Timing and genesis of the adakitic and shoshonitic intrusions in the Laoniushan complex, southern margin of the North China Craton: Implications for post-collisional magmatism associated with the Qinling Orogen

Lithos 126 (2011) 212–232

Contents lists available at ScienceDirect

Lithos j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / l i t h o s

Timing and genesis of the adakitic and shoshonitic intrusions in the Laoniushan complex, southern margin of the North China Craton: Implications for post-collisional magmatism associated with the Qinling Orogen Li-Xue Ding a, b, Chang-Qian Ma a, c,⁎, Jian-Wei Li a, d, Paul T. Robinson e, Xiao-Dong Deng a, d, Chao Zhang c, Wang-Chun Xu c a

State Key Laboratory of Geological Processes and Mineral Resources, China University of Geosciences, Wuhan 430074, China Wuhan Institute of Geology and Mineral Resources, Wuhan 430223, China c Faculty of Earth Sciences, China University of Geosciences, Wuhan 430074, China d Faculty of Earth Resources, China University of Geosciences, Wuhan 430074, China e Department of Earth Science, Dalhousie University, Halifax, Nova Scotia, Canada B3H 3J5 b

a r t i c l e

i n f o

Article history: Received 17 February 2011 Accepted 17 July 2011 Available online 23 July 2011 Keywords: Adakite Shoshonitic rocks Magma mixing Post-collision Qinling Orogen Southern margin of the North China Craton

a b s t r a c t The NWW-striking Qinling Orogen formed in the Triassic by collision between the North China and Yangtze Cratons. Triassic granitoid intrusions, mostly middle- to high-K, calc-alkaline, are widespread in this orogen, but contemporaneous intrusions are rare in the southern margin of the North China Craton, an area commonly considered as the hinterland belt of the orogen. In this paper, we report zircon U–Pb ages, elemental geochemistry, and Sr–Nd–Hf isotope data for the Laoniushan granitoid complex that was emplaced in the southern margin of the North China Craton. Zircon U–Pb dating shows that the complex was emplaced in the late Triassic (228 ± 1 to 215 ± 4 Ma), indicating that it is part of the post-collisional magmatism in the Qinling Orogen. The complex consists of, from early to late, biotite monzogranite, quartz diorite, quartz monzonite, and hornblende monzonite, which span a wide compositional range, e.g., SiO2 = 55.9–70.6 wt.%, K2O + Na2O = 6.6–10.2 wt.%, and Mg# of 24 to 54. The biotite monzogranite has high Al2O3 (15.5–17.4 wt.%), Sr (396–1398 ppm) and Ba (1284–3993 ppm) contents and relatively high La/Yb (mostly 14–30) and Sr/Y (mostly 40–97) ratios, but low Yb (mostly 1.3–1.6 ppm) and Y (mostly14–19 ppm) contents, features typical of adakitic rocks. The quartz monzonite, hornblende monzonite and quartz diorite have a shoshonitic affinity, with K2O up to 5.58 wt.% and K2O/Na2O ratios averaging 1.4. The rocks are characterized by strong LREE/HREE fractionation in chondrite-normalized REE pattern, without obvious Eu anomalies, and show enrichment in large ion lithophile elements but depletion in high field strength elements (Nb, Ta, Ti). The biotite monzogranite (228 Ma) has initial 87Sr/86Sr ratios of 0.7061 to 0.7067, εNd(t) values of −9.2 to −12.6, and εHf(t) values of −9.0 to −15.1; whereas the shoshonitic granitoids (mainly 217–215 Ma) have similar initial 87Sr/86Sr ratios (0.7065 to 0.7075) but more radiogenic εNd(t) (−12.4 to −17.0) and εHf(t) (−14.1 to −17.0). The Sr–Nd– Hf isotope data indicate that the rocks were likely generated by partial melting of an ancient lower continental crust with heterogeneous compositions, as partly confirmed by the widespread presence of early Paleoproterozoic inherited zircons. Mafic microgranular enclaves (MMEs), characterized by fine-grained igneous textures and an abundance of acicular apatites, are common in the Laoniushan complex. Compared with the host rocks, they have lower SiO2 (48.6–53.7 wt.%) and higher Mg# (51–56), Cr (122–393 ppm), and Ni (24–79 ppm), but equivalent Sr–Nd isotope compositions, indicating that the MMEs likely originated from an ancient enriched lithospheric mantle. The abundance of MMEs in the granitoid intrusions suggest that magma mixing plays an important role in the generation of the Laoniushan complex. Collectively, it is suggested that the Laoniushan complex was a product of post-collisional magmatism related to lithospheric extension following slab break-off. Formation of the adakitic and shoshonitic intrusions in the Laoniushan complex indicates that the Qinling Orogen had evolved into a post-collisional setting by about 230–210 Ma. © 2011 Elsevier B.V. All rights reserved.

1. Introduction ⁎ Corresponding author at: Faculty of Earth Sciences, China University of Geosciences, Lumo Road 388, Hongshan District, Wuhan, Hubei Province 430074, China. Tel.: + 86 27 67883139; fax: + 86 27 87515956. E-mail address: [email protected] (C.-Q. Ma). 0024-4937/$ – see front matter © 2011 Elsevier B.V. All rights reserved. doi:10.1016/j.lithos.2011.07.008

Adakites are sodium-rich igneous rocks that are derived from partial melting of subducted oceanic crust (Defant and Drummond, 1990). However, alternative mechanisms have also been proposed to

L.-X. Ding et al. / Lithos 126 (2011) 212–232

explain the generation of adakite-like rocks in various environments, and these mainly include (1) partial melting of thickened lower crust (Chung et al., 2003; Xu et al., 2007) or delaminated lower continental crust in the mantle lithosphere (Gao et al., 2004; Wang et al., 2004; 2006), (2) partial melting of peridotitic mantle metasomatized by slab melts (Martin et al., 2005), and (3) fractionational crystallization of mantle-derived magmas (Castillo et al., 1999; Li et al., 2008; 2009). Some adakites or adakite-like rocks may also have shoshonitic or high-K affinities (Castillo et al., 2002; Kepezhinskas et al., 1996). The shoshonite series commonly refer to volcanic rocks of basaltic to andesitic composition (Iddings, 1892; Joplin, 1965; Morrison, 1980), but recent studies show that their plutonic counterparts are also common (Duchesne et al., 1998; Eklund et al., 1998; Fowler and Henney, 1996; Jiang et al., 2002). Shoshonitic rocks typically occur in destructive plate margin settings, either during the initial stages of arc construction (Stern et al., 1988) or as the arc matures (Morrison, 1980), they may also form in extensional or post-collisional settings (Dostal et al., 2002; Duchesne et al., 1998; Hadi et al., 2003; Sun et al., 2008). Shoshonites constitute a unique potassium-rich suite of orogenic igneous rocks (Dostal et al., 2002; Duchesne et al., 1998; Kepezhinskas, 1995; Lange et al., 1993; Sun et al., 2008), and are of particular significance in reconstructing tectonic settings and under-

213

standing the genesis of associated mineral deposits (Ashley et al., 1994; Morrison, 1980; Redwood and Rice, 1997; Xie et al., 2006). The Qinling–Dabie–Sulu orogenic belt in east-central China marks the junction between the Yangtze and North China Cratons (Fig. 1a). This orogenic belt contains the largest collection of ultrahigh-pressure (UHP) metamorphic rocks in the world, and the age of the UHP metamorphism is thought to represent the time of continental collision. Available geochronological data and inferred P-T paths of metamorphic rocks show that the UHP metamorphism in the Dabie– Sulu belt occurred at 240 to 220 Ma (Hacker et al., 1998; Liu et al., 2004). The Qinling Orogen also contains a UHP metamorphic unit with isotopic ages of ~ 450–500 Ma, recording a middle Paleozoic collision between the North Qinling terrane and Yangtze Cratons (Yang et al., 2005a). These different ages have led to many controversies regarding the tectonic evolution of the Qinling Orogen (Jiang et al., 2010; Qin et al., 2009; Sun et al., 2000; Zhang et al., 2008). Certain types of magmatic activity can provide useful guides to specific tectonic regimes, e.g. the association of A-type granites, alkaline rocks and shoshonitic rocks with extensional environments (Seo et al., 2010; Yang et al., 2005b). Early Mesozoic granitoids, which have middle–high-K and calc-alkaline compositions, are widespread in the Qinling Orogen but their petrogenesis and tectonic implications

Fig. 1. (a) Simplified map of East China denoting the location of the Qinling orogenic belt. (b) Sketch map showing the distribution of Mesozoic granitoids in the Qinling orogenic belt. (c) Geological map of the Laoniushan complex showing the distinct intrusive units. Also shown are the sampling sites.

214

L.-X. Ding et al. / Lithos 126 (2011) 212–232

remain a matter of debate (Jiang et al., 2010; Qin et al., 2009; 2010; Sun et al., 2000; Wang et al., 2007a; Zhang et al., 2008). Hence, recognition of some special rocks, in particular those with significance in denoting tectonic setting, is critical for further understanding of the tectonic evolution of the Qinling orogenic belt. In the current study, we report a detailed account of the petrography, geochronology, and geochemistry of the Laoniushan intrusive complex on the southern margin of the North China Craton, which is composed of a rare association of adakitic and shoshonitic igneous rocks, attempting to provide new constraints on the timing and genesis of the rocks and tectonic evolution of the Qinling Orogen. 2. Regional geological setting and geology of the Laoniushan complex The North China Craton (NCC) consists of Archean to Paleoproterozoic basement that is variably overlain by Mesoproterozoic to Phanerozoic unmetamorphosed cover rocks (Zhao et al., 2000, and references therein). Based on age, lithological assemblages, tectonic evolution, and P-T-t paths, this craton is divided into an Eastern and Western Block, separated an approximately 300-km-wide central zone, termed the Trans-North China Orogen (Zhao et al., 1998), or Central Orogenic Belt (Kusky et al., 2007). During the Early Mesozoic, multiple tectono-thermal events occurred along the northern and eastern margins of the NCC, but coeval magmatism was rare in the interior and southern edge of the NCC. Numerous studies have shown that the Eastern Block experienced widespread lithospheric thinning during the Late Mesozoic (Xu et al., 2000; Yang et al., 2003; Zhai et al., 2005, 2007), as indicated by extensive volcanism, emplacement of voluminous granitoids, the development of rift basins, uplift, and formation of metamorphic core complexes. The Qinling Orogen is separated by the Shangdan suture in the north and the Mianlue suture in the south (Fig. 1b), and connects with the Kunlun and Qilian Orogens to the west and the Dabie UHP metamorphic belt to the east. Early Mesozoic granitoids, formed at 230 to 200 Ma, are widespread in the orogen. They can be broadly grouped into three compositional categories, i.e., adakite-like intrusions, middle-to-high-K calc-alkaline granitoids, and rapakivitextured plutonic rocks (Jin et al., 2005; Qin et al., 2008; Sun et al., 2000; Wang et al., 2007a; Zhang et al., 2005, 2007a). The Xiaoqinling–Xiong'ershan district is located in a zone between the southern margin of NCC and the North Qinling belt, northwest of the Dabie UHP orogenic belt (Fig. 1b). The geology of the district is dominated by Late Archean metamorphic rocks of the Taihua Group, which consist predominantly of amphibolite, felsic gneiss, migmatite, and metamorphosed supracrustal rocks. Mesoproterozoic marine sedimentary rocks and Paleoproterozoic volcanic rocks are variably exposed in the district (Zhao et al., 2004). Numerous granitoids, ranging in composition from monzogranite and alkali granite to quartz diorite and granodiorite, intrude the Archean and Proterozoic rocks. Mafic dykes, spatially associated with the granitoid intrusions (Wang et al., 2008a), as well as Early Cretaceous volcanic rocks (Xie et al., 2007), have also been reported in this area. Mao et al. (2010) have shown that the magmatism in this area occurred mostly in Late Jurassic to Early Cretaceous (157–112 Ma). The Laoniushan intrusive complex is a lenticular body that crops out over an area of about 440 km 2 (Fig. 1c). It intrudes various Precambrian metamorphic rocks, including the Archean Taihua Group gneiss, the Paleoproterozoic Tietonggou Group quartzite, the Xiong'er Group mafic volcanic rocks and quartzites, and the Gaoshanhe Group slates. Based on field relationships, four intrusive units are recognized, i.e. biotite monzogranite, quartz diorite, hornblende monzonite, and quartz monzonite. Lateral zoning is evident in the biotite monzogranite unit, as manifested by a coarse-grained central facies, a medium-grained transitional facies, and a fine-grained marginal facies (Fig. 1c), indicating prolonged cooling during emplacement.

The quartz diorite unit is exposed principally in the southwestern part of the complex, with a few small outcrops in the central and northern parts where it intrudes the biotite monzogranite (Fig. 1c). However, the quartz diorite is locally fractured and intruded by biotite monzogranite in some outcrops, indicative of coeval intrusion and cooling of the two magma types. The biotite monzogranite is also intruded by both the hornblende monzonite and quartz monzonite, whereas no clear intrusive relationship between the latter two units has been observed. Mafic microgranular enclaves (MMEs) are widespread in the Laoniushan complex, but are mainly hosted in the quartz monzonite and hornblende monzonite. They are mostly elliptical or oval in shape and range from a few centimeters to almost one meter in size. The contacts between the MMEs and the host rocks are generally sharp, but transitional contacts also occur locally. The enclaves are dioritic to monzodioritic in composition and show variable textures from equigranular to porphyritic, indicative of an igneous origin. Abundant angular to subangular wall-rock xenoliths also occur along the margins of the complex. 3. Petrography The four units differ in texture and mineral assemblages. The biotite monzogranite is pale gray and consists of quartz (20–30 vol.%), K-feldspar (~40 vol.%), plagioclase (~30 vol.%), biotite (3–5 vol.%) and minor amounts of amphibole. Some plagioclase grains host residual microcline, likely reflecting metasomatism (Fig. 2a). Microcline and perthitic microcline are primary species of alkali feldspar. Titanite, magnetite, zircon and apatite are the common accessory minerals. The quartz diorite is gray and porphyritic, with large euhedral plagioclase phenocrysts (An40–45, 8–10 vol.%) in a groundmass consisting of plagioclase (An35–40, 60 vol.%), K-feldspar (20–25 vol.%), quartz (~10 vol.%), amphibole (3–4 vol.%), biotite (1–2 vol.%) and accessory minerals (b1 vol.%) such as apatite and zircon. The quartz monzonite is equigranular and locally porphyritic in texture. It is composed of subequal amounts of plagioclase (An42–50, 50 vol.%) and K-feldspar (mainly microcline, 30 vol.%), with subordinate quartz (5–10 vol.%), amphibole (3–5 vol.%), biotite (b5 vol.%) and accessory minerals (b1 vol.%) including apatite, titanite, zircon and epidote. Plagioclase grains display euhedral zoning and polysynthetic twinning (Fig. 2b–c). The hornblende monzonite is similar in texture to the quartz monzonite, but contains much lesser quartz and more hornblende. The MMEs are monzodioritic or, less commonly, dioritic in composition, consisting of amphibole (30–35 vol.%), plagioclase (An60–65; 45–50 vol.%), K-feldspar (10–15 vol.%), and minor amounts of biotite, with or without quartz. Titanite, apatite, epidote, and allanite are the main accessory minerals. Both columnar and acicular apatites are present in the MMEs (Fig. 2d), but the latter is dominant. 4. Analytical methods Four samples, each from a different intrusive unit of the Laoniushan complex, were selected for LA-ICP-MS zircon U–Pb dating. Zircons were separated from a ca. 0.5–1 kg samples using standard crushing, grinding, heavy liquid, and magmatic separation methods. The grains were then handpicked under a binocular microscope based on size and morphology. The mineral separates were mounted in epoxy resin disks, polished, cleaned, and gold-coated. Cathodoluminescence (CL) images of the zircons were obtained prior to U–Pb analysis, using a JEOL JXA-8900 microprobe at the Chinese Academy of Geological Sciences, Beijing, China. U–Pb dating was performed using a laser ablation-inductively-coupled plasma mass spectrometer (LA-ICP-MS) at the State Key Laboratory of Geological Processes and Mineral Resources (GPMR), China University of Geosciences, Wuhan. Zircon 91500 was used as the standard, whereas the standard silicate glass NIST610 was used to optimize the machine. The laser beam was

L.-X. Ding et al. / Lithos 126 (2011) 212–232

215

Fig. 2. Photomicrographs showing the mineralogy and texture of distinct intrusions and enclosed MMEs of the Laoniushan complex. (a) Biotite monzogranite with plagioclase grain and enclosed residual microcline, likely reflecting metasomatism. (b) Quartz monzonite consisting of hornblende, biotite, plagioclase, K-feldspar, and quartz. Note the zoned plagioclase suggesting an unbalanced texture. (c) Plagioclase phenocrysts in quartz monzonite, with polysynthetic twin and oscillatory zones. Note the zoning of the plagioclase phenocryst is punctuated. (d) Columnar and acicular apatites in the MME. Mineral abbreviations: Pl—plagioclase, Kf—K-feldspar, Hb—hornblende, Bt—biotite, Qz—quartz, Apt—apatite.

