Neoproterozoic intraplate crustal accretion on the northern margin of ...

Precambrian Research 268 (2015) 97–114

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Neoproterozoic intraplate crustal accretion on the northern margin of the Yangtze Block: Evidence from geochemistry, zircon SHRIMP U–Pb dating and Hf isotopes from the Fuchashan Complex Lei Liu a , Xiaoyong Yang a,∗ , M. Santosh b,c , S. Aulbach d , Hongying Zhou e , Jianzhen Geng e , Weidong Sun f a CAS Key Laboratory of Crust-Mantle Materials and Environments, School of Earth and Space Sciences, University of Science and Technology of China, 96 Jinzhai Road, Hefei 230026, China b School of Earth Sciences and Resources, China University of Geosciences, 29 Xueyuan Road, Beijing 100083, China c Division of Interdisciplinary Science, Faculty of Science, Kochi University, Kochi 780-8520, Japan d Institut für Geowissenschaften, Johann Wolfgang Goethe-Universität, 60438 Frankfurt Am Main, Germany e Tianjin Institute of Geology and Mineral Resources, China Geological Survey, Tianjin 300170, China f CAS Key Laboratory of Mineralogy and Metallogeny, Guangzhou Institute of Geochemistry, 511 Kehua Street, Wushan, Guangzhou 510640, China

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Article history: Received 1 February 2015 Received in revised form 8 June 2015 Accepted 13 July 2015 Available online 22 July 2015 Keywords: Zircon U–Pb geochronology Yangtze Block Neoproterozoic Bimodal rift magmatism Fuchashan Complex

a b s t r a c t The South China Block preserves important imprints of Neoproterozoic intraplate magmatism in a continental setting associated with the amalgamation of the Yangtze and Cathaysia sub-blocks. Here we investigate the Fuchashan Complex exposed to the east of the NE to NNE trending Tan-Lu Fault (TLF) that dissects the North China Craton (NCC) and South China Block (SCB) in East China. We performed an integrated geochemical and geochronologic study on the basaltic to granitic rocks from the complex. Most of the zircons from one amphibolite, two diorites and one granitoid are characterized by oscillatory zoning with homogeneous cores, high Th/U ratios, variable negative Eu anomalies, and extremely HREE-enriched patterns, attesting to magmatic origin. Zircons from the four samples yield consistent Neoproterozoic ages (ca. 794–812 Ma), correlating with the widespread bimodal rift magmatism along the northern margin of SCB, possibly related to the Mid-Neoproterozoic breakup of the supercontinent Rodinia. The zircons from amphibolites show negative εHf (t) values of −16.4 and −8.3 and Hf model ages of 1699 ± 14 Ma and 2023 ± 10 Ma, indicating origin from partial melting of Paleoproterozoic enriched mantle. Their low SiO2 (48.28–51.75 wt%) contents also support mantle affinity. Furthermore, their high Mg numbers (54–60), and Cr (90–190 ppm) and Ni (31.6–51.1 ppm) content indicate primary and less evolved source magma. The markedly negative zircon εHf (t) values in the diorites and granitoids suggest partial melting of the Neoarchean to Paleoproterozoic lower crust. The extremely low zircon εHf (t) of −26.5 from a granitoid suggest reworking of Mesoarchean basement. A probable scenario for the generation of the extensive ca. 800 Ma bimodal suite in this region is the widespread Mid-Neoproterozoic intraplate magmatism triggered by asthenosphere upwelling, resulting in the partial melting of enriched subcontinental lithospheric mantle and underplating of mafic magma. Heat input from this magma led to further melting of lower continental crust at depths of 35–40 km, and the melt compositions controlled by fractionation of clinopyroxene and minor plagioclase under high fO2 conditions. © 2015 Elsevier B.V. All rights reserved.

1. Introduction Growth of juvenile crust in divergent plate margins and convergent plate boundaries has been well documented in many regions.

∗ Corresponding author. Tel.: +86 551 63606871. E-mail address: [email protected] (X. Yang). http://dx.doi.org/10.1016/j.precamres.2015.07.004 0301-9268/© 2015 Elsevier B.V. All rights reserved.

However, the tectonic and geophysical evidence indicates that rifted margins are likely the most productive areas for net growth of the continental crust (Courtillot et al., 1999; Hart et al., 1989; Peccerillo et al., 2003; Smedley, 1986; White and McKenzie, 1989; Zheng et al., 2006a). Thus it is important to understand the petrogenesis of magmatic suites formed in continental rift settings (Courtillot et al., 1999; Hart et al., 1989; Peccerillo et al., 2003; Smedley, 1986; White and McKenzie, 1989; Zheng et al., 2006a).

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L. Liu et al. / Precambrian Research 268 (2015) 97–114

The South China Block (SCB) has been central to the debate on the tectonic evolution of the Asian continental collage, as well as on the assembly, evolution and dispersal of the Neoproterozoic supercontinents Rodinia and Gondwana (e.g., Li et al., 2003a,b; Duan et al., 2011; Zhao and Cawood, 2012; Zhang and Zheng, 2013; Li et al., 2014; Kim et al., 2014; Nance et al., 2014; Santosh et al., 2014; Xu et al., 2014). Widespread Neoproterozoic continental intraplate magmatism has been recorded in the SCB, which is correlated to the dynamics associated with the amalgamation of the Yangtze Block (YB) and Cathaysia Block (CB) during the 1.0–0.86 Ga Sibao orogenesis (e.g., Li et al., 2002, 2003a, b, 2005, 2006, 2008a,b, 2009; Li et al., 2010; Ling et al., 2003; Zhang and Zheng, 2013), although some recent studies have assigned younger ages in the range of 0.86–0.82 Ga (Zhao et al., 2011; Wang et al., 2014a). The intraplate magmatism is commonly attributed to mantle plumes or superplume upwelling that broke up the supercontinent Rodinia (Li et al., 2003a,b; Li et al., 2008a). The Mid-Neoproterozoic magmatic suites display two age peaks of 830–800 Ma and 780–740 Ma (Zhang and Zheng, 2013). Zircon Hf isotope data suggest that the magmatism involved both addition of depleted mantle material to the juvenile Mesoproterozoic crust and remelting of preexisting old crustal components within rift setting (Li et al., 2008a; Zheng et al., 2006b; Zhang and Zheng, 2013). The volcanic and intrusive rocks from the SCB have been the subject of extensive tectonic, geochronological and geochemical investigations (e.g., Zhao and Cawood, 1999; Li et al., 2003a,b; Zhou et al., 2002; Zhao and Cawood, 2012). However, the petrogenesis and tectonic affiliations of these rocks remain equivocal. Two contrasting models have been proposed to interpret their formation, i.e. horizontal accretion in a collision-arc geodynamic environment (Zhao and Cawood, 1999; Zhou et al., 2002; Wang et al., 2004) vs. vertical crustal growth by mafic magma underplating/upwelling within a rift and plume-related environment (Li et al., 2003a,b; Li et al., 2008a). The main arguments supporting the former model are: (1) arc signatures from geochemical compositions (e.g., Nb, Ta and Ti depletion); and (2) the presence of arc assemblages showing identical Neoproterozoic ages (Zhou et al., 2002; Wang et al., 2004). The plume model is not only based on geochemical, geochronological and paleomagnetic data placing the SCB in the western part of Rodinia from intraplate basaltic rocks (Li et al., 2003a,b; Li et al., 2008a), but also other aspects of geological observations such as basin analysis and petrologic data (including anomalously high temperature suggesting komatiitic basalts) (Li et al., 2008b, 2014; Wang and Li, 2003; Wang et al., 2007a, 2012). These contrasting models require further evaluation in order to understand the geodynamics and crustal growth in the SCB. The Fuchashan Complex (also referred to as Fuchashan Pluton, Feidong Group/Terrane, and Fuchashan metamorphic zone in previous literature) has attracted considerable attention, particularly due to its proximity to the Tan-Lu Fault (TLF) (Lin et al., 2005, 2009; Zhang et al., 2007; Zhu et al., 2002, 2010; Shi et al., 2009; Kang et al., 2013). The complex has been traditionally considered as representing the initiation of the TLF as a sinistral transcurrent boundary in Early Triassic time (Zhang et al., 2007; Zhu et al., 2002, 2009, 2010), although the origin and formation age are poorly constrained. A precise understanding of the formation age and tectonic affinity of this complex is important in understanding the initiation and evolution of the TLF. Zircon is a highly robust mineral in most geological environments, and has been extensively used for dating as well as for geochemical tracing (Hanchar and Hoskin, 2003). Igneous zircons commonly record the multistage evolution of their host rocks (Schaltegger et al., 1999; Rubatto, 2002), and thus a combined study of in-situ U–Pb dating, cathodoluminescence (CL) imaging, trace element analysis and Hf isotope determination can be used to decipher the age and evolutionary history of their host rocks.

