are from Wang et al. (2016a); other symbols are as in Figure 6.

The ca. 496 Ma porphyritic monzogranites contain 70.1–71.8 wt% SiO2, 14.4–15.5 wt% Al2O3, 2.23–2.76 wt% total Fe2O3, 0.50–0.56 wt% MgO, 1.72–2.20 wt% CaO, 3.46–3.75 wt% Na2O, and 3.91–4.57 wt% K2O (Table 2), are enriched in light (L) REEs and LILEs, and depleted in heavy (H) REEs and HFSEs (Figs. 7C, 7D). The ca. 482 Ma biotite granodiorites are geochemically similar to the ca. 496 monzogranites, but contain lower concentrations of SiO2 (66.7–67.6 wt%) and K2O (2.28–2.41 wt%), higher concentrations of Fe2O3 (4.28–4.85 wt%), MgO (2.85–3.08 wt%) and the REEs (298–319 ppm), and have higher A/CNK values (Figs. 6 and 7; Table 2).

The ca. 475 Ma monzogranites and ca. 462–461 Ma biotite monzogranites and alkali-feldspar granites contain 62.8–74.7 wt% SiO2, 13.1–18.1 wt% Al2O3, 1.42–4.40 wt% total Fe2O3, 0.02–0.73 wt% MgO, 0.55–3.28 wt% CaO, 3.31–4.16 wt% Na2O, and 3.91–5.47 wt% K2O, and are characterized by LREE and LILE enrichments and HREE and HFSE depletions (Figs. 7E, 7F). The ca. 475 and 461 Ma granitoids show negligible Eu anomalies, contain lower concentrations of the HREEs (e.g., Yb = 0.88–1.45 ppm) and Y (8.55–16.6 ppm), and have higher Sr/Y ratios (24.4–57.4) than the ca. 462 Ma alkali-feldspar granites that are strongly depleted in Eu, Ba, Sr, Ti, and P (Figs. 7E, 7F).

The ca. 426 Ma tonalites are geochemically similar to the ca. 516 Ma tonalites (Table 2), although they contain higher concentrations of Na2O (5.87–6.51 wt%) and K2O (0.86–1.47 wt%), and have slightly fractionated REE patterns with higher REE abundances (14.0–14.8 ppm) and negligible Eu anomalies (Table 2). These tonalites are enriched in LREEs, Rb, Ba, Sr, U, and Pb, and depleted in HREEs, Th, Nb, Ta, Ti, and P (Figs.7G, 7H).

Zircon Hf Isotopes

The results of in situ Hf isotopic analyses of zircon grains obtained from seven samples during this study are given in Table 3.

TABLE 3.

ZIRCON Hf ISOTOPIC DATA FOR EARLY PALEOZOIC IGNEOUS ROCKS FROM THE ZHANGGUANGCAI RANGE

Primary zircon grains from the ca. 516 Ma tonalite (sample 11HNA7-1) in the Tadong area yielded 176Hf/177Hf ratios of 0.282268–0.282356, εHf(t) values from –7.22 to –3.72, and two-stage depleted-mantle model (TDM2) ages of 1936–1715 Ma. The captured zircon grains in this sample have εHf(t) values from –7.08 to –1.63 and TDM2 ages of 1940–1597 Ma. A magmatic zircon grain from the ca. 426 Ma tonalite (11HNA13-1) in the same area yielded 176Hf/177Hf ratio of 0.282321, εHf(t) value of –7.27, and TDM2 age of 1871 Ma. The captured zircon grains in this sample have εHf(t) values from –8.58 to –4.87 and TDM2 ages of 1998–1799 Ma.

The ca. 496 and 482 Ma primary zircon grains from the monzogranite (15XH30-1) and biotite granodiorite (HDL2-2) intrusions yielded 176Hf/177Hf ratios of 0.282516–0.282570, εHf(t) values from +1.08 to +3.58, and TDM2 ages of 1385–1238 Ma. In comparison, the ca. 563–512 Ma captured zircon grains in these samples yielded εHf(t) values from +1.16 to +5.26 and TDM2 ages of 1380–1171 Ma.

The ca. 475–461 Ma primary zircon grains from monzogranite samples 15XH32-1 and 15XH9-1, and alkali-feldspar granite sample 15XH10-1 yielded 176Hf/177Hf ratios of 0.282479–0.282644, εHf(t) values from –0.82 to +5.44, and TDM2 ages of 1490–1103 Ma. The ca. 554–495 Ma captured zircon grains in these rocks have εHf(t) values from +0.60 to +7.12 and TDM2 ages of 1430–1039 Ma.

