Formation Age and Evolution Time Span of the Koktokay No. 3 Pegmatite, Altai, NW China: Evidence from U Pb Zircon and 40 Ar 39 Ar Muscovite Ages

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1 bs_bs_banner doi: /rge Resource Geology Vol. 65, No. 3: Original Article Formation Age and Evolution Time Span of the Koktokay No. 3 Pegmatite, Altai, NW China: Evidence from U Pb Zircon and 40 Ar 39 Ar Muscovite Ages Qifeng Zhou, 1 * Kezhang Qin, 1 * Dongmei Tang, 1 Ye Tian, 1 Mingjian Cao 1 and Chunlong Wang 1,2,3 1 Key Laboratory of Mineral Resources, Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing, China, 2 Xinjiang Research Center for Mineral Resource, Xinjiang Institute of Ecology and Geography, Chinese Academy of Sciences, Urumqi, China and 3 University of Chinese Academy of Sciences, Beijing, China Abstract The Koktokay No. 3 pegmatite is the largest Li Be Nb Ta Cs pegmatitic rare-metal deposit of the Chinese Altai orogenic belt, and is famous for its concentric ring zonation pattern (nine internal zones). However, the formation age and evolution time span have been controversial. Here, we present the results of LA-ICP MS zircon U Pb dating and muscovite 40 Ar 39 Ar dating. Four groups of zircon U Pb ages ( 210 Ma, Ma, Ma and 172 Ma) for Zones II, V, VI, VII, and VIII, and a weighed mean 206 Pb/ 238 U age of 965 ± 11 Ma for Zone IV are identified. Also, Zones II, IV, and VI have muscovite 40 Ar 39 Ar plateau ages of ± 1.1 Ma, ± 1.0 Ma, and ± 1.1 Ma, respectively. Considering previous U Pb age studies (Zones I, V, and VII), the ages of emplacement, Li mineralization peak, hydrothermal stage of the No. 3 pegmatite are in ranges of Ma, Ma and Ma, with weighted mean 206 Pb 238 U ages of ± 2.3 Ma, ± 1.3 Ma and ± 3.9 Ma, respectively. The No. 3 pegmatite formed in the early Jurassic. The results of xenocrysts suggest that there is another pegmatite forming event of around 210 Ma in the mining district and the old zircon U Pb ages imply that Neoproterozoic crustal rocks pertain to sources of the No. 3 pegmatite. Including the previous muscovite 40 Ar 39 Ar age studies (Zones I and V), a cooling age range of Ma is considered as the time of hydrothermal stage and end of formation. The evolution process of the No. 3 pegmatite lasted 16 Ma. Therein, the magmatic stage continued for 9 11 Myr and the magmatic hydrothermal transition and hydrothermal stages were sustained at 5 7 Ma. These time spans are long because of huge scale, cupola shape, large formation depth, and complex internal zoning patterns and formation processes. Considering some pegmatite dikes in the Chinese Altai, there is an early Jurassic pegmatite forming event. Keywords: Altai, evolution time span, Muscovite 40 Ar 39 Ar age, the Koktokay No. 3 pegmatite, zircon U Pb age. 1. Introduction Pegmatite is characterized by diverse internal zonation patterns and various rare-metal mineralization types (e.g. Jahns & Burnham, 1969; Černý, 1991a, b; London, 1992, 2005, 2009; Linnen et al., 2012). A thorough understanding of the ages of different zones within a pegmatite as well as age analyses from a number of dating methods could supply time limits for cooling and evolution processes, and thereby clarify the Received 16 November Accepted for publication 22 April Corresponding authors: K. Qin, Key Laboratory of Mineral Resources, Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing , China. kzq@mail.iggcas.ac.cn. Q. Zhou, Key Laboratory of Mineral Resources, Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing , China. zhouqifeng85@163.com. 210

