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Article

Petrogenesis of Devonian and Permian Pegmatites in the Chinese Altay: Insights into the Closure of the Irtysh–Zaisan Ocean

by
Mengtao Wang
1,2,3 and
Xin Zhang
1,*
1
School of Geography and Planning, Ningxia University, Yinchuan 750021, China
2
Key Laboratory of High-Temperature and High-Pressure Study of the Earth’s Interior, Institute of Geochemistry, Chinese Academy of Sciences, Guiyang 550002, China
3
University of Chinese Academy of Sciences, Beijing 100049, China
*
Author to whom correspondence should be addressed.
Minerals 2023, 13(9), 1127; https://doi.org/10.3390/min13091127
Submission received: 29 June 2023 / Revised: 18 August 2023 / Accepted: 23 August 2023 / Published: 25 August 2023
(This article belongs to the Special Issue Petrogenesis, Geochronology, Mineralization and Geochemistry of Granite Rocks)
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Abstract

:
Owing to tectonic, magmatic, and metamorphic controls, pegmatites associated with different spatiotemporal distributions exhibit varying mineralisation characteristics. The petrogenesis of pegmatites containing rare metals can improve the understanding of geodynamic processes in the deep subsurface. In order to understand the difference of petrogenesis between Devonian and Permian pegmatites, zircon U-Pb geochronological and Hf-O isotope analyses were performed on samples of the Jiamanhaba, Amulagong, and Tiemulete pegmatites from the Chinese Altay. According to the results obtained, the Amulagong and Tiemulete pegmatites were formed during the Devonian, and samples that were analysed yielded zircon U-Pb ages of 373.0 ± 7.8 and 360 ± 5.2 Ma, respectively. Samples from these pegmatites produced εHf(t) values of 2.87–7.39, two-stage model ages of 900–1171 Ma and δ18O values of 9.55‰–15.86‰. These results suggest that the pegmatites were formed via an anatexis of mature sedimentary rocks deep in the crust. In contrast, the Jiamanhaba pegmatite was formed during the Permian, and its samples produced εHf(t) and δ18O values of 2.87–4.94 and 6.05‰–7.32‰, respectively, which indicate that the associated magma contained minor amounts of mantle/juvenile materials. The petrogenesis of pegmatites containing rare metals can reveal tectonic settings of their formation. A combination of data that were generated in the present study and existing geochronological and Hf-O isotope data for felsic igneous and sedimentary rocks in the Chinese Altay shows that the εHf(t) sharply increased while the δ18O suddenly decreased between Late Carboniferous and Early Permian. These changes highlight a tectonic transformation event during this critical period. This tectonic event promoted mantle–crustal interactions, and thus, it was probably linked to assemblages of the Altay orogen and the Junggar Block. The present study provides evidence that the Irtysh–Zaisan Ocean probably closed during the Late Carboniferous (~300 Ma).