Early Paleozoic Magmatism in the Eastern SZB

Previous research on the lithostratigraphic relationships and Rb-Sr and K-Ar geochronological data for intrusions in the eastern SZB suggested the presence of widespread early Paleozoic igneous rocks (Heilongjiang Bureau of Geology and Mineral Resources, 1993; Li and Zhao, 1991; Luan et al., 1991; Zhao et al., 1996). However, recent zircon LA-ICP-MS U-Pb dating suggests that some of the igneous rocks that were previously ascribed to the early Paleozoic actually formed during the late Paleozoic to early Mesozoic (Meng et al., 2011a; Wu et al., 2011). The spatial and temporal distributions of early Paleozoic igneous rocks in the southeastern SZB are also poorly defined. In this paper, combining our new data with the results of previous research indicates that the Zhangguangcai Range of the southeastern SZB contains widespread early Paleozoic (age peaks ca. 516, 497, 478–462, 451, and 425 Ma) igneous rocks (Figs. 5 and 8A; Wu et al., 2011; Wang et al., 2012a, 2014) consistent with the early Paleozoic magmatic events that are recorded in the northern SZB (age peaks ca. 505–499, 490, 471–450, and 433 Ma; Fig. 8B; Wilde et al., 2003; J.F. Liu et al., 2008; Wang et al., 2012a; Wei et al., 2013; Wang et al., 2016a). In addition, middle–late Cambrian (ca. 505–499 and ca. 490 Ma) igneous rocks are widely exposed in the Lesser Xing’an Range (Fig. 8B), whereas early Cambrian (ca. 516 Ma) magmatism and a minor ca. 496 Ma magmatic event have been identified in the Zhangguangcai Range (Fig. 8A; Wu et al., 2011). Ordovician magmatic events (ca. 478–450 Ma) have been identified throughout the eastern SZB (Figs. 8A, 8B; J.F. Liu et al., 2008; Wang et al., 2012a; Wang et al., 2016a), in addition to minor Silurian magmatic events (ca. 433–425 Ma) in the southeastern and northern SZB (Wei et al., 2013; Wang et al., 2012a, 2014). Similar early Paleozoic magmatic events have also been recognized in other parts of the eastern Central Asian Orogenic Belt. For example, early Cambrian magmatic events are recorded in the JB and Khanka block (541–513 Ma; Bi et al., 2014; Yang et al., 2014) and in the accretionary belt along the northern margin of the North China craton (ca. 518 Ma; Zhang et al., 2014). Voluminous late Cambrian to Ordovician magmatism is also recorded in this accretionary belt (ca. 493 and 467–445 Ma; Zhang et al., 2014; Pei et al., 2014, 2016) as well as in the Xing’an block (ca. 480–450 Ma; Ge et al., 2007b; Wu et al., 2011; She et al., 2012; Wu et al., 2015), the Erguna block (ca. 500–446 Ma; Ge et al., 2005, 2007a; Wu et al., 2005; Sui et al., 2006; Wu et al., 2011; She et al., 2012; Zhao et al., 2014), and the Russian part of the Bureya block (ca. 486–455 Ma; Kotov et al., 2009; Sorokin et al., 2010, 2011a, 2011b, 2011c; Smirnov et al., 2012); the JB also records extensive ca. 508 and ca. 490 Ma magmatic events (Wilde et al., 2003; Bi et al., 2014; Yang et al., 2014; Fig. 8C). However, only rare Silurian igneous rocks have been identified in the accretionary belt along the northern margin of the North China craton (ca. 438–419 Ma; Zhang et al., 2014; Pei et al., 2014, 2016; Wang et al., 2016b) and the Xing’an block (ca. 439 Ma; Guo et al., 2009). All of these suggest that early Paleozoic magmatic events are widespread not only in the eastern SZB, but also in other parts of the eastern Central Asian Orogenic Belt.

Figure 8.

Comparison of zircon age-probability plots. (A) The Zhangguangcai Range (this paper; Wu et al., 2011; Wang et al., 2012a, 2014). (B) The Lesser Xing’an Range of the northern Songnen–Zhangguangcai Range block (J.F. Liu et al., 2008; Wang et al., 2016a). (C) The Jiamusi block (Bi et al., 2014; Yang et al., 2014).

Figure 8.

Comparison of zircon age-probability plots. (A) The Zhangguangcai Range (this paper; Wu et al., 2011; Wang et al., 2012a, 2014). (B) The Lesser Xing’an Range of the northern Songnen–Zhangguangcai Range block (J.F. Liu et al., 2008; Wang et al., 2016a). (C) The Jiamusi block (Bi et al., 2014; Yang et al., 2014).

