In situ zircon U–Pb dating and O isotopes of the Neoarchean Hongtoushan VMS Cu–Zn deposit in the North China Craton: Implication for the ore genesis

Ore Geology Reviews 67 (2015) 354–367

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In situ zircon U–Pb dating and O isotopes of the Neoarchean Hongtoushan VMS Cu–Zn deposit in the North China Craton: Implication for the ore genesis Ming-Tian Zhu, Lian-Chang Zhang ⁎, Yan-Pei Dai, Chang-Le Wang Key Laboratory of Mineral Resources, Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing 100029, China

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Article history: Received 9 September 2014 Received in revised form 19 December 2014 Accepted 22 December 2014 Available online 24 December 2014 Keywords: Zircon U–Pb Oxygen isotopes Hydrothermal zircon Neoarchean Hongtoushan VMS deposit NCC

a b s t r a c t The Hongtoushan volcanogenic massive sulfide (VMS) deposit is the largest Archean Cu–Zn deposit in China, located in the Qingyuan greenstone belt on the northern margin of the North China Craton. The Cu–Zn mineralization was stratigraphically controlled by the interbeds (~100 m in thickness) of mafic–felsic volcanic sets and overlain by banded iron layers. However, the relationship between VMS deposits and associated volcanics has not been examined. This study ultimately clarifies the times and sources of the volcanics and mineralization. Based on in situ zircon U–Pb and O isotope on VMS-hosting mafic, felsic volcanic rocks, banded and massive sulfide ores and postmineralization pegmatite vein, we considered that there were two main formation stages for the Qingyuan Cu–Zn deposits; one was exhalative-hydrothermal sedimentation and another was further Cu– Zn enriched by later hydrothermal processes. The timing of the first stage occurred at 2571 ± 6 Ma based on the magmatic zircons in the VMS-hosting mafic volcanic rocks, from which the inherited zircons also indicate the existence of 2.65–3.12 Ga ancient supercrustal rocks in the Qingyuan district. A modern mantle-like δ18Ozircon value of 5.5 ± 0.1‰ (2SD) for this volcanism was well preserved in the inherited core domains of ore samples. It suggests that the mafic volcanics was most likely sourced from partial melting of juvenile crust, e.g., TTG granites. A large-scale metamorphic or hydrothermal event is documented by the recrystallized zircons in sulfide ores. The timing is tightly constrained by the hydrothermal zircon U–Pb ages. They are 2508 ± 4 Ma for the banded ore, 2507 ± 4 Ma for the massive ore and 2508 ± 2 Ma for the postmineralization pegmatite vein. These indistinguishable ages indicate that the 2507 Ma hydrothermal systems played a significant role in the upgrading of the VMS Cu–Zn orebodies. The weighted δ18O values of hydrothermal zircons show a successively increasing trend from 6.0 ± 0.1‰ (2σ) for the banded ore, 6.6 ± 0.2‰ (2σ) for the massive ore to 7.3 ± 0.2‰ (2σ) for the later pegmatite vein. This variation might be induced by gradual inputting of the δ18O-rich oceanic crust and/or oceanic sediment during the hydrothermal cycling system. Considering its modern mantle-like oxygen isotope composition of 2571 Ma volcanism, a submarine volcanic hydrothermal system involving mantle plumes is a preferred setting for the Neoarchean VMS Cu–Zn deposits in the Qingyuan greenstone belt. © 2014 Elsevier B.V. All rights reserved.

1. Introduction The Precambrian is an important metallogenic epoch in old cratons all over the world, especially in Canada, Australia and Russia. For example, the Abitibi greenstone belt comprises a well preserved submarine volcanic sequence that hosts a large number of volcanogenic massive sulfide (VMS) and important Au-rich VMS deposits, including the world-class Horne and RaRonde-Penna deposits (e.g., Claoue´-Long et al., 1990; Desrochers et al., 1993). The North China Craton (NCC) also has abundant Precambrian large-scale Cu–Zn, REE, and Fe ore deposits. Among them, Archean Cu–Zn deposits in China are primarily distributed in the Qingyuan greenstone belt, which hosts several important, economic VMS and Banded Iron Formation (BIF) deposits ⁎ Corresponding author.

http://dx.doi.org/10.1016/j.oregeorev.2014.12.019 0169-1368/© 2014 Elsevier B.V. All rights reserved.

that are spatially and temporally associated with mafic to felsic volcanic rocks (Zhai et al., 1985, 1990; Wang et al., 1987), such as Hongtoushan VMS Cu–Zn deposit and Xiaolaihe BIF deposit. The relationship between the VMS deposits, iron formation and volcanics in Qingyuan provides a key framework for interpreting ore genesis and studying the relationships between mineralization, crustal growth and tectonic background. Given the presence of granite–greenstone terranes, VMS and BIF deposits in Qingyuan, researches principally focused on dating of granites, volcanic geochemistry and petrology, microphysiography of ores (Zhai et al., 1984, 1985; Zhang et al., 1984; Gu et al., 2004a,b, 2007; Wan et al., 2005b,c; Zheng et al., 2008), but the timing for the volcanic rocks and VMS mineralization remained unknown. Precise age and isotope systematics will provide constraints on the temporal relationship between mineralization and magmatic/metamorphic events and hence potentially to the origin, genesis and setting of the ore-forming

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system. In situ zircon U–Pb is an accurate, effective tool, especially with regard to the application on hydrothermal zircons to date hydrothermal events and mineralization processes (e.g., Rubin et al., 1989; Hoskin, 2005). Exchange rates for oxygen are very slow in zircon (Watson and Cherniak, 1997), which may permit preservation of the protolith oxygen isotope composition through high-grade metamorphism (Valley et al., 1994). Hence, an understanding of the volcanic, mineralization times and magmatic hydrothermal sources in the Qingyuan district would offer a good insight into the sedimentary sequences and their relationship with subsequent metamorphism. In our contribution, we measured in situ zircon U–Pb and δ18O ratios for different rock types in the Hongtoushan VMS Cu–Zn deposit to determine the differences in the oxygen isotope ratios at specific times, to fingerprint magmatic, fluid history, and then to constraint ore genesis and tectonic setting. 2. Regional and local geology 2.1. Regional geology The Hongtoushan Cu–Zn deposit is situated in the Qingyuan greenstone belt of the NCC. The NCC is one of the oldest cratons in the world, consisting of Archean to Paleoproterozoic metamorphic basement overlain by Mesoproterozoic to Cenozoic unmetamorphosed cover. The metamorphic basement rocks are as old as 3.8–3.2 Ga (Liu et al., 1992; Song et al., 1996; Wan et al., 2005a); Neoarchean

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metamorphic rocks consist dominantly of 2.6–2.5 Ga tonalite– trondhjemite–granodiorite (TTG) gneisses, ~2.5 Ga syntectonic granites and a variety of supracrustal rocks that underwent greenschist to granulite facies regional metamorphism and polyphase deformation at about 2.5 Ga (Jahn and Zhang, 1984; Liu et al., 1985; Kröner et al., 1998; Grant et al., 2009; Wan et al., 2011). The NCC is divided into two major Archean to Paleoproterozoic blocks, the Eastern and Western Blocks, separated by a Paleoproterozoic orogen, the Trans-North China Orogen (Zhao et al., 2001; Zhao, 2014). The Qingyuan greenstone belt is situated at the northern margin of the Eastern Block (Fig. 1a). Traditionally, geologists divided the Qingyuan Archean basement into the Hunbei low-grade gneiss terrane and Hunnan high-grade granulite terrane separated by the Hunhe fault. Based on the latest researches, both terranes have similar features, from petrologic assemblages to geochronology (Wan et al., 2005a,c; Zhai, 2011). Thus, we collectively refer to them as the Qingyuan granite–greenstone belt, which is dominated by Neoarchaean TTG gneisses (60–70%) and greenstones (20–25%), with granulite–charnockite. The Archean greenstone belts (Fig. 1b) are further divided into the Shipengzi, Jinfengling, Hongtoushan and Nantianmen Formations (Liu, 1982; Zhai et al., 1984; Zhang et al., 1984). The Shipengzi Formation consists mainly of amphibolite, pyroxene amphibolite and fine-grained biotite gneiss; the Jinfengling Formation is comprised of amphibolites, locally with intercalations of plagioclase leptynite and biotite leptynite; the Hongtoushan Formation is composed primarily of interbedded meta-volcanics and meta-sedimentary rocks,

Fig. 1. Inset shows the distribution of basement in the North China Craton (a) and simplified geology and distribution of VMS mineral systems (b) in the Qingyuan granite–greenstone belts (after Zhang et al., 1984; Wan et al., 2005b; Gu et al., 2007; age's references are shown in the text).