2 Koktokay no. 3 pegmatite, Altai, NW China evolution time spans of different stages of a pegmatite. These time spans would help define whether a pegmatite formed through a slow cooling process (Jahns & Burnham, 1969) or fast cooling and crystallization (London, 1986, 2005; Chakoumakos & Lumpkin, 1990; Morgan & London, 1999; Webber et al., 1999; Sirbescu et al., 2008). Accurate age assessments of pegmatites could also contribute to determination of the mineralization ages, since different rare metals concentrated and deposited in different internal zones of a pegmatite dike. Although there have been many pegmatite dating studies (e.g. Clark, 1982; Romer & Wright, 1992; Dickin, 1995; Romer, 1995; Romer & Smeds, 1997; Wang et al., 2003; Smith et al., 2004; Zhu et al., 2006; Habler et al., 2007; Lupulscu et al., 2011; Ren et al., 2011), detailed dating of different zones of a pegmatite is scarce. Moreover, certain characteristics of pegmatites, such as the high mobility of 87 Sr (Clark, 1982; Wang et al., 2003), strong metamictization of zircons (Dickin, 1995; Romer et al., 1996), and alterations and exsolutions of columbite tantalite (Smith et al., 2004), limit application of some traditional dating methods. The Rb Sr isochron dating method (due to unreliable estimates of the 87 Sr 86 Sr ratios, Clark, 1982) and the microcline Ar Ar method (due to the loss of argon, Parsons et al., 1988 and the low closure temperature, Foland, 1994) are unsuitable for dating pegmatites. Zircon U Pb and columbite tantalite U Pb could be feasible methods to obtain emplacement age, but they necessitate appropriate care (Romer et al., 1996; Lupulscu et al., 2011), choosing large and clear grains or domains of grains and avoiding inclusions and cracks, for example. The muscovite 40 Ar 39 Ar dating method is reliable for pegmatites, but could just represent the cooling age on account of lower closure temperature ( 358 C, Hames & Bowring, 1994). Thus, application of suitable dating methods for pegmatite is significant for getting accurate ages and dating different internal zones of pegmatite would be conducive to constrain the formation age, evolution time spans and rare-metal mineralization ages. The Chinese Altai is a key part of the Central Asian Orogenic Belt (Sengör et al., 1993; Xiao et al., 2004), and is famous for hosting around 100,000 pegmatite dikes and dozens of pegmatitic rare-metal mineral deposits (Zou & Li, 2006). The Koktokay No. 3 pegmatite with concentric zoning and large amounts of multi raremetal reserves is the largest pegmatitic rare-metal deposit in the Chinese Altai and has been used as an example to illustrate magmatic hydrothermal evolution and mineralization process of pegmatites. Thus, accurate and credible dating of the Koktokay No. 3 pegmatite are of great significance in determining the formation age and evolution time span, which will in turn provide information about the pegmatiteformation stages of the Chinese Altai and shed light on formation process of pegmatitic rare metal mineral deposits. All of this information will be useful for future prospecting in this area. In the past few decades, a range of dating methods, including whole rock and mineral Rb Sr, zircon and uranimicrolite U Pb, and muscovite and microcline K Ar and Ar Ar methods, have been used and samples from different zones were analyzed to date the Koktokay No. 3 pegmatite. However, the age results are highly inconsistent, ranging from 120 to Ma (Zou et al., 1986; Chen et al., 2000; Zhu et al., 2006; Wang et al., 2007c; Ren et al., 2011), which makes the emplacement time uncertain. Among these ages, Rb Sr isochron ages (331.9 ± 1.5 Ma, Zou et al., 1986; 238 ± 2.5 Ma, Zhu & Zeng, 2002; ± 5.8 Ma, Zhu et al., 2006) are based on initial 87 Sr 86 Sr ratios with extremely high errors, muscovite and microcline K Ar ages ( Ma, Zou et al., 1986) are based on old techniques, and the microcline 40 Ar 39 Ar age (148 ± 1 Ma, Chen et al., 2000) has a large uncertainty caused by argon loss. These ages are probably unreliable and these dating methods are not appropriate for pegmatites. On the other hand, good ages obtained by appropriate means have been published, including 213 ± 2Ma (Wang et al., 2007c), ± 0.8 Ma and ± 0.5 Ma (Ren et al., 2011), but samples tested for these zircon U Pb ages are from other pegmatite dikes around the Koktokay No. 3 pegmatite and these ages could not exactly represent the formation age of the No. 3 pegmatite. On the basis of the unreliable ages, the estimated evolution time spans of 15 Ma for the late hydrothermal alteration stage (Wang et al., 2007c), 30 Ma (Chen et al., 2000) and 100 Ma (Zou et al., 1986) for the whole evolution process are consequently unreliable. All of this uncertainty makes the time limits of formation and the cooling process controversial. Thus, it is necessary to look at the previous data with some credulity in order to accurately and systematically date the Koktokay No. 3 pegmatite. To solve these problems, we present new data on the geochronology of the Koktokay No. 3 pegmatite using zircon LA ICP MS U Pb and muscovite 40 Ar 39 Ar dating methods. Samples from Zones II, IV, V, VI, VII, and VIII were analyzed using zircon LA ICP MS U Pb dating. Samples from Zones II, IV, and VI were analyzed using muscovite 40 Ar 39 Ar dating 211