Petrogenesis of Early Paleozoic Granitoids in the Zhangguangcai Range

The early Paleozoic granitoids in the Zhangguangcai Range formed ca. 516, 496–482, 475–461, and 426 Ma, and contain high concentrations of SiO2 and Al2O3 and low concentrations of TiO2, total Fe2O3, and MgO, suggesting that these intrusions were derived from magmas generated by the partial melting of crustal material (Zen, 1986; Barbarin, 1999; Nabelek et al., 2001; Xu et al., 2009; Koepke et al., 2007). In addition, the lack of coeval mafic and intermediate igneous rocks further confirms that they formed by partial melting rather than crystal fractionation (Lu and Xu, 2011). However, these early Paleozoic granitoids are also geochemically variable, suggesting they were derived from distinct source regions.

Early Cambrian (ca. 516 Ma) tonalites in the Tadong area are weakly to strongly peraluminous (A/CNK = 1.06–1.14), have high Na2O contents and CaO/Al2O3 values, low TiO2 contents, and low Rb/Sr (0.007–0.011) and Rb/Ba (0.025–0.031) ratios (Figs. 9A–9C). These features are geochemically similar to the compositions of experimental silicic melts derived from the anatexis of gabbroic and amphibolitic rocks (Sylvester, 1998; Patiño Douce, 1999; Koepke et al., 2007; France et al., 2010). The tonalites also have negative magmatic zircon εHf(t) values (–7.22 to –3.72) and ancient TDM2 ages (1936–1715 Ma), and relatively low REE abundances as well as extremely positive Eu and Sr anomalies, further suggesting that these intrusions formed from magmas generated by the partial melting of ancient accumulated gabbroic rocks (Fig. 9A; Koepke et al., 2007; France et al., 2010; Pei et al., 2014). This view is also supported by the resemblance of REE patterns for the ca. 516 Ma tonalites and the plagioclase in these samples (Figs. 7A, 7B). In addition, the ca. 516 Ma Na-rich tonalites show high Al2O3/(FeOT + MgO + TiO2) values and negative Ce anomalies (Ce/Ce* = 0.75–0.85; Table 2), and are characterized by enrichments in LREEs, LILEs (e.g., K, Rb, and Sr), and Pb, as well as by depletions in Th, Nb, Ta, and Ti, and high but variable Ba/La ratios (Figs. 7 and 9). These features suggest that the subducted-sediment–derived fluids were involved in the generation of their magmas (Figs. 9D, 9E; Hole et al., 1984; Hawkesworth et al., 1997; Kepezhinskas et al., 1997; Rosenbaum et al., 1997; Plank and Langmuir, 1998; Elliott, 2003; Hanyu et al., 2006; Guo et al., 2015). We conclude that these tonalitic magmas were generated by the partial melting of ancient accumulated gabbroic rocks with the involvement of subducted-sediment–derived fluids. The REE and trace element compositions of the ca. 516 Ma tonalites are similar to those of pegmatites in the Catalina Schist, a subduction zone terrane in southern California (Figs. 7A, 7B; Sorensen and Grossman, 1989), supporting this model for their origin.

Figure 9.

(A) TiO2–SiO2/50–K2O ternary plot for early Paleozoic igneous rocks from the Zhangguangcai Range, northeastern China (France et al., 2010). Other symbols are as in Figure 6. Fract. cryst.—fractional crystallization; MORB—mid-oceanic ridge basalt. (B) Rb/Ba versus Rb/Sr plot (Sylvester, 1998). (C) CaO/Al2O3 versus CaO + Al2O3 plot. (D) Al2O3/(FeOT + MgO + TiO2) versus Al2O3 + FeOT + MgO + TiO2 plot (Patiño Douce, 1999; Yu et al., 2017). (E) Th/Yb versus Ba/La (Hanyu et al., 2006) plot. (F) (La/Yb)N versus YbN (Martin, 1986) and Sr/Y versus Y (Defant and Drummond, 1990) plot.

Figure 9.

(A) TiO2–SiO2/50–K2O ternary plot for early Paleozoic igneous rocks from the Zhangguangcai Range, northeastern China (France et al., 2010). Other symbols are as in Figure 6. Fract. cryst.—fractional crystallization; MORB—mid-oceanic ridge basalt. (B) Rb/Ba versus Rb/Sr plot (Sylvester, 1998). (C) CaO/Al2O3 versus CaO + Al2O3 plot. (D) Al2O3/(FeOT + MgO + TiO2) versus Al2O3 + FeOT + MgO + TiO2 plot (Patiño Douce, 1999; Yu et al., 2017). (E) Th/Yb versus Ba/La (Hanyu et al., 2006) plot. (F) (La/Yb)N versus YbN (Martin, 1986) and Sr/Y versus Y (Defant and Drummond, 1990) plot.