32 μm in diameter. The detailed analytical procedures are described in Liu et al. (2010). Data reduction was conducted using an in-house program ICPMSDataCal (Liu et al., 2010). Common Pb correction was applied using the method of Andersen (2002). Calculations of weighted mean ages and concordia plots were made using ISOPLOT (Ludwig, 2003). The errors quoted in the U–Pb ages are at the 2σ (95% confidence) levels. Twenty-two fresh whole-rock samples were powdered to b200 mesh using an agate mill, which was cleaned between samples to avoid potential contamination. Major elements were determined using a Regaku 3080E1 XRF spectrometer at the Analytical Institute of the Hubei Bureau of Geology and Mineral Resources. The accuracies of the XRF analyses are estimated at ~ 2% for elements with concentrations greater than 0.5 wt.% and 5% for oxides present in concentrations N0.1 wt.%. Trace elements, including REE, were determined by inductively coupled plasma mass spectrometry (ICP-MS) coupled with an Agilent 7500a system, at the State Key Laboratory of Geological Processes and Mineral Resources (GPMR), China University of Geosciences, Wuhan. The detailed sample-digesting procedure for ICP-MS analyses and analytical precision and accuracy for trace elements are the same as described in Liu et al. (2008). Sr and Nd isotopic ratios were determined using a Finnigan Triton thermo-ion mass isotope spectrometer at GPMR (Wuhan), following the methods of Rudnick et al. (2004). The mass fractionation corrections for Sr and Nd isotopic ratios were based on 86 Sr/ 88Sr = 0.1194 and 146Nd/ 144Nd = 0.7219, respectively. The average 143Nd/ 144Nd ratio of the La Jolla standard measured during the

sample runs is 0.511862 ± 5 (2σ), whereas the NBS-987 standard gave 87Sr/ 86Sr as 0.710236 ± 16 (2σ). In situ zircon Hf isotope analysis was carried out at the Institute of Geology and Geophysics, Chinese Academy of Sciences (Beijing), using a Neptune MC-ICP-MS with a Geolas-193 laser. The signal of 180Hf of zircon 91500 was about 5 V using 15 J/cm2 of energy density, a 60 μm spot size, and 7 Hz ablation frequency. The detailed analytical procedures are described in Wu et al. (2006). The measured 176Lu/177Hf ratios and the 176Lu decay constant of 1.867 × 10 − 11 yr− 1 reported by Söderlund et al.(2004) were used to calculate initial 176Hf/177Hf ratios. For the calculations of εHf values, we used the chondritic values of 176Hf/ 177Hf and 176Lu/177Hf reported by Blichert-Toft and Albarede (1997). Single-stage model ages (TDM1) were calculated relative to the depleted mantle with present-day values (176Lu/177Hf)DM = 0.0384 and (176Hf/177Hf)DM = 0.28325 (Griffin et al., 2000). Two-stage model ages (TDM2) were calculated for the source rock of the magma by assuming a mean 176Lu/177Hf value of 0.015 for the average continental crust (Griffin et al., 2002). 5. Results 5.1. Zircon U–Pb ages Fig. 3 illustrates the typical morphologies and textures of zircons dated in this study. LA-ICP-MS U–Pb data are summarized in Table A1 (available as supplementary online material in Appendix A) and graphically illustrated in the concordia diagrams (Fig. 4).

216

L.-X. Ding et al. / Lithos 126 (2011) 212–232

Fig. 3. Cathodoluminescence (CL) images of representative zircons from the Laoniushan granitoids, together with 206Pb/238U ages and εHf(t) values. (a) biotite monzogranite (LN28); (b) quartz diorite (LN49); (c) quartz monzonite (LN44); (d) hornblende monzonite (LN9). Dashed and solid line circles indicate the areas of U–Pb and Hf isotope analyses, respectively.

Sample LN28 (109°33′27.6″E, 34°14′09.4″N) was taken from the biotite monzogranite. Zircons from this sample are short prismatic or stubby, and colorless or pale yellow. The crystals are 100–250 μm in diameter with length/width ratios of 1:1 to 3:1, and most show magmatic concentric zoning (Fig. 3a). U and Th concentrations range from 2207 to 436 ppm and 898 to 182 ppm, respectively, with Th/U ratios of 0.21–1.4. Twelve out of 14 analyses yielded a cluster of 206Pb/ 238U ages ranging from 223 ± 2 Ma to 230 ± 2 Ma, with a weighted mean of 228 ± 1 Ma (MSWD = 2.8) (Table A1; Fig. 4a). This is considered as the emplacement age of the biotite monzogranite unit. The remaining two analyses (LN28-4 and LN28-5) have apparently younger 206Pb/238U ages of ~213 Ma, possibly resulting from minor Pb loss due to a post-magmatic thermal disturbance, presumably associated with the emplacement of the quartz monzonite and/or hornblende monzonite (see below). Sample LN49 (109°32′26.8″E, 34°14′47.3″N) was collected from the quartz diorite unit. The analyzed zircons are euhedral, transparent, and colorless to light brown. The grains range in size from 150 to 450 μm, with length-to-width ratios of 2:1 to 3.5:1. Most of these zircons display concentric oscillatory zoning (Fig. 3b), indicative of a magmatic origin (Corfu et al., 2003). A few grains have dark cores, and two analyses of these inherited cores yielded relatively low U contents (268–282 ppm) and Th/U ratios (0.14–0.15), with 207Pb/ 206Pb ages of 2325 ± 15 Ma and 2283 ± 11 Ma, respectively. The other 22 analyses were made on the oscillatory bands of separate grains. With one exception (spot LN49-6), these analyses show high, variable U contents (232–4479 ppm) and Th/U ratios of 0.25 to 1.22. They yielded ages that cluster in two groups (Fig. 4b). The first group comprises 4 analyses with a weighted mean 206Pb/ 238U age of 238 ± 5 Ma. The second group comprises 17 analyses and yielded a weighted mean 206Pb/ 238U age of 227 ± 1 Ma (MSWD = 1.8), which is interpreted as the emplacement age of the quartz diorite.

Sample LN44 (109°32′22.4″E, 34°14′51.9″N) was collected from the quartz monzonite. Zircons are euhedral, short to long prismatic and colorless or light brown. The crystals are 100 to 200 μm long, with length/width ratios of 2:1 to 5:1. CL images reveal obvious oscillatory zoning and most zircons have a narrow overgrowth rim (Fig. 3c). A total of 22 analyses were made on 22 zircon grains, of which five were on inherited cores. The analyses on the oscillatory zones have high and variable U contents of 238 to 6319 ppm, with moderate Th/U ratios of 0.29 to 1.01, whereas those on the cores show lower and more restricted U contents (375–1313 ppm) and Th/U ratios (0.15– 0.40) (Table A1). The exception was grain LN44-10, which had a very low Th/U ratio (0.07). The five analyses from the inherited cores plot on a discordant line, with an upper intercept at 2253 ± 44 Ma (MSWD =7.2) (Fig. 4c), which is consistent with a weighted mean 207 Pb/206Pb age of 2258 ±14 Ma (MSWD =19). Among the other 17 analyses, two with younger discordant ages were rejected. The remaining 15 analyses yielded a weighed mean 206Pb/238U age of 217±2 Ma (MSWD =2.6). Sample LN9 (109°49′52.6″E, 34°23′23.3″N) was selected from the hornblende monzonite. Zircons are light brown (partly turbid due to high U content), translucent, and are characterized by well-developed prism and pyramid faces (Fig. 3d). The grains vary in length from 100 to 220 μm, with aspect ratios ranging from 1:1 to 6:1. CL images show that most of zircons have clear oscillatory zoning, and a few contain inherited cores. Eighteen analyses were made on 18 zircon grains. Compared to the other three samples, magmatic zircons in this sample have lower U (210–496 ppm) and Th (177–372 ppm), but higher Th/U ratios (0.8–1.4). Relative to the overgrowth rims, the inherited cores have higher, but more variable U (81–634 ppm) and Th (136– 2023 ppm) contents. Nine spots on the overgrowth rims are concordant or almost concordant (Fig. 4d). Of the 9 analyses, two with older ages (LN-3 and LN-5) were rejected as possibly representing early magmatic

L.-X. Ding et al. / Lithos 126 (2011) 212–232

217

Fig. 4. Zircon U–Pb concordia diagrams for (a) biotite monzogranite (LN28); (b) quartz diorite (LN49); (c) quartz monzonite (LN44); (d) hornblende monzonite (LN9).

activity. The remaining 7 analyses yielded a weighted mean 206Pb/ 238U age of 215 ± 4 Ma (MSWD = 1.1) (Fig. 4d), which is considered as the crystallization age for the hornblende monzonite. Nine spots on the inherited cores form a discordia, with an upper intercept at 2275± 57 Ma (MSWD= 18), consistent with a weighted mean 207Pb/206Pb age of 2193 ± 50 Ma (MSWD = 21) within analytical uncertainties. 5.2. Major and trace elements The major and trace element data for selected samples of both host rocks and MMEs are listed in Table 1. The host granitoids have an intermediate to felsic composition, with SiO2 contents ranging from 55.89 to 70.60 wt.%. These rocks have high total alkalis (K2O + Na2O = 6.97–10.18 wt.%) and high K2O/Na2O (0.8–1.8) ratios. They are characterized by low TiO2 (b1 wt.%), as well as high and variable Al2O3 (15.04 to 17.55 wt.%). Among the four intrusive units, the biotite monzogranite has the lowest TiO2 (0.16–0.25 wt.%), MgO (0.21–0.86 wt.%) and Mg # [atomic Mg/(Mg + Fe)] (24–44), FeO t (1.17–1.99 wt.%), MnO (0.04–0.09 wt.%), CaO (1.19–1.92 wt.%), P2O5 (0.05–0.19 wt.%), Cr (1.5–20.8 ppm) Ni (0.45–6.96 ppm), Y (13.9– 22.4 ppm) and Yb (1.34–2.30 ppm) concentrations, whereas the corresponding concentrations of the other rocks are relatively high, e.g. MgO = 1.62–4.03 wt.%, Mg # = 45–54. Moreover, the biotite monzogranite is high-K calc-alkaline to shoshonitic (Fig. 5) and weakly peraluminous whereas the other host rocks are shoshonitic (Fig. 5) and metaluminous. The MMEs are mafic to intermediate in

composition, with SiO2 of 48.57–53.71 wt.%. Compared with the host rocks, the MMEs have higher TiO2 (0.81–1.21 wt.%), FeO t (7.75– 10.83 wt.%), MnO (0.27–0.32 wt.%), MgO (4.77–6.88 wt.%), CaO (6.06–7.65 wt.%), P2O5 (1.12–1.30 wt.%), Cr (122–393 ppm), and Ni (24.4–78.8 ppm), but significantly lower Na2O (1.62–2.19 wt.%). On Harker diagrams (Fig. 5), with increasing SiO2 content, the Laoniushan granitoids become progressively depleted in FeO t, MgO, TiO2, CaO, P2O5, Y, V and Yb, possibly reflecting fractional crystallization of hornblende, apatite, and Fe–Ti oxides. Although all four units exhibit broadly similar, subparallel chondrite-normalized REE patterns, characterized by strong fractionation between LREE and HREE, with (La/Yb)N ratios of 11.6–30.3 (Table 2), the difference is clear, e.g. the biotite monzogranite possesses markedly negative P anomalies and strongly Ti negative anomalies, however, the other host rocks display obviously positive P anomalies and weakly Ti negative anomalies (Fig. 6). Compared with the host rocks, the MMEs have relatively low (La/Yb)N ratios of 4.7– 5.3. Both the host rocks and MMEs display a weakly concave middle REE pattern (Fig. 6), which is commonly considered to indicate the presence of residual amphibole in the source (Gromet and Silver, 1987). There are no marked negative Eu anomalies for most samples (Eu/Eu* = 0.80–1.1), except for sample LN21 and LN26 (Eu/Eu* = 0.59 and 0.68, respectively). As shown in Fig. 6, the majority of the host rocks, as well as the MMEs, have broadly similar primitive mantlenormalized multi-element spidergrams, among which, the biotite monzogranite has relatively lower HREE concentrations. Notably, all

218

Sample no. LN8

LN9

LN11

LN17

LN3

LN18

LN20

LN21

LN26

LN28

LN31

LN50

LN35

LN37

LN39

LN40

LN46

LN48

LN49

LN34

LN42

LN45

Rock type

HM

HM

HM

BMG

BMG

BMG

BMG

BMG

BMG

BMG

BMG

QM

QM

QM

QM

QM

QM

QD

MME

MME

MME

HM

Major oxides (wt.%) SiO2 56.65 TiO2 0.75 Al2O3 15.29 Fe2O3 2.02 FeO 3.82 MnO 0.12 MgO 3.69 CaO 6.02 Na2O 3.42 K2O 5.51 P2O5 1.17 H2O+ 0.74 CO2 0.07 Total 99.27 # Mg 53.85 A/CNK 0.68 K2O + 8.93 Na2O 1.61 K2O/Na2O Fe2O3/FeO 0.53 Trace elements (ppm) V 82.8 Cr 36.4 Co 11 Ni 12.3

60.11 0.5 17.1 1.76 2.68 0.09 2.1 4.17 4.11 5.45 0.56 0.52 0.07 99.22 46.76 0.84 9.56

55.89 0.8 15.04 2.33 4.25 0.16 4.03 6.12 3.06 5.58 1 0.88 0.15 99.29 53.1 0.68 8.64

58.56 0.59 16.55 1.78 3.08 0.11 2.61 5 3.99 5.36 0.87 0.58 0.15 99.23 49.85 0.77 9.35

67.39 0.2 17.36 0.74 0.89 0.04 0.54 1.58 5.25 4.93 0.07 0.39 0.07 99.45 38.23 1.03 10.18