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including amphibolite, biotite leptynite, biotite gneisses and quartz–feldspar gneisses; the Nantianmen Formation consists of biotite plagioclase leptynite and marble. Petrological and geochemical investigations indicate that the protolith of meta-volcanic rocks in the Jinfengling and Hongtoushan Formations is tholeiite to calc–alkalic volcanic rocks (Zhang et al., 1984; Zhai et al., 1985; Wan et al., 2005b). Previous geochronology studies using various methods indicate that the precursor rocks for the greenstones were older than 3.0 Ga and were intensely metamorphosed to upper amphibolite facies during the period of 2.9 to 2.8 Ga (Zhang et al., 1984; Wang et al., 1987; Shen et al., 1994). In recent years, Wan et al. (2005c) obtained SHRIMP zircon U– Pb ages of 2515 ± 6 Ma and 2510 ± 7 Ma for the Xiaolaihe and Tangtu hornblende leptynite, respectively, and thus, they considered that the Qingyuan supracrustal rocks mainly formed in the Late Neoarchean rather than Mesoarchean, similar to other greenstones in the NCC (e.g., Zhai and Santosh, 2011 and references therein). The emplacement period of TTG rocks is between 2560 and 2530 Ma and these rocks experienced a subsequent, extensive regional high-grade metamorphic event, recorded by the metamorphic overgrowth zircons, between 2517 and 2480 Ma (Fig. 2; Wan et al., 2005c; Grant et al., 2009). The 2505–2502 Ma K-granite bodies intruded the TTG and greenstone

rocks generated by partial melting of crustal rocks during this metamorphic event (Fig. 1; Li and Shen, 2000; Grant et al., 2009). The Qingyuan greenstone belt primarily contains three types of mineralization: VMS, BIF and ductile shear zone-hosted gold (Fig. 1b). Tens of Archean VMS deposits and smaller occurrences have been found within the greenstone belt with the Hongtoushan being the largest. The impermeable volcanic layers control the size and location of the Cu–Zn mineralization. All the VMS deposits are confined to a stratigraphic interval of about 100 m in thickness (Fig. 2). This interval, termed the ‘rhythmical member’ by local geologists, belongs to the upper portion of the Hongtoushan Formation and is composed of rhythmic ally interbedded of mica gneiss, quartz–feldspar gneiss, hornblende gneiss and amphibolite gneiss. The Archean submarine exhalative origins of the VMS deposits have a close genetic relationship with the volcanic rocks and are commonly overlain by banded quartz–iron layers with various thicknesses (Fig. 2). BIF deposits and occurrences are widely distributed in the Qingyuan greenstone belt from the bottom Jinfengling Formation to the top Nantianmen Formation. Few publications have reported on the BIF deposits, although several economic deposits occur in the Qingyuan area, such as the Xiaolaihe and Xiadianzigou deposits, the former occurs in the meta-volcanic rocks at

Fig. 2. Lithostratigraphic column of metamorphosed host rocks of the massive sulfide deposits (after Zhang et al., 1984; Gu et al., 2007).

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ca. 2515 Ma (Wan et al., 2005c) and the latter occurs in the Nantianmen biotite plagioclase leptynite. A dozen of Archean and Mesozoic ductile shear zone Au deposits occur in the Qingyuan greenstone belt (Fig. 1). They extend parallel to the strike of the greenstones in a NW direction. It is widely believed, no matter when the Au deposit formed, either in the Archean or the Mesozoic (Yanshanian event), that the gold is primarily sourced from the greenstones related with low salinity, CO2enriched metamorphic fluids (Dai and Liu, 1989; Zhang, 2013). These features are consistent with the orogenic-type gold deposit (Chen et al., 1998; Groves et al., 1998). The Nanlongwangmiao Au deposit is known to have formed in the Archean with a muscovite Rb–Sr isochron age of 2.41 Ga obtained by Fushun Geological Research Institute (quoted from Dai and Liu, 1989), whereas the Xiajiapu Au deposit occupies in the contact between Mesozoic granite and Archean supracrustal rocks and the granite emplaced at 234 Ma dated by zircon U–Pb method (Zhang, 2013). 2.2. Local geology The Hongtoushan deposit is the largest Archean VMS Cu–Zn deposit in China, with reserves of 0.5 Mt Cu with grades of 1.5–1.8%, 0.7 Mt Zn with grades of 2.0–2.5% and 20 t Au with grades of 0.5–0.8 g/t (Gu et al., 2007). The host rocks were metamorphosed to upper amphibolite facies, with temperatures of 600–650 °C and pressure of 0.8–1.6 GPa (Zhao and Cui, 1987). These rocks are interbedded amphibolite and biotite plagioclase gneisses, with minor of sillimanite biotite gneiss (Fig. 3a) and cordierite orthoamphibolite gneiss. Cordierite orthoamphibolite gneiss (Fig. 1 in Zheng et al., 2008) is widely distributed below banded ore bodies at 600–800 m, interpreted as metamorphosed alteration zone of a seafloor hydrothermal system, with enrichment in Fe and Mg and depletion in K (Zhang et al., 1984; Zheng et al., 2008). Petrological and geochemical investigations indicate that the protolith of gneisses was derived from mafic–felsic volcanic rocks intercalated with minor submarine sedimentary rocks (Zhai et al., 1985; Zheng et al., 2008). The gneissic foliation therein is roughly parallel to lithological boundaries and dips toward the southeast, defining a vertically plunging fold. The orebodies, although intensely deformed, are stratiform and are stratigraphically controlled by the ~100 m ‘rhythmic member’ in the upper part of the Hongtoushan Formation (Fig. 3a and b). This thin layer is composed mainly of rhythmically interbedded biotite plagioclase gneisses and amphibolite gneisses. An individual layer ranges from several cm to 5 m in thickness and is gradually thickened along with the depth (Fig. 4). Massive sulfide orebodies occur in the upper part of the biotite plagioclase gneiss that is characterized by cordierite, sillimanite and anthophyllite. The ore bodies at all levels above −467 m exhibit a Y-shaped form, with the two branches opening toward east. These two branches merge into a single entity toward west, attaining the maximum thickness and the highest grade in Cu, Au and Ag, and thus form a substantial, vertical ore pillar. Below the depth of 467 m, the orebodies occur mainly in the limbs of the fold. Sulfide-bearing veins are ubiquitous throughout the deposit running through the wall rocks and orebodies in variable directions with sharp or relatively straight boundaries. Most of the veins are not folded and the minerals are not deformed (Gu et al., 2007). No granite intrusions have been found in the mine. Pegmatoidal quartz– amazonite veins are sporadically distributed in the contact zone between the ore body and gneiss, and locally crosscuts the gneiss (Zhang et al., 1984). Numerous diabase dykes penetrated into the massive sulfide orebodies and the metamorphic rocks (Figs. 3 and 4) and most extend consistently in approximately N–S directions, with local branching. Xenoliths of the ores and metamorphic rocks can be found within the dykes. These dykes are tholeiitic in composition (Gu et al., 2007). Most of the orebodies are mainly stratiform or stratiform-like in shape and are massive. Banded and disseminated orebodies occur locally and contain synchronously folded sulfide and gneiss layers. Yet the sequences of various orebodies are difficult to define because of complex deformation. Sulfide minerals amount to 60 to 80 vol.% of the