The ca. 496–482 Ma granitoids (including the monzogranites and biotite granodiorites) in the study area have P2O5 and Al2O3 concentrations that negatively correlate with SiO2 concentrations, suggesting that they are I-type granites (Chappell, 1999). The dominantly peraluminous affinity, variable SiO2 concentrations, high CaO/Al2O3 values, low Rb/Sr (0.20–0.68), Rb/Ba (0.17–0.23), and Al2O3/(FeOT + MgO + TiO2) ratios (2.24–5.20) of these I-type granites suggest that they formed from magmas derived from an intermediate to mafic source (e.g., amphibolite; Figs. 9B–9D; Sylvester, 1998; Patiño Douce, 1999), consistent with the general model for the generation of I-type granite magma by the dehydration melting of igneous rocks or their metamorphic equivalents (Wolf and Wyllie, 1994; Chappell, 1999). The ca. 496–482 Ma granitoids have relatively flat but high HREE patterns and low Sr/Y ratios (Figs. 7C and 9F). These results, combined with the zircon εHf(t) values (+1.08 to +3.58) and TDM2 ages (1385–1238 Ma) for these granitoids, suggest they were derived from magmas generated by the partial melting of Mesoproterozoic lower crustal material at pressures below the garnet stability field (Rapp et al., 1991; Martin, 1999). The ca. 482 Ma granodiorites contain lower concentrations of SiO2, higher concentrations of MgO, Cr, and Ni, and have higher Mg# values than the ca. 496 Ma monzogranites (Table 2), suggesting that the magmas that formed these intrusions contained minor amounts of mantle-derived material (Yang et al., 2007; Tang et al., 2014), consistent with the elevated TZr values (863–865 °C; Watson et al., 2006) of these granodiorites.

Early–Middle Ordovician (ca. 475–461 Ma) granitoids in the study area are dominated by weakly peraluminous I-type biotite monzogranites and alkali-feldspar granites. In particular, the ca. 475 and 461 Ma biotite monzogranites in the Tadong and Shuguang areas contain relatively low concentrations of HREEs and Y (Figs. 7E, 7F), and have high Sr/Y ratios, suggesting that these intrusions have adakitic affinities (Fig. 9F; Martin, 1986; Defant and Drummond, 1990; Q. Wang et al., 2006). The absence of coeval mafic igneous rocks and the slightly positive Eu anomalies of these biotite monzogranites preclude their derivation by the fractional crystallization of basaltic magmas (Castillo et al., 1999). These monzogranites are also relatively K-rich, have low Mg# values, low concentrations of Cr, Co, and Ni, and contain magmatic zircon crystals with variable but low εHf(t) values (+3.78 to +5.44 and –0.82 to +1.48) and ancient TDM2 ages (1209–1103 Ma and 1490–1344 Ma), indicating that they were derived from magmas generated by the partial melting of thickened Mesoproterozoic lower continental crustal material under pressures and temperatures where garnet ± amphibole are stable (Petford and Atherton, 1996; Martin, 1999; Chung et al., 2003; Garrison and Davidson, 2003; Xiong et al., 2003; Cao et al., 2013; Li et al., 2015), rather than by the partial melting of subducted oceanic crust (Defant and Drummond, 1990; Defant et al., 1991) or delaminated thickened crustal material (Xu et al., 2002; Xu et al., 2006; Gao et al., 2008). In contrast, the ca. 462 Ma I-type alkali-feldspar granites in the Shuguang area have flat but elevated HREE patterns (e.g., Yb concentrations of 1.90–3.37 ppm; Fig. 7E) and low Sr/Y ratios, suggesting that they were derived from a garnet-free source. These alkali-feldspar granites also contain zircon grains with εHf(t) values of +1.08 to +4.79 that yield ancient TDM2 ages (1385–1135 Ma). These signatures suggest that their magmas were generated by the partial melting of Mesoproterozoic lower crustal material at pressures below the garnet stability field. The ca. 462 Ma alkali-feldspar granites also show relatively high Na2O and K2O concentrations, but are strongly depleted in Eu, Ba, Sr, P, and Ti (Figs. 7E, 7F), implying that these intrusions recorded the fractional crystallization of plagioclase, K-feldspar, apatite, and Fe-Ti oxides prior to emplacement.