Furthermore, Lu–Hf isotope in zircons has been used to trace the origin and the evolution of crust and mantle over time (Amelin et al., 1999, 2000; Griffin et al., 2000, 2002; Zheng et al., 2006a,b). In this paper, we report new sensitive high-resolution ion microprobe (SHRIMP) zircon U–Pb ages, trace element compositions and Lu–Hf isotopes from Neoproterozoic magmatic rocks on the northern margin of YB. The results suggest that the protoliths of the Fuchashan Complex belong to bimodal magmatism through asthenospheric upwelling within a rift setting, and that these rocks were derived through crystal fractionation of dominantly clinopyroxene and minor plagioclase. 2. Geological setting The Fuchashan Complex in Feidong County, Anhui Province, is exposed to the east of the TLF as a NNE-trending narrow and linear zone of ca. 35 km length and ca. 4 km width (Fig. 1). The major rock types in this complex are biotite–plagioclase gneiss and hornblende–plagioclase gneiss, associated with slivers of metasedimentary-metavolcanic rocks (Zhang et al., 2007; Tong and Xu, 2000). Minor amphibolites occur as blocks or lenses within the gneisses (Fig. 2). All these rocks were subjected to loweramphibolite facies metamorphism during the Early Cretaceous as inferred from the metamorphic mineral paragenesis of hornblende + plagioclase + titanite (Zhang et al., 2007; Zhu et al., 2002, 2010). Estimates of the metamorphic conditions show that the peak P–T conditions for the Fuchashan Complex are in the range of 610–690 ◦ C and 0.61–0.81 GPa, and that the average P–T conditions are 656 ± 25 ◦ C and 0.71 ± 0.06 GPa, corresponding to the amphibolite facies (Kang et al., 2013). Their protoliths are considered to be intermediate-mafic plutonic rocks (Jing et al., 1991; Anhui Geological Survey, 1998). The gneisses form part of a NNE–SSW trending antiform with a brittle normal fault corresponding to the TLF located along the western margin of the antiform (Lin et al., 2005). The contact relationship between amphibolite and granitic gneiss (Fig. 2c and d) suggests that the protoliths of the amphibolites intruded into the granitic gneiss in a semi-solid state. These features clearly suggest that the mafic and felsic magmas were coeval, and in places co-magmatic. The whole-rock, biotite and amphibole 40 Ar–39 Ar ages reported in previous studies from these rocks show metamorphic ages of 100–145 Ma, and were regarded as the time of the TLF had evolved into normal faulting (Zhu et al., 2002, 2010; Zhang et al., 2007). The U–Pb age (809 ± 7 Ma) of zircons from an amphibolite from this complex is thought to constrain its formation age, correlating with the YB (Kang et al., 2013). Based on petrographic observation, Kang et al. (2013) divided the complex into three units, composed of granitic gneiss enclosing minor amphibolite lenses, biotite–plagioclase gneiss enclosing banded or lentoid amphibolites, and granodioritic gneisses and hornblende–plagioclase gneisses. In contrast, Xu et al. (2002) proposed that the complex consists of two parts based on field observation and lithological characteristics: a lower fine-grained monzonitic gneiss interlayered with hornblende or biotite plagioclase gneiss, and an upper fine-grained monzonitic gneiss, plagioclase gneiss and hornblende–biotite plagioclase gneiss. In this study, the rock types are designated as amphibolite, diorite and granitoid based on their major mineral compositions. 3. Sample description Four representative rocks from the Fuchashan Complex were selected for dating (Fig. 1). Based on their mineral assemblage, sample FC02 is a plagioclase amphibolite, FC01 corresponds to a monzonitic gneiss, KJ01 is a biotite–plagioclase gneiss and FC11


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Fig. 1. Sketch geological map of the Fuchashan Complex on the northern margin of the Yangtze Block and sample locations (modified after Niu et al., 2010). The inset shows the tectonic framework for the study area.

Fig. 2. Field photographs of Neoproterozoic plutonic rocks in this study. (a) Amphibolite interlayered with felsic vein (below the red dashed line); (b) hornblende–plagioclase gneiss enclosed by fine-grained monzonitic gneiss, along the eastern side, foliation in hornblende-plagioclase gneiss dips to the east at a high angle of >70◦ ; (c) amphibolite lens within a high-K granite vein enclosed by biotite–plagioclase gneiss; (d) the lithological contacts between dioritic and granitic gneiss is gradual and transitional, indicating chemical reaction between them.

Table 1 Modal abundance of minerals (vol.%) in representative rocks from the Fuchashan Complex on the northern margin of the Yangtze Block. Sample

Mineral composition (%)

Accessory mineral

Rock name

14FC01 14FC02 14FC05 14FC11 14FC12 14FC16 TL9901 14KJ01

Qtz(30), Pl(25), Kfs(15), Bt(15), Amp(10) Amp(50), Pl(30), Qtz(10), Bt(5), EP(5) Amp(55), Pl(20), Qtz(10), Bt(15) Pl(25), Kfs(25), Qtz(35), Bt(10) Pl(30), Kfs(30), Qtz(30), Bt(10) Pl(40), Kfs(20), Qtz(30), Bt(10) Pl(50), Kfs(20), Qtz(20), Bt(5) Pl(50), Kfs(20), Qtz(15), Bt(10), Amp(5)

Zir, Mgt, Mnz Aln, Zir Zir, Mgt Ep, Zir, Apt Zir, Apt, Mgt Zir, Apt, Mnz Zir, Apt, Mgt Zir, Apt, Ttn

Monzonitic gneiss Plagioclase amphibolite Plagioclase amphibolite Granitic gneiss Granitic gneiss Granitic gneiss Granitic gneiss Biotite–plagioclase gneiss

Mineral abbreviations are after Kretz (1983).