3 Q. Zhou et al. to complement the systematic age results of different zones. In light of these data and previous work, we attempt to determine the formation age and to constrain the time spans of the evolution processes of the Koktokay No. 3 pegmatite, as well as to correlate the Koktokay No. 3 pegmatite to other pegmatite dikes in the Chinese Altai. 2. Geological setting The Altai orogenic belt, containing Mongolian Altai, Chinese Altai and Russian Altai (from east to west), is situated between the Sayan and Gorny Altai of southern Siberia to the north and the Junggar block to the south (Xiao et al., 1992). The Chinese Altai, composed of variably deformed and metamorphosed Vendian to Paleozoic sedimentary, volcanic, and granitic rocks (Xiao et al., 2009), is divided into six fault-bounded integral parts that are the Altaishan terrane, the NW Altaishan terrane, the Central Altaishan terrane, the Qiongkuer-Abagong terrane, the Erqis terrane, and the Perkin-Ertai terrane (Windley et al., 2002). Each terrane shows different stratigraphy, metamorphism, deformation patterns, and age relations (He et al., 1990; Qu & Zhang, 1991) (Fig. 1). The northern and southern parts of the Chinese Altai were formed in a arc setting (island arc or forearc basin), while the central parts were deposited in a continental environment, with many granitoid intrusions (Qin, 2000; Windley et al., 2002; Xiao et al., 2004; Qin et al., 2005; Wang et al., 2006c). The Chinese Altai underwent a complex process of subduction and accretion in the Paleozoic (Xiao et al., 1992, 2004; Sengör et al., 1993; He et al., 1994; Windley et al., 2007) and finally progressed into relatively stable continent development with alternating moderate tension and compression (Li & Poliyangsiji, 2001), after finishing the formation of its basic tectonic framework and no later than the middle Carboniferous (He et al., 1994; Windley et al., 2002; Li et al., 2003; Xiao et al., 2004; Wang et al., 2005). About 200 granitoid plutons occupy at least 40% of the Chinese Altai (Zou et al., 1989; Cai et al., 2011b), indicating that magmatism played an important role in the development of the Chinese Altai (Zou et al., 1989; Wang et al., 1998). The emplacement of these granites took place at approximately five distinct time periods: Ma (Wang et al., 2006c; Cai et al., 2011b; Lv et al., 2012); Ma (climax age interval) (Zou et al., 1989; Liu, 1993); Ma (Zou et al., 1989; Liu, 1990; Zhang et al., 1996; Yuan et al., 2007b); Ma (Zhu et al., 2006; Wang et al., 2007c; Liu et al., 2014) and approximately 151 Ma (Windley et al., 2002). The granitic intrusions, which were emplaced earlier than 300 Ma, are metaluminous to peraluminous in composition (Wang et al., 2006c; Yuan et al., 2007b; Sun et al., 2008; Cai et al., 2011b). The parental magmas of metaluminous granites (I-type) and peraluminous granites (S-type) were generated by dehydration melting of hornblende-bearing mid-crustal source and mica-bearing mid-upper crustal source, respectively (Yuan et al., 2007a; Cai et al., 2011b), and some magma chambers experienced composite assimilation and fractional crystallization processes (Liu et al., 1997). Geochemical studies suggest that the metaluminous and peraluminous granites were derived from a mixture of continental sources and mantle-derived components, although the relative proportions of the mixture remains controversial (Zhao et al., 1993; Jahn et al., 2000; Chen & Jahn, 2002; Wang et al., 2006c, 2009b; Yuan et al., 2007b; Sun et al., 2008, 2009; Cai et al., 2011a). A-type granites with emplacement ages younger than 300 Ma are dominantly alkalifeldspar granites and were interpreted as post-tectonic granitoids (e.g. Tong et al., 2006; Shen et al., 2011). There are more than 100,000 pegmatite dikes in the Chinese Altai. Some of these dikes are mineralized enough to be rare metal deposits, have been explored, and have been found to have a variety of mineralization associations: muscovite, Li, Be, Nb, Ta and Cs (Wang et al., 1981; Zou et al., 1986; Zou & Li, 2006; Zhu, 2007). These pegmatite deposits, divided into nine pegmatite fields (Fig. 1) (Zou & Li, 2006), are concentrated in the Central Altaishan terrane and the Qiongkuer-Abagong terrane and are hosted in metagabbros, granites and metasediments. There are three primary types of pegmatite deposits: muscovite pegmatite ( Ma, orogenic stage, Wang et al., 2001; Ren et al., 2011), muscovite-rel (rare-element) pegmatite (369 Ma, post-orogenic stage, Wang et al., 2003, 2004), and REL (rare-element) pegmatite ( Ma, non-orogenic stage, Chen et al., 2000; Wang et al., 2000, 2003, 2007c; Ren et al., 2011; Lv et al., 2012; Qin et al., 2013). Among the REL pegmatites, Be/Be Nb Ta pegmatites occur through the periods of Ma, while complex pegmatites (Li Be Nb Ta Cs) occur later. From Permian to Jurassic, the REL pegmatitic deposits host more mineralization elements, become larger on scales and more zoned and show more evolved and fractionated signatures (Wang et al., 2004). The pegmatites formed at the same time or later than the emplacement episodes of granites. The pegmatites might have a common source with the coeval granites (Zhu et al., 212

4 Koktokay no. 3 pegmatite, Altai, NW China Fig. 1 Geological sketch map of the Chinese Altai terranes, showing the geological setting of the Koktokay No. 3 pegmatite (modified from Windley et al., 2002; Wang et al., 2006c, 2007c; Luan et al., 1995; Cai et al., 2011a). I, Altaishan terrane; II, NW Altaishan terrane; III, Central Altaishan terrane; IV, Qiongkuer-Abagong terrane; V, Erqis terrane; VI, Perkin-Ertai terrane. 1, Qinghe pegmatite district; 2, Keketuohai pegmatite district; 3, Kuwei-Jiebiete pegmatite district; 4, Kelumute- Jideke pegmatite district; 5, Kalaeerqisi pegmatite distrct; 6, Dakalasu-Kekexier pegmatite district; 7, Xiaokalasu-Qiebielin pegmatite district; 8, Hailiutan-Yeliuman pegmatite district; 9, Jiamanhaba pegmatite district. 2006; Cao et al., 2013), but according to Goodenough et al. (2014), the relationship of pegmatite and granite needs further research. The Koktokay No. 3 pegmatite is the largest pegmatite deposit with abundant raremetal reserves and complex internal zonations in this region. 3. Geology of the Koktokay No. 3 pegmatite The Koktokay No. 3 pegmatite is located near Keketuohai town, which is about 50 km away from Fuyun county, in Xinjiang, NW China (Fig. 1). The Koktokay No. 3 pegmatite intruded a metagabbro pluton (2 km 2 km, with a depth of more than 1300 m) (Zou & Li, 2006), which is divided into two parts: eastern part composed of meta amphibole gabbro and western part consisting of plagioclase amphibolite with a little amphibolitic rock. The metagabbro pluton belongs to the Keketuohai mafic complex, which was dated at 408 ± 6 Ma (zircon SHRIMP U Pb, Wang et al., 2006c) and 409 ± 5Ma (zircon LA ICP MS U Pb, Cai et al., 2012). The petrogenesis and geochemical studies imply that the 213