The middle Silurian (ca. 426 Ma) low-K tonalites are geochemically similar to the ca. 516 Ma tonalites, as exemplified by their high SiO2 and Na2O contents, high CaO/Al2O3 values, low TiO2 contents, and Rb/Sr (0.013–0.021) and Rb/Ba (0.031–0.037) ratios (Figs. 9A–9C), weakly peraluminous affinities (A/CNK = 1.07–1.09), high Al2O3/(FeOT + MgO + TiO2) values (Fig. 9D) and slightly negative Ce anomalies (Ce/Ce* = 0.79–0.84; Table 2), enrichments in LREEs, LILEs, and Pb, depletions in Th, Nb, Ta, and Ti (Figs. 7G, 7H), and the negative εHf(t) values (–8.58 to –4.87) of both captured and primary zircon grains (Table 3). However, the ca. 426 Ma tonalites have negligible Eu anomalies. These observations suggest that these tonalitic magmas were generated by the partial melting of ancient gabbroic or amphibolitic rocks with the involvement of subducted-sediment–derived fluids (Hole et al., 1984; Hawkesworth et al. 1997; Kepezhinskas et al., 1997; Plank and Langmuir, 1998; Hanyu et al., 2006; Guo et al., 2015). This view is supported by the extreme Sr enrichments and negligible Eu anomalies of these samples (Figs. 7G, 7H) that reflect the sourcing of Sr from altered oceanic crust in the downgoing slab (Elliott, 2003; Staudigel, 2003; Niu et al., 2013).

Tectonic Setting of Early Paleozoic Magmatism and Tectonic Evolution of the SZB and JB

The early Cambrian (ca. 516 Ma) magmatism in the southeastern SZB formed a series of Na-rich tonalite intrusions. Such Na-rich granitoids can form in a subduction zone setting as a result of the partial melting of accumulated gabbroic rocks or subducted oceanic crust or thickened lower continental crust (Drummond and Defant, 1990; Petford and Atherton, 1996; Koepke et al., 2007; Pei et al., 2014), or during continent-continent collision by the high-pressure hydrous melting of metasedimentary or metabasaltic rocks (Patiño Douce and Harris, 1998; Gao et al., 2009). The origin of ca. 516 Ma tonalites incorporating ancient accumulated gabbroic rocks as well as subducted-sediment–derived fluids indicates that they formed in a subduction zone environment (Fig. 10A; Sorensen and Grossman, 1989; Woodhead, 1998; Hanyu et al., 2006; Jian et al., 2008) rather than in a continent-continent collisional setting. This view is consistent with the exposure of these units near the Jiayin-Mudanjiang belt in an area that consists of a series of subducted accretionary complexes and has been considered to represent the suture belt between the SZB and the JB (Fig. 2A; Zhang, 1992; Li et al., 1999; Xie et al., 2008a; Wang et al., 2012a; Zhang et al., 2015). In addition, the ca. 516 Ma tonalites have low whole-rock FeOT / MgO values (Fig. 6D) and relatively high zircon Ce/Ce* values (Fig. 11) that suggest that they formed in a highly oxidized environment, consistent with their formation near the subduction zone (Berndt et al., 2004; Liang et al., 2006; Koepke et al., 2007; Trail et al., 2011, 2013; Liu et al., 2016, and references therein).

Figure 10.

Simplified tectonic models for the early Paleozoic evolution of the eastern Songnen–Zhangguangcai Range block (SZB) and the adjacent Jiamusi block (JB). LXR––Lesser Xing’an Range.

Figure 10.

Simplified tectonic models for the early Paleozoic evolution of the eastern Songnen–Zhangguangcai Range block (SZB) and the adjacent Jiamusi block (JB). LXR––Lesser Xing’an Range.

Figure 11.

(A) Variations in bulk-rock Sr/Y versus formation ages for early Paleozoic igneous rocks in the eastern Songnen–Zhangguangcai Range block (SZB), from north to southeast. Gray open rhombuses and squares indicate data from Wang et al. (2012a); other symbols are as in Figure 6. (B) Magmatic zircon Hf TDM2 (depleted mantle model) ages. (C) Zircon Ce/Ce*.

Figure 11.

(A) Variations in bulk-rock Sr/Y versus formation ages for early Paleozoic igneous rocks in the eastern Songnen–Zhangguangcai Range block (SZB), from north to southeast. Gray open rhombuses and squares indicate data from Wang et al. (2012a); other symbols are as in Figure 6. (B) Magmatic zircon Hf TDM2 (depleted mantle model) ages. (C) Zircon Ce/Ce*.

The late Cambrian to Early Ordovician (ca. 496–482 Ma) medium- to high-K calc-alkaline I-type biotite granodiorite and monzogranite intrusions in the study area are associated with minor amounts of metamorphosed basaltic andesites (our data). These ca. 496–482 Ma granitoids are enriched in LREEs and LILEs, and are depleted in HREEs and HFSEs (e.g., Nb, Ta, and Ti; Figs. 7C, 7D). The geochemistry of these associated intrusions is indicative of their derivation from magmas that formed in an active continental margin setting (Gill, 1981; Grove and Donnelly-Nolan, 1986; Grove et al., 2003; Deering et al., 2007; Lu and Xu, 2011; Wang et al., 2012a; Tang et al., 2014, 2016). Previous research suggested that the oceanic crust between the southern SZB and the southern JB existed during the Neoproterozoic to Ordovician (Li et al., 1999; Xie et al., 2008a). In addition, Ordovician calc-alkaline igneous rocks occurred in the eastern SZB but not in the southern JB (Fig. 8; Wilde et al., 2003; Wang et al., 2012a; Bi et al., 2014; Yang et al., 2014; Wang et al., 2016a). A deep reflection seismic profile has also identified a westward-dipping fossil subduction zone beneath the eastern SZB (Zhang et al., 2015). These results suggest that the early Cambrian to Early Ordovician magmatism in the southeastern SZB occurred in an active continental margin setting associated with the northwestward subduction of an oceanic plate between the southern SZB and the southern JB (Figs. 10A, 10B; Li et al., 1999; Wang et al., 2012a; Xu et al., 2012).