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Fig. 3. Outcrop photos and photomicrographs of characteristic textures and mineral assemblages in samples FC01, FC02, FC11 and KJ01. (a, c, e and g) Field photos of samples FC01, FC02, FC11 and KJ01, red and blue shadow zone in (c) and (g) represent the lithological contact zone; (b) photomicrograph in plane polarized light of amphibolite (FC02) with amphibole, plagioclase, quartz and biotite; (d) diorite (FC01) associated with granitoid, showing elongated crystals oriented parallel to each other, composed of plagioclase, quartz and amphibole with accessory biotite and epidote, plane polarized light; (f) granitoid (FC11), mainly consisting of plagioclase, quartz and K-feldspar with minor biotite; the quartz grains are found in different habits, as discrete, commonly subhedral grains and as anhedral to subhedral granoblastic aggregates, cross polarized light; (h) Diorite (KJ01), mainly consisting of plagioclase, K-feldspar, quartz and amphibole with minor epidote and biotite, cross polarized light.

is a granitic gneiss. Compositionally, they cover basaltic, dioritic to granitic. The summary of mineral modal abundance in these samples is given in Table 1. In general, the rock types show gradational and transitional contact, although in some cases, discrete boundaries can be observed (Fig. 3). The plagioclase amphibolite (FC02, Fig. 3a) occurs as bands (10–50 cm thick) or as blocks within the felsic gneiss. The rock

consists of hornblende (60–70%), plagioclase (20–25%) and quartz (5%). Hornblende grains are anhedral to subhedral, and are oriented with grain size varying from 0.2 to 1.5 mm. Chlorite is commonly developed around their rims. Plagioclase is anhedral with irregular shape, ranging in sizes from 0.1 to 0.4 mm. The accessory minerals are mainly biotite, apatite and magnetite, as well as minor titanite and epidote (Fig. 3b).


The monzonitic gneiss (FC01, Fig. 3c), exposed along the western margin of the complex, is composed of quartz, plagioclase, K-feldspar, biotite and hornblende with monazite and magnetite as minor accessories. The quartz grains in the matrix are generally equant, but also occur as slightly elongate ribbons (Fig. 3d). Grain size of plagioclase and K-feldspar ranges from ∼0.5 mm to ∼5 mm. Biotite occurs as subhedral to euhedral laths. Retrograde chlorite is commonly developed around the edge of biotite. The granitic gneiss (FC11, Fig. 3e) mainly consists of plagioclase (20–25%), K-feldspar (25–30%), quartz (30–35%), biotite (5–10%) and magnetite (∼5%). Plagioclase and K-feldspar grains are subhedral to anhedral, and range in size from 0.4 to 0.8 mm. Biotite is subhedral to euhedral with smaller sizes of 0.3–0.5 mm, and chlorite is commonly developed around the rim of biotite (Fig. 3f). The biotite–plagioclase gneiss (KJ01, Fig. 3g) constitutes the main unit in the Fuchashan Complex. Samples of this rock are composed of quartz (25–30%), plagioclase (25–30%), K-feldspar (10–15%), biotite (15–20%), hornblende ( granitoid (32–36, mean = 34) (Table S3). With respect to transition metal contents, e.g., V, Cr, Ni and Sc, generally, the amphibolites have higher abundances than diorites and granitoids (Table S3). All samples have low LOI contents of 0.14–0.83 wt%, suggesting only weak hydrothermal alteration. The amphibolites have similar chondrite-normalized REE patterns with OIB (ocean island basalt), displaying moderately enriched LREE (light rare earth element), weakly fractionated HREE and slight negative Eu anomalies (Eu/EuN * = 0.93–1.03, N denotes chondrite normalization) (Fig. 8a). Their primitive mantlenormalized trace element patterns show broadly similar shapes, characterized by a marked enrichment in large ion lithophile elements (LILE) such as Rb, Ba, Th, U and K, but pronounced negative anomalies of Ta, P, Zr, Hf and Ti (Fig. 8b). When compared with typical OIB, the amphibolites display more enrichment in Ba, but depletion in Nb and Ta (Fig. 8b), which is characteristic of intraplate basalts.

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Fig. 6. SHRIMP U–Pb zircon dating results for samples FC02 (a), FC01 (b), KJ01 (c) and FC11 (d) from the Fuchashan Complex on the northern margin of the Yangtze Block.

The diorite samples have variable REE contents but similar chondrite-normalized REE patterns with moderately enriched LREE, weakly fractionated HREE and slight positive Eu anomalies (Eu/EuN * = 1.00–1.48) (Fig. 8c, in addition to 4 samples from this study, trace compositions from Tong and Xu (2000) are shown). In the primitive mantle-normalized spider diagrams (Fig. 8d), the rocks exhibit consistent Nb, Ta, P and Ti negative anomalies, suggesting a magmatic arc signature. The LILE show higher enrichment than those in the amphibolites. Four samples (LP15Y61, 2058Y1, 2139Y1 and TP14Y51-1 in Tong and Xu, 2000) have high Sr (825–1302 ppm) and low Y (8.7–13.2 ppm) contents with high Sr/Y ratios (82–130), suggesting adakitic composition (Defant and Drummond, 1990). The other five samples display high Sr (495–1230 ppm) and high Y (17.5–19.9 ppm) contents with slightly low Sr/Y ratios (25–69), typical of the compositional field of arc rocks. When compared with average upper or lower continental crust, these rocks display trace element patterns and contents similar to upper continental crust (UCC). The granitoid samples exhibit strong enrichment in LREE and significantly fractionated HREE (relatively high (La/Yb)N = 18–49) with marked negative Eu anomalies (Eu/EuN * = 0.71–0.85) (Fig. 8e). In contrast to the diorites, they display higher enrichment in LILE and LREE, but more depletion in HFSE (such as P, Ti, Nb and Ta) (Fig. 8f), suggesting that they are less fertile with regard to the magma evolution history. All the granitoids display low Sr/Y ratios (17–30) with variably high Y contents (13.8–19.5 ppm), locating them in the typical arc field. In addition, two samples (FC13 and FC14) show very similar patterns and contents to UCC, whereas the other samples display much more enrichment in LREE and LILE than UCC.

5.3. Zircon trace elements Table S2 (supplementary data) lists the REE and trace elements in zircons that were analyzed for U–Pb dating. The chondritenormalized REE patterns of zircons are presented in supplementary data Fig. S1. Zircons from the amphibolite sample FC02 generally show similar REE patterns, with significant enrichment of HREE, positive Ce and negative Eu anomalies (Fig. S1a). LREE contents are low with a slight variation (0.10 < Pr/PrN * < 0.35, Pr/PrN * = PrN /SQRT(CeN × NdN ), where N denotes the normalization to chondrite values after Sun and McDonough (1989)). Ti-in-zircon temperatures were calculated using the thermometry of Watson et al. (2006), and the results are shown in Fig. S2a. The REE patterns of zircons from diorite sample FC01 are characterized by a clear positive Ce anomaly and a weakly negative Eu anomaly (Fig. S1b). The calculated Ti-in-zircon temperatures are remarkably similar to those in the amphibolite (Fig. S2b), indicating that these rocks might have evolved during the same magmatic event. The HREE contents show limited variation, whereas abundances of the LREE are more variable. Zircons from another diorite sample KJ01 generally show similar REE patterns to those of sample FC01 except for their strong negative Eu anomalies. Furthermore, they have more homogenous HREE and more scatter in LREE (Fig. S1c). The higher LREE enrichment is commonly present within and around the bright portion of the crystals in CL images (Fig. 5c). Alternatively, the LREE enrichment is also as a result of the presence of microscale LREE-bearing mineral inclusions (such as apatite, monazite or epidote), especially in metamorphic zircons (Xia et al., 2010; Wang et al., 2014b). The calculated Ti-in-zircon temperatures