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massive sulfide ore, consisting dominantly of pyrite, pyrrhotite, sphalerite and chalcopyrite in approximate proportions of 5:5:1:1. Less abundant are cubanite, electrum, chalcocite and magnetite; galena is only locally observed. A notable feature in various ores is the welldeveloped pyrite porphyroblast, the fractures of which are filled with chalcopyrite, pyrrhotite and sphalerite. Based on detailed microscope investigations, Gu et al. (2004a) considered that the ores underwent extensive mechanical remobilization and recrystallization in the course of metamorphism and a considerable portion of the cubic pyrite phenocrysts are porphyroblasts formed during retrograde metamorphism. Gangue minerals are dominated by quartz, with subordinate plagioclase, garnet, hornblende, actinolite, biotite, muscovite and gahnite. Zn/Cu ratios generally increase with depth, manifesting an enrichment of Cu and Zn in the shallow and the deep levels, respectively. δ34S of sulfide minerals varies from −0.7 to 3.1‰, implying a sulfur composition of magmatic origin. δ30Si and δ18O of quartz vary from − 0.8 to 0.4‰ and 8.5 to 9.5‰, respectively, similar to the isotope characteristics of submarine volcanics (Hou et al., 2006). 3. Sample selection and analytical techniques 3.1. Sample selection and description All the samples were collected at the underground mining level of −827 m (Fig. 4), including amphibolite gneiss and biotite plagioclase gneiss, pegmatite and massive, banded ores. The samples of amphibolite gneiss (HTS1-1) and biotite plagioclase gneiss (HTS2-1) are interbedded. Amphibolite gneiss consists mainly of amphibole and plagioclase, with geochemical affinity of tholeiite series (Zhai et al., 1985); biotite plagioclase gneiss is composed of plagioclase, quartz and biotite, and is considered sourced from felsic volcanic rocks (Zhai et al., 1985; Zheng et al., 2008). Pegmatite sample (HTS3-1) was taken from the pegmatoidal quartz– amazonite vein, consisting primarily of quartz, amazonite and plagioclase, with minor of biotite, muscovite and sericite alterated from plagioclase. The light green amazonite (Fig. 5a), a variant of microcline containing ~1.4% Rb2O and ~0.2% Cs2O, forms graphic intergrowth with quartz (Fig. 5b). The sizes of amazonite and quartz range from 1–10 cm to 0.5–5 cm, respectively. Chalcopyrite is the only sulfide disseminated within these grains. This green pegmatoidal vein has been dated by the amazonite Rb–Sr system at 2360 Ma (quoted from Yu, 2006) and was interpreted as generation of late hydrothermal processes. Sulfide minerals amount to 60 vol.% of the massive sulfides ore sample (HTS3-4), consisting of pyrite, pyrrhotite, sphalerite, and chalcopyrite (Fig. 5c). Sulfide grains are commonly 0.5–1 cm in size except pyrite porphyroblasts, which are much larger, normally between 2 and 5 cm, formed during retrograde metamorphism (Gu et al., 2007). Some pyrite grains are granular, most likely representing a primary sulfide, whereas the sphalerite fills in the spaces of pyrite grains. Gangue minerals consist dominated of quartz, plagioclase, hornblende, muscovite, biotite, garnet, sillimanite and gahnite to form globular textures. Banded ore sample (HTS3-9) in metamorphic rocks mainly occurs in the gradational boundary of massive sulfide and gneiss. Pyrite, pyrrhotite, chalcopyrite and sphalerite are uniformly distributed along the gneissic schistosity (Fig. 5d). These zones, tens of meters thick and more than 100 m in length, have Cu and Zn contents below the economic grade, but locally contain exploitable ore pods. 3.2. Analytical techniques The samples for U–Pb and O isotope analyses were broken into fistsized pieces. The zircon grains were separated via a combination of heavy liquid and magnetic techniques, handpicked and subsequently mounted in epoxy resin and polished to remove the upper one third of the grains. The mounts were vacuum-coated with high-purity gold prior to geochronology analyses. Cathodoluminescence (CL) was

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Fig. 3. Geology at level −167 m (a) and section map of No. 14 prospecting line (b) of the Hongtoushan deposit (after Deng, 1994).

obtained based on LEO1450VP SEM to identify internal structures and choose the potential target sites for U–Pb and O analyses. Zircon U–Pb analyses for the amphibolite gneiss sample (HTS1-1) were performed using a laser ablation-inductively-coupled plasma mass spectrometer (LA-ICP-MS) at the state Key Laboratory of Continental Dynamics, Northwest University, Xi'an, China. A pulsed (Geolas) 193 nm ArF Excimer laser energy of 50 mJ/pulse was used for ablation at a repetition rate of 10 Hz. The diameter of the laser spot was 32 μm. The detailed analytical procedures used follow those described by Yuan et al. (2007). Zircon 91500 was used as an external standard for U–Pb dating and was analyzed twice every five unknown analyses. The NIST SRM610 glass was used as an external standard to calculate U, Th and Pb concentrations of unknowns. Uncertainty of preferred values for the external standard 91500 was propagated to the ultimate results of the samples. Concordia diagrams and weighted mean calculations were made using Isoplot/Ex_ver3 (Ludwig, 2003). Zircon U–Pb analyses for the biotite plagioclase gneiss (HTS2-1), pegmatite (HTS3-1) and ore samples (HTS3-4 and HTS3-9) were conducted using a Cameca IMS-1280 ion microprobe at the Institute of Geology and Geophysics, CAS (IGGCAS). The detailed analytical procedures can be found in Li et al. (2010a). The analyzed ellipsoidal spot size is approximately 20 × 30 μm in size. During the analysis, the standard zircon 91500 was used as an external standard for U–Pb dating. Uncertainty of 1% for individual analyses (ratio and ages) was propagated to the unknowns, whereas calculated weighted mean ages are within 95% confidence limits. The data were processed using the Isoplot/Ex_ver3 (Ludwig, 2003). In situ zircon O isotope analyses were conducted on zircons that were previously dated, using the Cameca IMS-1280 ion microprobe at the IGGCAS. After U–Pb dating, the sample mount was repolished to ensure that any oxygen implanted in the zircon surface used for U–Pb analysis was removed. One analysis consists of 20 cycles, with an internal precision generally better than 0.4‰ on the 18O/16O ratio. The detailed analytical procedures are similar to those reported by Li et al. (2010b). Measured 18O/16O values were normalized to the Vienna Standard Mean Ocean Water composition (VSMOW, 18O/16O = 0.0020052). The instrumental mass fractionation factor (IMF) was corrected using the Penglai zircon standard with δ18O value of 5.3‰ (Li et al., 2010c). A second zircon standard (Qinghu) was analyzed as an unknown to ascertain the veracity of the IMF. Nine measurements of Qinghu zircon during the course of this study yield a weighted mean δ18O = 5.5 ± 0.3‰ (2SD; standard deviation), which is consistent with the reported value of 5.4‰ ± 0.2‰ (2SD) (Li et al., 2013).