66.51 0.25 17.03 0.74 1.27 0.06 0.86 1.92 4.46 5.56 0.17 0.42 0.07 99.32 44.2 1.01 10.02

67.78 0.25 16.93 0.86 1.22 0.05 0.73 1.79 4.65 4.56 0.12 0.37 0.07 99.38 39.5 1.07 9.21

69.35 0.2 16.73 0.86 0.8 0.09 0.32 1.41 3.77 5.42 0.08 0.57 0.11 99.71 26.61 1.14 9.19

69.57 0.16 17.12 0.64 0.6 0.06 0.21 1.58 3.97 5.22 0.05 0.4 0.07 99.65 24.15 1.14 9.19

69.71 0.25 15.48 0.9 1.03 0.06 0.79 1.88 4.31 4.28 0.18 0.47 0.3 99.64 43.36 1.02 8.59

68.15 0.21 16.64 0.68 1 0.05 0.7 1.81 4.63 5.03 0.19 0.41 0.07 99.57 43.64 1.02 9.66

70.6 0.18 16.16 0.76 0.68 0.06 0.32 1.19 4.42 4.53 0.06 0.46 0.07 99.49 29.49 1.13 8.95

64.81 0.38 16.44 1.16 2.12 0.1 1.62 2.98 3.76 4.77 0.3 0.65 0.44 99.53 47.72 0.98 8.53

60.95 0.61 16.13 1.37 3.65 0.12 2.62 4.37 3.47 4.25 0.62 1.19 0.15 99.5 48.89 0.88 7.72

62.23 0.54 17.55 1.22 2.42 0.08 1.64 3.49 3.8 5.41 0.37 0.63 0.11 99.49 45.39 0.95 9.21

59.25 0.75 15.35 1.82 4.48 0.15 3.35 4.77 3.17 4 0.66 1.65 0.15 99.55 49.4 0.84 7.17

58.96 0.74 16.2 1.73 4.17 0.13 3.23 4.49 3.32 4.35 0.42 1.58 0.22 99.54 50.14 0.88 7.67

62.6 0.53 16.63 1.37 2.87 0.1 2.17 3.59 3.62 4.32 0.3 1.39 0.04 99.53 48.53 0.97 7.94

60.17 0.72 15.85 2.13 3.72 0.13 3 4.9 3.77 3.2 0.64 1.15 0.11 99.49 48.69 0.85 6.97

53.71 0.81 14.46 2.86 5.18 0.27 4.77 6.06 2.19 5.71 1.18 1.79 0.67 99.66 52.31 0.7 7.9

50.64 0.98 12.67 2.96 6.92 0.32 6.88 7.65 1.62 5.1 1.12 2.22 0.52 99.6 56.14 0.57 6.72

48.57 1.21 13.98 3.65 7.55 0.29 6.27 7.26 2.1 4.46 1.3 2.62 0.41 99.67 50.78 0.65 6.56

1.33 0.66

1.82 0.55

1.34 0.58

0.94 0.83

1.25 0.58

0.98 0.7

1.44 1.08

1.31 1.07

0.99 0.87

1.09 0.68

1.02 1.12

1.27 0.55

1.22 0.38

1.42 0.5

1.26 0.41

1.31 0.41

1.19 0.48

0.85 0.57

2.61 0.55

3.15 0.43

2.12 0.48

22.5 11.3 2.81 3.72

30 20.8 4.12 6.96

27.5 14.7 3.53 5.1

10.4 1.52 1.5 0.45

8.9 1.91 1.14 0.52

30.5 15.1 3.62 4.44

25.6 14.7 3.37 4.54

14.7 2.97 1.23 0.98

56.2 31.4 7.42 9.58

59.3 31.9 8.01 6.88

119 72.1 16.6 20.3

127 72.7 19.4 23.5

94.9 36.2 13.1 14.6

132 218 20.8 43.7

111 64.2 15.7 19.6

106 63.2 15.9 16.4

70 40.4 10.5 10

105 52.7 13.2 14

132 218 20.8 43.7

183 393 28.6 78.8

204 122 26.7 24.4

L.-X. Ding et al. / Lithos 126 (2011) 212–232

Table 1 Data of major and trace elements for Laoniushan granitoids and enclaves.

Table 1 (continued) LN9

LN11

LN17

LN3

LN18

LN20

LN21

LN26

LN28

LN31

LN50

LN35

LN37

LN39

LN40

LN46

LN48

LN49

LN34

LN42

LN45

Rock type

HM

HM

HM

HM

BMG

BMG

BMG

BMG

BMG

BMG

BMG

BMG

QM

QM

QM

QM

QM

QM

QD

MME

MME

MME

Ga Rb Sr Y Zr Nb Cs Ba Hf Ta Pb Th U La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu ∑ REE Sr/Y (La/Yb)N δEu

19.8 18.5 19.7 20 20.9 19.6 19.2 24.5 24.8 20.3 21.3 20.3 25.6 20.5 20.1 22.7 20.4 19.1 20.3 25.6 21.5 26.5 110 115 109 103 115 131 106 248 210 155 161 139 231 165 152 114 155 106 98.9 231 140 127 1917 1575 1520 1761 1284 1377 1398 396 600 816 929 897 416 1105 1214 1087 1076 1123 1289 416 580 582 22.2 27.8 34 25.2 13.9 16.6 14.4 21.4 19.3 17.3 15 22.4 41.7 22.3 18.8 32.5 21.1 19.2 29.3 41.7 35.4 50.2 289 422 411 446 286 269 294 286 227 186 193 178 420 239 302 434 129 117 290 420 290 376 15.2 16.8 20.9 15 22.2 23.2 21.6 53.7 24.7 18.9 16.5 30.9 38.2 20.2 16.3 29.2 16.5 16.1 22.4 38.2 25 39.4 1.38 1.27 1.26 1.88 1.26 1.9 2.82 3.2 2.16 6.33 4.71 2.89 2.61 8.02 3.61 1.58 1.6 1.35 4.28 2.61 1.08 1.68 4538 4073 4278 4410 2978 3993 3221 1284 1698 1579 1960 2595 1395 2227 2676 2107 2126 2120 2291 1395 2009 1398 6.99 10.2 10.1 10.5 7.66 7.15 7.8 8.61 6.24 5.82 6.29 5.16 13.1 6.98 8.24 12.1 3.9 3.54 8.22 13.1 8.07 10.6 0.73 0.65 0.72 0.62 1.62 1.81 1.55 3.34 1.55 1.21 1.04 1.89 1.72 1.3 0.8 1.63 0.81 0.97 1.16 1.72 1.06 2.24 56.8 56.5 54.6 61.2 72.2 73.4 59.3 42.6 45.2 92.2 88.5 98 204 72 84.9 72.4 62.9 68.3 75.9 204 94.3 85.7 9.56 8.65 8.68 9.54 55.7 36.9 42.8 33.1 25.1 18.1 13.1 18.6 19 12.5 10.4 15.3 12.1 9.72 13.5 19 4.33 8.37 1.91 2.41 1.97 2.42 4.92 4.16 4.35 22.6 4.12 5.94 5.64 4.96 5.13 6.29 3.28 4.42 3.22 2.7 6.49 5.13 4.13 3.46 63.3 72.9 77.9 75.2 58.1 65.4 60 63.3 55.5 30.5 27.2 39.4 34.9 39.7 35.4 54.1 44.3 34.8 50.2 34.9 27 34.2 109 132 156 134 85.9 97.7 89.6 110 96.2 50.9 43.5 63.2 79.8 72.1 66.6 106 78.3 64.2 101 79.8 74.6 90 12.1 15.3 18.6 14.7 8.21 9.47 8.59 11.5 9.98 5.7 4.92 6.72 10.3 8.25 7.35 12.5 8.77 7.41 11.9 10.3 10.9 13 43.9 56.9 69.2 53.9 26.1 31 27.3 38.6 33.7 21.2 17.6 23.5 43.3 31.1 27.9 48.6 32.3 27.8 46.5 43.3 48.1 58.3 7.31 9.54 11.7 8.71 3.7 4.58 4.04 6.18 5.3 3.74 3.09 4.01 8.94 5.61 4.76 9 5.47 4.91 8.5 8.94 9.83 12.7 2.11 2.51 2.81 2.43 1.13 1.39 1.22 1.05 1.04 1.02 0.89 1.09 2.16 1.48 1.55 2.18 1.56 1.38 2.16 2.16 2.68 2.96 5.35 7.19 8.44 6.47 2.7 3.33 2.84 4.31 3.7 2.98 2.45 3.29 7.49 4.38 3.76 7.13 4.29 3.7 6.54 7.49 7.72 10.6 0.73 0.93 1.17 0.85 0.37 0.44 0.38 0.6 0.51 0.42 0.35 0.49 1.05 0.63 0.54 0.99 0.6 0.53 0.88 1.05 1.07 1.53 3.8 5.05 6.11 4.61 2.01 2.53 2.1 3.23 2.62 2.39 1.98 2.87 6.09 3.46 3.05 5.63 3.45 3.02 5.02 6.09 5.97 8.59 0.73 0.96 1.19 0.89 0.4 0.5 0.42 0.63 0.5 0.5 0.39 0.62 1.3 0.7 0.63 1.1 0.71 0.61 0.97 1.3 1.22 1.75 2 2.66 3.2 2.39 1.12 1.43 1.17 1.74 1.44 1.44 1.15 1.89 3.92 1.97 1.85 3.12 1.87 1.65 2.72 3.92 3.36 4.76 0.28 0.36 0.44 0.33 0.18 0.22 0.18 0.28 0.21 0.21 0.17 0.28 0.56 0.27 0.26 0.43 0.26 0.24 0.36 0.56 0.48 0.69 1.91 2.52 3.08 2.35 1.34 1.56 1.34 2.06 1.57 1.64 1.35 2.3 4.49 2.09 1.92 3.1 1.9 1.73 2.7 4.49 3.51 4.9 0.3 0.37 0.45 0.35 0.21 0.25 0.19 0.32 0.23 0.24 0.22 0.35 0.69 0.3 0.3 0.5 0.29 0.24 0.43 0.69 0.52 0.72 252.82 309.19 360.29 307.18 191.47 219.8 199.37 243.8 212.5 122.88 105.26 150.01 204.99 172.04 155.87 254.38 184.07 152.22 239.88 204.99 196.96 244.7 86 57 45 70 92 83 97 20 31 47 62 40 10 50 65 33 51 58 44 10 16 12 22.4 19.55 17.09 21.62 29.3 28.33 30.26 20.76 23.89 12.57 13.62 11.58 5.25 12.84 12.46 11.79 15.76 13.59 12.56 5.25 5.2 4.72 0.99 0.89 0.83 0.95 1.05 1.04 1.05 0.59 0.68 0.9 0.96 0.89 0.79 0.88 1.08 0.8 0.95 0.95 0.85 0.79 0.91 0.76

BMG—biotite monzogranite; HM—hornblende monzonite; QM—quartz monzonite; QD—quartz diorite; MME—microgranular enclave.

L.-X. Ding et al. / Lithos 126 (2011) 212–232

Sample no. LN8

219

220

L.-X. Ding et al. / Lithos 126 (2011) 212–232

the samples are enriched in Sr (396–1917 ppm) and Ba (1284– 4538 ppm) and depleted in Nb (15–53 ppm), Ta (0.62–3.34 ppm) and Ti (Table 2; Fig. 8). 5.3. Isotopic compositions 5.3.1. Sr–Nd isotopes The Sr–Nd isotopic data of the host rocks and MMEs are listed in Table 2. The initial 87Sr/ 86Sr ratios and εNd(t) values were calculated using an average emplacement age (220 Ma) for the four intrusive units. Depleted mantle Nd model ages (TDM) were calculated using the model of DePaolo (1981). The biotite monzogranite has a radiogenic Nd isotopic composition with εNd(t) values ranging from −9.2 to − 12.6 and initial 87Sr/ 86Sr ratios ranging from 0.7061 to 0.7067, corresponding to a TDM2 in the range of 1.7 to 2.0 Ga. The other host rocks have more negative εNd(t) values ranging from −12.4 to −17.0, and higher ( 87Sr/ 86Sr)i values ranging from 0.7065 to 0.7075, corresponding to a TDM2 in the range of 2.0 to 2.4 Ga. The MMEs have ( 87 Sr/ 86 Sr) i = 0.7065–0.7070, ε Nd (t) = − 13.1 to − 15.9, and TDM2 = 2.1–2.3 Ga. 5.3.2. Zircon Hf isotopes Haflium isotopes of zircon have recently been used as a powerful tool to constrain sources of magmatic rocks (Andersen, 2002; Nebel et al., 2007). Zircons from the four dated samples were selected for insitu Hf isotope analysis, which was done on the same domains as the U–Pb dating. The analytical results are listed in Table 3. Twenty-two spot analyses were made on sample LN49. With two exceptions from inherited zircons ( 176Hf/ 177Hf = 0.281305–0.281387; εHf(t) = −1.2 to + 1.3), they give relatively high and significantly variable 176 Hf/ 177Hf ratios ranging from 0.282078 to 0.282428, and show a single distribution in initial Hf isotope ratio with εHf(t) values ranging from − 7.6 to − 19.7 (Table 3). The weighted mean εHf(t) value is −16.7 ± 0.7, corresponding to a weighted mean two-stage Hf model age (TDM2) of 2.3 ± 0.1 Ga (Fig. 7). Sample LN28 shows a single distribution, with 176Hf/ 177Hf ratios of 0.282198 to 0.282385 and εHf(t) values of −9.0 to −15.4. Thirteen spots yielded a weighted mean εHf(t) value of −12.5 ± 1.4 corresponding to a mean TDM2 of 2.0 ± 0.2 Ga (Fig. 7). Sample LN44 also shows a single distribution, but has consistent 176Hf/ 177Hf ratios of 0.282121 to 0.282224 and more negative initial Hf isotope values of −14.7 to −18.6. With one exception (Spot 10), nine analyses yielded a weighted mean εHf(t) value of −17.0 ± 0.9, with a mean TDM2 of 2.2 ± 0.1 Ga (Fig. 7). Fifteen analyses were made on sample LN9, including six analyses from the inherited cores. Spots on the overgrowth bands have 176Hf/177Hf ratios ranging from 0.282194 to 0.282362 and display a single distribution. The calculated initial Hf isotope values range from −10.0 to −16.0 and with a weighted mean εHf(t) value of −14.1 ± 1.2, corresponding to a mean TDM2 of 2.1 ± 0.1 Ga (Fig. 7). By contrast, the spots from the inherited cores, have relatively lower 176Hf/177Hf ratios (0.281312–0.281665), higher εHf(t) values (−0.1 to +8.7), with TDM1 of 2.3 to 2.9 Ga (Table 3). 6. Discussion 6.1. Genetic classification Samples of the biotite monzogranite are characterized by high Sr (396–1398 ppm) and Ba contents (1284–3993 ppm) but low Y (mostly b18 ppm) and Yb (mostly b 1.8 ppm) concentrations, and high Sr/Y (20–97) and La/Yb (12–30) ratios (Table 2). The rocks have high SiO2 (N60 wt.%) and low MgO (b1 wt.%) contents, low Mg #s (b45), and low Cr and Ni concentrations. These geochemical signatures are consistent with high-silica adakites (HAS, Martin et al., 2005). In the Sr/Y versus Y and (La/Yb)N versus YbN diagrams, most biotite monzogranite samples fall within the adakite field (Fig. 8a–b). On the other hand, the biotite monzogranite samples plot in the high