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Fig. 7. Harker variation diagrams for the different rock types from the Fuchashan Complex.

are systematic lower by 50–70 ◦ C than the estimates from samples FC01 and FC02 (Fig. S2c). The REE patterns of zircons from granitoid sample FC11 are characterized by a clear positive Ce anomaly and a slight negative Eu anomaly (Fig. S1d). HREE show limited variation, whereas LREE are more variable. Only three spots displaying LREE enrichment may be due to the presence of microscale LREE-bearing mineral inclusions. Except for analysis 6.1c, which yields an abnormally high Ti content of 205 ppm, the other 9 analyses have Ti contents of 21.8–44.5 ppm, and the calculated Ti-in-zircon temperatures are indistinguishable from the data obtained from samples FC02 and FC01 (Fig. S2d). 5.4. Zircon Lu–Hf isotopes In-situ Lu–Hf isotope analyses were carried out on the zircon grains from the four dated rocks of the Fuchashan Complex. The results are given in supplementary data Table S4 and shown in Fig. 9. Initial 176 Hf/177 Hf ratios and εHf (t) values were calculated at the corresponding 206 Pb/238 U ages obtained from the same zircon domains, which registers the timing of zircon growth from the primary magmas.

Twelve Lu–Hf spot analyses were performed on 11 zircon grains from sample FC02. The results show variable 176 Lu/177 Hf ratios of 0.0004–0.0023, but 176 Hf/177 Hf ratios in the range of 0.281856–0.282069 are similar. According to their corresponding U–Pb ages, the calculated εHf (t) values are in the range of −16.4 to −8.3 with a mean of −13.3 (Fig. 9). The results show that TDM1 ages range from 1699 ± 14 Ma to 2023 ± 10 Ma (2), defining two peaks of zircon Hf model ages at 1.80 and 1.91 Ga (Fig. 10a). Twelve Lu–Hf analyses were obtained for 9 zircon grains in diorite sample FC01. Their 176 Lu/177 Hf ratios vary from 0.0006 to 0.0014, and 176 Hf/177 Hf ratios are in the range of 0.281879–0.282020, thus showing an even more limited variation as compared to those in sample FC02 (Fig. 9). They have εHf (t) values of −15.7 to −9.5, corresponding to two-stage Hf model ages of 2289 ± 18 Ma to 2637 ± 12 Ma (2), and yielding three peaks of TDM2 at 2.42, 2.53 and 2.59 Ga (Fig. 10b). Twenty analyses for Lu-Hf isotope composition were performed on 20 zircon grains from diorite sample KJ01, which yielded variable 206 Pb/238 U ages of 647–868 Ma (Table S4). The results show similar 176 Lu/177 Hf ratios (0.0005–0.0020) and 176 Hf/177 Hf ratios (0.281857–0.282026) compared to the other diorite sample FC01. Their εHf (t) values are in the range of −18.4 to −9.3, and corresponding TDM2 are 2290 ± 14 Ma to

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Fig. 8. Chondrite-normalized REE patterns and primitive mantle-normalized spidergrams for the different rock types from the Fuchashan Complex. (a and b) Amphibolites, (c and d) diorites and (e, f) granitoids. The normalizing values of chondrite and primitive mantle are from Sun and McDonough (1989) and McDonough and Sun (1995), respectively. Trace element concentrations of MORB, continental arc basalts (CAB), oceanic arc basalts (CAB) and continental arc andesites (CAA) are from Kelemen et al. (2003). Trace element concentrations of the upper continental crust (UCC) and lower continental crust (LCC) are from Rudnick and Gao (2003).

2740 ± 11 Ma (2), defining two peaks of TDM2 at 2.47 and 2.57 Ga (Fig. 10c). Twelve analyses were performed for 12 zircon grains from granitoid sample FC11. They have variable 176 Lu/177 Hf ratios varying from 0.0008 to 0.0024, but similar 176 Hf/177 Hf ratios in the range of 0.281535–0.281685 (Table S4). The 176 Hf/177 Hf ratios are the

lowest of all four samples, yielding the lowest εHf (t) values are in range of −26.5 to −21.2 (Fig. 9). The calculated TDM2 are also the oldest, and vary from 3034 ± 11 Ma to 3359 ± 10 Ma (2), yielding two peaks of TDM2 at 3.23 and 3.33 Ga (Fig. 10d).

6. Discussion

Fig. 9. Zircon Hf-isotope evolution diagram for sample FC02, FC01, KJ01 and FC11. The shadow in blue and red represent the range of northern margin of the Yangtze Block and the Huangling terrane within the Yangtze Block (data from Zhang and Zheng (2013) and references therein).

Petrographic observations indicate minor alteration in all the samples, consistent with their low LOI values ranging from 0.14 to 0.83 wt%, far lower than the 1.33 wt% alteration criterion as defined by Polat and Hofmann (2003). In addition, three rock types (amphibolite, diorite and granitoid) exhibit fairly homogenous major and trace compositions (Table S3), which also suggest insignificant alteration effects. Polat et al. (2002) demonstrated that metavolcanic rocks having Ce/CeN * ratios between 0.9 and 1.1 display limited LREE mobility, whereas those with Ce/CeN * < 0.9 and Ce/CeN * > 1.1 are characterized by large LREE mobility. In this study, the Ce/CeN * ratios of all samples are fall between 0.96 and 1.07, thereby showing limited LREE mobility (Table S3). In general, the major elements Al, Ti, Fe, P, the HFSE (Th, Nb, Ta, Zr, Hf), REE (excepting Ce and Eu), transition metals (Cr, Ni, Sc, V) and Y are relatively immobile and insensitive to alteration, whereas other major elements, such as Na, K and Ca, and LILE (Cs, Rb, Ba, Sr) and Pb tend to be mobile during alteration and metamorphism (Arndt, 1994; Manikyamba et al., 2014; Polat and Hofmann, 2003; Said et al., 2010). In the following sections, principally the features of immobile elements are applied for rock classification and further petrogenetic discussions.


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Fig. 10. Histograms of zircon εHf (t) values for samples FC02 (a), FC01 (b), KJ01 (c) and FC11 (d). The numbers in red and black denote TDM1 and TDM2 , repectively. See Table S4 for detail.

6.1. Petrogenesis 6.1.1. Mafic rocks Six amphibolites with ages of ca. 800 Ma from different locations within the Fuchashan Complex belong to the sub-alkaline series with compositional features of typical intraplate basalts (Fig. 8). These amphibolites have a limited range of major and trace elements, and display uniform REE and trace element patterns (Fig. 8a and b). As Mg# decreases, Al2 O3 and Cr decrease (Table S3), whereas Fe2 O3 T and TiO2 register an increase. These chemical variations are consistent with crystal fractionation of dominantly clinopyroxene associated with a minor amount of plagioclase (Fig. S3). The lack of any prominent negative Eu anomalies (Eu/EuN * = 0.93–1.03) in these amphibolites also supports this interpretation, because the minor amount of plagioclase fractionation and resulting negative Eu anomalies would be balanced by clinopyroxene fractionation causing Eu positive anomalies. HREE are insignificantly fractionated from LREE in the whole-rock (Fig. 8a), particularly as reflected in the extreme enrichment of HREE in zircons from the amphibolite (Fig. S1a), and preclude the presence of garnet in residue, or its fractionation from the mafic magma, thus suggesting a depth of less than 40 km. With the exception of the sample FC10, all the studied amphibolites exhibit immobile trace element patterns similar to intraplate basalts, rather than to volcanic arc basalts and MORB (Fig. 11). Based on the highest Mg# and the lowest Zr content, the sample FC10 is considered to be the least evolved; thus, its composition may represent the closest analog to the primary magma. The Zr/Y ratios relative to the index of fractionation-Zr (Fig. 11c) is a potential method to distinguish between MORB, volcanic arc basalt and