4. Analysis results 4.1. Results of in situ zircon U–Pb dating 4.1.1. Amphibolite gneiss Zircons extracted from sample amphibolite HTS1-1 range in length from 100 to 300 μm, with length to width ratios of 1:1 to 3:1. Some zircons have homogeneous luminescence. Considering to their U–Pb systems, we consider that they are of detrital origin. The majority of grains have typical oscillatory zoning with variable CL intensity, indicating their magmatic affinity (Fig. 6a). A small domain of blocky sectorzones is locally observed in some zircons. A total of 21 analyses on 21 grains (Table 1) were determined from the amphibolite gneiss. They cluster into three populations on concordia diagram with different levels of Pb loss. (1) The oldest 4 zircon grains have relative concentrated Th/U ratios of 0.30–0.72; (2) the Th/U ratios of 6 zircon grains range from 0.09 to 0.57; and (3) the remaining 11 grains gave relatively concentrated apparent 207Pb/206Pb ages and have variable concentration of U and Th, ranging from 11 to 1183 ppm and 17 to 351 ppm, respectively; the Th/U ratios range from 0.17 to 1.87, with an average ratio of 0.79. All analyses show variable discordance (Fig. 7a) and a range of apparent 207Pb/206Pb ages from 2565 to 3125 Ma. Four oldest analyses define a weighted mean 207Pb/206Pb age of 3115 ± 12 Ma (MSWD = 1.4; not shown). Six apparent 207Pb/206Pb ages of the middle population are between 2623 and 2703 Ma, with a weighted mean 207Pb/206Pb age of 2646 ± 34 Ma (MSWD = 2.6; not shown). These zircons contained in amphibolite gneiss reflect the existence of ancient supercrustal rocks of 2.65–3.12 Ga in this area. The remaining 11 analyses in the discordia line gave an upper intercept age of 2571 ± 6 Ma (MSWD = 0.1). This age might represent the formation age of mafic volcanics.

4.1.2. Biotite plagioclase gneiss Zircons separated from sample HTS2-1 are relatively small but of good enough size for the Cameca IMS 1280 analysis. They are 40–80 μm in length with aspect ratios of approximately 1:1.5. These grains are relatively rounded and are purple–brown in color. The majority of grains have typical core–rim texture with different luminescence. The rim domains have relatively blurry, dark luminescence, suggesting metasomatism, whereas most core domains show light luminescence. But the sizes of core domains are only several micrometers, which are too narrow for further studies using the Cameca IMS-1280 (Fig. 6b).

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Fig. 4. Geology at level −827 m of the Hongtoushan deposit. Sampling locations as yellow star showing; Symbols as in Fig. 3.

Therefore, we consider that these rim domains might be due to recrystallization. Because of small size of the zircon grains, 15 U–Pb analyses were only conducted on the homogenous grains (Table 1). The U and Th contents range from 126 to 558 ppm (mean = 206 ppm) and from 40 to 158 ppm (mean = 70 ppm), respectively. Therefore, the Th/U ratios concentrate on 0.21–0.48 (Fig. 8a), with an average ratio of 0.35. All analyses show concordant apparent 207Pb/206Pb ages from 2495 to 2523 Ma and form a perfect-defined cluster (Fig. 7b), almost without Pb loss. These grains define a weighted mean 207Pb/206Pb age of 2507 ± 5 Ma (MSWD = 2.0), which is interpreted as representing the zircon recrystallized time. 4.1.3. Pegmatite Zircons from the pegmatite sample (HTS3-1) range in length from 100 to 200 μm, with length to width ratios of 1:1–2:1. The majority of grains are prismatic with blunt pyramidal terminations and all grains are purple–brown in color. CL images show that some grains contain luminescent ‘cores’ surrounded by weakly luminescent, low contract ‘rims’ (Fig. 6c). However, when comparing with their undistinguished Th/U ratios, 207Pb/206Pb ages and δ18Ozircon values, we consider that these two domains are the same in origin and could not represent core–rim texture. Eleven spots on 11 grains (Table 1) were analyzed. These grains contain slightly high concentrated U, ranging from 622 to 2287 ppm, and low Th contents, ranging from 53 to 145 ppm. The Th/U ratios are significantly lower and range from 0.04 to 0.11 (mean = 0.09; Fig. 8a). All analyses show concentrated apparent 207Pb/206Pb ages from 2503 to 2511 Ma and form a perfect-defined cluster (Fig. 7c), almost without any Pb loss. These grains define a weighted mean 207Pb/206Pb age of 2508 ± 2 Ma (MSWD = 0.8), which is interpreted as representing the formation age of the pegmatite vein. 4.1.4. Massive sulfide ore Zircons extracted from massive sulfide ore sample (HTS3-4) range in length from 100 to 150 μm, with length to width ratios of 1:1–2:1. The majority of grains are semi-prismatic with a blunt pyramidal termination and all grains are purple–brown in color. CL images show that some grains contain luminescent cores surrounded by weakly luminescent, low contrast rims and some grains have uniform luminescent domains (Fig. 6d). Most of cores are surrounded by less luminescent rim, and resulting in core domains that were too narrow for analyses using the Cameca IMS-1280.

A total of 13 analyses on 13 grains (Table 1) were determined. Among them, only 2 cores were analyzed because of their small width. The concentration of U in all 11 rims ranges from 123 to 3194 ppm, consistently higher than in 2 cores (120 and 189 ppm). Thorium contents in the rims are extremely low, ranging from 2 to 51 ppm (mean = 11 ppm), whereas the content in 2 cores ranges from 71 to 105 ppm. The Th/U ratios of 2 core analyses have values typical of magmatic zircons, between 0.55 and 0.59, in contrast to the 11 rim analyses, which are all significantly lower and range from 0.01 to 0.04 (mean = 0.03) due to a marked decrease in the Th content (Fig. 8a). Although the zircon U–Pb results for core domains are just two analyses, these analyses spread along the concordia between 2560 and 2573 Ma (Fig. 7d), which should represent the crystallization age of the core, within error of the zircon age from amphibolite gneiss sample (HTS1-1). All analyses for the rim domains plot close to concordia and define a weighted mean 207Pb/206Pb age of 2507 ± 4 Ma (MSWD = 7.8). It is interpreted as a time of metamorphism relative with the formation of the massive sulfide ores. 4.1.5. Banded ore Zircons extracted from sample HTS3-9 range in length from 100 to 200 μm, with length to width ratios between 2:1 and 3:1. The majority of grains are prismatic with blunt pyramidal terminations and all grains are purple–brown in color. CL images of zircons show strongly luminescent cores surrounded by weakly luminescent, low contract rims (Fig. 6d). The cores have faint oscillatory zoning and are locally transgressed by less luminescent rim in different levels with discordant boundaries. The size of remaining cores is variable due to the levels of transgression by the rims, indicating either hydrothermal or metamorphism. Within the dark rims of zircons, nonuniform structures are discernible as either irregular-shaped patches or sector zones of low but variable CL intensity. The dark CL response and low contrast of the rims inhibits the recognition of fine details, but euhedral banding can be observed in some rims, and it is parallel to both the outer grain margins and the core–rim boundary. A total of 29 analyses on 26 grains (Table 1) were determined from the banded ore. Of the 29 analyses collected, 13 analyses were located in core domains and 16 analyses in rim domains. The U content of the cores ranges from 57 to 137 ppm (mean = 95 ppm), whereas the content in the rims ranges from 162 to 705 ppm (mean = 297 ppm). Thorium concentrations in the cores ranging from 27 to 94 ppm are similar to the contents in the rims from 21 to 69 ppm. The average Th content of core analyses is 55 ppm; nearly equal the average value of the rims (46 ppm), but the U shows a general increase from core to rim. The

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Fig. 5. Hand specimen photographs of the Hongtoushan deposit. Pegmatite vein consists mainly of white quartz (Q) and green amazonite (Amz), with sizes of 2–5 cm and 2–10 cm, respectively (a). Under cross-polarized light, amazonite shows cross hatched twin and form graphic texture with quartz; plagioclase (Pl) was altered to sericite (Srt) (b). Massive sulfide ore is composed of pyrite (Py), pyrrhotite, sphalerite (Sph), chalcopyrite (Ccp), quartz etc. (c). Veined sulfides distribute in the gneiss schistosity (d).