5 Q. Zhou et al. Fig. 2 Geological map of the Koktokay No. 3 pegmatite (modified from Zou et al., 1986; Zhu et al., 2000). (a) SHRIMP zircon U Pb ages of granite and metagabbro from Wang et al. (2006c); (b) LA ICP MS zircon U Pb age of metagabbro from Cai et al. (2012); (c) SHRIMP zircon U Pb age of pegmatite dyke from Wang et al. (2007c); (d) LA ICP MS zircon U Pb age of pegmatite dyke from Ren et al. (2011). parental magma of the Keketuohai mafic complex was high-mg tholeiitic basaltic melt produced by partial melting of the lithospheric mantle, facilitated by ridge subduction (Cai et al., 2012). Moreover, in the mining district, there are three types of granites: biotite granite, two-mica granite and muscovite granite (Fig. 2). The biotite granite is divided into undeformed biotite granite in the east with a zircon SHRIMP U Pb age of 409 ± 7 Ma (Wang et al., 2006c) and gneissic biotite granite in the southwest (Fig. 2). The two-mica granites and the muscovite granites occur within the undeformed biotite granite and the metagabbro pluton. The compositions of apatite indicate a source correlation between the muscovite granites and the No. 3 pegmatite (Cao et al., 2013). Zhu et al. (2006) suggested that the biotite granites and two-mica granites outside of the mining district, which display Rb Sr isochron ages of ± 7.5 Ma and ± 6.3 Ma, respectively, were derived from a common magma source with the No. 3 pegmatite, as indicated by similar initial ε Nd(T) values. The relationship of the No. 3 pegmatite and the different types of granites around inferred by the similar geochemical characteristics might reflect the common features of this region or the magma sources. The stratigraphic sequence of the mining district, exposed in the east and south, mainly belongs to the Habahe group and is composed of biotite-plagioclase-quartz schist with staurolite, biotitequartz schist with andalusite, and quartz-biotite schist, metamorphosed to greenschist and amphibole facies locally (Fig. 2) (Zou & Li, 2006). The Habahe group, distributed in the northwest and central to southeast regions of the Chinese Altai, is m thick (Windley et al., 2002; Cai et al., 2011a), hosting detrital zircons with U-Pb ages of Ma (Long et al., 2007, 2010; Yuan et al., 2007a), and mainly came from juvenile materials and minor evolved continental crust (Long et al., 2007, 2008, 2010; Cai et al., 2011a). The Koktokay No. 3 pegmatite is located in the intersection of the E W trend and NNW trend of joints and fractures which might be related with the Fuyun fault. Also, in the mining district, there are other pegmatite 214

6 Koktokay no. 3 pegmatite, Altai, NW China Fig. 3 (A) photograph of the mining pit, showing the sample locations; (B) shape of the No. 3 pegmatite, modified from (Zou & Li, 2006); (C) cross section of the No. 3 pegmatite at an elevation of 1186 m, showing the distribution of the nine concentric ring zones from rim to core, with previous geochronological results; (D) diagrammatic section of the No. 3 pegmatite, displaying the distribution of the mineral assemblages and the sample locations. C and D are modified from (Zou et al., 1986). a1, muscovite K Ar ages from Zou et al. (1986); a2, microcline K Ar ages from Zou et al. (1986); b1, muscovite 40 Ar 39 Ar plateau ages from Chen et al. (2000); b2, K-feldspar 40 Ar 39 Ar plateau ages from Chen et al. (2000); c1, uranmicrolite U Pb ages from Zou et al. (1986); c2, zircon SHRIMP U Pb ages from Wang et al. (2007c); d1, whole rock Rb Sr ages from Zou et al. (1986); d2, whole rock-muscovite Rb Sr ages from Zhu and Zeng (2002); d3, whole rockmuscovite-apatite Rb Sr ages from Zhu et al. (2006). dikes with different scales and mineralization types (Fig. 2). Some of these pegmatite dikes were dated at 213 ± 2 Ma using the SHRIMP zircon U Pb method and ± 0.8 Ma using the LA ICP MS zircon U Pb method (Wang et al., 2007c; Ren et al., 2011). The Koktokay No. 3 pegmatite is regularly described as a straw hat in form and consists of two main parts: a steeply dipping cupola (Fig. 3A, B) and a gently dipping plate (Fig. 2). The cupola is 250 m long, 150 m wide, 250 m deep, and plunges toward the NE at (Zou et al., 1986). The gently dipping plate, usually described as a fan, is m thick, extending along a strike of for 2160 m, and dipping down to the SW at an angle of for 1660 m (Zou & Li, 2006). The Koktokay No. 3 pegmatite is highly fractionated and is strongly internally diversified. The cupola part, showing a perfect concentric ring structure, consists of nine mineralogical-textural zones, numbered I to IX from rim to core (Fig. 3B D) (Wang et al., 1981; Zou et al., 1986; Zou & Li, 2006). These nine internal zones are mostly annular with the exception of the lenticular Zone VIII and Zone IX (Fig. 3B D). The border zone, between wall rock and Zone I, has various thicknesses ( 10 cm), and is formed by medium-fine grained quartz-muscovite with accessory garnet and 215