intraplate basalt (Pearce and Norry, 1979). In this study, the Zr/Y ratios are positively correlated with Zr contents, falling within the intraplate basalt field (Fig. 11c), and indicating that the Zr/Y ratios increased as crystallization differentiation proceeded. The only major phases with a significant effect on the Zr/Y ratios are clinopyroxene, amphibole and garnet (Pearce and Norry, 1979). As mentioned above, garnet was probably absent in the magma source. The partition coefficients of elements between mineral and melt (expressed as mineral/melt Delement , abbr. mineral Delement ) reflect the degree of compatible and incompatible elements. Thus, if mineral Delement > 1, the element is considered compatible in the mineral, whereas mineral Delement < 1 indicates that the element is incompatible in the mineral. In the basic rocks, the Amphibole DY = 1 and the Amphibole DZr = 0.5 (Pearce and Norry, 1979), indicating that Y contents remain constant relative to the primary rock if the amphibole dominates the fractionation process. In this respect, Zr contents in the residual melt increase as fractionation proceeds, and the slope coefficient of Zr/Y compared to Zr would be close to 1/Y (e.g., 1/27, assuming Y content of 27 ppm in sample FC10 represents Y content of the primary melt). However, the trend in Fig. 11c argues against this probability. In contrast, the clinopyroxene dominated fractionation agrees well with the trend (the fractionation direction for clinopyroxene is determined by clinopyroxene DZr = 0.1 and clinopyroxene DY = 0.5, Pearce and Norry, 1979). Therefore clinopyroxene fractionation is inferred to have controlled the melt compositions during Neoproterozoic intraplate magmatism on the northern margin of YB. Widespread Neoproterozoic intraplate magmatism throughout the South China region (including the Yangtze and Cathaysia blocks) marks continental rift magmatism, and these events were suggested to be most

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Fig. 11. Geochemical discrimination diagrams of (a) Ti–Zr (Pearce, 1996), (b) Ti–Zr–Y (Pearce and Cann, 1973) and (c) Zr/Y–Y (Pearce and Norry, 1979) for the Fuchashan Neoproterozoic amphibolites. The slope coefficient of Zr/Y compared to Zr for amphibolite dominant fractionation is set as 1/27 based on Zr and Y partition coefficients of amphibole as well as Y content in sample FC10; the clinopyroxene fractionation direction depends on clinopyroxene DY and clinopyroxene DZr . WPB – within plate basalt; VAB – volcanic arc basalt; IAT – island arc tholeiite; CAB – calcalkaline basalt; M + I, MORB and island arc tholeiite; M + W-MORB and within plate basalt; M + V-MORB and volcanic arc basalt. Arrows indicate fractionation vectors for clinopyroxene and amphibolite. Symbols are as in Fig. 4.

likely related to the mantle superplume upwelling that led to the breakup of the supercontinent Rodinia (Li et al., 2003a,b, 2005, 2006, 2008a,b, 2009, 2010; Ling et al., 2003; Zheng and Wu, 2009). Zircon Hf isotopes from the amphibolite display variable negative values (εHf (t) = −8.3 to −16.4), apparently reflecting an enriched mantle rather than depleted mantle as the major source involved in the basaltic melt generation. In the Hf isotope evolutional diagram (Fig. 9), the enriched mantle is probably metasomatized subcontinental lithospheric mantle (SCLM) as a result of crust-mantle interaction. If so, their arc geochemical signatures such as depletion in Ti, Nb and Ta (Fig. 8a and b) are probably inherited by fluid/melt interaction with the overlying SCLM. However, their REE patterns show OIB-nature (Fig. 8a). The inherited feature is also present in granitoid rocks in the YB, e.g., Li et al. (2002), which is interpreted as the arc-like signature of the Guandaoshan granitoid pluton as inherited from lower crustal protoliths. Neoproterozoic rocks on the northern margin of the YB were mainly considered as protoliths of metamorphic rocks along the Qinling-Tongbai-Hong’an-Dabie-Sulu orogenic belt. Several studies were carried out on the zircon U–Pb ages and Lu–Hf isotopes of these rocks (Zheng et al., 2005; Zheng and Wu, 2009; Zhao et al., 2008; Chen et al., 2014; Liu et al., 2012; Tang et al., 2008; Wu et al., 2008; Xia et al., 2009). However, their zircon U–Pb ages primarily cluster at 740–780 Ma, and zircon εHf (t) values fall into two groups of 1.1–10.1 and −9.1 to −2.7 (Zheng and Wu, 2009 and references therein), which is different from the present results (Fig. 9). Interestingly, our results are analogous to those reported from the Huangling Complex (Fig. 9). The Huangling Complex lies in the Yangtze Gorge area near the Yichang city (located in the interior of the YB). Zircon U–Pb dating and Lu–Hf analyses of the granitic batholith and the mafic dykes occurring within the intrusion yielded consistent ages of 800–820 Ma and are characterized by pronounced negative εHf (t) values with Archean Hf model ages (Zhang et al., 2008, 2009), which are indistinguishable from our results (Figs. 7 and 12). These features suggest the presence of ancient Archean continental lithosphere in the northern Yangtze in the Neoproterozoic, which further support the hypothesis proposed by Zheng et al. (2006b) on the presence of widespread Archean basement beneath the YB. Zircon trace element data provide important insights on melt compositions and temperatures (Rubatto, 2002; Rubatto and Hermann, 2003, 2007; Rubatto et al., 2009; Watson et al., 2006). In this study, the steep MREE to HREE distribution patterns with high (Lu/Gd)N ratios of 13.6–19.1 from amphibolite are typical of magmatic zircon formed in the absence of garnet in the source (Fig. S1a, e.g., Hoskin and Black, 2000; Rubatto, 2002; Whitehouse and Platt, 2003; Hoskin and Schaltegger, 2003). Positive Ce anomalies are also very prominent (Ce/CeN * 1, Fig. S1a), which is ascribed to high oxidation state (indexed as high fO2 ) during zircon crystallization (Peck et al., 2001). The relatively small negative Eu anomalies (Eu/EuN * = 0.38–0.45) (Fig. S1a) in the zircons indicates that they were a late crystallizing phase that formed after minor plagioclase crystallization (Rubatto and Hermann, 2003). Ti contents in these zircon grains range from 26.1 to 37.2 ppm, yielding zircon crystallization temperatures of 833–871 ◦ C with a mean of 851 ◦ C, which are close to the zircon saturation temperature (Fig. S2a). The zirconium concentrations in basic rocks do not readily approach the saturation level (Watson and Harrison, 1983), and this may be accounted for metasomatism by fluid/melt derived from subducted continental crust. In summary, the Neoproterozoic intraplate magmatism on the northern margin of YB broadly involved the growth of juvenile continental crust and reworking of the Paleoproterozoic-Neoarchean ancient crust. These crustal rocks are characterized by OIB and arclike signature in terms of their trace element compositions. The asthenosphere upwelling is considered to have supplied the heat