1. Introduction

Granitic pegmatites (hereafter referred to as pegmatites) are igneous rocks comprising skeletal crystals or graphic zones and/or large mineral grains, and these are associated with systematic variations [1]. Pegmatites are often enriched in Li, Be, Nb, Ta, Rb, Cs, Hf, and other rare metals, and thus, these rocks are associated with several mineral deposits around the world. Prominent pegmatite-linked deposits include the following: the Keketuokay No. 3 Li-Be-Nb-Ta-Rb-Cs-Hf deposit in Xinjiang, China, Tanco Li-Ta-Cs deposit in Canada and Bikita Li deposit in Zimbabwe [2]. The distribution of pegmatites in the world and continental convergence events exhibit a strong spatiotemporal correlation, and this highlights the significance of tectonic activities on the formation of pegmatites [3,4,5]. Therefore, numerous studies on the petrogenesis of pegmatites have been undertaken in different parts of the world.
The classification of pegmatite is as complex as the pegmatite itself. Others classification schemes focus on the internal structure, paragenetic relationship, bulk composition, petrogenetic aspects, nature of parent medium, and geochemical features. However, all schemes ignore differences in the geological environment, which is important for understanding pegmatites. Ginsburg et al. [6] and Černy [7] proposed a new classification scheme: according to the petrogenetic of pegmatites, three classes can be distinguished here including the NYF family, LCT family, NYF + LCT family. The NYF family has a prominent accumulation of Nb, Ta, Ti, Y, Sc, REE, Zr, U, Th and F (Nb > Ta). The parent granites are mainly subaluminous to metaluminous A- to I-types, which are correlated with anorogenic setting [8,9]. The LCT family is enriched in Li, Rb, Cs, Be, Sn, Ta, Nb (Ta > Nb), B, P and F. The parent granites are mainly midly to peraluminous S-, I-type which are usually strongly fractionated within individual bodies [10,11]. They are correlated with orogenic setting (syn-orogenic and post-orogenic) [8,9]. The NYF + LCT family pegmatites display mixed geochemical and mineralogical characteristics [12]. Jiamanhaba, Amulagong, Tiemulete—the three pegmatites we studied—are all the LCT-type. Two petrogenetic models have been advanced regarding pegmatites, and based on the genetic relationship between LCT-type pegmatites and S-type granites, Černy and Meintzer [10], Černý [8] and Selway et al. [13] indicated that pegmatites represent consolidation products of residual magmas from the late fractional crystallisation of granitic magmas. The J-B model was proposed by Jahns and Burnham [14] based on the experimental data. They considered water-saturated magma playing a very important role in the forming of internal zoning (in response to fractional crystallisation) in pegmatite. The coarse grain texture is ascribed to the effects of the rapid transfer of materials. Supercritical fluids, which contain H2O and volatiles, are crucial for the transportation of materials that promote the crystallisation of minerals. Therefore, large crystals usually form in the fluid phase, whereas small crystals develop directly from the silicate melt. In contrast, London et al. [15,16] and London [17] proposed the London model, in which pegmatite formation was associated with low H2O contents of granitic magmas. The zonations in pegmatites were attributed to the presence of boundary layers, where volatiles and incompatible elements cause crystallisation from the exterior to the interior. Furthermore, based on experimental studies and the observation of miarolitic cavities in pegmatites, a high content of H2O was indicated to inhibit the growth rate of crystals, and this accounts for the absence of large crystals. According to another model, the petrogenesis of pegmatites is related to partial melting deep in the crust instead of the fractional crystallisation of granitic magmas [18,19,20,21,22]. In fact, a genetic relationship between pegmatites and granites is commonly lacking. The Tanco (Canada), Greenbush (Australia) and Keketuokay No. 3 (China) pegmatites exhibit no genetic relationship with the host granites. However, previous studies suggest that pegmatites are commonly associated with anatexis [21,22,23,24,25,26,27,28,29,30].