7 Q. Zhou et al. tourmaline. Graphic pegmatite zone (Zone I), with a ring circumference (RC) of 665 m, ring thickness (RT) of 3 7 m and extension (E) of 220 m, is almost graphic intergrowth of microcline and quartz with coarsemedium grain sizes. Saccharoidal albite zone (Zone II, with a little smaller scale than Zone I) consists of massive microcline (55%) accompanied with graphic pegmatite (10%), irregular body of saccaroidal albite (29%), and a muscovite-quartz rim (6%). Massive microcline zone (Zone III: RC, 580 m; RT, 0 35 m; E, 185 m) includes massive microcline and huge graphic microcline-quartz, which hosts coarse-grained muscovite-quartz pockets. Muscovite-quartz zone (Zone IV: RC, 520 m; RT, 4 13 m; E, 150 m) is composed of coarse-huge muscovite-quartz assemblages (60%) rich in bluish beryl and tourmaline, massive microcline (30%), and cleavelandite-spodumene (10%). Zone IV is the main economic source of beryl and mica. The cleavelandite spodumene zone (Zone V: RC, 400 m; RT, 3 30 m; E, 130 m) consists of cleavelandite spodumene assemblages (65%), quartz spodumene assemblages (30%), and muscovite quartz aggregates (5%), together with columbite tantalite and elbaite. Quartz spodumene zone (Zone VI: RC, 350 m; RT, 3 5 m; E, 100 m) follows Zone V and occurs by increasing frequency of quartz at the expense of the proportion of cleavelandite. Zones V and VI are the main Li-minerals source with huge spodumen crystals. Muscovite-slice albite zone (Zone VII: RC, 280 m; RT, 5 7 m, 50 m; E, 70 m) mostly consists of mediumcoarse muscovite-slice albite with a little quartzspodumene and quartz block. Lepidolite-slice albite zone (Zone VIII, with a length of 50 m, width of 3 7 m, and thickness of 15 m) is lens-shape, dipping to the E and NE with an inclination of 75 above Zone IX and crosscutting Zones VII and VI. It is formed by medium-fine grained lepidolite-slice albite, assembled with elbaite and columbite-tantalite, and a few muscovite-slice albite aggregates. Quartz core (Zone IX, with a length of m, width of 5 40 m, and thickness of 80 m) is composed of two segments which are quartz block (79%) and massive microcline (21%). This zone is difficult to observe now due to intensive mining. On the other hand, the plate part consists of seven zones: three continuous zones (the graphic pegmatite zone, saccharoidal albite zone, and aplite zone) and four discontinuous zones (the blocky microcline zone, muscovite quartz zone, cleavelandite quartz spodumene zone, and lepidoliteslice albite zone). Černý and Ercit (2005) considered the Koktokay No. 3 pegmatite as the spodumene subtype of the rare-element class and the Li Cs Ta (LCT) family. On the basis of the previous studies of mineralogy, inclusions and stable isotopes (Zou et al., 1986; Wu et al., 1994, 1995; Lu et al., 1997; Zhu et al., 2000; Zhang, 2001; Zhang et al., 2004a, b, 2008a, b; Liu & Zhang, 2005; Wang et al., 2006b, 2007b, 2009a; Zou & Li, 2006; Zhou, 2013; Zhou et al., 2013, 2015 accepted) as well as our studies, the Koktokay No. 3 pegmatite formed from fertile hydrous silicate melt enriched in fluxing elements (H 2O, B, and F) and rare-metal elements (Li, Be, Nb, Ta, and Cs), and mainly evolved in three stages (magmatic stage, magmatic-hydrothermal transition stage, and hydrothermal stage). Zones I, III, and to some degree Zones II and IV, were crystallized directly from the pegmatite magma in magmatic stage. Zones V, VI, VII were derived from the rare-metal rich residual silicate melt during the magmatic-hydrothermal transition stage. Zones VIII and IX are regarded as metasomatic units formed in late hydrothermal stage. The homogenization temperature ranges of inclusions for these stages are approximately C, C, and C, respectively (Wu et al., 1994, 1995; Lu et al., 1997; Zhu et al., 2000; Zhou, 2013). In light of pegmatite types, lithium mineral assemblages, and CO 2 H 2O inclusions, the formation pressures of the No. 3 pegmatite vary from 1.2 to 6.5 kbar and the formation depth is around 11 km (Lu et al., 1997; Zhu et al., 2000; Zhou, 2013). The H O C isotopic results show that there was not obvious incorporation of foreign material during the formation of the No. 3 pegmatite (Zou et al., 1986; Zou & Li, 2006). Also, the formation ages of the Koktokay No. 3 pegmatite range from 120 Ma to Ma (Fig. 3C) (Zou et al., 1986; Chen et al., 2000; Zhu et al., 2006; Wang et al., 2007c; Ren et al., 2011; Liu et al., 2014), resulting in confused emplacement ages and various evolution time spans. 4. Samples The pegmatite samples investigated here come from the huge open pit of the Koktokay No. 3 pegmatite dike (from the cupola) (Fig. 3A, D). Samples for zircon LA ICP MS U Pb dating are D-1 (Zone II), D-3 (Zone VII), D-4 (Zone V), D-5 (Zone IV), D-7 (Zone VIII) and D-8 (Zone VI). Samples for muscovite 40 Ar 39 Ar dating are KH-42 (Zone II), KH-52 (Zone IV), and KH-62 (Zone VI). Samples from Zone II consist of saccharoidal albite (70 80%), fine quartz (15 20%), and garnet (mainly spessartine) (2 5%). Sample from Zone IV consists of coarse translucent quartz (55 65%), muscovite 216