In contrast, the collision between the northern SZB and the northern JB may have occurred during the Cambrian, with subsequent postcollisional extension occurring during the late Cambrian (J.F. Liu et al., 2008; Wang et al., 2016a). This view is supported by the presence of middle–late Cambrian (ca. 505–496 Ma) K-rich adakitic monzogranites and late Cambrian (ca. 491–486 Ma) A-type granitoids in the northern SZB (Figs. 10A, 10B; J.F. Liu et al., 2008; Wang et al., 2016a), and by the ca. 508 and ca. 490 Ma collisional and postcollisional granitoids in the northern JB (Bi et al., 2014). In addition, the increase of Sr/Y ratios in the ca. 505–496 Ma monzogranites over time in the northern SZB (Fig. 11A; Wang et al., 2016a) is consistent with crustal thickening (Chapman et al., 2015) associated with the Cambrian continent-continent collision. The decrease in Sr/Y recorded by the ca. 490 Ma A-type granitoids coincides with a crustal thinning event related to postcollisional extension (Figs. 10A, 10B; J.F. Liu et al., 2008; Wang et al., 2016a). Furthermore, combining field observations with our unpublished detrital zircon U-Pb age data for strata in the northwestern part of the northern SZB indicates the presence of a significant unconformity between the early Cambrian trilobite-bearing carbonate units (i.e., the Wuxingzhen Formation) and the late Cambrian siliciclastic units (depositional age between ca. 506 and 496 Ma), an unconformity that was probably generated by the Cambrian continent-continent collision.

The Early–Middle Ordovician igneous rocks (ca. 475–461 Ma) in the southeastern SZB consist of a series of high-K calc-alkaline biotite monzogranites and alkali-feldspar granites (Fig. 6). The biotite monzogranites in the Tadong and Shuguang areas have adakitic affinities, and were generated by the partial melting of the thickened lower crust associated with the onset of compressional tectonism. The similar Early–Middle Ordovician igneous rocks (ca. 470 Ma) in the northern SZB consist of a suite of calc-alkaline intermediate to felsic intrusive rocks and minor rhyolites (Wang et al., 2016a). The Sr/Y values of the ca. 470 Ma granitoids increase over time (Fig. 11A), suggesting that crustal thickening also occurred in the northern SZB. In addition, the Early–Middle Ordovician magmatic rocks (barring the 462 Ma alkali-feldspar granites) have Sr/Y ratios that increase from north (the Lesser Xing’an Range) to southeast (the Tadong area) in the eastern SZB (Fig. 11A). This indicates that the eastern SZB underwent compressional tectonism during the Early to Middle Ordovician, with crustal thickening being more intense in the southeastern than the northern SZB. These rock associations and geochemical features suggest that the Early–Middle Ordovician igneous rocks formed in a compressional active continental margin setting (Ducea et al., 2015). In addition, the residual oceanic crust between the southern SZB and the southern JB still existed during the Ordovician, with no coeval igneous rocks or unconformity identified in the western JB (Heilongjiang Bureau of Geology and Mineral Resources, 1993; Li et al., 1999; Xie et al., 2008a). This precludes the possibility of collision between the southern parts of these two blocks during the Early–Middle Ordovician, and further confirms that the active continental margin setting recorded in the study area could be related to the northwestward subduction of an oceanic plate between the southern SZB and the southern JB (Fig. 10C; Li et al., 1999; Wang et al., 2016a).