K calc-alkaline and shoshonitic fields on the SiO2 versus K2O diagram, a feature distinctly different from slab-derived adakites that are sodic. Potassic to ultra-potassic adakite-like rocks have also been recognized elsewhere, such as the early Cretaceous high-K calc-alkaline adakitic granites in the Northern Dabie complex central China (Wang et al., 2007a), late Triassic high-K calc-alkaline adakitic monzogranite, NW margin of the South China Block (Qin et al., 2010), and Cenozoic potassic adakites in Tibet and the Andes (Chung et al., 2003; Hou et al., 2004; Kay et al., 2005). Thus, the geochemical features, together with the radiogenic Sr–Nd isotope compositions, indicate that the Laoniushan biotite monzogranite can be considered as “potassiumrich adakite” (Rapp et al., 2002). The remaining three units differ from the biotite monzogranite in that they have high Y and Yb contents. Specifically, they are characterized by low TiO2 contents, high Sr and Ba contents, K2O/Na2O and Fe2O3/FeO ratios, moderate Th/U ratios, LILE and LREE enrichment and HFSE depletion without obvious Eu anomalies. These features are compatible with typical shoshonitic rocks (Eklund et al., 1998; Jiang et al., 2002; Liégeois et al., 1998; Nardi, 1986). The hornblende monzonite, quartz monzonite, and quartz diorite all plotted in the shoshonitic field of the SiO2 vs. K2O diagram (Peccerillo and Taylor, 1976; Fig. 5). In terms of the rock series, the I- and S-type granitoids of the traditional typology are of (high K) calc-alkaline series, whereas the M-type granites are of the tholeiitic series. They are distinguished from calc-alkaline and tholeiitic series on the Ce/Yb vs. Ta/Yb and Th/Yb vs. Ta/Yb diagrams (Fig. 8c–d). In addition, albeit these rocks are K-rich and enriched in incompatible elements, they are not alkaline in character, as indicated by their agpaitic indices (A.I.= a + K/Al, molar ratio), which are less than 0.8 (not shown), much lower than those of alkali granites (A.I. = 0.87; Liégeois and Black, 1987). The ubiquitous biotite and amphibole in the rocks also preclude their classification as A-type granite. Based on these characteristics, we suggest that the Laoniushan granitoids belong to the shoshonitic series. 6.2. The nature and origin of the MMEs MMEs coexisting with felsic host rocks provide a potential window for understanding the process of magma interaction between mafic and felsic end-members (Kaygusuz and Aydinçakir, 2009) and the genesis of granitoid rocks (Barbarin and Didier, 1991; Didier, 1973). However, the origin of MMEs is still debated, and several petrogenetic hypotheses have been proposed; (1) refractory restites derived from the source region (Chappell and White, 1992); (2) cognate fragments of cumulate minerals (Chappell et al., 1987; Noyes et al., 1983); (3) interdiffusion of elements between the MMEs and enclosing melts (Didier and Barbarin, 1991); and (4) mafic liquids or hybrid magmas involving mafic-felsic magma mixing (Gerdes et al., 2000; Perugini et al., 2003). The low (La/Yb)N ratios (4.7–5.3) of the Laoniushan MMEs relative to those of the host rocks (11.6–30.3) may suggest that the MMEs represent solid residues of partial melting. However, the typical igneous textures and lower A/CNK values (0.57–0.70) relative to the host rocks (0.68–1.14) clearly point to an igneous origin. The cognate model suggests that MMEs form by crystal-liquid differentiation within a single, parental granitoid magma (Dahlquist, 2002; Fershtater and Borodina, 1991). This hypothesis could account for the similarities in mineralogy and chemical and isotopic compositions between the Laoniushan enclaves and their hosts. However, a cumulate model involving crystal settling at the bottom of the magmatic chamber is not compatible with the small grain size of most enclaves. Donaire et al. (2005) proposed a rapid cooling crystallization process, operating at the margins of magma conduits, to account for both the textural and chemical features of microgranular enclaves from the Los Pedroches granodiorite. Such a cognate model could account for the Laoniushan MMEs if the enclaves were restricted to the margins of the plutons where cooling was rapid

L.-X. Ding et al. / Lithos 126 (2011) 212–232

221

Fig. 5. Representative binary variation plots for the Laoniushan granitoids and enclosed MMEs. The K2O vs. SiO2 diagram are from Rickwood (1989). The data of GD, BM, QM and HM are from this paper, whereas the data of NG, AG and RG are from Li et al. (2004); Qin et al. (2007); Wang et al. (2008b); Yan and Zhang (2005); Zhang et al. (2005); Zhang et al. (2007a); Zhou et al. (2008); and Zhu et al. (1998). Abbreviations: GD, quartz diorite; BM, biotite monzogranite; QM, quartz monzonite; HM, hornblende monzonite; NG, normal granite; AG, adakite-like granitoid; RG, Rapakivi-textured granitoid.

because they are all systematically finer-grained than their hosts. However, in Laoniushan complex the exposed MMEs are randomly distributed in the plutons, a feature difficult to be explained by such a model. Moreover, the low differentiation indices of the MMEs (DI = 45–56) suggest that they are not highly evolved rocks. Interdiffusion of elements between the MMEs and enclosing melts has also been proposed as a mechanism generating plastic MMEs

within granitoids (Didier and Barbarin, 1991). Actually, this model is consistent with the so-called “Soret Effect”. With this effect, the total contents of Mg and Fe will increase in the mafic melt of the magma chamber, but Mg/(Mg + Fe) ratios might be constant overall. In contrast, Na/(Na + Ca) ratios of the mafic and felsic melts would be markedly different due to separation of Na and Ca. The Laoniushan MMEs are different from the host rocks both in Mg/(Mg + Fe) and

222

L.-X. Ding et al. / Lithos 126 (2011) 212–232

Table 2 Sr and Nd isotopic data for Laoniushan granitoids and MMEs. Sample no.

Rock types

Rb (ppm)

Sr (ppm)

87

Rb/86Sr

LN18 LN34 LN31 LN9 LN28 LN45 LN50 LN17 LN49 LN37 LN46

BMG Enclave BMG HM BMG Enclave BMG HM QD QM QM

66.5 131 231 161 115 155 127 139 103 98.9 98

576 1377 416 929 1575 816 582 897 1761 1289 1221

0.2752 1.6072 0.5015 0.2113 0.5496 0.6314 0.4484 0.1692 0.2220 0.2322 0.4168

87

Sr/86Sr

2σ

Isr

Sm (ppm)

Nd (ppm)

147

0.706947 0.711488 0.708218 0.707197 0.708359 0.708951 0.708064 0.707200 0.708193 0.708043 0.708154

8 8 4 3 8 15 6 8 18 7 8

0.706086 0.706459 0.706649 0.706536 0.706639 0.706975 0.706661 0.706671 0.707498 0.707316 0.706850

4.58 8.94 3.09 9.54 3.74 12.7 4.01 8.71 8.5 8.04 5.47

31 43.3 17.6 56.9 21.2 58.3 23.5 53.9 46.5 41.8 32.3

0.0893 0.1248 0.1061 0.1014 0.1067 0.1317 0.1032 0.0977 0.1105 0.1163 0.1024

Na/(Na + Ca) ratios, indicating that element interdiffusion is a less likely mechanism to form the MMEs. Additionally, the exposed MMEs are not concentrated at the margin but are randomly distributed in the plutons. Thus, interdiffusion of elements is not sufficient to account for the formation of the present MMEs.

Sm/144Nd

143

Nd/144Nd

0.511930 0.511863 0.511864 0.511861 0.511866 0.511730 0.512031 0.511862 0.511643 0.511678 0.511632

2σ

εNd(t)

TDM2 (Ga)

3 2 5 7 4 3 4 3 3 2 3

− 10.8 − 13.1 − 12.6 − 12.5 − 12.5 − 15.9 − 9.2 − 12.4 − 17.0 − 16.5 − 17.0

1.9 2.1 2.0 2.0 2.0 2.3 1.7 2.0 2.4 2.3 2.4

Compared with their enclosing granitic rocks, the MMEs have relatively higher Mg-numbers (51–56) and lower SiO2 contents (48–53%). In principle, there are three main possibilities to account for these characteristics: (1) melting of mafic lower crust; (2) melting of upper mantle lithosphere previously enriched by subduction, and

Fig. 6. Chondrite-normalized REE pattern and Primitive mantle normalized trace-element spiderdiagrams for the Laoniushan shoshonitic granitoids (a–f). Normalizing values for REE and trace elements are from Taylor and McLennan (1985) and Sun and McDonough (1989), respectively. The reference values of lower continental crust and bulk continental crust are from Rudnick and Gao (2003). Symbols as in Fig. 5.

L.-X. Ding et al. / Lithos 126 (2011) 212–232 Table 3 Hf-isotope data from zircons of Laoniushan granitoids. Spot no.

176

176

176

177

177

177

Yb/ Hf

Lu/ Hf

Hf/ Hf

2σ

Age (Ma)

εHf(t)

TDM1 (Ma)

TDM2 (Ma)

LN49 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22

quartz diorite 0.039027 0.001351 0.093698 0.003201 0.026814 0.001068 0.120650 0.004260 0.032980 0.001287 0.033106 0.001311 0.030338 0.001166 0.015485 0.000566 0.029156 0.001006 0.041349 0.001424 0.042763 0.001625 0.045758 0.001715 0.041273 0.001423 0.032709 0.001070 0.054080 0.002058 0.044880 0.001715 0.048999 0.001826 0.044744 0.001528 0.043968 0.001598 0.044852 0.001450 0.022817 0.000861 0.015377 0.000545

0.282123 0.282428 0.282184 0.282206 0.282178 0.282177 0.282195 0.282154 0.282078 0.282154 0.282124 0.282203 0.282104 0.282170 0.282167 0.282165 0.282181 0.282171 0.282152 0.282138 0.281305 0.281387

0.000019 0.000016 0.000015 0.000015 0.000016 0.000014 0.000015 0.000015 0.000019 0.000017 0.000016 0.000014 0.000016 0.000016 0.000018 0.000015 0.000016 0.000016 0.000014 0.000015 0.000016 0.000017

231 229 228 228 228 239 242 227 227 228 226 235 227 226 228 226 235 249 227 226 2325 2283

− 18.1 − 7.6 − 15.9 − 15.7 − 16.2 − 16 − 15.3 − 17 − 19.7 − 17.1 − 18.2 − 15.2 − 18.9 − 16.5 − 16.7 − 16.8 − 16.0 − 16.0 − 17.2 − 17.7 − 1.2 1.3

1607 1238 1509 1615 1527 1530 1498 1532 1655 1566 1617 1509 1636 1529 1575 1562 1545 1546 1577 1589 2709 2576

2400 1741 2263 2243 2279 2276 2231 2327 2500 2333 2403 2224 2444 2296 2311 2312 2274 2285 2341 2370 2938 2755

LN28 1 2 3 4 5 6 7 9 10 11 12 13

biotite monzogranite 0.036481 0.001178 0.039975 0.001407 0.039761 0.001455 0.035315 0.001143 0.043965 0.001387 0.030085 0.001013 0.033550 0.001096 0.051869 0.001707 0.045644 0.001611 0.027831 0.000962 0.043530 0.001387 0.038764 0.001338

0.282248 0.282368 0.282321 0.282300 0.282287 0.282198 0.282207 0.282221 0.282326 0.282224 0.282263 0.282385

0.000015 0.000011 0.000014 0.000016 0.000014 0.000015 0.000014 0.000015 0.000012 0.000013 0.000014 0.000014

230 227 227 214 213 229 229 228 230 228 228 223

− 13.7 − 9.5 − 11.2 − 12.2 − 12.7 − 15.4 − 15.1 − 14.7 − 11.0 − 14.5 − 13.2 − 9.0

1424 1264 1331 1350 1377 1487 1478 1483 1330 1450 1412 1238

2121 1859 1963 2016 2047 2231 2212 2187 1953 2174 2092 1823

LN44 1 2 3 4 5 6 7 8 9 10

quartz monzonite 0.053868 0.001571 0.049142 0.001715 0.050575 0.001547 0.040741 0.001284 0.047465 0.001416 0.061922 0.002193 0.036990 0.001173 0.039185 0.001337 0.032346 0.000931 0.051553 0.001790

0.282224 0.282139 0.282188 0.282196 0.282121 0.282135 0.282139 0.282148 0.282180 0.281542

0.000028 0.000015 0.000013 0.000015 0.000017 0.000024 0.000017 0.000014 0.000016 0.000024

224 219 217 217 209 217 218 217 217 2232

− 14.7 − 17.8 − 16.1 − 15.8 − 18.6 − 18.1 − 17.8 − 17.5 − 16.3 3.7

1473 1600 1523 1501 1612 1626 1576 1571 1509 2445

2181 2375 2266 2244 2417 2388 2370 2353 2278 2563

LN9 hornblende monzonite 1 0.022114 0.000861 2 0.019879 0.000734 3 0.034162 0.001274 4 0.030174 0.001113 5 0.059417 0.002053 6 0.029303 0.001117 7 0.021558 0.000818 8 0.016850 0.000602 9 0.015054 0.000568 10 0.044147 0.001482 11 0.031891 0.001173 12 0.022498 0.000765 13 0.019989 0.000709 14 0.041564 0.001428 15 0.020926 0.000719

0.282194 0.282301 0.282362 0.282203 0.282255 0.282251 0.282267 0.282197 0.282234 0.281665 0.281610 0.281565 0.281312 0.281393 0.281562

0.000022 0.000027 0.000029 0.000020 0.000026 0.000019 0.000020 0.000016 0.000015 0.000027 0.000025 0.000028 0.000019 0.000018 0.000018

207 231 214 248 217 217 218 218 217 2237 2154 2156 2353 2460 2169

− 16 − 11.7 − 10 − 14.9 − 13.8 − 13.8 − 13.2 − 15.6 − 14.3 8.7 5.3 4.4 − 0.1 4 4.7

1487 1334 1267 1485 1448 1418 1384 1473 1421 2254 2312 2349 2689 2628 2350

2252 1999 1877 2210 2122 2122 2083 2237 2155 2264 2405 2464 2892 2725 2459

(3) melting of mantle lithosphere followed by crustal contamination. Experimental studies (Helz, 1976; Rapp et al., 1991; Wolf and Wyllie, 1994) and theoretical considerations (Roberts and Clemens, 1993) suggest that dehydrating basaltic lower crust can produce large volumes of mafic melts, particularly in regions with high heat flow that facilitates higher degree partial melting (40–60%) to form mafic