8 Koktokay no. 3 pegmatite, Altai, NW China (20 30%), albite (10 15%), green columnar beryl (1 2%), and in some cases trace columbite tantalite. The sample from Zone V consists of pink coarse columnar spodumene (30 35%), cleavelandite (45 50%), coarse translucent quartz (10 15%), and minor muscovite (1 5%). Samples from Zone VI are composed of pink-green coarse columnar spodumene (20 25%), coarse translucent quartz (30 40%), cleavelandite (20 25%), muscovite (10 15%), and occasionally tourmaline (mainly elbaite). The sample from Zone VII consists of middle-coarse muscovite (20 25%), slice albite (60 65%), minor quartz (5 10%), and trace columbite tantalite. The sample from Zone VIII consists of middle-coarse purple lepidolite (10 15%), slice albite (50 65%), minor muscovite (5 10%), elbaite (5 10%) and minor columbite tantalite (<5%). 5. Analytical methods 5.1 LA ICP MS zircon U Pb dating Zircons were separated using standard heavy-liquid and magnetic techniques for LA ICP MS studies. Representative grains were handpicked using a binocular microscope, mounted in an epoxy resin disc and then polished to be exposed. Prior to LA ICP MS analysis, zircons were photographed in transmitted and reflected light and imaged by cathodoluminescence (CL) using a LEO 1450VP SEM. These images were used to choose potential target sites for U Pb dating and to characterize the internal features of the zircons, including growth zones and inclusions. U Pb isotopic analyses were carried out on an Agilent 7500a Q-ICPMS with a 193 nm laser ablation system hosted at the Institute of Geology and Geophysics, Chinese Academy of Sciences in Beijing. A detailed description of the Q-ICPMS can be found in Xie et al. (2008). During analysis, a primary laser beam of 10kv acceleration voltage with a 10 Hz pulse rate is focused on a spot approximately 44 μm in diameter. The dwell time for each isotope was set at 6 ms for Si, Ti, Nb, Ta, Zr, and REE, 15 ms for 204 Pb, 206 Pb, 207 Pb and 208 Pb, and 10 ms for 232 Th and 238 U. Each spot analysis consisted of approximately 30 s of background acquisition and 40 s of sample data acquisition. To monitor the stability of the instrument and to ensure the reliability of the measured results, the zircon and NIST SRM 610 standards were measured once, respectively, for every five sample points. 207 Pb 206 Pb, 206 Pb 238 U, 207 U 235 U( 235 U = 238 U/137.88), and 208 Pb 232 Th ratios were corrected using the zircon as the external standard. The fractionation correction and results were calculated using GLITTER 4.0 (Macquarie University). All of the measured isotope ratios of during the process of sample analysis were regressed and corrected using reference values. The relative standard deviations of reference values for were set at 2%. The weighted mean U Pb ages were processed using ISOPLOT/Ex Version 3.0 (Ludwig, 2003). 5.2 Muscovite 40 Ar 39 Ar dating Samples of muscovite were crushed and sieved, and muscovites with grain sizes between mm were separated by magnetic and gravimetric means. About 100 mg were further purified by hand-picking under a binocular microscope to remove all visible impurities. 40 Ar 39 Ar measurements were performed at the State Key Laboratory of Lithospheric Evolution, Institute of Geology and Geophysics, Chinese Academy of Sciences (IGGCAS) in Beijing. Details of the analyses are given in Wang et al. (2006a, 2007a). Aliquots of mineral separates were wrapped in aluminum foil and stacked in quartz vials (at IGGCAS) along with the neutron flux monitor GA1550 biotite which has a K Ar age of 98.5 ± 0.8 Ma (Spell & McDougall, 2003) (at IGGCAS). The samples were irradiated at the H8 position in the 49-2 Reactor at the China Institute of Atomic Energy in Beijing, with a neutron flux of 6.5*10 12 ncm 2 s 1. Samples were heated stepwise with a double vacuum resistance furnace. Following an additional 5 minutes of released gas purification on Zr Al getters, the isotopic data were measured using a MM5400 mass spectrometer (at IGGCAS). After correcting for mass discrimination, system blanks, and radiometric interference, 40 Ar 39 Ar ages were calculated according to 40 Ar* 39 Ar K ratios, and the J value was obtained by analyses of the monitors. Plateau ages were determined from three or more contiguous steps, comprising >50% of the 39 Ar released, revealing concordant ages at the 95% confidence level (2σ). The raw data were processed using the ArAr CALC-software of Koppers (2002). The correction factors of interfering isotopes produced during irradiation were determined by analysis of irradiated pure CaF 2 and K 2SO 4. Mass discrimination was monitored using an on-line air pipette from which multiple measurements are made before and after each incremental-heating experiment. The decay constant used is λ= year 1 (Steiger & Jäger, 1977). All 37 Ar abundances were corrected for radiogenic decay (half-life 35.1 days). The uncertainty for each 217

9 Q. Zhou et al. Fig. 4 Cathodoluminenscenece (CL) images of dated zircon grains from Zones II, IV, V, VI, VII, and VIII of the Koktokay No. 3 pegmatite. Analyzed spots are circled and the codes correspond to the results in Table 1. apparent age is given at one standard deviation. The inverse isochrones were calculated from the plateau steps using the York regression algorithm (York, 1969). 6. Results 6.1 Zircon characteristics The zircon populations from the Koktokay No. 3 pegmatite are mainly opaque, dark, spongy, and morphologically complex, ranging from equant euhedral grains to round or anhedral grains in CL imaging. The zircon grains in samples D-1 (Zone II), D-4 (Zone V), D-7 (Zone VIII) and D-8 (Zone VI) are mostly dark, spongy, euhedral to anhedral grains with variable sizes (Fig. 4). Some zircon grains have porous and blurred cores with weakly thin or small oscillatory rims (e.g. grain 24 in sample D-1 and grain 18 in sample D-8) (Fig. 4). The zircons in sample D-5 (Zone IV) are mainly euhedral grains with columnar shapes and have black cores with irregular mantles (Fig. 4). The zircon populations from sample D-3 (Zone VII) are subeuhedral to euhedral grains, hosting black or irregular mantles and regular oscillatory rims (Fig. 4). These zircon characteristics, especially the darkness, the spongy and porous parts, and the small oscillatory rims on CL images suggest that zircons from the Koktokay No. 3 pegmatite have suffered strong metamictization or recrystallization, likely due to high U concentrations or deuteric fluid. This is common in rare-element pegmatitic deposits due to high fluxes and the presence of magmatic-hydrothermal transition in the systems. Thus, the obtained data have been checked carefully and both of the concordant and discordant age data points are considered and discussed here. 6.2 LA ICP MS zircon U Pb ages Sample D-1 (Zone II) Analyses of 11 spots of 11 zircon grains from sample D-1 (Zone II) were obtained (Table 1; Fig. 4). These analyses illustrate that U concentrations range from 2341 to 5174 ppm, Th contents vary from to ppm and Th/U ratios are in the range of 0.01 and U Pb isotopic results form a weighted mean 206 Pb 238 U age of ± 2.0 Ma with a mean sum weighted deviation (MSWD) of 2.1 (Fig. 5). Some data deviate from the concordia line, indicating a Pb loss on account of metamictization and high U contents (Fig. 5) Sample D-5 (Zone IV) Three analyses of three zircons from sample D-5 (Zone IV) were obtained (Table 1; Fig. 4). Zircons analyzed from sample D-5 have variable U concentrations ( ppm), Th contents ( ppm) and Th/U ratios ( ). The 206 Pb 238 U ratios for these three spots yielded a weighted mean age of 965 ± 11 Ma with a MSWD of (Fig. 5). 218