The Altay Orogenic Belt in the north of the Xinjiang Uygur Autonomous Region is an important component of the Central Asian Orogenic Belt (CAOB) [31,32,33,34,35,36,37,38,39,40] (Figure 1). In this prominent metallogenic belt, more than 100,000 pegmatite veins are distributed across 38 fields [41,42,43]. These rare metal-rich pegmatite veins are primarily associated with Permian and Triassic Periods. Among these fields, major deposits include the following: super large Li-Be-Ta-Nb-Cs in the Keketuohai No. 3, super large Li in Kaluan 650, 805, 806, and 807, large Li-Be-Nb-Ta in Kelumute No. 112, small Be-Li-Nb-Ta in Xiaokalasu, small Be-Nb-Ta Dakalasu and Yelaman, and small Be in Jiamanhaba pegmatites [44,45,46,47,48,49,50,51]. Relatedly, Devonian and Carboniferous rare metal deposits include the following: small Li-Be-Ta-Nb deposit in Amulagong and small Li-Be-Ta-Nb-Cs deposit in Talati pegmatites [23]. Previous studies on the pegmatites in Chinese Altay mainly focus on the chronology of the pegmatites and the geochemical evolution of the magma. The main dating methods used in the Institute of Chronology are muscovite Ar-Ar dating [42,52,53,54], zircon U-Pb dating [21,45,46,49,50,51,53,54,55,56,57,58,59], and niobium-tantalum U-Pb dating, which has appeared in recent years [48,54]. Because muscovite is easily disturbed by magmatic hydrothermal, the dating results of muscovite are quite different from those of zircon and niobium-tantalite. The petrogenesis of pegmatite in Chinese Altay has always been controversial. Predecessors have speculated that pegmatite may be the product of the late evolution of granite batholith, but no clear evidence has been presented to prove the genetic evolution relationship between the two. In the present study, Hf and O isotope data were utilised to evaluate source and tectonic setting differences between the Permian Jiamanhaba pegmatite and Devonian Amulagong and Tiemulete pegmatites which show different distribution, quantity and mineralisation of them during two periods.
Minerals 13 01127 g001
Figure 1. Geological map of the Chinese Altay (modified after Xiao et al. [60]; Cai et al. [61,62]; Xiao and Santosh, [63]; the age data are from Zhang et al. [21]; Lv et al. [22,23,50]; Wang et al. [45]; Chen, [46]; Ma et al. [49]; Ren et al. [51]; Wang et al. [64]). The age data of pegmatites are shown in Table 1.
Figure 1. Geological map of the Chinese Altay (modified after Xiao et al. [60]; Cai et al. [61,62]; Xiao and Santosh, [63]; the age data are from Zhang et al. [21]; Lv et al. [22,23,50]; Wang et al. [45]; Chen, [46]; Ma et al. [49]; Ren et al. [51]; Wang et al. [64]). The age data of pegmatites are shown in Table 1.
Minerals 13 01127 g001
Table 1. The age of pegmatites in Chinese Altay.
Table 1. The age of pegmatites in Chinese Altay.
NumberPegmatite DepositMineralisation TypePositionScaleMethodAge (Ma)Reference
1AmulagongLi-Be-Ta-NbUnit 3smallLA-ICP-MS373.0 ± 7.8 this study
358[23]
2TalatiP-Cs-Nb-Be-LiUnit 3smallLA-ICP-MS386[23]
3BaichengLi-Be-Ta-NbUnit 3smallLA-ICP-MS274[22]
4Tiemulete1(F)-Nb-P-B-BeUnit 3smallLA-ICP-MS333[23]
5Tiemulete2Mica-BeUnit 3smallLA-ICP-MS360.0 ± 5.2this study
6BuluketeBeUnit 3smallLA-ICP-MS275.5 ± 4.2[51]
7AsikaerteBe-Nb-MuUnit 2largeLA-ICP-MS229.0 ± 3.0 [56]
8Kokotokay No. 3Li-Be-Ta-Nb-Cs-Rb-HfUnit 2super largeLA-ICP-MS208.1 ± 0.8[51]
LA-ICP-MS212.7 ± 0.5[51]
LA-ICP-MS220[45]
LA-ICP-MS220–209 [46]
LA-ICP-MS190.6 ± 1.2[51]
LA-ICP-MS214.9 ± 2.1[51]
LA-ICP-MS180.7 ± 0.5[51]
9HusiteBe-Nb-TaUnit 2middleLA-ICP-MS244.3 ± 1.1[51]
10QunkuBe-Nb-TaUnit 2smallLA-ICP-MS206.8 ± 1.6[51]
11Azubai 01Be-Nb-TaUnit 2middleLA-ICP-MS191.6 ± 2.0 [21]
12JiamukaiLi-Be-Ta-Nb-CsUnit 2middleLA-ICP-MS212.2 ± 1.7[51]
192.0 ± 2.3 [21]
13KelumuteLi-Be-Nb-TaUnit 2largeLA-ICP-MS202.9 ± 0.8[51]
238.3 ± 2.0[50]
14KaluanLi-Nb-TaUnit 2super largeLA-ICP-MS223.7 ± 1.8[49]
221 ± 15[49]
216 ± 2.6[49]
224.6 ± 2.3 [21]
15KukalagaiLi-Nb-TaUnit 2super largeLA-ICP-MS211.3 ± 2.4[49]
16DakalasuBe-Nb-TaUnit 3smallLA-ICP-MS270.1 ± 1.7[51]
272.5 ± 1.4[51]
263.7 ± 4.4[64]
258[22]
17HulugongBe-Nb-TaUnit 2smallLA-ICP-MS246.8 ± 1.2[51]
18XiaokalasuBe-Li-Nb-TaUnit 3smallLA-ICP-MS267.5 ± 3.5[64]
19QiemuerqiekeREEUnit 3smallLA-ICP-MS253[22]
20TaerlangBe-Nb-Ta-REEUnit 3smallLA-ICP-MS256[22]
21QiebielinNb-P-B-BeUnit 3smallLA-ICP-MS403[23]
22AkebasitawuBe-Nb-Ta-REEUnit 3smallLA-ICP-MS249.7 ± 0.7[51]
253[22]
23SaerjiakeLi-Be-Ta-NbUnit 3smallLA-ICP-MS253[22]
24HailiutanREEUnit 3smallLA-ICP-MS254[22]
25YelamanBe-Nb-TaUnit 3smallLA-ICP-MS267.8 ± 1.4[51]
262.8 ± 3.1[64]
263[22]
26JiamanhabaBe-Ta-NbUnit 3smallLA-ICP-MS269.4 ± 1.6[51]
260.6 ± 2.5this study
260[22]
27Jiamanhaba02Nb-P-B-BeUnit 3smallLA-ICP-MS395[23]