Th/U ratios of the core analyses have values typical of magmatic zircons, ranging from 0.31 to 0.69 (mean = 0.57), in contrast to the rim analyses, which are all slightly lower and range from 0.04 to 0.34 (mean = 0.20) due to a marked increase in the U content (Fig. 8a). When integrating the isotope data with the CL patterns, it is evident that two separate events are recorded in zircons from the banded ore (Fig. 7e). All analyses plot close to concordia with core and rim analyses forming two clusters about 60 Ma apart. An older group of 13 core analyses defines a weighted mean 207Pb/206Pb age of 2568 ± 5 Ma (MSWD = 0.7), interpreted as the crystallization age of zircon cores, within error of the age of 11 zircons from amphibolite gneiss sample (HTS1-1). The rim analyses define a weighted mean 207Pb/206Pb age of 2508 ± 4 Ma (MSWD1.9); it is interpreted as a metamorphic or hydrothermal event recorded by the banded ores, and within errors of the ages obtained from the felsic volcanic rock (HTS2-1), pegmatite (HTS3-1) and massive sulfide ore samples (HTS3-4). In summary, the zircon U–Pb results for the banded ore show two distinct ages recorded from the core (~2568 Ma) and rim (~2508 Ma) domains, respectively.

4.2. Results of in situ zircon oxygen isotope Sixty five in situ O isotope analyses were conducted on the same zircon zones of the U–Pb analyses (Table 1). The δ18O values of magmatic zircons from the biotite plagioclase gneiss (HTS2-1) range from 5.3 to 5.8 (mean = 5.5 ± 0.2‰; 2SD), a range that is very similar to modern mantle-like δ18O value (δ18O = 5.3 ± 0.6‰; Valley et al., 1998). The zircons from the pegmatite (HTS3-1) have a homogenous, δ18O-rich isotope composition, with δ18Ozircon values ranging from 6.7‰ to 7.7‰ (mean = 7.3 ± 0.3‰; 2SD). The δ18Ozircon values of the massive sulfide ore (HTS3-4) and the banded ore (HTS3-9) occur as two groups according to their core–rim structure. The δ18Ozircon values of cores have a typical characteristic of modern mantle-like, ranging from 5.3‰ to 5.8‰ (mean = 5.5‰ ± 0.1‰; 2SD), which is distinct with the values of rims, ranging from 5.6‰ to 7.0‰ (mean = 6.2 ± 0.4‰; 2SD). The

preservation of oxygen isotope homogeneity in inherited cores proves that diffusion of oxygen was extremely slow in these zircon crystals. In summary, the δ18Ozircon values occur as two groups, representing two different sources (Fig. 8b). One is the mantle origin, with δ18Ozircon of 5.3‰–5.8‰ and forming a Gaussian distribution with a mean δ18Ozircon of 5.5 ± 0.2‰ (2SD; not shown), and another metamorphic or hydrothermal origin has variable δ18Ozircon and is relatively oxygen-enriched, with δ18Ozircon of 5.6‰–7.7‰.

5. Discussion 5.1. Result interpretations Generally, zircon can be divided into two types: growth of new zircon and modification of existing zircon. The new zircon can grow from a magmatic melt and is regarded as magmatic zircon or from an aqueous fluid as hydrothermal zircon (e.g. Ahrens, 1965; Rubin et al., 1989; Vavra, 1990), while the modification of existing zircons is regarded as metamorphic or hydrothermal zircon (e.g., Hoskin, 2005). Therefore, the hydrothermal zircon could refer to the new zircon growing in hydrothermal environments and overgrowing on pre-existing zircon. The zircons from Hongtoushan show complex features and have magmatic and hydrothermal signatures. Oxygen isotope systematics have been extensively employed to trace magma source and constrain the mixing history with crustal material. Magmas with no supracrustal input generally have a uniform oxygen isotope ratio that is distinct from magmas that assimilated or were generated directly from supracrustal sources. The oxygen isotope value of the normal mantle, regardless of the lithosphere mantle (4.90–5.5‰; Mattey et al., 1994; Zhang et al., 2001) or asthenosphere (5.2–5.8‰ of Mid-Ocean-Ridge Basalt; Pearson et al., 1991; Bindeman et al., 2012) or mantle-derived zircon (4.8–5.6‰; Page et al., 2007), is about 5.3 ± 0.6‰ (2SD; Valley et al., 1998), whereas the values of Archean sedimentary rocks vary from 9‰ to 12‰ (Shieh and Schwarcz, 1978). Thus, even

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361

Fig. 6. Representative zircon cathodoluminescence images of various samples from the Hongtoushan deposit. The numbers show their 207Pb/206Pb age (Ma) and δ18O value (‰). In situ zircon O isotope analyses were conducted on zircons that were previously dated.

small inputs of subducted sediments or continental crust into a magma with mantle-like oxygen will generally raise the δ18O of the magma. 5.1.1. The host rock From zircon petrographic features, the zircons in the VMS-hosting mafic volcanics could be divided in to a inherited older population and a younger concentrated population, the latter showing magmatic affinity, which is compatible with the U/Th ratios and oxygen isotope composition (mean = 5.5 ± 0.2‰; 2SD). Therefore, we constrain the formation age of the mafic volcanics at 2571 Ma, which also represent the formation age of the VMS Cu–Zn mineralization and related banded iron layers. This late Archean period of volcanic activity and polymetallic mineralization as well as Algoma-type BIF is well developed in the NCC, their formation reaching a peak during the late Archean (e.g., Zhai and Windley, 1990; Zhang et al., 2012). In the Anshan– Benxi area, the BIF-hosting rocks were primarily formed in the late Archean, such as the 2551 ± 10 Ma chlorite–quartz schist in the Chentaigou BIF (Dai et al., 2013), the 2541–2553 Ma plagioclase amphibolite in the Shirengou BIF (Zhang et al., 2012) and the 2554 ± 12 Ma plagioclase–hornblende schist at the Gongchangling BIF (Han et al., 2014). On the contrary, the zircons from the felsic volcanics have distinct core–rim texture (Fig. 6d). The core domains show up remaining serrated shape, while the rim domains irregularly surround the cores. Therefore, the rims should be a modification origin based on the inherited grains. Additionally, they have a limited range of Th/U ratios between 0.21 and 0.48 and a normal mantle range of δ18Ozircon value between 5.2‰ and 5.8‰ (Fig. 8). This is clear evidence that the modified fluid was magmatic in origin. Hence, when combined with their identical weighted mean 207Pb/206Pb ages, their 2507 Ma age represents metamorphic time of the felsic volcanic rocks.

When comparing the identical ages of the amphibolite gneiss (2571 ± 6 Ma for HTS1-1) and the core domains of the ore samples (2560 and 2573 Ma for HTS3-4 and 2568 ± 5 Ma for HTS3-9), a reasonable interpretation is that the zircon cores of ore samples were inherited from the grains of the amphibolite gneiss. Therefore, the oxygen isotope composition of core domains (a modern mantle-like δ18Ozircon value of 5.5 ± 0.1‰; 2SD) should present δ18Ozircon values of the mafic. Both the presence of 2.6–2.5 Ga TTG intrusions in Qingyuan (Wan et al., 2005b; Grant et al., 2009) and their mantle-like δ18Ozircon of 5.5 ± 0.4‰ (e.g., King et al., 1998) for TTG lithologies from the Superior Province support that the TTG granitoids should be a preferred alternative for the source of the core domains. It shows that the 2571 Ma volcanic rocks were most likely derived from partial melting of juvenile crust (TTG sets). 5.1.2. The Cu–Zn ore and pegmatite Zircons selected from the hydrothermal alteration and/or mineralization veins have been broadly reported and interpreted to be either hydrothermal in origin (Rubin et al., 1989; Kerrich and King, 1993; Hu et al., 2004; Hoskin, 2005; Pettke et al., 2005; Lawrie et al., 2007; Pelleter et al., 2007; Fu et al., 2009; Wan et al., 2012) or magmatic relicts (e.g., Fu et al., 2009; Zhu et al., 2013). The core–rim texture observed in most of the zircons from the Cu–Zn ore samples (HTS3-4 and HTS3-9) shows different luminescence (Fig. 6d) and has continuous inner textures. Their disparity includes: (1) increase in U concentration and decrease in Th/U ratio in the rims compared to the cores; (2) disparate 207 Pb/206Pb ages; and (3) distinct higher δ18Ozircon value in the rims (mean = 6.2 ± 0.4‰; 2SD) than that in the cores (mean = 5.5 ± 0.1‰; 2SD), supporting an interpretation that most of the rim domains were formed by recrystallization of early zircon grains under fluid conditions rather than modification of existing zircon that should have

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Table 1 Zircon U–Pb and O isotopic data from the Hongtongshan VMS Cu–Zn deposit. No.