10 Koktokay no. 3 pegmatite, Altai, NW China Table 1 LA ICP MS zircon U Pb isotopic data for the Koktokay No. 3 pegmatite Sample spot 206 Pb (ppm) 207 Pb (ppm) 208 Pb (ppm) Pb (ppm) Th (ppm) U (ppm) Th U 207 Pb 235 U σ 206 Pb 238 U σ 207 Pb 235 U σ 206 Pb 238 U σ Age (Ma) Age (Ma) D-1, Zone II (Saccharoidal albite) D D D D D D D D D D D D-5, Zone IV (Muscovite quartz) D D D D-4, Zone V (Cleavelandite-spodumene) D D D D D-8, Zone VI (Quartz-spodumene) D D D D D D D D D D D D D-3, Zone VII (Muscovite-slice albite) D D D D D D-7, Zone VIII (Lepidolite-slice albite) D D D Pb indicates the contents of radiogenic lead; the rations of 206 Pb 238 U and 207 Pb 235 U indicate the radiogenic lead; errors are 1σ. 219

11 Q. Zhou et al. Fig. 5 Concordia and weighed mean prism diagrams showing the LA ICP MS U Pb ages of zircon grains from Zones II, IV, V, VI, VII and VIII of the Koktokay No. 3 pegmatite. 220

12 Koktokay no. 3 pegmatite, Altai, NW China Sample D-4 (Zone V) Four analyses of four zircons from sample D-4 (Zone V) were obtained (Table 1; Fig. 4). Three analyses from sample D-4 have variable U concentrations ( ppm), Th contents ( ppm) and Th/U ratios ( ). The 206 Pb 238 U ratios for these three spots yielded a weighted mean age of ± 6.3 Ma with a MSWD of 1.6 (Fig. 5). The result of another zircon from sample D-4 has 1429 ppm U, ppm Th and a Th/U ratio of 0.26, with a 206 Pb 238 U age of 186 ± 2 Ma (Table 1; Fig. 4) Sample D-8 (Zone VI) Twelve analyses of 12 zircons from sample D-8 (Zone VI) were obtained (Table 1; Fig. 4). Zircons analyzed from sample D-8 have highly variable U concentrations ( ppm), Th contents ( ppm) and Th/U ratios ( ). The 206 Pb 238 U ratios for these 12 points yielded a weighted mean age of ± 2.7 Ma with a MSWD of 1.7 (Fig. 5) Sample D-3 (Zone VII) Five analyses were obtained for zircon grains from sample D-3 (Zone VII) (Table 1; Fig. 4). Four of these analyses illustrate that U concentrations range from to 2981 ppm, Th contents vary from to ppm, and that the Th/U ratios are in the range of 0.16 and U Pb isotopic results form a weighted mean 206 Pb 238 U age of ± 6.4 Ma with a mean sum weighted deviation (MSWD) of 4.1 (Fig. 5). The other data point defines a 206 Pb/ 238 U age of 214 ± 2 Ma with ppm U, ppm Th and a Th/U ratio of 1.12 (Table 1; Fig. 4) Sample D-7 (Zone VIII) Three analyses of three zircons from sample D-7 (Zone VIII) were obtained (Table 1; Fig. 4). Two of these analyses have variable U concentrations ( ppm), Th contents ( ppm) and Th/U ratios ( ), giving a weighted mean 206 Pb 238 U age of ± 4.2 Ma with a MSWD of 0.5 (Fig. 5). The result of the other zircon from sample D-7 has 2712 ppm U, ppm Th and a Th U ratio of 0.14, yielding a 206 Pb 238 U age of 172 ± 3 Ma (Table 1; Fig. 4). The distinctly variable U and Th contents and Th/U ratios are indicative of metamictization, which leads to formation of opaque, dark and spongy zircons without apparent oscillatory or sector zones on CL images (Fig. 4). 6.3 Muscovite 40 Ar/ 39 Ar ages The 40 Ar 39 Ar dating of muscovite from samples, including KH-42 (Zone II), KH-52 (Zone IV) and KH-62 (Zone VI), were performed. The results, obtained from a temperature interval of C, show similar plateau ages of Ma with % of 39 Ar K released. The inverse isochron ages, in the range of Ma, are coincident with corresponding plateau ages. The muscovite 40 Ar 39 Ar dating of sample KH-42, obtained at temperatures between 750 and 1290 C, displays a fairly flat age spectrum with a welldefined plateau, giving a plateau weighted age of ± 1.1 Ma (2σ) (MSWD = 3.89) and an inverse isochron age of ± 1.1 Ma (MSWD = 3.94) with over 99.41% of 39 Ar K released (Table 2; Fig. 6a, b). Similarly, the 40 Ar 39 Ar step-heating results ( C) of muscovite in sample KH-52 provide a plateau weighted age of ± 1.0 Ma (2σ) (MSWD = 2.67) and an inverse isochron age of ± 1.0 Ma (MSWD = 2.80) with 98.36% of 39 Ar K released (Table 2; Fig. 6c, d). The muscovite 40 Ar/ 39 Ar dating of sample KH-62 step heated from 750 to 1200 C yield a plateau weighted age of ± 1.1 Ma (2σ) (MSWD = 3.15) and an inverse isochron age of ± 1.1 Ma (MSWD = 3.44) with 90.68% of 39 Ar K released (Table 2; Fig. 6e, f). As shown in Table 2, there are two older ages in sample Kh-42 and one older age in sample Kh-62, obtained from temperature intervals of C and C. The cumulative argon release during these stages only accounts for 0.31% and 0.58% of the total gas released, respectively. These older ages are not accurate, probably due to excess Ar. Nevertheless, analyses at temperatures from C are relatively coincident, forming remarkably flat age plateaus with 99.10% and 90.10% of the 39 Ar K released, indicating the absence of excess argon or any diffusive argon loss. Thus, the ages of ± 1.1 Ma, ± 1.0 Ma and ± 1.1 Ma for samples KH-42, KH-52 and KH-62 are reliable estimates for the crystallization age of muscovite from Zones II, IV, VI according to the criteria of Dalrymple and Lamphere (1971). 7. Discussion 7.1 Interpretation and geological significance of zircon U Pb ages In this study, we found weighed mean 206 Pb 238 U ages of ± 2.0 Ma for Zone II, 965 ± 11 Ma for Zone IV, and ± 2.7 Ma for Zone VI (Fig. 5). Also, some samples host two sets of ages for one zone (Fig. 5). 221