223

liquids of basaltic and basaltic andesite compositions (Rapp and Watson, 1995). However, these mafic liquids are mostly marked by low MgO and high Al2O3, which are not the case of our MMEs. MgO concentrations of the MMEs are significantly higher than the values of experimental melts derived by melting low K-basaltic, alkali basaltic, and high-Al basaltic sources (Fig. 9a). Although Jung et al. (2002) showed that melting mafic lower crust could generate relatively highMgO (b5%) dioritic to basaltic melts, the extremely high MgO (N6%) and K2O/Na2O ratios (N3) of the Laoniushan MMEs (Fig. 9) do not favor this possibility. Furthermore, partial melts of metabasaltic rocks at highest pressure should have lowest Y and HREE but highest Al2O3 and Sr, corresponding to the highest Sr/Y ratios. Nevertheless, the MME samples with lowest Y and highest Sr/Y ratio have the lowest Al2O3 content. All these observations indicate that the MMEs are unlikely to originate from a mafic lower crust. Instead, high Nb/Ta (18–24) ratios indicate that they were most likely derived from a mantle source, as partly confirmed by the Ni vs. Cr binary variation diagram (Fig. 10; Tsuchiya et al., 2005). The MMEs have evolved Nd isotopic compositions [εNd(t) = −13.1 to −15.9]. The negative εNd(t) values could be acquired by melting either of an enriched or a depleted mantle source, followed by significant contamination by crustal materials during magma evolution. In order to further identify the nature of the mantle source of the MMEs, we conducted an isotopic simulation using the mixing model of Langmuir et al. (1978). The depleted mantle (Nd = 19 ppm, εNd = +9, Sr = 190 ppm, ISr = 0.703; Zindler et al., 1984) and the lower continental crust (Nd = 38 ppm, εNd = − 28, Sr = 300 ppm, and ISr = 0.709; Jahn et al., 1987) are considered as end-members for the modeling. The calculated result indicates that the MMEs would require 30–40% input from continental crustal magma if they originated from a depleted mantle. However, such a high degree of mixing or assimilation would significantly modify the major element composition of the MMEs, which is not observed in our samples. The Sr–Nd isotope compositions of the MMEs ( 87Sr/ 86Sr = 0.7090–0.7115, 143 Nd/ 144Nd = 0.5117–0.5119) are clearly different from those of the depleted mantle ( 87Sr/ 86Sr = 0.7015–0.7025; 143Nd/ 144Nd = 0.5133– 0.5136; Rollinson, 1993), providing additional argument against melting of a deplete mantle source for their genesis. The evolved Sr–Nd isotopic features, in combination with enrichment of LILE (e.g. Rb, Ba, Sr and K) and depletion of HFSEs (e.g. Nb, Ta, Ti), favor the derivation of the Laoniushan MMEs from partial melting of ancient enriched subcontinental lithospheric mantle. 6.3. Magma mixing/mingling and origin of the shoshonitic granitoids In terms of present petrological, geochemical, and isotopic data, we propose that generation of the Laoniushan granitoids and MMEs therein resulted from mixing/mingling of crustal- and mantle-derived magmas. Several lines of evidence support this interpretation, which are addressed below. The MMEs are randomly scattered in the different rock units and have rounded to ovoid shapes, with no evidence of solid-state deformation. In most cases, the contacts between the enclaves and host rocks are sharp (Fig. 11a), although a few transitional boundaries are also observed (Fig. 11b), reflecting hybridization of mafic and felsic magmas. K-feldspar megacrysts occurring in host granitoids sometimes cross-cut the contacts between the host rocks and enclaves, and even enter the MMEs. In addition, the K-feldspar megacrysts in the host rocks are generally euhedral to subhedral, but tend to be round in the MMEs (Fig. 11c and d). The systematic variations in the size and morphology of the K-feldspar megacrysts in MMEs and hosts reflect mechanical transfer of mineral grains from a felsic magma into a mafic magma during magma mixing (Vernon et al., 1988). Furthermore, plagioclase phenocrysts in the quartz monzonite show marked punctuated zoning (Fig. 2b and c), which can form by infiltration of melt into partially crystalline felsic materials

224

L.-X. Ding et al. / Lithos 126 (2011) 212–232

Fig. 7. Histograms of zircon εHf(t) and TDM2 for the granitoids of the Laoniushan complex.

(Campos et al., 2002; Silva et al., 2000). Abundance of acicular apatite in the MMEs further confirms magma mixing as an important process in the formation of the MMEs and their host rocks (Barbarin and Didier, 1992). On most variation diagrams of element abundance versus SiO2 (Fig. 5), the MMEs and host rocks form a roughly continuous trend without a chemical gap, consistent with a mixing trend between two

end-members (Castro, 1990; Perugini et al., 2003). The two also display typical mixing trends on the diagrams of (Fe2O3 + FeO)/SiO2 vs. K2O/CaO and Na2O/CaO vs. Al2O3/CaO (not shown; Neves and Vauchez, 1995; Zhou, 1994). The shoshonitic granitoids from the Laoniushan complex have lower SiO2 (56–63 wt.%), higher MgO (1.62–4.03 wt.%), Cr (31.4–72.7 ppm)

L.-X. Ding et al. / Lithos 126 (2011) 212–232

225

Fig. 8. Plots of (a) Sr/Y vs. Y and (b) (La/Yb)N vs. YbN for the Laoniushan adakitic biotite monzogranite (modified after Defant and Drummond, 1990). Adakite fields from Martin (1999). (c) Ce/Yb vs. Ta/Yb (Pearce, 1982) and (d) Th/Yb vs. Th/Yb diagrams (Müller and Groves, 1995) for the Laoniushan shoshonitic granitoids. Symbols as in Fig. 5.

and Ni (6.88–23.5 ppm) contents, and more radiogenic Sr–Nd–Hf isotopic compositions. Available data have shown that a majority of shoshonitic granitoids are produced by partial melting of continental lithospheric mantle (Miller et al., 1999; Turner et al., 1996; Xie et al., 2006). However, the existence of inherited zircons in these granitoids strongly indicates that the granitoids originated from partial melting of an ancient lower continental crust (Fig. 3). Furthermore, they plot in the field of metabasaltic to metatonalitic sources on the diagram of Al2O 3 /(MgO + FeO*) versus CaO/(MgO + FeO*)(not shown), supporting the hypothesis. Nevertheless, Fig. 10 shows that the shoshonitic granitoids have Mg #s and K2O/Na2O ratios significantly higher than the values of experimental melts of basaltic materials (Sen and Dunn, 1994; Rapp and Watson, 1995) and the high-Mg quartz diorites produced by melting mafic lower crust in the Damara Orogen of Namibia (Jung et al., 2002). This indicates that they are unlikely to be produced merely by partial melting of lower continental crust. Instead, mafic, mantle-derived components must have been involved in their formation. This view is partly supported by the wide range of zircon Hf isotope compositions (Table 3; Fig. 7). It is suggested that the Laoniushan potassic granitoids may have resulted from binary hybridization between a potassium-rich mafic end-member and a felsic end-member (Clemens et al., 2009, and references therein). The inflections in plots of K2O, Ba and Sr vs. SiO2 (Fig. 5) indicate that some fractional crystallization was involved. It has been shown that Ba has relatively high distribution coefficients in biotite and K-feldspar (Gan, 1993). As shown in Fig. 12, the diagrams of Rb/Sr vs. Sr and Ba vs. Rb display a trend consistent with fractional crystallization of potassium feldspar, respectively, confirming the presence of fractional crystallization during magma evolution. However, the MgO versus FeO Total diagram (Fig. 13a) further demonstrates that magma mixing, rather than fractional crystallization, was a major mechanism in the formation of the shoshonitic granitoids. Lastly, the relationship between MgO and Ni (Fig. 13b) for the Laoniushan rocks mimics the late Triassic Mishuling pluton in the

Qinling Orogen, which has been demonstrated to be a product of magma mixing (Qin et al., 2009). 6.4. Origin of the adakitic biotite monzogranite The origin of adakites has been hotly debated in the last two decades (e.g. Moyen, 2009; Richards and Kerrich, 2007). The proposed models mainly include (1) melting of subducted oceanic crust (Defant and Drummond, 1990; Martin, 1999); (2) partial melting of peridotitic mantle metasomatized by slab melts (Martin et al., 2005); (3) fractionational crystallization of mantle-derived basalts (Castillo et al., 1999; Li et al., 2009); (4) melting of thickened (Wang et al., 2007b; Xu et al., 2007) or delaminated basaltic lower continental crust (Gao et al., 2004). The adakitic biotite monzogranite in the Laoniushan complex has relatively high SiO2 (65–71 wt.%), lower MgO (0.21–0.86 wt.%), Cr (1.52–20.8 ppm) and Ni (0.45–6.96 ppm), and evolved Sr–Nd isotopic compositions with (87Sr/ 86Sr)i of 0.7061 to 0.7067 and εNd(t) of −9.2 to −12.6, showing continental crust signatures. In addition, their high K2O/Na2O ratios, negative Nb and Ta anomalies, and negative εNd(t) values are distinctive from oceanic crust-derived adakitic melts (Defant and Drummond, 1990). Thus, melting of subducted oceanic crust and mantle peridotite can be ruled out. As mentioned before, fractional crystallization processes were dominant in the Laoniushan magmatic system, although they could have occurred locally. Therefore, the model (3) can be ruled out as well. Furthermore, the Laoniushan adakitic biotite monzogranite is also unlikely to have formed by partial melting of ancient mafic lower continental crust (eclogite) that foundered into the convecting mantle. Commonly, the adakites formed by this model have characteristically high Mg#, Cr and Ni (Gao et al., 2004; Zhang et al., 2010), which are distinctly different from the Laoniushan adakitic rocks. In the SiO2 versus MgO diagram, the biotite monzogranites also deviate from the field of adakitic rocks derived from partial melting of delaminated lower crust (Fig. 14). Crustal thickening most likely occurred during the Triassic collision between North China and Yangtze

226

L.-X. Ding et al. / Lithos 126 (2011) 212–232

Zircons from the biotite monzogranite have two-stage Hf model ages of 1.8 to 2.2 Ga, which are consistent with the whole-rock, twostage Nd model ages (TDM2) of 1.7 to 2.0 Ga (Table 2). These ages are in broad agreement with the U–Pb ages of inherited zircons from the biotite monzogranite, ranging from 2.0 to 2.5 Ga. The implication is that the biotite monzogranite was generated by melting the ancient crustal rocks. The Nd and Hf model ages are close to that of the Taihua Group metamorphic rocks in the Xiaoqinling area (2.26–2.14 Ga; Wan et al., 2006;) implying that the Taihua Group could be a possible source for the biotite monzogranite. However, the biotite monzogranite shows more continental isotopic signatures than the Taihua Group metamorphic rocks (ISr = 0.703; εNd(t) = +7.1, unpublished data), indicating that the Taihua Group cannot be the source of the Laoniushan adakitic rocks. In the ( 87Sr/ 86Sr)i vs. εNd(t) diagram (Fig. 15), the εNd(t) values for the Laoniushan biotite monzogranite samples overlap the Late Mesozoic granitoids from the NCC (Hong et al., 2003, and references therein), but are distinctly different from the Early Mesozoic granitoids in the Qinling Orogen (Wang et al., 2007a; Wang et al., 2008a; Zhang et al., 2005; Zhang et al., 2007b; Zhu et al., 1998). This indicates that the Laoniushan biotite monzogranite may share similar source with the Late Mesozoic granitoids in the NCC. As illustrated in Fig. 15, they lie between the EM1-type enriched mantle and lower continental crust (LCC), but far from the upper continental crust (UCC), confirming their derivation from ancient lower crustal materials. 6.5. Timing of magmatism and tectonic implications

Fig. 9. (a) Plot of SiO2 vs. Mg# (Mg# = molar Mg/(Mg + Fe)); (b) Plot of SiO2 vs. K2O/ Na2O. The shadow field represents the experimental melts of basalts (Rapp and Watson, 1995; Sen and Dunn, 1994) and quartz diorites from melting of mafic lower crust (Jung et al., 2002).The symbols of the Laoniushan samples as in Fig. 5.

Cratons along the Mianlue suture (Fig. 1b). Thus, having ruled out the possibilities stated above, we suggest that the Laoniushan adakitic biotite monzogranite was produced by partial melting of thickened lower continental crust.

Fig. 10. Ni vs Cr binary variation diagram for the enclaves. The mantle-derived melts are from Tsuchiya et al. (2005).

Previous geochronological studies yielded K–Ar and whole-rock Rb–Sr isochron ages ranging from 428 to 95 Ma (Nie and Fan, 1989; Zhang et al., 2007a; Zhou et al., 1987; Zhu, 1995). The large spread of these ages suggests that they cannot provide precise constraint on the timing of magmatism. More recently, Zhu et al. (2008) reported a LAICP-MS zircon U–Pb age of 146.4 ± 0.6 Ma for a biotite monzogranite sample (E 109°57′48.3″; N 34°20′15.8″) from the eastern margin of the Laoniushan complex. This age led them to suggest that the monzogranite formed during the Late Mesozoic intracontinental extension that has been commonly invoked for the whole of eastern China. Zircons from our samples are generally characterized by spectacular oscillatory zoning, and have relatively high Th/U ratios and REE contents, strong negative Eu anomalies, as well as enrichment in HREE, indicative of a magmatic origin (Corfu et al., 2003). The U–Pb ages of these samples therefore demonstrate that the Laoniushan complex formed mainly by two magmatic pulses at 228–227 Ma and 217–215 Ma in the Late Triassic. This indicates that the complex was linked to the Triassic orogeny of the Qinling Orogen, and was unrelated to the Late Mesozoic intracontinental extension widely recognized in North China Craton and the whole of eastern China (Zhu et al., 2008). Actually, the monzogranite reported by Zhu et al.(2008) and the studied monzogranite samples in this paper have distinctly different CL images of zircons and geochemical charateristics, indicating that they could be from different intrusions. Thus, the U– Pb age reported by Zhu et al.(2008) may represent a rock unit emplaced in the Late Mesozoic within the Laoniushan complex or some other younger body. The adakitic biotite monzogranite is temporally equivalent to many adakite-like intrusions in the South Qinling terrane emplaced in the 233–215 Ma interval (Qin et al., 2005; Wang et al., 2008a; Wu et al., 2009), whereas the shoshonitic granitoids are coeval with the rapakivi-textured granitoids (217–200 Ma; Gong et al., 2009; Lu et al., 1999) and mafic dikes (Wang et al., 2007b) in the western Qinling Orogen. Early Mesozoic granitoids are widespread in the Qinling Orogen, however, their petrogenetic processes and tectonic implications have long been debated. Zhang et al.(1994) and Sun et al.(2000) contended

L.-X. Ding et al. / Lithos 126 (2011) 212–232

227

Fig. 11. Photographs showing various hybrid textures in the Laoniushan complex. (a) Mafic enclaves having sharp contact with the host rocks; (b) Transitional contact between the host intrusion and mafic enclave. Note that feldspar crystals enclosed in the enclave have been corroded into fragments; (c) Mafic enclaves in biotite monzogranite. K-feldspar megacrysts in the enclave are similar to those in the host, with some megacrysts cross-cutting the contact between the two or enter into the enclaves; (d) A fine-grained mafic enclave enclosed in the biotite monzogranite. Note that the K-feldspar xenocryst is corroded and rounded.

that the widespread granitic magmatism in the Qinling Orogen were formed in a syn-collisional setting; Zhang et al.(2001) considered that the Triassic granitoids were caused by the northward subduction of the Paleo-Tethys oceanic crust; Jiang et al. (2010) argued that the early granitic plutons in the Qinling Orogen were emplaced in a continental arc environment coupled with the northward subduction of the Paleo-Tethyan oceanic crust, whereas the later plutons were emplaced during continental collision between the South Qinling terrane and Yangtze Craton; Zhang et al.(2008) and Qin et al.(2009, 2010) suggested that the granitic magmatism was formed in a postcollisional setting. The recognition of Early Mesozoic adakitic granitoids in the Qinling Orogen is of importance to constrain the evolution of the orogen. Commonly, adakitic magmatism could occur in the following tectonic

settings: (1) island arc (Defant and Drummond, 1990); (2) active continental margin (Feeley and Davidson, 1995; Kay and Kay, 2002); (3) post-collisional setting (Qin et al., 2008; 2010; Zhang et al., 2008); and (4) intracontinent setting (Chung et al., 2003; Stevenson et al., 2005; Zhang et al., 2010). In terms of previous researches, the Early Mesozoic adakitic rocks exposed in the western Qinling Orogen are mostly high-K calc-alkaline, with K2O/Na2O ≤1. Compared with typical adakitic rocks, they also have relatively low Mg, Cr and Ni contents. The Laoniushan adakitic biotite monzogranites are high-K calc-alkaline to shoshonitic, with K2O/Na2O ≥1, belonging to potassic adakites. Experimental studies have shown that potassic adakites could be formed by low degrees of partial melting of a source rich in water and LILE under extremely high pressure (Rapp et al., 2002). This means that the Laoniushan adakitic biotite monzogranites are likely

Fig. 12. (a) Diagram of Rb/Sr vs. Sr; (b) Diagram of Ba vs. Rb. Symbols as in Fig. 5.