Combining new data with the results of previous research suggests that Late Ordovician igneous rocks (ca. 460–445 Ma) crop out widely in the eastern SZB (Wang et al., 2012a; Wang et al., 2016a; L.-Y. Zhang, 2017, personal commun.). The Late Ordovician igneous rocks in the southeastern and central SZB (i.e., in the Tadong and Shuguang areas) consist of a calc-alkaline suite that includes diabase, diorite, tonalite, monzogranite, and rhyolite units, all of which have arc-type igneous rock affinities (Wang et al., 2012a; L.-Y. Zhang. 2017, personal commun.). The Late Ordovician magmatism in the southeastern SZB probably occurred in an active continental margin setting associated with the northwestward subduction beneath the southeastern SZB (Fig. 10D; Li et al., 1999; Wang et al., 2012a; Xu et al., 2012). The Late Ordovician igneous rocks in the northern SZB are composed of a series of A-type rhyolites and alkali-feldspar granites that formed in locations distal from the subduction zone, implying an extensional setting most likely associated with backarc extension (Fig. 10D; Li et al., 1999; Wang et al., 2012a; Wang et al., 2016a). The backarc extension is also indicated by the presence of Middle–Late Ordovician Xiaojingou Formation shallow-marine facies carbonate and siltstone sedimentary units interlayered with minor amounts of andesite in the inboard region of the southeastern SZB (Fig. 10D; Heilongjiang Bureau of Geology and Mineral Resources, 1993; Li et al., 1999).

The ca. 426 Ma Na-rich tonalites of the Tadong area are similar geochemically to the ca. 516 Ma tonalites; they formed from magmas generated by the partial melting of ancient gabbroic or amphibolitic rocks with the addition of subducted-sediment–derived fluids. This suggests that the northwestward subduction of the oceanic plate beneath the southeastern SZB could have continued until at least the middle Silurian. The formation of coeval (ca. 420 Ma) calc-alkaline igneous rocks in the southeastern SZB, including basalts, basaltic andesites, and andesites (L.-Y. Zhang, 2017, personal commun., and our data), also supports the continuation of subduction until this time. In contrast, the presence of ca. 425 Ma collision-related peraluminous monzogranites in the central SZB (i.e., the Shuguang area; Wang et al., 2012a), combined with the absence of the Silurian sedimentary units in the northern and central SZB as well as the JB, suggests that the collision between the central SZB and the central JB could have occurred during the middle Silurian. This indicates that the central to southern SZB progressively collided with the JB before final collision occurred at the end of the early Paleozoic (Fig. 10E). This was followed by the formation of the Devonian bimodal volcanic rocks (ca. 386 Ma) and stable sedimentary cover sequences in both two blocks (Heilongjiang Bureau of Geology and Mineral Resources, 1993; Li et al., 1999; Meng et al., 2010, 2011a; Xu et al., 2012).

Crustal Accretion and Reworking in the SZB

It remains unclear whether Precambrian crust existed beneath the SZB and when crustal growth and reworking occurred in this area. Zircon Hf isotopes can be used to track the history of the chemical differentiation of crust and mantle as well as the heterogeneity of the continental crust (Kinny et al., 1991; Yang et al., 2007; Kemp et al., 2009; Dhuime et al., 2012; Tang et al., 2016; Vervoort and Kemp, 2016). The new zircon Hf isotopic data for the early Paleozoic igneous rocks presented in this study can provide key insights into the crustal accretion and reworking processes recorded in the eastern SZB.

The zircon Hf isotopic compositions of the early Paleozoic igneous rocks in the southeastern SZB indicate that crustal accretion in this area occurred between the Paleoproterozoic and Mesoproterozoic (TDM2 age peaks at 1926–1715 and 1484–1209 Ma, as well as a secondary peak at 1091 Ma), which is comparable to the northern SZB (Fig. 11B; Wang et al., 2016a), suggesting that deeper parts of the crust beneath the eastern SZB could have contained ancient Precambrian basement material. This inference is supported by the presence of many Paleoproterozoic and Mesoproterozoic detrital zircon grains in Paleozoic sediments in this area (Meng et al., 2010; Wang et al., 2012b, 2014; Zhou et al., 2012). Recent research also identified similar crustal accretion events in other microcontinents in the eastern Central Asian Orogenic Belt, including the adjacent JB (ca. 1.7–1.2 Ga; Fig. 12B; Bi et al., 2014; Yang et al., 2014) and the Erguna block (major peaks ca. 1.8 and 1.4–1.2 Ga; Fig. 12C; Ge et al., 2007a). This suggests that the microcontinents in the eastern Central Asian Orogenic Belt (including the SZB, the JB, and the Erguna block) share a common Paleoproterozoic and Mesoproterozoic continental crustal basement that is coupled with the coeval subcontinental lithospheric mantle underlying the eastern Central Asian Orogenic Belt (Re depletion age peaks ca. 1.8 and 1.3 Ga; Fig. 12A; Wu et al., 2003; Zhou et al., 2007, 2010; Zhang et al., 2011; Guo et al., 2017).

Figure 12.