Isotopic ratios 207

Pb/206Pb

Age (Ma) ±σ (%)

207

1σ

207

1σ

206

0.4903 0.4889 0.4878 0.4642 0.3969 0.4668 0.4735 0.4236 0.4192 0.4138 0.4351 0.5019 0.5002 0.5022 0.5189 0.6201 0.6156 0.5264 0.5049 0.5084 0.6203

0.87 0.94 1.26 0.85 1.09 0.64 0.58 1.04 0.56 0.84 1.09 0.56 0.77 0.64 0.63 1.00 1.30 1.03 0.60 1.04 0.60

2572 2572 2569 2571 2572 2571 2571 2570 2569 2571 2565 2624 2623 2626 2703 3114 3097 3096 2645 2655 3125

25 26 34 24 31 19 18 29 18 24 30 18 22 19 19 23 35 31 18 32 17

2572 2569 2566 2521 2376 2526 2539 2435 2425 2414 2457 2623 2619 2625 2699 3113 3095 2943 2640 2653 3120

12 13 18 12 16 8 6 15 6 11 16 6 10 8 7 13 14 10 7 10 7

HTS2-1, biotite plagioclase gneiss analyzed by Cameca IMS-1280 1 0.16525 0.47 11.02046 1.57 0.4837 2 0.16395 0.39 10.98878 1.55 0.4861 3 0.16428 0.43 10.58598 1.58 0.4673 4 0.16459 0.36 10.39285 1.56 0.4580 5 0.16411 0.41 10.79685 1.56 0.4772 6 0.16498 0.39 10.49260 1.58 0.4613 7 0.16374 0.39 10.73110 1.56 0.4753 8 0.16465 0.50 10.48313 1.59 0.4618 9 0.16651 0.26 10.81813 1.56 0.4712 10 0.16441 0.53 10.40800 1.59 0.4591 11 0.16504 0.43 10.59560 1.57 0.4656 12 0.16449 0.46 10.48823 1.58 0.4625 13 0.16433 0.27 10.50281 1.56 0.4635 14 0.16604 0.37 10.68336 1.55 0.4667 15 0.16498 0.41 10.84268 1.56 0.4767

1.50 1.51 1.52 1.51 1.50 1.53 1.51 1.51 1.54 1.51 1.51 1.51 1.54 1.51 1.51

2510 2497 2500 2503 2498 2507 2495 2504 2523 2502 2508 2502 2501 2518 2507

8 7 7 6 7 7 7 8 4 9 7 8 5 6 7

2525 2522 2487 2470 2506 2479 2500 2478 2508 2472 2488 2479 2480 2496 2510

HTS3-1, pegmatoidal quartz–amazonite vein analyzed by Cameca IMS-1280 1 0.16519 0.24 10.46021 1.52 0.4593 1.50 2 0.16479 0.25 10.48734 1.53 0.4616 1.51 3 0.16515 0.20 10.49322 1.53 0.4608 1.51 4 0.16460 0.15 10.23451 1.51 0.4510 1.50 5 0.16511 0.39 10.34946 1.55 0.4546 1.50 6 0.16513 0.19 10.57113 1.51 0.4643 1.50 7 0.16517 0.26 10.50750 1.53 0.4614 1.50 8 0.16529 0.10 10.96320 1.52 0.4810 1.52 9 0.16538 0.19 10.54597 1.52 0.4625 1.51 10 0.16545 0.14 10.75773 1.51 0.4716 1.50 11 0.16510 0.32 10.51210 1.54 0.4618 1.51

2509 2505 2509 2503 2509 2509 2509 2511 2511 2508 2509

4 4 3 3 7 3 4 2 3 2 5

HTS3-4, massive sulfide ore analyzed by Cameca IMS-1280 (R—Rim; C—core) R1 0.16436 0.11 10.87299 1.51 0.4798 1.50 R2 0.16507 0.19 10.65225 1.51 0.4680 1.50 R3 0.16503 0.18 10.87126 1.51 0.4778 1.50 R4 0.16635 0.30 10.56934 1.58 0.4608 1.55 R5 0.16402 0.35 10.75262 1.55 0.4755 1.51 R6 0.16583 0.40 10.84194 1.56 0.4742 1.51 R7 0.16556 0.50 10.80382 1.64 0.4733 1.56 R8 0.16450 0.23 11.46527 1.64 0.5055 1.63 R9 0.16496 0.43 10.79433 1.59 0.4746 1.53 R10 0.16395 0.43 11.29808 1.57 0.4998 1.51 R11 0.16432 0.59 11.42709 1.61 0.5044 1.50 C1 0.17161 0.33 11.59229 1.54 0.4899 1.50 C2 0.17023 0.42 11.50556 1.62 0.4902 1.56

2505 2508 2508 2521 2498 2516 2513 2502 2507 2497 2520 2573 2560

HTS3-9 banded sulfide ore analyzed by Cameca IMS-1280 (R—Rim; C—core) R1 0.16529 0.43 10.19587 1.56 0.4474 1.51 R2 0.16341 0.50 10.57349 1.59 0.4693 1.51 R3 0.16408 0.39 10.57549 1.55 0.4675 1.50 R4 0.16520 0.24 10.83480 1.52 0.4757 1.50

2510 2491 2498 2510

±σ (%)

207

Pb/235U

±σ (%)

HTS1-1, amphibolite gneiss analyzed by LA-ICP-MS 1 0.1715 1.49 11.5951 1.27 2 0.1715 1.59 11.5592 1.40 3 0.1712 2.06 11.5139 1.97 4 0.1714 1.46 10.9688 1.24 5 0.1715 1.87 9.3831 1.72 6 0.1714 1.16 11.0293 0.81 7 0.1714 1.08 11.1864 0.68 8 0.1713 1.76 10.0019 1.60 9 0.1712 1.06 9.8920 0.65 10 0.1713 1.46 9.7762 1.24 11 0.1708 1.82 10.2456 1.69 12 0.1769 1.06 12.2459 0.63 13 0.1768 1.33 12.1918 1.06 14 0.1771 1.16 12.2616 0.80 15 0.1855 1.13 13.2697 0.76 16 0.2391 1.45 20.4430 1.29 17 0.2366 2.20 20.0750 1.49 18 0.2364 1.98 17.1533 1.02 19 0.1791 1.10 12.4692 0.72 20 0.1802 1.98 12.6373 1.04 21 0.2408 1.07 20.5949 0.69

206

Pb/238U

Pb/206Pb

Concentration (ppm) Pb/235U

Pb/238U

O isotope (‰) δ18O

2σ

0.34 0.31 0.31 0.37 0.35 0.34 0.29 0.48 0.45 0.30 0.44 0.38 0.21 0.36 0.33

5.59 5.33 5.61 5.37 5.67 5.53 5.29 5.56 5.31 5.45 5.67 5.33 5.71

0.29 0.34 0.17 0.49 0.37 0.31 0.25 0.25 0.25 0.33 0.28 0.30 0.38

5.35

0.24

53 68 86 145 60 63 108 85 84 126 89

0.08 0.10 0.11 0.11 0.09 0.04 0.11 0.04 0.07 0.09 0.11

7.04 7.54 7.29 7.33 7.20 7.51 7.47 7.42 7.19 7.71 6.65

0.27 0.27 0.33 0.30 0.37 0.28 0.31 0.25 0.29 0.26 0.24

3194 1114 980 344 182 137 243 340 169 123 127 189 120

51 10 8 8 6 5 11 8 7 4 2 105 71

0.02 0.01 0.01 0.02 0.04 0.04 0.04 0.02 0.04 0.04 0.02 0.55 0.59

6.49 6.66 6.49 6.55

0.22 0.20 0.17 0.40

6.98 6.54 6.18

0.28 0.38 0.28

6.47 6.74 5.51 5.79

0.29 0.34 0.21 0.36

166 177 214 535

40 48 63 21

0.24 0.27 0.29 0.04

5.66 5.88 5.86 6.10

0.27 0.21 0.37 0.41

1σ

U

Th

Th/U

2572 2566 2561 2458 2155 2470 2499 2277 2257 2232 2329 2622 2615 2623 2694 3110 3092 2726 2635 2650 3111