2. Geological Setting

2.1. Regional Geology

The Chinese Altay Orogenic Zone extends from Mongolia in the east to Kazakhstan and Russia in the west. It is situated between the South Siberian Craton in the north and the Junggar Block in the south. The Chinese Altay was partitioned into four terranes by the Hongshanzui–Nort, Kurti–Abagong, Fuyun–Xibodu and Irtysh faults [32,42,45,61,62,65]. The North Altay Terrane (Unit 1) is bounded in the south by the Hongshanzui–Nuoerte fault [32,66], and it contains mainly clastic sediments (shales, siltstones, greywackes, and sandstones), limestones, granites, and minor greenschist facies island-arc metavolcanic rocks that were formed in the Late Devonian to Early Carboniferous [67,68,69]. The Central Altay Terrane (Unit 2) is sandwiched between the Hongshanzui–Nuoerte fault in the north and the Kurti–Abagong fault in the south [31,32]. This terrane contains mainly metamorphosed sedimentary rocks, early to middle Paleozoic granites [33,38,39,49,61,62,66,69,70,71,72,73] and volcanic rocks (Habahe Group). In this terrane, granite commonly occurs as intrusions into sedimentary rocks, and this indicates frequent magmatism. Relatedly, the Qiongkuer Terrane (Unit 3) is sandwiched between the Kurti–Abagong fault in the north and the Fuyun–Xibodu fault in the south. It comprises the Silurian Kulumuti Group and Early Devonian Kangbutiebao and Middle-Late Devonian Altay formations [32,74], and these strata are coveal with exposed granites [38,66,70,71,72,75,76,77,78]. Lastly, the South Altay Terrane (Unit 4) is sandwiched between the Fuyun–Xibodu fault in the north and the Irtysh fault in the south. The northwestern part is composed by sediments. The southeastern part contains mainly Devonian fossil-bearing sedimentary rocks (Kangbutiebao Formation) and the overlain Carboniferous Formations. In addition, a few Carboniferous–Permian volcaniclastic rocks are exposed in this area [32,74,79].
Geochronology data show that granitic intrusions were common in the Chinese Altay from the Early Palaeozic to the Early Mesozoic, and these attained a peak in the Devonian [35,37,38,44,50,51,61,62,72,80,81,82,83,84]. The Palaeozoic granites, which are dominantly I- and S-types, are widely distributed in the Central Altay Terrane [61,85]. The peak of granitic magmatism in the Permian was coeval with pegmatites and pyrometamorphism [21,50,51,70,77,78,80,86]. The Permian granites involve I, A, I-A, and I-S types, and these are common across the Qiongkuer Terrane [78,87,88,89]. In the Chinese Altay, pegmatites are dominantly found in the Central Altay and the Qiongkuer terranes. The over 100,000 pegmatite veins, which are distributed in 38 fields, occur mainly in the Halong–Qinghe and Jiamanhaba–Dakalasu metallogenic subzones. Owing to tectonic, magmatic, and metamorphic controls, pegmatites associated with different spatiotemporal distributions show varying mineralisation characteristics. Devonian–Carboniferous pegmatites are concentrated in the Halong–Qinghe subzone, whereas Permian pegmatites are abundant in the Jiamanhaba–Dakalasu subzone in the Qiongkuer Terrane. Triassic pegmatites occur in both the Qiongkuer and the Central Altay terranes, whereas most Jurassic pegmatites are found in the Central Altay Terrane, with subordinate occurrences in the North Altay Terrane. Minerals that are common in the Permian and Devonian–Early Carboniferous pegmatites in the Qiongkuer Terrane include the following types: Be, Nb, Ta, and Li. Minerals and mineral assemblages that are dominant in Triassic pegmatites in the Qiongkuer and the Central Altay terranes are the following types: Be, Li-Nb-Ta, Li-Be-Nb-Ta, and Be-Li-Nb-Ta (-Cs-Rb-Hf).