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Fig. 12. Nb–Yb diagram (a) and Rb–Y + Nb diagram (b) of granitoids and diorites from the Fuchashan Complex (after Pearce et al., 1984). (c) R1–R2 diagram (after Batchelor and Bowden, 1985). The numbers in panel c are from Batchelor and Bowden (1985). (d) (Zr + Ce + Nb + Y) vs. 10,000 × Ga/Al diagram (after Whalen et al., 1987) showing an A-type granite affinity. See text for the abbreviations. Symbols are as in Fig. 4.

to induce melting of the enriched SCLM in an extensional environment. The melt compositions are controlled by crystal fractionation of dominant clinopyroxene associated with minor amounts of plagioclase under high fO2 . The absence of garnet in the residue and in the fractionation from the primary magma suggests that the depth of magmatism was less than 40 km. 6.1.2. Intermediate to felsic rocks The granitoid and diorite samples in this study show close correspondence with felsic and intermediate rocks in compositions. Rubidium, Y and Nb are commonly employed to discriminate the tectonic settings of granitic rocks (Pearce et al., 1984). Both the granitoids and diorites fall within the field of volcanic arc granite (VAG) + syn-collision granite (syn-COLG) in Fig. 12a. Furthermore, they plot within the VAG field in the Rb–(Y + Nb) diagram (Fig. 12b). In the R1–R2 discrimination diagram (Batchelor and Bowden, 1985; Fig. 12c), however, all granitoids fall in the syn-COLG field. The contrasting results from the two classification schemes suggest that their compositional features were influenced by source rock characteristics, formation and evolution of magma, and in turn the complex tectonic environments (Li et al., 2003a; Frost et al., 2001). In contrast, the basalts representing primary composition are more likely to reflect the tectonic setting (Li et al., 2008a, 2012). Therefore, we correlate the amphibolite to within plate extension setting. In addition, the coeval nature of the basaltic, intermediate to felsic rocks suggests a coherent geologic system. With the exception of the samples FC13 and FC14, all the studied granitoids and diorites fall within the A-type granite field following Whalen et al. (1987) (Fig. 12d). A-type granites are characterized by their alkalinity and anhydrous, compositionally high

SiO2 , Na2 O + K2 O, Fe/Mg, Ga/Al and low CaO and Sr (Whalen et al., 1987). They are commonly formed in anorogenic tectonic setting (Loiselle and Wones, 1979) or continental extension and/or rifting environment, where the crust tends to be thin and magmatic advection of heat can approach the shallow level of the crust (Whalen et al., 1987; Eby, 1990; Frost et al., 2001; Li et al., 2008a). In this study, all felsic and intermediate samples show features analogous to A-type granites, in particular their highly alkaline nature, resembling intraplate granite (Fig. 4b; Pearce et al., 1984). These results are also in accordance with those from the mafic rocks as mentioned above. The petrogenesis of A-type granites is controversial, and mechanisms involving melting of crustal and mantle materials, or fractional crystallization of mafic magmas plus assimilation of crustal rocks, are often suggested (Whalen et al., 1987; Eby, 1990; Frost et al., 2001; Li et al., 2008a). It is noted that all the felsic rocks are characterized by systematically lower zircon εHf (t) values between −26.5 and −21.2 than those in the intermediate rocks (−18.4 to −9.3) (Table S4), suggesting that granitoids were generated by remelting of ancient crust, and precluding any mantle sources. The remelting at the relatively shallow level of the crust requires the temperature to be >900 ◦ C (Li et al., 2008a). This requires transfer of heat from mantle-derived or hot mafic magmas, and most likely heat source was the coeval, extensive intraplate mafic magmas. The extremely negative zircon εHf (t) values in the granitoid sample FC11 suggest the presence of crustal basement older than 3.2 Ga, the most likely candidate being TTG gneisses, which have been identified from the Kongling terrane in the northwestern of the YB (Jiao et al., 2009; Gao et al., 2011). Most of the A-type granitoids over the world are metaluminous, and peralkaline and peraluminous compositions are rare (Frost

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et al., 2001). In this study, all the granitoids are weakly peraluminous with A/CNK ratios from 1.00 to 1.04 and A/NK ratios from 1.28 to 1.37 (Fig. 4c). Neoproterozoic anorogenic two-mica granites from Colorado in North America also belong to A-type peraluminous category (Anderson and Thomas, 1985). They were generated from an oxidized peraluminous quartzofeldspathic source at P ≥ 10 kbar (≥35 km) as constrained by their minimum melt compositions in the granite system (Anderson and Thomas, 1985). They display high abundances of LILE, LREE and other incompatible elements (Anderson and Thomas, 1985). These features are also apparent in our granitoid samples (Table S3; Fig. 8e and f). Furthermore, strong positive Ce anomalies are also evident (Ce/CeN * 1, Fig. S1d) in zircon grains, which are ascribed to high fO2 during zircon crystallization (Peck et al., 2001). This is in agreement with an oxidized source as suggested by Anderson and Thomas (1985). Therefore, we interpret that the A-type peraluminous granites in this region were generated in a similar setting at a depth ≥35 km. Compared to basaltic rocks, the granitic rocks have higher K2 O/Na2 O ratio, in particular the ratios of all granitoids are ≥1.21 (Table S3), suggesting potassium enrichment during late magmatic stage. Melting experiments demonstrate that late K-rich magmatism is related to remelting of the preexisting crustal rocks (Xiao and Clemens, 2007; Watkins et al., 2007), which is also supported by the zircon Hf isotope data (Fig. 9). Zircons from diorite samples FC01 and KJ01 share similar REE pattern with granitoid sample FC11, although they have distinct whole-rock major element compositions (Table S3; Fig. S1b, c and d). The large variation of the LREE in sample KJ01 indicates that some of the magmatic zircons experienced solid-state recrystallization (Hoskin and Black, 2000), consistent with their internal structure in CL images (broad bright overgrowth rim in Fig. 5c). This can also explain the unusually low Ti contents in zircon but high Zr contents in the whole-rock, corresponding to the Ti-in-zircon temperatures that are lower than zircon saturation temperatures (Fig. S2c), whereas the other three samples show the opposite trend (Fig. S2a, b and d). The largest negative Eu anomalies (Eu/EuN * = 0.02–0.07, Fig. S1c) in this sample indicate the highest degree of plagioclase fractionation during zircon growth. The lowest Mg# (34) and the highest Zr contents (353 ppm) in this sample also support this inference. In contrast, samples FC01 and FC11 show a consistent Ti-in-zircon temperature with a mean of 845 ◦ C, which is analogous to the result from the amphibolite (Table S2, Fig. S2a, b and d), revealing indistinguishable temperatures between the basic and felsic magmas. The temporal distributions of basaltic and felsic rocks in the Fuchashan Complex are displayed in Fig. S4, where a comparable age distribution is obvious from the age peaks. Their close temporal and spatial association indicates bimodal magmatism at ca. 800 Ma. The intermediate rocks (diorites) are likely the products of post-magmatism as discussed in following section, the time gap must be small due to similar formation ages between amphibolite and diorite. The intermediate rocks also could be the result of magma mixing (e.g., Li et al., 2003b). However, evidence from zircon Hf isotopes apparently argue against this model, with a substantial overlap of zircon εHf (t) values between the diorite and amphibolite, whereas the zircon εHf (t) values of the diorite are consistently higher than that of the granitoid sample FC11 (Table S4; Fig. 9). Alternatively, the diorites can be generated from the partial melting of basaltic rocks. The similarity of Hf isotopes and trace element features between amphibolites and diorites (Figs. 8 and 9) seem to support this possibility. In this case, the TDM2 of the diorites should be similar to the TDM1 of the amphibolites, but this is not evident in Fig. 10. Moreover, as melt phase, the incompatible elements (e.g., Ti, P and Nb) of the diorites should be systematically higher than those of amphibolites, which is also not observed in Table S3. However, individual high-Mg diorite (FC03) with high Mg#, Cr,