18 20 27 17 20 13 12 20 11 16 21 12 17 14 14 25 32 23 13 23 15

65 60 19 80 121 226 317 106 1183 308 44 1303 256 138 228 143 858 946 224 590 197

82 50 17 85 44 39 142 110 351 136 82 111 85 45 131 103 259 85 122 179 123

1.28 0.83 0.92 1.07 0.36 0.17 0.45 1.04 0.30 0.44 1.87 0.09 0.33 0.33 0.57 0.72 0.30 0.56 0.55 0.30 0.62

15 15 15 15 15 15 15 15 15 15 15 15 15 15 15

2543 2554 2472 2431 2515 2445 2507 2447 2489 2436 2464 2450 2455 2469 2513

32 32 31 31 31 31 31 31 32 31 31 31 31 31 31

172 126 142 166 126 145 146 204 350 197 171 149 558 246 192

59 40 44 62 44 50 43 99 158 58 75 57 117 87 63

2476 2479 2479 2456 2467 2486 2481 2520 2484 2502 2481

14 14 14 14 14 14 14 14 14 14 14

2436 2446 2443 2399 2416 2459 2446 2532 2451 2491 2447

31 31 31 30 30 31 31 32 31 31 31

622 677 788 1280 651 1434 1001 2287 1172 1409 820

2 3 3 5 6 7 8 4 7 7 10 5 7

2512 2493 2512 2486 2502 2510 2506 2562 2506 2548 2559 2572 2565

14 14 14 15 14 15 15 15 15 15 15 14 15

2526 2475 2518 2443 2507 2502 2498 2637 2504 2613 2633 2570 2572

31 31 31 32 31 31 32 35 32 32 33 32 33

7 8 6 4

2453 2486 2487 2509

15 15 14 14

2384 2480 2472 2508

30 31 31 31

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363

Table 1 (continued) No.

Isotopic ratios 207

Pb/206Pb

Age (Ma) ±σ (%)

207

Pb/235U

±σ (%)

206

Pb/238U

±σ (%)

HTS3-9 banded sulfide ore analyzed by Cameca IMS-1280 (R—Rim; C—core) R5 0.16584 0.23 10.94247 1.52 0.4786 1.50 R6 0.16514 0.45 10.56173 1.61 0.4638 1.54 R7 0.16480 0.39 10.52481 1.55 0.4632 1.51 R8 0.16483 0.42 10.42341 1.57 0.4586 1.51 R9 0.16612 0.33 10.22320 1.54 0.4463 1.51 R10 0.16504 0.44 10.37600 1.57 0.4560 1.51 R11 0.16422 0.41 10.63072 1.56 0.4695 1.50 R12 0.16381 0.36 10.48105 1.57 0.4640 1.53 R13 0.16555 0.33 10.52964 1.55 0.4613 1.51 R14 0.16435 0.48 10.63825 1.57 0.4695 1.50 R15 0.16460 0.36 10.62380 1.55 0.4681 1.51 R16 0.16617 0.28 11.10352 1.53 0.4846 1.50 C1 0.16995 0.56 11.26175 2.06 0.4806 1.99 C2 0.17127 0.48 11.20924 1.58 0.4747 1.51 C3 0.17104 0.62 11.04320 1.66 0.4683 1.53 C4 0.17095 0.56 11.37887 1.60 0.4828 1.50 C5 0.17139 0.71 11.56227 1.68 0.4893 1.52 C6 0.17188 0.66 11.10346 1.64 0.4685 1.50 C7 0.17122 0.64 11.39119 1.65 0.4825 1.52 C8 0.17247 0.62 11.21147 1.62 0.4715 1.50 C9 0.17073 0.69 11.24619 1.65 0.4777 1.50 C10 0.16958 0.55 10.84950 1.68 0.4640 1.59 C11 0.17223 0.60 11.67221 1.62 0.4915 1.50 C12 0.17034 0.74 11.54106 1.68 0.4914 1.51 C13 0.17120 0.61 11.31986 1.62 0.4796 1.50

207

Concentration (ppm)

Pb/206Pb

2516 2509 2505 2506 2519 2508 2500 2495 2513 2501 2504 2515 2557 2570 2568 2567 2571 2576 2570 2582 2565 2554 2579 2561 2569

similar U concentration, oxygen isotope composition and a correlation between U content and age (Tomaschek et al., 2003). Considering the preservation of rim–core zircons in the Cu–Zn ores, these grains thus represent growth of new zircon from a fluid. Hence, we consider that these zircons are hydrothermal zircons that overgrew from the inherited older zircon grains. Moreover, abundant fluid inclusions captured in the rim domains (Fig. 7d and e) also suggest a hydrothermal origin. Thereby, the ages of hydrothermal zircons define the modification age of the Hongtoushan VMS Cu–Zn deposit at 2507 Ma that is nearly synchronous with the metamorphic processes of the felsic volcanic rocks. The pegmatoidal quartz–amazonite vein was once considered to represent a later hydrothermal process, with no apparent genetic relationship with the Cu–Zn mineralization (Zhang et al., 1984; Yu, 2006). The zircons in the pegmatite most likely are hydrothermal in origin based on (1) the Th/U ratio is low (0.09 ± 0.02), reflecting depletion in Th, consistent with formation by hydrothermal involvement; (2) the U–Pb age of 2508 ± 2 Ma is consistent with the ages of the felsic volcanic rocks and Cu–Zn mineralization, indicating their synchronous precipitation processes; (3) the mean δ18Ozircon value 7.3 ± 0.3‰ (2SD) is greater than that of magmatic origin (Fig. 9); and (4) fluid inclusions were captured in the zircon grains (Fig. 7c). Therefore, the aforementioned features support an interpretation that the formation of the pegmatite was sourced from post hydrothermal solution. As mentioned above, the hydrothermal fluids responsible for the modification of the Cu–Zn ores and the formation of pegmatite were both derived from the hydrothermal process and had δ18Ozircon values more elevated than volcanic rocks (Fig. 9). Considering almost no δ18Ozircon influence for the mafic and felsic volcanics, this increase of δ18Ozircon should indicate an input of a high δ18O source rather than exchange of protolith with surface waters at low temperature (Hoskin, 2005). The submarine sedimentary environment of Hongtoushan VMS deposit supports the oceanic crust and/or oceanic sediment as a high-δ18O inputting source during metamorphic hydrothermal cycling processes. 5.2. Implications for tectonic setting at Qingyuan In Qingyuan, high-precise in situ zircon U–Pb analyses allow us to build an approximate geochronology sequence of various units. It

1σ

207

Pb/235U

4 8 6 7 6 7 7 6 6 8 6 5 9 8 10 9 12 11 11 10 11 9 10 12 10

2518 2485 2482 2473 2455 2469 2491 2478 2483 2492 2491 2532 2545 2541 2527 2555 2570 2532 2556 2541 2544 2510 2578 2568 2550

1σ

206

Pb/238U

14 15 15 15 14 15 15 15 14 15 14 14 19 15 16 15 16 15 15 15 16 16 15 16 15

2521 2457 2454 2434 2379 2422 2481 2457 2445 2481 2475 2547 2530 2504 2476 2539 2568 2477 2538 2490 2517 2457 2577 2577 2525