2.2. Geology of the Jiamanhaba, Tiemulete and Amulagong Pegmatites

2.2.1. Jiamanhaba Pegmatite

The Jiamanhaba pegmatite field is in the northwest of the Chinese Altay, approximately 50 km from Habahe County in the south and just 10 km from the border between China and Kazakhstan in the west [51]. It is in the west terminal of the Jiamanhaba–Dakalasu metallogenic subzone in the Qiongkuer Terrane, which belongs to the west region of the Kabaer–Xibodu magmatic arc. These pegmatites are associated with many intrusive rocks in the area, including biotite granites and Caledonian plagiogranites. In the south of the field, minor outcrops of Late Variscan two-mica granites are also found. Pegmatites are scattered around this field, and these are characterised by lengths of <500 m and widths of <10 m. These pegmatites exhibit a low degree of fractionation, simple zonation, and common Be mineralisation.
The Jiamanhaba pegmatite is hosted in biotite granites, and based on outcrops that measure 50 m in length and 7 m in width, it is associated with an ESE strike of 98° and a dip of 20°. Across the upper to the lower section, the pegmatite vein can be divided into the following internal zones: graphic, saccharoidal albite, massive microcline, and quartz–muscovite nested. Major minerals in the Jiamanhaba pegmatite are muscovite, microcline, quartz, and albite, whereas spessartite, apatite, tourmaline, and zircon are dominant accessory minerals. Prominent rare metal minerals in this pegmatite are beryl and columbite-tantalite. The Be mineralisation is abundant in the graphic and massive microcline texture zones.

2.2.2. Amulagong Pegmatite

The Amulagong pegmatite field is in the southeast of the Chinese Altay, approximately 35 km from Qinghe County in the east. It occurs between the Mieriteke and Baleqige–Keketuomusike faults, and this field was intruded by numerous Variscan porphyritic biotite granites (strike = WNW 315°) that contain medium-coarse crystals. Variscan diorites (strike = WNW 350°) are also scattered in the north and south of the Nuogan River over a length and a width of 3.9 and 1.4 km, respectively. The 116 pegmatite veins in this field cover an area of >3 km2, and these veins, which intruded diorites along fissures, display a NW strike and dip of 45°. The majority of these veins are <100 m long, whereas some are >200 m long, but all are <20 m wide. Several pegmatite veins are characterised by low Be and Be-Nb-Ta mineralisation, whereas others show high Be-Li-Nb-Ta mineralisation.
The studied Amulagong pegmatite is hosted in medium Variscan diorites, and it is characterised by a length of approximately 70 m and an average width of 8 m. Across the upper to the lower section, pegmatite veins can be partitioned into the following zones: graphic, saccharoidal albite, massive microcline, quartz–muscovite nested, and cleavelandite albite. Major minerals in this pegmatite are muscovite, microcline, quartz, and albite, whereas spessartite, apatite, tourmaline, and zircon constitute accessory minerals. Ore minerals in this deposit are beryl, columbite–tantalite, spodumene, and pollux. Beryllium mineralisation is concentrated in the massive microcline and quartz–muscovite nested zones. Conversely, Nb and Ta are hosted in manganotantalites mainly in the massive microcline zone, whereas spodumene crystallised in both the massive microcline and cleavelandite albite zones.