Comparisons of magmatic (black line) and captured (gray shadow area) zircon Hf TDM2 (depleted mantle model) age probability plots for early Paleozoic igneous rocks. (A) Eastern Songnen–Zhangguangcai Range block (SZB) (this paper; Wang et al., 2012a; Wang et al., 2016a). CAOB—Central Asian Orogenic Belt. (B) Jiamusi block (JB) (Bi et al., 2014; Yang et al., 2014). (C) Erguna block (Ge et al., 2007a). Dark gray line in A indicates Re depletion ages (TRD ages) of mantle-derived rocks from the eastern CAOB (Wu et al., 2003; Zhou et al., 2007, 2010; Zhang et al., 2011; Guo et al., 2017).

Figure 12.

Comparisons of magmatic (black line) and captured (gray shadow area) zircon Hf TDM2 (depleted mantle model) age probability plots for early Paleozoic igneous rocks. (A) Eastern Songnen–Zhangguangcai Range block (SZB) (this paper; Wang et al., 2012a; Wang et al., 2016a). CAOB—Central Asian Orogenic Belt. (B) Jiamusi block (JB) (Bi et al., 2014; Yang et al., 2014). (C) Erguna block (Ge et al., 2007a). Dark gray line in A indicates Re depletion ages (TRD ages) of mantle-derived rocks from the eastern CAOB (Wu et al., 2003; Zhou et al., 2007, 2010; Zhang et al., 2011; Guo et al., 2017).

Zircon Hf isotopic compositions can also be used to reveal the crustal reworking processes, and our new data indicate that coeval granitoids and rhyolites in different areas and granitoids and rhyolites that formed at different times in the same area have different zircon Hf isotopic compositions, suggesting the heterogeneity of lower crust beneath the eastern SZB.

Widespread early Cambrian to middle Silurian granitoid and rhyolite magmatism is recorded in the eastern SZB. However, the ca. 490 Ma and ca. 470 Ma granitoids in the Tadong area have much younger zircon Hf TDM2 ages (peaks ca. 1.2 Ga) than coeval granitoids and rhyolites from the Shuguang area and the Lesser Xing’an Range (main peaks ca. 1.5–1.4 Ga and a secondary peak ca. 1.8 Ga; Fig. 11B; Wang et al., 2016a). It suggests that the reworked lower crustal material increases in age from the southeastern SZB to the northern SZB, supporting the presence of laterally heterogeneous deep crustal material beneath the eastern SZB.

The large variations in zircon Hf TDM2 ages between the different stages of early Paleozoic granitoids and rhyolites in the same area are exemplified by the gradual increase in zircon Hf TDM2 ages of the Early Ordovician to middle Silurian granitoids in the central SZB (i.e., the Shuguang area; Fig. 11B). The zircon Hf TDM2 ages of the majority of the ca. 490 and ca. 470–450 Ma granitoids and rhyolites in the Lesser Xing’an Range also decrease over time (Fig. 11B; Wang et al., 2016a). These temporal variations in zircon Hf isotopic data, combined with the spatial variations discussed here, suggest that the deep crust beneath the eastern SZB is vertically and laterally heterogeneous. The reworking of the heterogeneous lower crustal material beneath the eastern SZB generated the widespread early Paleozoic granitoid and rhyolite magmatism in the study area.

  1. Widespread early Cambrian to middle Silurian (age peaks ca. 516, 497, 478–462, 451, and 425 Ma) magmatic events occurred in the southeastern SZB.

  2. The early Cambrian to middle Silurian magmatism consisted of low-K tholeiitic to calc-alkaline granitoids. They formed in an active continental margin setting associated with the northwestward subduction of an oceanic plate between the southern SZB and the southern JB.

  3. Subduction beneath the southeastern SZB occurred between the early Cambrian and middle Silurian, whereas the Cambrian tectonic evolution of the northern SZB was dominated by collision-related processes associated with the northern JB, and the northwestward subduction had affected the northern SZB since the Early Ordovician.

  4. Crustal growth in the eastern SZB occurred mainly between the Paleoproterozoic and Mesoproterozoic. Early Paleozoic subduction- and collision-related processes in this region caused significant reworking of the ancient crust in the eastern SZB.

We thank the staff of the State Key Laboratory of Geological Processes and Mineral Resources, China University of Geosciences, Wuhan, and the Geological Laboratory Centre, China University of Geosciences, Beijing, for their advice and assistance during zircon laser ablation–inductively coupled plasma–mass spectrometry U-Pb dating, major and trace element analyses, and zircon Hf isotope analysis. We also thank Damian Nance for editorial handing, Wenjiao Xiao and two anonymous reviewers for constructive comments, and Zhuang Li (Peking University) for helpful discussions. This work was supported by the National Natural Science Foundation of China (grants 41572043 and 41330206), the National Basic Research Program of China (grant 2013CB429802), and the Opening Foundation of the State Key Laboratory of Geological Processes and Mineral Resources, China University of Geosciences (Wuhan) (grant GPMR201503).

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