and Ni contents can result by the melting of mafic crustal rocks. The likely explanation for most diorites is that they were derived from an evolved mafic magma after fractionation of the amphibolite. Moreover, the degree of clinopyroxene fractionation increases during magma evolution. This is demonstrated by the variation of whole-rock Eu/EuN * with Zr contents and Mg# (Fig. S5a and b), and correspondingly, zircon Eu/EuN * variation would follow the same pattern (Fig. S5c and d). However, sample KJ01 does not conform to this trend (Fig. S5c and d), and the strong negative Eu anomalies (Eu/EuN * = 0.02–0.07) suggest that during zircon growth, the dominantly fractionating mineral is plagioclase instead of clinopyroxene, and this may represent the last stage of magma evolution, consistent with the lowest Mg# and the highest Zr content in this sample. 6.2. Tectonic significance 6.2.1. Geodynamic model Models on the timing and tectonics of assembly of the SCB through the amalgamation of the Yangtze and Cathaysia blocks are diverse, with some workers assigning this as during the Sibao orogenesis at 1.0–0.86 Ga, correlating with the global Grenvillian event (e.g., Li et al., 2002, 2008a,b, 2009, 2010, 2012). However, Precambrian stratigraphic sequences and magmatic records suggest that the Sibao or Jiangnan Orogen lasted until 830 Ma, or even up to 800 Ma, and is not connected with the main Grenvillian orogeny (e.g., Wang et al., 2007b; Yao et al., 2013; Zhao and Cawood, 2012; Zhao, 2014). The existence of ca. 860 Ma arcs in both west and east segments of the Jiangnan Orogen has also been identified in some recent works with the possibility that the amalgamation of the Yangtze and Cathaysia Blocks in the eastern part occurred later than ca. 871–864 Ma, or even at ca. 830 Ma (Yao et al., 2011, 2013, 2014). The voluminous pre-850 Ma igneous rocks are generally considered as products of the Sibao magmatic arc, whereas the younger 830–740 Ma igneous suite are believed to be the products of anorogenic magmatism in an intracontinental rift setting related to superplume activity during the breakup of Rodinia (Li et al., 2002, 2003a,b, 2008a,b; Cawood et al., 2013). Alternate models consider that these Mid-Neoproterozoic igneous rocks formed in an active continental margin of Rodinia related to an arc system (Zhao and Cawood, 1999; Zhou et al., 2002; Wang et al., 2004; Zhao et al., 2011). Although these younger 830–740 Ma magmatic suites are widespread throughout the SCB, their tectonic setting remains controversial and competing models have been proposed (Li et al., 2002, 2003a,b, 2008a,b; Cawood et al., 2013; Zhao and Cawood, 1999; Zhou et al., 2002; Wang et al., 2004; Zhao et al., 2011; Zhao, 2014). The major aspect of the controversy is the interpretation of the arc-like geochemical signatures such as depletions in Nb, Ta and Ti relative to LREE in the voluminous rocks along the western margin of the YB (e.g., Kangding gneissic complex and the Gongcai gneissic complex). These are considered as the critical evidence for the petrogenesis of Mid-Neoproterozoic igneous rocks related to arc magmatism. In this model, garnet is necessary as a residual phase in the source to account for the depleted Y and HREE, and presumes the source also had residual rutile and amphibole to account for the strong depletion of Nb, Ta and Ti (Zhou et al., 2006a,b). Although all samples exhibit arc-like signatures, the extreme enrichment in HREE and raised HREE patterns in their zircons (Fig. S1) does not favor garnet in the residue or fractionation from primary magma. Amphibole is also excluded as the Amphibole D = 1 and Amphibole D = 0.5 compared to the slope coefY Zr ficient of Zr/Y close to 1/Y in the basaltic rocks (Pearce and Norry, 1979), which is not apparent in Fig. 11c. Rutile controls more than 90% of the whole-rock Nb and Ta budget in eclogites, retaining the HFSE during arc magmatism in subduction zone, whereas other


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Fig. 13. Geodynamic cartoon showing continental crust growth and evolution under supercontinental rifting setting within the Yangtze Block during ca. 800 Ma. The arrow denotes direction of terrane motion. SCLM = subcontinental lithospheric mantle, ESCLM = enriched subcontinental lithospheric mantle, LCC = lower continental crust. Amphibolite forms by crystallization from the underplated melt of the ESCLM and diorite forms from the residual melt. Attendant heating causes partial melting of older lower crustal basement, from which granitoids are generated. See text for detailed description.

incompatible elements are partitioned into the melt or fluid causing further metasomatism in the overlying mantle wedge (Aulbach et al., 2008; Rudnick et al., 2000; Zack et al., 2002; Liu et al., 2014a,b). Most of the experimental studies have proved that rutile has variable degree of high Nb and Ta partition coefficients with consistent rutile DNb /rutile DTa < 1 (Foley et al., 2000; Klemme et al., 2005; Schmidt et al., 2004; Xiong et al., 2005, 2011), and thus a rutile-bearing residue would elevate Nb/Ta in the melts. Nevertheless, a number of samples with Nb/Ta < 13 (the continental crust Nb/Ta: 12–13, Barth et al., 2000) in our suite, as well as in the other areas within the YB (Zhou et al., 2002, 2006a,b; Li et al., 2003a, b, 2008a,b), suggests that rutile is absent in the residue at least for the generation of low Nb/Ta rocks. Moreover, the formation of rutile is strictly limited at P > 1.5 GPa (>50 km, e.g., Xiong et al., 2005, 2011), which is at variance with the shallower depth required by the absence of garnet in the magma genesis. On the other hand, a temperature of 800 ◦ C is not readily attained even at 40 km based on the general thermal gradients of arc systems in supra-subduction setting, which is also in conflict with the magma melting temperature >871 ◦ C or 900 ◦ C estimated in this study. Therefore, we suggest that the arc-like signatures in the amphibolites might have been inherited from fluid/melt derived from subducted continental crust. The possible continental crust growth scenario during ∼800 Ma in this region is shown in Fig. 13. In the Rodinia rifting setting, asthenospheric upwelling caused the partial melting of enriched SCLM at shallow depth (e.g., 35–40 km), and the protoliths of the amphibolites witnessed variable degree of crystal fractionation involving dominant clinopyroxene associated with minor amount of plagioclase. The remaining magmas after amphibolite extraction further evolved into diorites. Their consistently higher Mg#, Cr, Ni in whole rock compositions and zircon εHf (t) values than those in the granitoids suggest that they were derived from SCLM rather than ancient mafic crustal rocks. The underplated mafic magmas led to partial melting of the lower continental crust and produced the voluminous granitoids in this region. 6.3. Tectonic constraints The NE- to NNE-striking TLF along the East China, with a length of 2400 km, is the largest fault within the continental margin of East China. Several models have been proposed to explore its formation, onset and evolution (Li, 1994; Xu et al., 1987; Zhang et al., 2007; Zhu et al., 2002, 2009, 2010; Shi et al., 2009). The main controversy surrounds the timing of formation of the TLF, whether syn-collisional or postdating the formation of the Dabie-Sulu orogen. The common occurrence of P-bearing rocks (such as phosphorite) in both