2.2.3. Tiemulete Pegmatite

This pegmatite in the southeast of the Chinese Altay is approximately 25 km from Qinghe County in the southeast. It is within the Qinghe pegmatite ores area in the Qiongkuer Terrane, which represents the core of the Buleke Anticline. Few muscovite- and beryl-bearing pegmatites intrude the medium-sized Early Variscan diorite body with an east–west strike located east of the pegmatite field. In the pegmatite field, strata involve those of the Early–Middle Ordovician Habahe Formation, including garnet biotite gneisses, sillimanite–two mica gneisses, augen gneisses, striped, enterolithic and augen migmatites, as well as biotite–plagioclase gneisses, which exhibit a strike of 130° and dip of 40° [23]. In the west of this field, 40 muscovite-bearing pegmatite veins occur as intrusions, and 28 of these contain rare metals. These veins are scattered across the field, and most are <200 m long and <5 m wide.
The Tiemulete pegmatite is hosted in biotite–plagioclase and plagioclase hornblende gneisses, and it measures 190 m in length and 2.21 m in width, with a strike of WNW 330° and a dip of 78°. The veins are shallow and exhibit characteristics of expansion and contraction. Across the upper to the lower section, the pegmatite vein can be divided into the following zones: graphic, saccharoidal albite, massive microcline, and quartz–muscovite nested. The quartz–muscovite nested zone is characterised by a random mixture of smoky grey quartz and scaly muscovite. Major minerals in this pegmatite are muscovite, microcline, quartz, and albite, whereas spessartite, apatite, tourmaline, and zircon are accessory minerals. Ore minerals in this deposit include beryl and columbite–tantalite. The Be mineralisation is concentrated in the saccharoidal albite zone, and the beryl crystals are small and dispersed. Muscovite is abundant in the quartz–muscovite nested zone, but this is non-uniformly distributed in the pegmatite vein.

2.3. Materials

Jiamanhaba pegmatite can be divided into four mineral assemblage zones: graphic assemblage; saccharoidal albite assemblage; massive microcline assemblage and nested quartz–muscovite assemblage. Graphic assemblage is composed of quartz (40%), potassium feldspar (40%) and muscovite (20%). It is located in the upper wall of the gently inclined pegmatite vein and accounts for 10% of the total volume of the whole vein (Figure 2a). The saccharoidal albite assemblage is mainly composed of quartz (45%) and albite (55%), accounting for 10% of the total volume of the whole vein (Figure 2a,c). The massive microcline assemblage is mainly composed of block microcline (98%) and muscovite (2%), accounting for 40% of the total volume of the whole vein. The nested quartz–muscovite assemblage is mainly composed of quartz (30%), albite (50%) and muscovite (20%), accounting for 40% of the total volume of the whole vein (Figure 2b).
Amulagong pegmatite can be divided into five mineral assemblage zones: graphic assemblage; saccharoidal albite assemblage; massive microcline assemblage; nested quartz–muscovite assemblage and cleavelandite albite–spodumene assemblage. The graphic assemblage is located on the edge of the upper wall of the inclined pegmatite vein, consisting of medium coarse-grained quartz (40%) and potassium feldspar (60%), with rare metal mineralisation. It accounts for 20% of the total volume of the whole vein (Figure 2e). The saccharoidal albite assemblage consists of fine-grained quartz (30%) and albite (70%), accounting for 30% of the total vein volume (Figure 2d,f). The massive microcline assemblage is mainly composed of microcline (60%), quartz (30%) and muscovite (10%). Niobium–tantalite crystals in the shape of chicken feet can be seen in albitisation rocks, which accounts for 20% of the total volume of the whole vein. The nested quartz–muscovite assemblage, located in the dike footwall, about 0.5 m thick. It is mainly composed of quartz (30%), albite (40%) and muscovite (30%). It accounts for 10% of the total volume of the whole vein. The cleavelandite albite–spodumene assemblage consists mainly of foliated albite (60%), spodumene (20%) and muscovite (20%). It accounts for 20% of the total volume of the whole vein.
The Tiemulete pegmatite can be divided into four mineral assemblage zones: graphic assemblage; saccharoidal albite assemblage; massive microcline assemblage and nested quartz–muscovite assemblage. The graphic assemblage, located in the upper wall of the gently inclined pegmatite vein, is mainly composed of feldspar (80%), quartz (16%) and muscovite (4%), with beryl crystals. It accounts for 30% of the total volume of the whole vein (Figure 2h). The saccharoidal albite assemblage is mainly composed of albite (70%), quartz (27%) and muscovite (3%). It accounts for 10% of the total volume of the whole vein (Figure 2i). The massive microcline assemblage is mainly composed of feldspar (80%), quartz (13%) and muscovite (7%). It accounts for 20% of the total volume of the whole vein. The nested quartz–muscovite assemblage is an irregularly shaped aggregate of quartz (20%) and muscovite (80%). The colour of mica is brown, white, light green, and it is in the form of a wedge, plate, laminated composite. It accounts for 40% of the total volume of the whole vein (Figure 2g).