228

L.-X. Ding et al. / Lithos 126 (2011) 212–232

Fig. 13. (a) Diagram of FeOt–MgO (after Zorpi et al., 1989) for the Laoniushan MMEs and their host granitoids; (b) MgO versus Ni. The dark shadow field represents the Mishuling pluton (Qin et al., 2009), which is formed by magma mixing.

related to thickened continental crust, which is consistent with the foregoing conclusion. Thickening of continental crust commonly could have been caused by continent–continent collision or slab subduction in the active continental margin setting. However, there is no evidence for a Mesozoic back-arc setting in the north Qinling during the Mesozoic (Meng and Zhang, 2000). Therefore, the active continental margin setting is less likely. A regional unconformity in the Lower Triassic (ca. 259–245 Ma) was recently identified along the northern margin of the Yangtze Block, south of the Qinling orogenic belt (Shen et al., 2010), indicating that the North China and Yangtze Cratons were amalgamated in the Early Triassic. Thus, we propose that the Laoniushan adakitic rocks, together with the Early Mesozoic adakitic granitoids in the western Qinling Orogen, were formed in a post-collisional setting. The occurrence of the Laoniushan shoshonitic granitoids further supports this hypothesis and indicates an extensional tectonic setting at ~210 Ma. Based on the discussions above, we propose the following scenario to explain the tectonic evolution of the Qinling Orogen and generation of the Laoniushan intrusive complex (Fig. 16). During the late

Fig. 14. (a) MgO vs. SiO2 diagrams of Laoniushan biotite monzogranite. The experimental melts of metabasalt or eclogite at pressures of 1.0–4.0 GPa are from Rapp and Watson (1995), Rapp et al. (1991), Sen and Dunn (1994), Skjerlie and Patiño Doudce (2002), and Springer and Seck (1997); adakitic rocks derived from partial melting of lower crust are from Atherton and Petford (1993), Johnson et al. (1997), Muir et al. (1995), and Petford and Atherton (1996); adakites derived from slab melting are from Defant and Drummond (1993), Drummond et al. (1996), Kay and Kay (1993), Martin (1999), Sajona et al. (2000), Stern and Kilian (1996), and Yogodzinski et al. (2001); adakitic rocks derived from partial melting of delaminated lower crust are from Xu et al., 2002.

Paleozoic to early Triassic, Paleo-Tethyan oceanic crust was subducted beneath the South Qinling terrane at a low angle. The Paleo-Tethyan ocean closed, and continental collision between the Yangtze and North China Cratons along the Qinling–Dabie orogen caused the Yangtze continental lithosphere to be subducted beneath the North China Craton in the early to middle Triassic. During and prior to the collision, no magmatism occurred. Afterwards, slab break-off caused rising of the asthenosphere, which led to partial melting of the relic continental slab to form the adakite-like intrusions mostly along the Mianlue subduction zone. The upwelling asthenosphere also provided sufficient heat energy to melt the enriched lithospheric mantle, resulting in voluminous mafic magmas. Underplating of the mafic magmas induced partial melting of thickened lower crust, and led to the formation of voluminous felsic magmas. The felsic magmas were emplaced in the hinterland of the orogen, far from the subduction zone, and formed the Laoniushan adakitic biotite monzogranite. Subsequently, further lithospheric extension and asthenospheric upwelling triggered partial melting of more mafic materials of the lower crust. The resultant felsic magmas mixed with mafic magmas derived from enriched lithospheric mantle to form the Laoniushan

Fig. 15. Initial 87Sr/86Sr versus εNd(t) diagram for the Laoniushan granitoids and MMEs (t = 220 Ma), with comparison to the Early Mesozoic granitoids of the Qinling Orogen (t = 220 Ma) (Wang et al., 2007b; Wang et al., 2008b; Zhang et al., 2005, 2007a; Zhu et al., 1998) and Mesozoic granitoids from the North China Craton (Hong et al., 2003). Upper continental crust (UCC) and lower continental crust (LCC) are from Jahn et al. (1988) and Wang et al. (2007a), respectively. Symbols as in Fig. 5.

L.-X. Ding et al. / Lithos 126 (2011) 212–232

229

Acknowledgments This study was supported by the National Nature Science Foundation of China (grants 90814004, 40334037, and 40821061), China Geological Survey Bureau (grants 11212010918002), the Fundamental Research Funds for the Central Universities (CUG090102), and the MOST special fund from the State Key Laboratory of Geological Processes and Mineral Resources (MSFGPMR201005). We thank Tim Kusky, Jinyang Zhang and Zhenbing She whose comments helped to greatly improve the manuscript. We are also grateful to Profs. Yongsheng Liu, Lian Zhou and Yueheng Yang for assistance with the analytical work.

Appendix A. Supplementary data Supplementary data to this article can be found online at doi:10. 1016/j.lithos.2011.07.008.

References

Fig. 16. Cartoons of tectonic scenarios showing the tectonic evolution of the Qinling Orogen and generation of the Laoniushan intrusive complex. (a) Northward subduction of the Paleo-Tethyan oceanic crust beneath the South Qinling terrane at a low angle during the late Paleozoic to early Triassic times. (b) Continental collision between the Yangtze and North China Cratons after the closure of the Paleo-Tethyan ocean during the early to middle Triassic; (c) Slab break-off caused upwelling of the asthenosphere in the late Triassic and partial melting of enriched lithospheric mantle and lower crust, contributing to the formation of voluminous granitoids in the Qinling Orogen, including the Laoniushan complex. More details are given in the text.

shoshonitic granitoids, together with rapakivi-textured granitoids and mafic dikes in the Qinling Orogen. 7. Conclusions (1) The Laoniushan complex consists of four intrusive units, namely the biotite monzogranite, quartz diorite, quartz monzonite, and hornblende monzonite. The biotite monzogranite is similar to potassium-rich adakite, whereas the other three units have a shoshonitic affinity. LA-ICP-MS zircon U–Pb dating indicates that the adakitic intrusion formed at ca. 228 Ma, whereas the shoshonitic intrusions were mainly emplaced between 217 and 215 Ma. (2) Both the adakitic and shoshonitic rocks contain abundant MMEs. The host rocks were produced mainly by partial melting of ancient thickened lower continental crust of the North China Craton, but the shoshonitic rocks were derived from more mafic source than the adakitic equivalents. The MMEs were mainly derived from an ancient enriched lithospheric mantle. Magma mixing played an important role in the generation of the Laoniushan complex. (3) The petrological, geochronological, and geochemical data, when combined with regional data, document that the Laoniushan complex was generated in a post-collisional setting following slab break-off. The identification of the Laoniushan complex confirms the existence of Early Mesozoic magmatism along the southern margin of the North China Craton, and indicates that post-collisional extension in the Qinling Orogen began at about 230 Ma.

Andersen, T., 2002. Correction of common lead in U–Pb analyses that do not report 204 Pb. Chemical Geology 192, 59–79. Ashley, P.M., Cook, N.D.J., Hill, R.L., 1994. Shochonitic lamprophyre dykes and their relation to mesothermal Au–Sb veins at Hillgrove, New South Wales, Australia. Lithos 32, 249–272. Atherton, M.P., Petford, N., 1993. Generation of sodium-rich magmas from newly underplated basaltic crust. Nature 362, 144–146. Barbarin, B., Didier, J., 1991. Conclusions: enclaves and granite petrology. In: Didier, J., Barbarin, B. (Eds.), Enclaves and Granite Petrology. Elsevier, Amsterdam, pp. 545–549. Barbarin, B., Didier, J., 1992. Genesis and evolution of mafic microgranular enclaves through various types of interaction between coexisting felsic and mafic magmas. Transactions of the Royal Society of Edinburgh, Earth Sciences 83, 145–153. Blichert-Toft, J., Albarede, F., 1997. The Lu–Hf geochemistry of chondrites and the evolution of the mantle-crust system. Earth and Planetary Science Letters 148, 243–258. Campos, T.C.F., Neiva, A.M.R., Nardi, L.V.S., 2002. Geochemistry of the Rio Espinharas hybrid complex, northeastern Brazil. Lithos 64, 131–153. Castillo, P.R., Janney, P.E., Solidum, R.U., 1999. Petrology and geochemistry of Camiguin Island, southern Philippines: insights to the source of adakites and other lavas in a complex arc setting. Contributions to Mineralogy and Petrology 134, 33–51. Castillo, P.R., Solidum, R.U., Punongbayan, R.S., 2002. Origin of high field strength element enrichment in the Sulu Arc, southern Philippines, revisited. Geology 30, 707–710. Castro, A., 1990. Microgranular enclaves of the Quintana granodiorite (Los Pedroches batholith). Petrogenetic significance. Revista Sociedad Geologica de España 3, 7–21. Chappell, B.W., White, A.J.R., 1992. I- and S-type granites in the Lachlan Fold Belt. Transactions of the Royal Society of Edinburgh, Earth Sciences 83, 1–26. Chappell, B.W., White, A.J.R., Wyborn, D., 1987. The importance of residual source material (restite) in granite petrogenesis. Journal of Petrology 28, 1111–1138. Chung, S.L., Liu, D.Y., Ji, J.Q., Chu, M.F., Lee, H.Y., Wen, D.J., Lo, C.H., Lee, T.Y., Qian, Q., Zhang, Q., 2003. Adakites from continental collision zones: melting of thickened lower crust beneath southern Tibet. Geology 31, 1021–1024. Clemens, J.D., Darbyshire, D.P.F., Flinders, J., 2009. Sources of post-orogenic calcalkaline magmas: the Arrochar and Garabal Hill–Glen Fyne complexes, Scotland. Lithos 112, 524–542. Corfu, F., Hanchar, J.M., Hoskin, P.W.O., Kinny, P., 2003. Atlas of zircon textures. Reviews in Mineralogy and Geochemistry 53, 469–495. Dahlquist, J.A., 2002. Mafic microgranular enclaves: early segregation from metaluminous magma (Sierra de Chepes), Pampean Ranges, NW Argentina. Journal of South American Earth Sciences 15, 643–655. Defant, M.J., Drummond, M.S., 1990. Derivation of some modern arc magmas by melting of young subducted lithosphere. Nature 347, 662–665. Defant, M.J., Drummond, M.S., 1993. Mount St. Helens: potential example of the partial melting of the subducted lithosphere in a volcanic arc. Geology 21, 547–550. DePaolo, D.J., 1981. Neodymium isotopes in the Colorado Front Range and crust-mantle evolution in the Proterozoic. Nature 291, 193–196. Didier, J., 1973. Granites and their enclaves. Developments in Petrology. Elsevier, Amsterdam. 393 pp. Didier, J., Barbarin, B., 1991. The different types of enclaves in granites—nomenclature. In: Didier, J., Barbarin, B. (Eds.), Enclaves and Granite Petrology. Elsevier, Amsterdam, pp. 19–23. Donaire, T., Pascual, E., Pin, C., Duthou, J.L., 2005. Micro-granular enclaves as evidence of rapid cooling in granitoid rocks: the case of the Los Pedroches granodiorite, Iberian Massif, Spain. Contributions to Mineralogy and Petrology 149, 247–265. Dostal, J., Caby, R., Keppie, J.D., Maza, M., 2002. Neoproterozoic magmatism in Southwestern Algeria (Sebkha el Melah inlier): a northerly extension of the TransSaharan orogen. Journal of African Earth Science 35, 213–225.