the Feidong and Susong Complexes have been taken to suggest that they belong to the same metamorphic complex that was separated into two parts through displacement of ca. 300 km by the TLF sinistral slip, thus supporting the model that the TLF formed after the Dabie-Sulu orogeny (Xu et al., 1987). However, based on the regional geological analysis, alternate models suggest that the TLF formed during the collision between the NCC and SCB (Li, 1994; Zhu et al., 2009, 2010). The protolith ages of 782 ± 13 Ma for a metagranite from the Susong Complex (Xia et al., 2009) are indistinguishable from the ages for the Fuchashan granitoid (part of the Feidong Complex), suggesting that they belong to the SCB (Zheng, 2008). Nevertheless, the Susong Complex witnessed Triassic high-pressure metamorphism associated with the collision between the NCC and SCB (Xia et al., 2008, 2009, 2010), whereas the Feidong Complex did not experience this metamorphic event (Kang et al., 2013; Shi et al., 2009; Zhang et al., 2007). The Triassic high-pressure metamorphism significantly modified the Neoproterozoic igneous zircons through a variety of fluid and/or melt migration including supercritical fluid, generating a large population of metamorphic zircons including metamorphic overgrowth and recrystallization (Xia et al., 2008, 2009, 2010). The spongy texture of some of the zircons are consistent enrichment of LREE, HREE, Th, U and HFSE relative to Neoproterozoic magmatic zircons, and Xia et al. (2010) suggested that a supercritical fluid was involved that transported the LREE, HREE, Th, U and HFSE in the accessory minerals to recrystallized zircons. This mechanism of dissolution and recrystallization is responsible for the spongy texture and the very high concentration of trace elements in this type of metamorphic zircons. In this study, the CL images of granitoid zircons display very distinct features as compared to those of metagranite zircons from the Susong Complex (Fig. 5d vs. Fig. 4 in Xia et al., 2009 and Fig. 5 in Xia et al., 2010), although they have similar protolith ages. Furthermore, our granitoid zircons fall on or close to the concordia (Fig. 7d), showing little metamorphic effects, in contrast with the metagranite zircons that record Triassic lower-intercept ages (Xia et al., 2009, 2010). These differences indicate that the Feidong Complex did not witness the Triassic high-pressure metamorphism, whereas the Susong Complex did, providing additional evidence to support the formation of the TLF subsequent to the Triassic collision between the NCC and SCB. 7. Conclusions Zircon SHRIMP U–Pb dating of an amphibolite, two diorites and a granitoid from the Fuchashan Complex on the northern margin of YB shows that the protoliths of these rocks formed at

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794 ± 11 Ma, 796 ± 12 Ma, 812 ± 8 Ma and 803 ± 11 Ma, respectively. The age data define rift-related magmatism (∼800 Ma) along the northern margin of the SCB, possibly in response to MidNeoproterozoic breakup of the supercontinent Rodinia. Coeval and spatially related mafic and felsic rocks suggest that they formed by bimodal magmatic activity. The significant arc-like signatures in amphibolites probably reflect inheritance from enriched SCLM, which were metasomatized by fluid/melt derived from subducted continental crust. The diorites may represent products of the further evolution of the mafic magmas. The extremely low zircon εHf (t) value of −26.5 indicates the presence of crustal rocks older than 3.2 Ga in the basement of the YB. Absence of garnet and generation of A-type peraluminous granite constrain a depth range of 35–40 km for the Neoproterozoic bimodal magmatism, with the melt compositions controlled by crystal fractionation of dominant clinopyroxene associated with minor amount of plagioclase under high fO2 conditions. Comparisons between zircons from the Feidong and Susong Complexes suggest that the TLF formed after the Triassic collision between the NCC and SCB. Acknowledgements We are indebted to Prof. Chang, Y. F., Yao, Z. B., Shi, Y. H. and Dr. Chen Y. X. for their kind guidance during this research. Li, N., Hu, Z. L. and Chu, G. are acknowledged for their help with the zircon SHRIMP U–Pb analyses. Comments from Associate editor Dr. Wan Y. S. and two anonymous reviewers are gratefully acknowledged. This study was supported by the State Key Basic Research Development Program of China (2012CB416602), Natural Science Foundation of China (41090372, 41373053), the Fundamental Research Funds for the Central Universities and the China Postdoctoral Science Foundation (2014M561832). Appendix A. Supplementary data Supplementary data associated with this article can be found, in the online version, at http://dx.doi.org/10.1016/j.precamres.2015. 07.004 References Amelin, Y., Lee, D.C., Halliday, A.N., Pidgeon, R.T., 1999. Nature of the earth’s earliest crust from hafnium isotopes in single detrital zircons. Nature 399, 252–255. Amelin, Y., Lee, D.C., Halliday, A.N., 2000. Early-middle Archean crustal evolution deduced from Lu–Hf and U–Pb isotopic studies of single zircon grains. Geochim. Cosmochim. Acta 64, 4205–4225. Anderson, J.L., Thomas, W.M., 1985. Proterozoic anorogenic two-mica granites: Silver Plume and St. Vrain batholiths of Colorado. Geology 13, 177–180. Anhui Geological Survey, 1998. Geology map of Liangyuanzhen, Anhui Province: Anhui Geological Survey, scale 1:50000., pp. 1–122 (in Chinese). Arndt, N.T., 1994. Archean komatiites. In: Condie, K.C. (Ed.), Archean Crustal Evolution. Elsevier, Amsterdam, pp. 11–44. Aulbach, S., O’Reilly, S.Y., Griffin, W.L., Pearson, N.J., 2008. Subcontinental lithospheric mantle origin of high niobium/tantalum ratios in eclogites. Nat. Geosci. 1, 468–472. Barth, M.G., McDonough, W.F., Rudnick, R.L., 2000. Tracking the budget of Nb and Ta in the continental crust. Chem. Geol. 165, 197–213. Batchelor, R.A., Bowden, P., 1985. Petrogenetic interpretation of granitoid rock series using ulticationic parameters. Chem. Geol. 48, 43–55. Black, L.P., Kamo, S.L., Allen, C.M., Aleinikoff, J.N., Davis, D.W., Korsch, R.J., Foudoulis, C., 2003. TEMORA 1: a new zircon standard for Phanerozoic U–Pb geochronology. Chem. Geol. 200, 155–170. Blichert-Toft, J., Albarède, F., 1997. The Lu–Hf isotope geochemistry of chondrites and the evolution of the mantle–crust system. Earth Planet. Sci. Lett. 148, 243–258. Cawood, P.A., Wang, Y., Xu, Y., Zhao, G., 2013. Locating South China in Rodinia and Gondwana: a fragment of greater India lithosphere? Geology 41, 903–906. Chen, Y.-X., Zheng, Y.-F., Li, L., Chen, R.-X., 2014. Fluid-rock interaction and geochemical transport during protolith emplacement and continental collision: a tale from Qinglongshan ultrahigh-pressure metamorphic rocks in the Sulu orogen. Am. J. Sci. 314, 357–399. Compston, W., Williams, I.S., Kirschvink, J.L., Zichao, Z., Guogan, M.A., 1992. Zircon U–Pb ages for the early Cambrian time-scale. J. Geol. Soc. 149, 171–184.

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