3. Methods

We distinguish the mineral assemblage zoning and production sequence of each pegmatite, and we identify the graphic assemblage as the first sequence mineral assemblage. Rocks with a fresh structural plane of graphic mineral assemblage in pegmatite are taken as samples. The contact zone between pegmatite and country rock often develops UST, which is mainly composed of potassium feldspar, tourmaline and muscovite-oriented growth perpendicular to the contact zone, and the crystal particles are fine. The saccharoidal albite assemblage located in the inner side of graphic assemblage presents transitional contact with the graphic assemblage. Since the genetics of the saccharoidal albite assemblage are not clear, samples from the middle of the graphic assemblage are taken to avoid the change of minerals and lithology in the rocks located in the inner and outer parts of the graphic assemblage. In order to ensure the selection of enough zircon grains, the weight of the sample is >10 kg.
Zircon targeting and cathodoluminescence (CL) imaging were performed in the State Key Laboratory of Mineral Deposit Geochemistry, Institute of Geochemistry, Chinese Academy of Sciences. Samples were pulverised and zircon grains were selected in a pollution-free environment. Zircons were selected through magnetic separation and flotation, and these were then sequentially pasted to target sites using epoxy resin. A scanning electron microscope was used to produce cathodoluminescence images for subsequent zircon U-Pb geochronology and Hf and O isotope measurements.
In situ zircon U-Pb isotope analysis was performed using an NWR193 laser ablation microprobe (Elemental Scientific Lasers LLC) that was attached to an Analytikjena M90 ICP-MS in the LA-ICP-MS laboratory of the State Key Laboratory of Mineral Deposit Geochemistry, Institute of Geochemistry, Chinese Academy of Sciences. Helium was used as a compensation gas to adjust the sensitivity of the laser ablation system before entry into the ICP-MS. The energy density of the ultraviolet beam that was generated by the laser was 10 J/cm2, whereas the beam spot diameter was 30 μm. The time-resolved analysis for each sample involved durations of approximately 20–30 and 40 s for the blank and sample signals, respectively. The U-Pb data were corrected for isotope fractionation using zircon 91500 as an external standard [90,91]. Ten points in each sample were analysed in duplicate, and a point for Plesovice was used as the monitor. The NIST SRM 610 and Si were correspondingly used as internal and external standards to correct the Pb concentration. The instrument drift sensitivity calibration and calculation of element concentrations and U-Th-Pb isotope ratios were achieved using offline processing software. The Isoplot 3 offline software was used to calculate the weighted average ages and to plot these ages in the Concordia diagram [90,92,93]. The common Pb adjustment was conducted using Andersen’s software [94].
The zircon Hf isotope analysis was also performed in situ using the NWR193 laser-ablation microprobe (Coherent GMBH) attached to a Neptune multicollector ICP-MS in the LA-ICP-MS laboratory of the State Key Laboratory of Mineral Deposit Geochemistry, Institute of Geochemistry, Chinese Academy of Sciences. The laser beam diameter was 32 µm and the beam energy was 10 J/cm2. Helium was used as the carrier gas, and it was mixed with argon using a Y-shaped transfer interface. The gas mixture then transported the laser-denudated sample to the mass spectrometer for analysis. To correct for interferences of 176Lu and 176Y on 176Hf, the 176Lu/175 Lu = 0.02658 and 176Yb/173Yb = 0.796218 ratios were determined [95]. Analysis of the zircon 91500 and Plesovice standards was interspersed with that of samples to monitor the stability of the instrument. The initial 176Hf/177Hf ratios (εHf(t)) were calculated using the chondritic uniform reservoir (CHUR) that was proposed by Blichert-Toft and Albarede [96] as the reference. The TDM1 of Hf was calculated based on the depleted mantle, which involves present-day 176Hf/177Hf and 176Lu/177Hf ratios correspondingly of 0.28325 and 0.0384 [97].
A CAMECA IMS 1280-HR secondary ion mass spectrometer (France) was used for in situ zircon O isotope analysis in the State Key Laboratory of Isotope Geochemistry, Guangzhou Institute of Geochemistry, Chinese Academy of Sciences. A primary beam of Cs+ involving an impact potential of 10 kV and a beam current of 7 nA was spread over 10 μm2 to generate a homogeneous beam density. A normal-incidence electron flood gun was used to compensate for the charging of samples during the analyses, while negative secondary ions were extracted using a 10 kV electric field. The O isotope analysis was conducted based on the multicollector technique using two Faraday cups [98]. Isotopic ratios of elements that are determined using the ion microprobe are often impacted by drift. The instrumental mass fractionation (IMF) between light and heavy ions is the ratio of measured to real values [99,100]. In the present study, the IMF correction was performed using the 91500 zircon standard, which is assigned a δ18O value of 9.9‰ [101].