230

L.-X. Ding et al. / Lithos 126 (2011) 212–232

Drummond, M.S., Defant, M.J., Kepezhinskas, P.K., 1996. The petrogenesis of slab derived trondhjemite–tonalite–dacite/adakite magmas. Transactions of the Royal Society of Edinburgh, Earth Science 87, 205–216. Duchesne, J.C., Berza, T., Liégeois, J.P., Vander Auwera, J., 1998. Shoshonitic liquid line of descent from diorite to granite: the Late Precambrian post-collisional Tismana pluton (South Carpathians, Romania). Lithos 45, 281–303. Eklund, O., Konopeklo, D., Rutanen, H., Fröjdö, S., Shebanov, A.D., 1998. 1.8 Ga Svecofennian post-collisional shoshonitic magmatism in the Fennoscandian shield. Lithos 45, 87–108. Feeley, T.C., Davidson, J.P., 1995. Petrology of calc-alkaline lavas at Volcan Ollague and the origin of compositional diversity at Central Andean Stratovolcanoes. Journal of Petrology 35, 1295–1340. Fershtater, G.B., Borodina, N.S., 1991. Enclaves in the Hercynian granitoids of the Ural Mountains, USSR. In: Didier, J., Barbarin, B. (Eds.), Enclaves and Granite Petrology. Elsevier, Amsterdam, pp. 83–94. Fowler, M.B., Henney, P.J., 1996. Mixed Caledonian appinite magmas: implications for lamprophyre fractionation and high Ba–Sr granite genesis. Contributions to Mineralogy and Petrology 126, 199–215. Gan, G.L., 1993. Mineral-melt element partition coefficients: data and major variation regularities. Acta Petrologica et Mineralogica 12, 144–180 (in Chinese). Gao, S., Rudnick, R.L., Yuan, H.L., Liu, X.M., Liu, Y.S., Xu, W.L., Ling, W.L., Ayers, J., Wang, X.C., Wang, Q.H., 2004. Recycling lower continental crust in the North China Craton. Nature 432, 892–897. Gerdes, A., Wörner, G., Finger, F., 2000. Hybrids, magma mixing and enriched mantle melts in post-collisional Variscan granitoids: the Rastenberg Pluton, Austria. In: Franke, W., Haak, V., Oncken, O., Taner, D. (Eds.), Orogenic Process: Quantification and Modellig in the Variscan Belt, 179. Geological Society of London, Special Publications, pp. 415–431. Gong, H.J., Zhu, L.M., Sun, B.Y., Li, B., Guo, B., Wang, J.Q., 2009. Zircon U–Pb ages and Hf isotope characteristics and their geological significance of the Shahewan, Caoping and Zhashui granitic plutons in the South Qinling Orogen. Acta Petrologica Sinica 25, 248–264 (in Chinese with English abstract). Griffin, W.L., Pearson, N.J., Belousova, E., Jackson, S.E., van Achterbergh, E., O'Reilly, S.Y., Shee, S.R., 2000. The Hf isotope composition of cratonic mantle: LAM-MCICPMS analysis of zircon megacrysts in kimberlites. Geochimica et Cosmochimica Acta 64, 133–147. Griffin, W.L., Wang, X., Jackson, S.E., Pearson, N.J., O'Reilly, S.Y., Xu, X., Zhou, X., 2002. Zircon chemistry and magma mixing, SE China: in-situ analysis of Hf isotopes, Tonglu and Pingtan igneous complexes. Lithos 61, 237–269. Gromet, L.P., Silver, L., 1987. REE variations across the Peninsular Ranges Batholith: implications for batholithic petrogenesis and crustal growth in magmatic arcs. Journal of Petrology 28, 75–125. Hacker, R.B., Ratschbacher, L., Webb, L., 1998. U–Pb zircon ages constrain the architecture of the ultrahigh-pressure Qinling–Dabie Orogen. China: Earth and Planetary Science Letters 161, 215–230. Hadi, E.H., Tahiri, A., Reddad, A., 2003. Les granitoides hercyniens post-collisionnels du Maroc oriental: une province magmatique calco-alcaline a shoshonitique. Comptes Rendus Geosciences 335, 959–967. Helz, R., 1976. Phase relations of basalt in their melting ranges at PH2O = 5 kb. Part 2. Melt compositions. Journal of Petrology 17, 139–193. Hong, D.W., Wang, T., Tong, Y., Wang, X.X., 2003. Mesozoic granitoids from North China Block and Qinling–Dabie–Sulu orogenic belt and their deep dynamic process. Earth Science Frontiers 10, 231–256 (in Chinese with English abstract). Hou, Z.Q., Gao, Y.F., Qu, X.M., Rui, Z.Y., Mo, X.X., 2004. Origin of adakitic intrusives generated during mid-Miocene east–west extension in southern Tibet. Earth and Planetary Science Letters 220, 139–155. Iddings, J.P., 1892. The origin of igneous rocks. Bulletin of Philosophical Society 12, 89–213. Jahn, B.M., Auvray, B., Cornichet, J., Bai, Y.L., Shen, Q.H., Liu, D.Y., 1987. 3.5 Ga old amphibolites from eastern Hebei province, China: field occurrence, petrology, Sm– Nd isochron age and REE geochemistry. Precambrian Research 34, 311–346. Jahn, B.M., Auvray, B., Shen, Q.H., Liu, D.Y., Zhang, Z.Q., Dong, Z.Q., Ye, X.J., Zhang, Q.Z., Cornichet, J., Macé, J., 1988. Archaean crustal evolution in China: the Taishan Complex, and evidence for juvenile crustal addition from long-term depleted mantle. Precambrian Research 38, 381–403. Jiang, Y.H., Jiang, S.Y., Ling, H.F., Zhou, X.R., Rui, X.J., Yang, W.Z., 2002. Petrology and geochemistry of shoshonitic plutons from the western Kunlun orogenic belt, Xinjiang, northwestern China: implications for granitoids geneses. Lithos 63, 165–187. Jiang, Y.H., Jin, G.D., Liao, S.Y., Zhou, Q., Zhao, P., 2010. Geochemical and Sr–Nd–Hf isotopic constraints on the origin of Late Triassic granitoids from the Qinling Orogen, central China: implications for a continental arc to continent–continent collision. Lithos 117, 183–197. Jin, W.J., Zhang, Q., He, D.F., Jia, X.Q., 2005. SHRIMP dating of Adakites in western Qinling and their implications. Acta Petrologica Sinica 2 (l), 959–966 (in Chinese with English abstract).. Johnson, K., Barnes, C.G., Miler, C.A., 1997. Petrology, geochemistry, and genesis of highAl tonalite and trondhjemites of the Cornucopia stock, Blue Mountains, Northeastern Oregon. Journal of Petrology 38, 1585–1611. Joplin, G.A., 1965. The problem of potash-rich basaltic rocks. Mineralogical Magazine 34, 266–275. Jung, S., Hoernes, S., Mezger, K., 2002. Synorogenic melting of mafic lower crust: constraints from geochronology, petrology and Sr, Nd, Pb and O isotope geochemistry of quartz diorites (Damara Orogen, Namibia). Contributions to Mineralogy and Petrology 143, 551–566. Kay, R.W., Kay, S.M., 1993. Delamination and delamination magmatism. Tectonophysics 219, 177–189.

Kay, R.W., Kay, S.M., 2002. Andean adakites: three ways to make them. Acta Petrologica Sinica 18, 303–311. Kay, S.M., Godoy, E., Kurtz, A., 2005. Episodic arc migration, crustal thickening, subduction erosion, and magmatism in the south-central Andes. Geological Society of America Bulletin 117, 67–88. Kaygusuz, A., Aydinçakir, E., 2009. Mineralogy, whole-rock and Sr–Nd isotope geochemistry of mafic microgranular enclaves in Cretaceous Dagbasi granitoids, Eastern Pontides, NE Turkey: evidence of magma mixing, mingling and chemical equilibration. Chemie der Erde 69, 247–277. Kepezhinskas, P., 1995. In: Smelie, J.L. (Ed.), Diverse shoshonite magma series in the Kamchatka Arc: relationships between intra-arc extension and compositon of alkaline magmas, 81. Geological Society of London, pp. 249–264. Kepezhinskas, P.K., Defant, M.J., Drummond, M.S., 1996. Progressive enrichment of island arc mantle by melt–peridotite interaction inferred from Kamchatka xenoliths. Geochemica et Cosmochimica Acta 60, 1217–1229. Kusky, T.M., Windley, B.F., Zhai, M.G., 2007. Tectonic evolution of the North China Block: from orogen to craton to orogen. Geological Society of London, Special Publications 280, 1–34. Lange, R.A., Carmichael, I.S.E., Renne, P.R., 1993. Potassic volcanism near Mono basin, California: evidence for high water and oxygen fugacities inherited from subduction. Geology 21, 949–952. Langmuir, C.H., Vocke, R.D., Hanson, G.N., 1978. A general mixing equation with applications to Icelandic basalts. Earth and Planetary Science Letters 37, 380–392. Li, Y.J., Li, Z.C., Ding, S.P., Liu, Z.W., Li, M.X., Li, J.B., 2004. Petrology features and magma mingling marks of the Wenquan granite from western Qinling. Journal of Earth Sciences and Enviroment 26, 7–21 (in Chinese with English abstract). Li, J.W., Zhao, X.F., Zhou, M.F., Vasconcelos, P., Ma, C.Q., Deng, X.D., de Souza, Z.S., Zhao, Y.X., Wu, G., 2008. Origin of the Tongshankou porphyry-skarn Cu–Mo deposit, eastern Yangtze craton, Eastern China: geochronological, geochemical, and Sr–Nd– Hf isotopic constraints. Mineralium Deposita 43, 315–336. Li, J.W., Zhao, X.F., Zhou, M.F., Ma, C.Q., de Souza, Z.S., Vasconcelos, P., 2009. Late Mesozoic magmatism from the Daye region, eastern China: U–Pb ages, petrogenesis, and geodynamic implications. Contributions to Mineralogy and Petrology 157, 383–409. Liégeois, J.P., Black, R., 1987. Alkaline magmatism subsequent to collision in the PanAfrican belt of the Adrar des Iforas. In: Fitton, J.G., Upton, B.G.J. (Eds.), Alkaline Igneous Rocks. The Geological Society. Blackwell, pp. 381–401. Liégeois, J.P., Navez, J., Hertogen, J., Black, R., 1998. Contrasting origin of post-collisional high-K calc-alkaline and shoshonitic versus alkaline and peralkaline granitoids: the use of sliding normalization. Lithos 45, 1–28. Liu, J., Xu, Z., Liou, J.G., Song, B., 2004. SHRIMP U–Pb ages of ultrahigh-pressure and retrograde metamorphism of gneisses, south-western Sulu terrane, eastern China. Journal of Metamorphic Geology 22, 315–326. Liu, Y.S., Zong, K.Q., Kelemen, P.B., Gao, S., 2008. Geochemistry and magmatic history of eclogites and ultramafic rocks from the Chinese continental scientific drill hole: subduction and ultrahigh-pressure metamorphism of lower crustal cumulates. Chemical Geology 247, 133–153. Liu, Y.S., Gao, S., Hu, Z.C., Gao, C.G., Zong, K.Q., Wang, D.B., 2010. Continental and oceanic crust recycling-induced melt–peridotite interaction in the Trans-North China Orogen: U–Pb dating, Hf isotopes and trace elements in zircons of mantle xenoliths. Journal of Petrology 51, 537–571. Lu, X.X., Wei, X.D., Xiao, Q.H., 1999. Geochronological studies of rapakivi granites in Qinling and its geological implications. Geological Journal of China Universities 15, 373–377 (in Chinese with English abstract). Ludwig, K.R., 2003. ISOPLOT 3.0: a geochronological toolkit for Microsoft excel, Berkeley Geochronology Centre. Special publication no. 4. Mao, J.W., Xie, G.Q., Pirajno, F., Ye, H.S., Wang, Y.B., Li, Y.F., Xiang, J.F., Zhao, H.J., 2010. Late Jurassic–Early Cretaceous granitoids magmatism ini Eastern Qinling, centraleastern China: SHRIMP zircon U–Pb ages and tectonic implications. Australian Journal of Earth Sciences 57, 51–78. Martin, H., 1999. Adakitic magmas: modern analogues of Archaean granitoids. Lithos 46, 411–429. Martin, H., Smithies, R.H., Rapp, R., Moyen, J.F., Champion, D., 2005. An overview of adakite, tonalite–trondhjemite–granodiorite (TTG), and sanukitoid: relationships and some implications for crustal evolution. Lithos 79, 1–24. Meng, Q.R., Zhang, G.W., 2000. Geologic framework and tectonic evolution of the Qinling Orogen, central China. Tectonophysics 323, 183–196. Miller, C., Schuster, R., Klötzli, U., Frank, W., Purtscheller, F., 1999. Post-collisional potassic an ultrapotassic magmatism in SW Tibet: geochemical and Sr–Nd–Pb–O isotopic constraints for mantle source characteristics and petrogenesis. Journal of Petrology 40, 1399–1424. Morrison, G.W., 1980. Characteristics and tectonic setting of the shoshonite rock association. Lithos 13, 97–108. Moyen, J.F., 2009. High Sr/Y and La/Yb ratios: the meaning of the “adakitic signature”. Lithos 112, 556–574. Muir, R.J., Weaver, S.D., Bradshaw, J.D., Eby, G.N., Evans, J.A., 1995. Geochemistry of the Cretaceous Separaton Plint Batholith, New Zealand: granitoid magmas formed by melting of mafic lithosphere. Journal of the Geological Society 152, 689–701. Müller, D., Groves, D.I., 1995. Potassic Igneous Rocks and Associated Gold–Copper Mineralization. Springer Verlag, Berlin, pp. 1–144. Nardi, L.V.S., 1986. As rochas granitoides da série shoshonitica. Revista Brasileira de Geociencias 16, 3–10. Nebel, O., Nebel-Jacobsen, Y., Mezger, K., Berndt, J., 2007. Initial Hf isotope compositions in magmatic zircon from early Proterozoic rocks from the Gawler Craton, Australia: a test for zircon model ages. Chemical Geology 241, 23–37. Neves, S.P., Vauchez, A., 1995. Successive mixing and mingling of magmas in a plutonic complex of Northeast Brazil. Lithos 34, 275–299.

L.-X. Ding et al. / Lithos 126 (2011) 212–232 Nie, F.J., Fan, J.T., 1989. A Study on REE Geochemistry of the Molybdenum-Bearing Granitoids in the Jinduicheng–Huanglongpu Area, Shanxi. Acta Petrologica Et Mineralogica 8, 24–32. Noyes, H.J., Frey, F.A., Wones, D.R., 1983. A tale of two plutons: geochemical evidence bearing on the origin and differentiation of the Red Lake and Eagle Peak plutons, Central Sierra Nevada, California. Journal of Geology 91, 487–509. Pearce, J.A., 1982. Trace element characteristics of lavas from destructive plate boundaries. In: Thorpe, R.S. (Ed.), Andesites. Wiley, pp. 525–548. Peccerillo, A., Taylor, S.R., 1976. Geochemistry of Eocene calcalkaline volcanic rocks from the Kastamonu area, Northern Turkey. Contributions to Mineralogy and Petrology 58, 68–81. Perugini, D., Poli, G., Christofides, G., Eleftheriadis, G., 2003. Magma mixing in the Sithonia Plutonic Complex, Greece: evidence from mafic microgranular enclaves. Mineralogy and Petrology 78, 173–200. Petford, N., Atherton, M., 1996. Na-rich partial melts from newly underplated basaltic crust: the Cordillera Blanca Batholith, Peru. Journal of Petrology 37, 1491–1521. Qin, J.F., Lai, S.C., Li, Y.F., Bai, L., Wang, J., 2005. Zircon saturation temperatures of Yangba granodiorite and mafic microgranular enclaves in northern margin of Yangtze Block and its significance. Northwestern Geology 38, 1–5. Qin, J.F., Lai, S.C., Li, Y.F., 2007. Genesis of the Indosinian Guangtoushan adakitic biotite plagiogranite in the Mianxian–Lueyang (Mianlue) suture, South Qinling, China, and its tectonic implications. Geological Bulletin of China 26, 466–471 (in Chinese with English abstract). Qin, J.F., Lai, S.C., Li, Y.F., 2008. Slab breakoff model for the Triassic post-collisional adakitic granitoids in the Qinling Orogen, central China: zircon U–Pb ages, geochemistry, and Sr–Nd–Pb isotopic constrains. International Geology Review 50, 1080–1104. Qin, J.F., Lai, S.C., Grapes, R., Diwu, C.R., Ju, Y.J., Li, Y.F., 2009. Geochemical evidence for origin of magma mixing for the Triassic monzonitic granite and its enclaves at Mishuling in the Qinling Orogen (central China). Lithos 112, 259–276. Qin, J.F., Lai, S.C., Diwu, C.R., Ju, Y.J., Li, Y.F., 2010. Magma mixing origin for the postcollisional adakitic monzogranite of the Triassic Yangba pluton, Northwestern margin of the South China Block: geochemistry, Sr–Nd isotopic, zircon U–Pb dating and Hf isotopic evidences. Contributions to Mineralogy and Petrology 159, 389–409. Rapp, R.P., Watson, E.B., 1995. Dehydration melting of metabasalt at 8–32 kbar: implications for continental growth and crust–mantle recycling. Journal of Petrology 36, 891–932. Rapp, R.P., Watson, E.B., Miller, C.F., 1991. Partial melting of amphibolite/eclogite and the origin of Archaean trondhjemites and tonalites. Precambrian Resources 51, 1–25. Rapp, R.P., Xiao, L., Shimizu, N., 2002. Experimental constraints on the origin of potassium-rich adakite in east China. Acta Petrologica Sinica 18, 293–311. Redwood, S.D., Rice, C.M., 1997. Petrogenesis of Miocene basic shoshonitic lavas in the Bolivian Andes and implications for hydrothermal gold, silver and tin deposits. Journal of South American Earth Science 10, 203–221. Richards, J.R., Kerrich, R., 2007. Special paper: Adakite-like rocks: their diverse origins and questionable role in metallogenesis. Economic Geology 102, 537–576. Rickwood, P.C., 1989. Boundary lines within petrologic diagrams which use oxides of major and minor elements. Lithos 22, 247–264. Roberts, M.P., Clemens, J.D., 1993. Origin of high-potassium, calc-alkaline, I-type granitoids. Geology 21, 825–828. Rollinson, H.R., 1993. Using Geochemical Data: Evalution, Presentation, Interpretation. Prentice Hall, pp. 1–35. Rudnick, R., Gao, S., 2003. The role of lower crustal recycling in continent formation. Geochimica Et Cosmochimica Acta 67 (Suppl.1) 403 pp. Rudnick, R.L., Gao, S., Ling, W.L., 2004. Petrology and geochemistry of spinel peridotite xenoliths from Hannuoba and Qixia, North China Craton. Lithos 77, 609–637. Sajona, F.G., Naury, R.C., Pubellier, M., Leterrier, J., Bellon, H., Cotton, J., 2000. Magmatic source enrichment by slab-derived melts in a young post-collision setting, central Mindanao (Philippines). Lithos 54, 173–206. Sen, C., Dunn, T., 1994. Dehydration melting of a basaltic composition amphibolite at 1.5 and 2.0 GPa: implications for the origin of adakites. Contributions to Mineralogy and Petrology 117, 394–409. Seo, J., Choi, S.G., Oh, C.W., 2010. Petrology, geochemistry, and geochronology of the post-collisional Triassic mangerite and syenite in the Gwangcheon area, Hongseong Belt, South Korea. Gondwana Research 18, 479–496. Shen, Z.Y., Xiao, A.C., Wang, L., Guo, J., Wei, G.Q., Zhang, L., 2010. Unconformity in the Lower Triassic of Micangshan Area, northern Sichuan Province: its discovery and significance. Acta Petrologica Sinica 26, 1313–1321 (in Chinese with English abstract). Silva, M.M.V.G., Neiva, A.M.R., Whitehouse, M.J., 2000. Geochemistry of enclaves and host granites from the Nelas area, central Portugal. Lithos 50, 153–170. Skjerlie, K.P., Patiño Doudce, A.E., 2002. The fluid-absent partial melting of a zoisite-bearing quartz eclogite from 1.0 to 3.2 GPa: implications for melting in thickened continental crust and for subduction-zone processes. Journal of Petrology 43, 291–314. Söderlund, U., Patchett, P.J., Vervoort, J.D., Isachsen, C.E., 2004. The Lu-176 decay constant determined by Lu–Hf and U–Pb isotope systematics of Precambrian mafic intrusions. Earth and Planetary Science Letters 219, 311–324. Springer, W., Seck, H.A., 1997. Partial fusion of basic granulites at 5 to 15 kbar: implications for the origin of TTG magmas. Contributions to Mineralogy and Petrology 127, 30–45. Stern, C.R., Kilian, R., 1996. Role of the subducted slab, mantle wedge and continental crust in the generation of adakites from the Austral Volcanic Zone. Contributions to Mineralogy and Petrology 123, 263–281. Stern, R.A., Bloomer, S.H., Lin, P.N., Ito, E., Morris, J., 1988. Shoshonitic magmas in nascent arcs: new evidence from submarine volcanoes in the northern Marianas. Geology 16, 426–430.