1σ 31 32 31 31 30 31 31 31 31 31 31 32 42 31 32 32 32 31 32 31 31 33 32 32 31

U 705 162 231 196 309 310 203 277 292 191 165 617 88 137 80 107 57 94 82 114 120 108 98 85 61

Th 47 47 22 55 39 28 69 57 37 60 46 56 52 94 49 69 32 49 50 56 80 70 57 27 33

O isotope (‰) Th/U

δ18O

2σ

0.07 0.29 0.10 0.28 0.13 0.09 0.34 0.21 0.13 0.31 0.28 0.09 0.59 0.69 0.61 0.64 0.56 0.52 0.61 0.49 0.67 0.65 0.59 0.31 0.54

5.77 6.08 5.95 5.82 5.82 6.06 6.29 6.09 6.04 6.39 6.04 5.62 5.36 5.36 5.59 5.29 5.40 5.30 5.46 5.34 5.61 5.35 5.45 5.54 5.62

0.31 0.37 0.26 0.29 0.15 0.36 0.20 0.24 0.30 0.24 0.42 0.29 0.24 0.31 0.26 0.32 0.33 0.28 0.22 0.31 0.35 0.20 0.22 0.31 0.29

reveals that the widespread intrusions of TTG granitoids concentrated at 2560–2530 Ma that are considered as a major continental growth epoch (Zhai and Windley, 1990; Zhai et al., 1990; Zhai and Santosh, 2011; Zhao and Cawood, 2012), followed by eruption of the volcanics at 2515–2507 Ma (limited to the Hongtoushan, Xiaolaihe and Tangtu greenstone) and emplacement of syenogranite of 2505–2502 Ma (Fig. 1). It shows that the Qingyuan plutons and volcanics, ranging in composition from diorite to potassic granite, were emplaced over a short time interval, less than 60 Ma, signifying that they were related to the same major tectonothermal episode. This younger episode (2515–2507 Ma) in modification of Cu–Zn mineralization closely corresponds with volcanism during 2.50 to 2.45 Ga mantle plume breakout event of global extent (Heaman, 1997; Isley and Abbott, 1999; Barley et al., 2005). In the eastern NCC, the Neoarchean magmatism extends over a width of more than 800 km, which is difficult to explain by subduction processes taking place over one short time-span (Zhao et al., 2001). Therefore, Zhao et al. (2001) proposed the mantle plume model as an alternative explanation, which can reasonably explain the extensive exposure of 2.6–2.5 Ga TTG granite, granitoid plutons and later volcanics in the eastern NCC (e.g., Jahn et al., 1988; Wu et al., 2005; Smithies et al., 2007; Yang et al., 2008; Grant et al., 2009; Van Kranendonk et al., 2014). Moreover, geochronology and oxygen isotopic data in our study show that the 2571 magmatic rocks were the results of melting mantlederived magma. Therefore, we suggest that the Neoarchean magmatism and mineralization in Qingyuan were the result of mantle plume activity. 5.3. Cu–Zn–Fe ore genesis at Qingyuan VMS mineralization universally requires a submarine heat source in an extensional tectonic setting, including oceanic environment or at least marine basins developed on rifted crust (e.g., Franklin et al., 2005). The VMS Cu–Zn deposits in the Qingyuan district formed under extensional condition in submarine volcanics probably developed under hot spots driven by mantle plume (Zhai et al., 1985; Grant et al., 2009). We consider that there were two main epochs for the formation of the Qingyuan Cu–Zn deposits; one from primary exhalative processes at 2571 Ma and another by subsequent reworking of the ore at 2507 Ma. The first is that emplacement of sufficient volumes of mantle-derived

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Fig. 7. Concordia diagram of SIMS U–Pb data of various samples from the Hongtoushan deposit. Under cross-polarized light, zircon grains from pegmatite (c), massive (d) and banded (e) ore samples contain fluid inclusions (FI; amplified in yellow circles) and thus are interpreted as hydrothermal in origin. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

magma through underplating led to large-scale melting of juvenile crust with mantle-like δ18Ozircon values and induced extensive eruption of mafic–felsic volcanic rocks. Thus, the volcanic hydrothermal event played a significant role in formation of a typical VMS Cu–Zn in Qingyuan. The 2507 Ma reworking of the primary VMS changed the nature and structure of the orebodies. The most obvious changes undoubtedly include migration and enrichment of metallic elements, as well as deformation of Cu–Zn orebodies. Ohmoto et al. (2006) and Galley et al. (2007) pointed out that Algoma-type BIFs are typically distal to VMS deposits. The largest VMS deposits are genetically linked to bimodal arc volcanism (Franklin et al.,

2005). Significantly, Isley and Abbott (1999) have suggested that plume-driven hydrothermal activity enhances the deposition of BIF by increasing the Fe flux to the global ocean through submarine hydrothermal processes. Hence, hydrothermal systems that produce VMS deposits also deliver iron and manganese to the deep ocean (e.g., German and Von Damm, 2006). In Qingyuan, Algoma-type BIF layers normally overlie at the upper sections of the volcanics and VMS Cu–Zn deposits and are also distally precipitated. It reflects that the VMS hydrothermal systems and Algoma-type BIF have a genetic relationship linked to submarine hydrothermal system. The massive sulfides commonly formed near the volcanic vents, whereas the iron is derived by hydrothermal leaching of

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resulted in the escaping vapor phase rich in hydrogen and HCl (Foustoukos and Bekker, 2008), and thus, transporting water particularly rich in Fe and poor in O (Isley and Abbott, 1999), leading to the deposition of iron formations overlying the volcanics.

6. Conclusions Our study shows that hydrothermal zircon U–Pb system is an effective tool to constrain the timing of the metallic mineralization and hydrothermal processes, even for the Archean rocks, which were subjected to extensive and high grade metamorphism. Zircon oxygen isotopes of volcanics were still preserved during amphibolite facies high-pressure and subsequent high-temperature events, as well as a greenschist-facies overprint, but the oxygen composition of hydrothermal zircons was elevated by input of oceanic crust and/or oceanic sediment. The new geochronological data presented in this study provide temporal constraints on intrusion, volcanic rocks and VMS mineralization. Possibly associated with the mantle plume activity, the ca. 2.57 Ga VMS-hosting volcanic rocks have mantle-like δ18Ozircon value and were most likely derived from partial melting of juvenile crust, such as TTG granites, and have a genetic relationship with the original exhalative formation of the Cu–Zn mineralization. This submarine Neoarchean volcanism was widespread in the Qingyuan district, especially in the north of Hunhe fault and was genetically related to the Cu–Zn–Fe metallogenesis. The orebodies were further enriched by the 2507 Ma hydrothermal processes linked with metamorphism.

Acknowledgment

Fig. 8. δ18Ozircon vs. Th/U (a) and δ18Ozircon vs. 207Pb/206Pb (b) relational diagrams. Zircons scatter as two assemblages; one is mantle in origin and another is hydrothermal source. Mantle δ18O value of 5.3 ± 0.6‰ is after Valley et al. (1998).

submarine volcanic rocks in the deeper part of the ocean resulting in the chemical sediments represented by the iron formation layers. Precipitation of Cu–Zn in the vicinity of volcanic vents consumed a mass of S and

The authors are grateful to the mining geologists Drs. Mingran Huang and Weichun Zhang of Hongtoushan Mining Company and to Dr. Yuwang Wang of Beijing Institute of Geology for Mineral Resources, Beijing, for their help during field work and to Xianhua Li, Qiuli Li and Xiaoxiao Ling for their assistance during analyses. Dr. Bo Wan's suggestions have greatly elevated the article's value. Great thanks are also given to Dr. Franco Pirajno for his text improvement and constructive suggestions. This research was financially supported by the National

Fig. 9. Zircon weighted mean δ18O values (2σ) of various samples from the Hongtoushan deposit, implying the δ18Ozircon evolution during the magmatic hydrothermal convective processes. δ18Ozircon values are successively elevated from banded ore, massive sulfide ore to pegmatite samples, indicating the exchange of protoliths with surface waters at low temperature. Mantle δ18O value of 5.3 ± 0.6‰ is after Valley et al. (1998).

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