13 Q. Zhou et al. Table 2 40 Ar 39 Ar step heating data for muscovite from the No. 3 pegmatite Temperature ( C) 40 Ar 39 Ar 37 Ar 39 Ar 36 Ar 39 Ar 40 Ar* 39 Ar k 40 Ar*(%) 39 Ark(%) Age (Ma) ±2σ Sample name KH-42 (Zone II, Saccharoidal albite) J = ± C ± C ± C ± C ± C ± C ± C ± C ± C ± C ± C ± C ± C ± C ± Sample name KH-52 (Zone IV, Muscovite-quartz) J = ± C ± C ± C ± C ± C ± C ± C ± C ± C ± C ± C ± C ± C ± C ±10.84 Sample name KH-62 (Zone VI, Quartz-spodumene) J = ± C ± C ± C ± C ± C ± C ± C ± C ± C ± C ± C ± C ± C ± C ± C ± C ± C ±8.17 Zone V displays a weighed mean 206 Pb 238 U age of ± 6.3 Ma and a 206 Pb 238 U age of 186 ± 2 Ma. Zone VII shows a weighed mean 206 Pb 238 U age of ± 6.4 Ma and a 206 Pb 238 U age of 214 ± 2 Ma. Zone VIII displays a weighed mean 206 Pb 238 U age of ± 4.2 Ma and a 206 Pb 238 U age of 172 ± 3 Ma. Several 206 Pb 238 U ages of ca. 965 Ma (Zone IV), ca. 210 Ma (Zones V and VII), ca Ma (Zones II, VII and VIII), ca. 186 Ma (Zones V and VI) and ca. 172 Ma (Zone VIII) are on the concordia line, while some 206 Pb 238 U ages of ca Ma for Zones II and VI deviate from the concordia line in different degrees (Fig. 5). We focus on those 222

14 Koktokay no. 3 pegmatite, Altai, NW China Fig. 6 Plateau and inverse isochron Ar Ar ages of muscovite for sample KH-42 (Zone II), KH-52 (Zone IV), and KH-62 (Zone VI). (A, C, E) Plateau ages of KH-42, KH-52, and KH-62, respectively; (B, D, F) inverse isochron ages of KH-42, KH-52, and KH-62, respectively. 206 Pb 238 U ages on the concordia line and discuss their geological significance. The zircons of ca. 965 Ma seem to be relic zircons, and the ages ranging from 172 to 213 Ma are related to formation of pegmatite Zircon weighed mean 206 Pb 238 U age of 965 ± 11 Ma The zircons from Zone IV, yielding a weighed mean 206 Pb 238 U age of 965 ± 11 Ma, are possibly xenocrysts from country rocks or the magma source. The direct country rock of the Koktokay No. 3 pegmatite is a metagabbro pluton which formed in Ma (Wang et al., 2006c; Cai et al., 2012). The nearby granites formed in the lower to middle Devonian (Wang et al., 2006c) or in the lower Triassic (Zhu et al., 2006). The detrital zircon U Pb ages of the Habahe group mainly vary from 540 to 439 Ma (Long et al., 2007, 2010; Yuan et al., 2007a). The ages of the country rocks around the No. 3 pegmatite are not in accordance with ca. 965 Ma; hence, these old zircons do not come from the country rocks, including the metagabbro, granites, and the Habahe group. The other possibility is that these old zircons from Zone IV may possibly stem from the 223

15 Q. Zhou et al. Fig. 7 Four groups of zircon U Pb ages of the Koktokay No. 3 pegmatite. (A) plot of zircon U Pb ages of different zones with error bars. Data are from Wang et al. (2007c) (gray solid circles) and this study (white solid circles). (B) there are four weighed mean zircon U Pb ages corresponding to the four regions (age groups I, II, III and IV) in Figure 7a, respectively. magma source of the Koktokay No. 3 pegmatite. If so, the old zircons reflect that some of the rocks which were melted to form the Koktokay No. 3 pegmatite, are Neoproterozoic rocks. Since pegmatites are the result of crustal rocks melting, at least some of the source rocks of the No. 3 pegmatite are Neoproterozoic crustal rocks. The source rocks should be researched further Zircon 206 Pb 238 U age range of Ma Based on the previous zircon U Pb analyses (Wang et al., 2007c) and the zircon U Pb results in this study for Zones I, II, V, VI, VII and VIII, there are different sets of zircon U Pb ages for a single zone (Figs 5, 7a) and four groups of zircon U Pb ages for the Koktokay No. 3 pegmatite as a whole: Ma, Ma, Ma and Ma, are identified, with corresponding weighed mean 206 Pb 238 U ages of ± 3.0 Ma, ± 2.3 Ma, ± 1.3 Ma and ± 3.9 Ma, respectively (Fig. 7b). These age groups reflect different events (emplacement and mineralization) and evolution stages which are related to the continuous formation process of the Koktokay No. 3 pegmatite Age group I: Ma (212.0 ± 3.0 Ma). The zircon grains with the age range of Ma are interpreted as xenocrysts for the following reasons. Wang et al. (2007c) reported that the SHRIMP zircon U Pb age of the 10 m-thick Be pegmatite dike above 224