231

Stevenson, J.A., Daczko, N.R., Clarke, G.L., Pearson, N., Klepeis, K.A., 2005. Direct observation of adakites melts generated in the lower continental crust, Fiordland, New Zealand. Terra Nova 17, 73–79. Sun, S.S., McDonough, W.F., 1989. Chemical and isotope systematics of oceanic basalts; implication for mantle compositions and processes. In: Saunders, A.D., Nory, M.J. (Eds.), Magmatism in the Ocean Basins, 42. Geological Society of London, Special Publications, pp. 313–345. Sun, W.D., Li, S.G., Chen, Y.D., Li, Y.J., 2000. Zircon U–Pb dating of granitoids from South Qinling, Central China and their geological significance. Geochimica 29, 209–216 (in Chinese with English abstract). Sun, L.H., Wang, Y.J., Fan, W.M., Zi, J.W., 2008. Post-collisional potassic magmatism in the Southern Awulale Mountain, western Tianshan Orogen: petrogenetic and tectonic implications. Gondwana Research 14, 383–394. Taylor, S.R., McLennan, S.M., 1985. The Continental Crust: Its Composition and Evolution. Blackwell, Oxford, pp. 57–114. Tsuchiya, N., Suzuki, S., Kimura, J.I., Kagami, H., 2005. Evidence for slab melt/mantle reaction: petrogenesis of Early Cretaceous and Eocene high-Mg andesites from the Kitakami Mountains, Japan. Lithos 79, 179–206. Turner, S., Arnaud, N., Liu, J., Rogers, N., Hawkesworth, C., Harris, N., Kelley, S., Valstern, P., Deng, W., 1996. Post-collision, shoshonitic volcanism on the Tibetan Plateau: implications for convective thinning of the lithosphere and source of ocean island basalts. Journal of Petrology 37, 45–71. Vernon, R.H., Etheridge, M.A., Wall, V.J., 1988. Shape and microstructures of microgranitoid enclaves: indicators of magma mingling and flow. Lithos 22, 1–11. Wan, Y.S., Wilde, S.A., Liu, D.Y., Yang, C.X., Song, B., Yin, X.Y., 2006. Further evidence for − 1.85 Ga metamorphism in the Central Zone of the North China Craton: SHRIMP U–Pb dating of zircon from metamorphic rocks in the Lushan area, Henan Province. Gondwana Research 9, 189–197. Wang, Q., Zhao, Z.H., Bao, Z.W., Xu, J.F., Liu, W., Li, C.F., Bai, Z.H., Xiong, X.L., 2004. Geochemistry and petrogenesis of the Tongshankou and Yinzu adakitic intrusive rocks and the associated porphyry copper–molybdenum mineralization in southeast Hubei, east China. Resource Geology 54, 137–152. Wang, Q., Xu, J.F., Jian, P., Bao, Z.W., Zhao, Z.Z., Li, C.F., Xiong, X.L., Ma, J.L., 2006. Petrogenesis of adakitic porphyries in an extensional tectonic setting, Dexing, South China: implications for the genesis of porphyry copper mineralization. Journal of Petrology 47, 119–144. Wang, Q., Wyman, D.A., Xu, J.F., Jian, P., Zhao, Z.H., Li, C.F., Xu, W., Ma, J.L., He, B., 2007a. Early Cretaceous adakitic granites in the Northern Dabie Complex, central China: implications for partial melting and delamination of thickened lower crust. Geochimica et Cosmochimica Acta 71, 2609–2636. Wang, X.X., Wang, Tao, Jahn, B.M., Hu, N.G., Chen, W., 2007b. Tectonic significance of Late Triassic post-collision lamprophyre dykes from the Qinling Mountains (China). Geological Magazine 144, 837–848. Wang, J., Jin, Q., Lai, S.C., Qin, J.F., Li, X., 2008a. Research of geochemistry and petrogenesis of the Wulong granites in the Fuping region, South Qinling mountains. Journal of Mineralogy and Petrology 28, 79–87. Wang, T.H., Mao, J.W., Xie, G.Q., 2008b. Petrochemical research on intermediate-basic dyke rocks in Xiaoqinling, Xiong ershan goldfield, Henan province, Central China (China University of Geosciences, Beijing; Peking University) Earth Science Frontiers 15, 250–266 (in Chinese with English abstract). Wolf, M.B., Wyllie, P.J., 1994. Dehydration-melting of amphibolite at 10 kbar: the effects of temperature and time. Contributions to Mineralogy and Petrology 115, 369–383. Wu, F.Y., Yang, Y.H., Xie, L.W., Yang, J.H., Xu, P., 2006. Hf isotopic compositions of the standard zircons and baddeleyites used in U–Pb geochronology. Chemical Geology 234, 105–126. Wu, F.H., Liu, S.W., Li, Q.G., Wang, Z.Q., Su, L., Yang, K., Zhang, F., Yan, Q.R., Yan, Z., 2009. Zircon U–Pb geochronology and geological significance of Guangtoushan granitoids from Western Qinling, central China. Acta Scientiarum Naturalium Universitatis Pekinensis 45, 811–818. Xie, C.F., Zhu, J.C., Ding, S.J., 2006. Identification of Hercynian shoshonitic intrusive rocks in central Hainan Island and its geotectonic implications. Chinese Science Bulletin 51, 2507–2519. Xie, G.Q., Mao, J.W., Li, R.L., Ye, H.S., Zhang, Y.X., Wan, Y.S., Li, H.M., Gao, J.J., Zheng, R.F., 2007. U–Pb dating for volcanic rocks of the Daying Formation from Baofeng basin in eastern Qinling, China and its implications. Acta Petrologica Sinica 23, 2387–2396 (in Chinese with English abstract). Xu, X.S., O'Reilly, S.Y., Griffin, W.L., Zhou, X.M., 2000. Genesis of young lithospheric mantle in southeastern China—an LAM-ICPMS trace element study. Journal of Petrology 41, 111–148. Xu, J.F., Shinjio, R., Defant, M.J., Wang, Q., Rapp, R.P., 2002. Origin of Mesozoic adakitic intrusive rocks in the Ningzhen area of east China: partial melting of delaminated lower continental crust. Geology 12, 1111–1114. Xu, H.J., Ma, C.Q., Ye, K., 2007. Early Cretaceous granitoids and their implications for the collapse of the Dabie orogen, eastern China: SHRIMP zircon U–Pb dating and geochemistry. Chemical Geology 240, 238–259. Yan, Y.X., Zhang, C.L., 2005. Geochemical characteristics and tectonic setting of Zhangjiaba and Xinyuan plutons in South Qinling. Geological Journal of China Universities 11, 425–433 (in Chinese with English abstract). Yang, J.H., Wu, F.Y., Wilds, S.A., 2003. A review of the geodynamic setting of large-scale Late Mesozoic gold mineralization in the North China Craton: an association with lithospheric thinning. Ore Geology Reviews 23, 125–152. Yang, J.H., Chung, S.L., Wilde, S.A., Wu, F.Y., Chu, M.F., Lo, C.H., Fan, H.R., 2005a. Petrogenesis of post-orogenic syenites in the Sulu Orogenic Belt, East China: geochronological, geochemical and Nd–Sr isotopic evidence. Chemical Geology 214, 99–125.

232

L.-X. Ding et al. / Lithos 126 (2011) 212–232

Yang, J.S., Liu, F.L., Wu, C.L., Xu, Z.Q., Shi, R., Chen, S., Deloule, E., Wooden, J., 2005b. Two ultrahigh-pressure metamorphic events recognized in the central orogenic belt of China: evidence from the U–Pb dating of coesite-bearing zircons. International Geology Review 47, 323–343. Yogodzinski, G.M., Lees, J.M., Churikova, T.G., Doreneorf, F., Woerner, G., Volynets, O.N., 2001. Geochemical evidence for the melting of subducting oceanic lithosphere at plate edges. Nature 409, 500–504. Zhai, M.G., Fan, Q.C., Zhang, H.F., Sui, J.L., 2005. Lower crust processes during the lithosphere thinning in eastern China: magma underplating, replacement and delamination. Acta Petrological Sinica 21, 1509–1526 (in Chinese with English abstract). Zhai, M.G., Fan, Q.C., Zhang, H.F., Sui, J.L., Shao, J.A., 2007. Lower crustal processes leading to Mesozoic lithospheric thinning beneath eastern North China: underplating, replacement and delamination. Lithos 96, 36–54. Zhang, B.R., Luo, T., Gao, S., Ouyang, J., Chen, D., Ma, Z., Han, Y., Gu, X., 1994. Geochemical Study of the Lithosphere, Tectonism and Metallogenesis in the Qinling–Dabieshan Region. Chinese University of Geoscience Press, Wuhan, pp. 110–122 (in Chinese with English abstract). Zhang, Z.Q., Zhang, G.W., Tang, S.H., Wang, J.H., 2001. On the age of metamorphic rocks of the Yudongzi Group and the Archean crystalline basement of the Qinling Orogen. Acta Geologica Sinica 75, 198–204. Zhang, H.F., Xiao, L., Zhang, L., 2005. Geochemical and Pb–Sr–Nd isopotic compositions of granitoids from western Qinling belt: constraints on basement nature and tectonic affinity. Science in China Series D: Earth Sciences 50, 184–196. Zhang, H.F., Xiao, L., Zhang, L., Yuan, H.L., Jin, L.L., 2007a. Geochemical and Pb–Sr–Nd isotopic compositions of Indosinian granitoids from the Bikou Block, northwest of the Yangtze plate: constraints on petrogenesis, nature of deep crust and geodynamics. Science in China (Series D): Earth Sciences 50, 972–983. Zhang, Z.Q., Zhang, G.W., Liu, D.Y., 2007b. Geochronology and Geochemistry of Ophiolites, Granitoids and Clastic Rocks in Qinling Orogen. Geological Publishing House, Beijing, pp. 176–302 (in Chinese). Zhang, C.L., Wang, T., Wang, X.X., 2008. Origin and tectonic setting of the Early Mesozoic granitoids in Qinling orogenic belt. Geological Journal of China Universities 14, 304–316 (in Chinese with English abstract).

Zhang, C., Ma, C.Q., Holtz, F., 2010. Origin of high-Mg adakitic magmatic enclaves from the Meichuan pluton, southern Dabie orogen (central China): implications for delamination of the lower continental crust and melt–mantle interaction. Lithos 119, 467–484. Zhao, G.C., Wilde, S.A., Cawood, P.A., Lu, L.Z., 1998. Thermal evolution of Archean basement rocks from the eastern part of the North China carton and its bearing on tectonic setting. International Geology Review 40, 706–721. Zhao, G.C., Cawood, P.A., Wilde, S.A., Sun, M., Lu, L.Z., 2000. Metamorphism of basement rocks in the Central Zone of the North China Craton: implications for Paleoproterozoic tectonic evolution. Precambrian Research 103, 55–88. Zhao, G.C., Sun, M., Wilde, S.A., Guo, J.H., 2004. Late Archaean to Palaeoproterozoic evolution of the Trans-North China Orogen: insights from synthesis of existing data from the Hengshan–Wutai–Fuping belt. Geological Society of London 226, 27–55. Zhou, X.R., 1994. Hybridization in the genesis of granitoids. Earth Science Frontiers 1, 77–97. Zhou, X.R., Song, X.H., Zheng, J.L., 1987. Research on the rare-earth element existence in some granites. Earth Science-Journal of Wuhan College of Geology 12, 277–284 (in Chinese with English abstract). Zhou, B., Wang, F.Y., Sun, Y., 2008. Geochemistry and tectonic affinity of Shahewan orogenic rapakivi from Qinling. Acta Petrological Sinica 24, 1261–1272 (in Chinese with English abstract). Zhu, M., 1995. The ages of K–Ar isochron and 39Ar-40Ar of granites from Qinling area and their geological significance. Acta Petrologica Sinica 11, 179–192 (in Chinese with English abstract). Zhu, M.X., Luo, T.C., Zhang, H.F., 1998. Pb, Sr and Nd isotopic geochemical characteristics of Dongjiangkou plutons in Southern Qinling Belt and constrains on their sources. Geology–Geochemistry 1, 30–35 (in Chinese with English abstract). Zhu, L.M., Zhang, G.W., Guo, B., 2008. U–Pb (LA-ICP-MS) zircon dating for the large Jinduicheng porphyry Mo deposit in the East Qinling, China, and its metallogenetic geodynamical setting. Acta Geologica Sinica 82, 204–220 (in Chinese with English abstract). Zindler, A., Staudigel, H., Batiza, R., 1984. Isotope and trace element geochemistry young Pacific seamounts: implications for the scale of upper mantle heterogeneity. Earth and Planetary Science Letters 70, 175–195. Zorpi, M.J., Coulon, C., Orsini, J.B., Cocirta, C., 1989. Magma mingling, zoning and emplacement in calc-alkaline granitoid plutons. Tectonophysics 157, 315–329.