4. Results

4.1. Zircon U-Pb Dating

4.1.1. Jiamanhaba Pegmatite

Approximately 60 zircon grains measuring 50–200 μm in length were selected from Jiamanhaba pegmatite samples, and 30 were analysed. Data for a representative subset of all the zircons analysed are shown in Figure 3a. Most of the zircons analysed were short, columnar, and euhedral, and these produced aspect ratios varying from 2:3 to 3:4, whereas subordinate elongated grains yielded aspect ratios ranging from 1:2 to 1:3. The zircons that were analysed produced dark to almost black CL images (Figure 3a), but these also exhibited abundant spots and alteration zones.
Among the 30 zircons that were analysed for U-Pb isotopes, the concentrations of Th and U for the zircon grains that were analysed ranged from 106.64 to 1851.54 and 77.86 to 6619.34 ppm, respectively, and these yielded Th/U ratios that vary from 0.035 to 1.37, whereas the 206Pb/238U ages range from 191 to 422 Ma. Sixteen yielded 206Pb/238U ages that vary from 191 to 422 Ma, which obviously deviate from the concordant line. These results suggest that because of their high Th/U of 0.07–1.37 (Table 2), the samples were likely xenocrysts, whereas the other 14 samples, which produced an average weighted age of 260.6 ± 2.5 Ma, probably reflect the age of emplacement of the Jiamanhaba pegmatite (Figure 4a).

4.1.2. The Amulagong Pegmatite

Regarding this pegmatite, approximately 60 zircon grains measuring 50–150 μm were separated, and 15 were analysed. Data for a representative subset of the zircons that were analysed are displayed in Figure 3b. The analysed zircons are equiaxial, and these produced comparable aspect ratios of approximately 1:1. These zircons also yielded dark to almost black CL images (Figure 3b), and many crystals exhibited weak magmatic oscillatory annular zones.
Among the 15 zircons that were analysed for U-Pb isotopes, the concentrations of Th and U for the samples range from 0.83 to 10.46 and 1120.51 to 3837.11 ppm, respectively, and these produced Th/U ratios that vary from 0.001 to 0.003, whereas the 206Pb/238Uages range from 368 to 430 Ma (Table 2). One yielded an age of 430 Ma, and this was considered a xenocryst because the associated age was evidently distinct from those obtained for other samples. Data for the other 14 samples produced a weighted average age of 373.0 ± 7.8 Ma, and this likely represents the age of emplacement of the Amulagong pegmatite (Figure 4b).

4.1.3. Tiemulete Pegmatite

Approximately 60 zircon grains measuring 100–200 μm were separated from samples of this pegmatite, and 20 were analysed. Figure 3c shows data for a representative subset of the analysed zircons. The zircons were short and columnar, and they exhibited aspect ratios that vary from 2:3 to 3:4. These zircons also displayed dark to almost black CL images (Figure 3c), and several crystals were characterised by weak magmatic oscillatory annular zones. The concentrations of Th and U for the zircons analysed range from 2.58 to 696.37 and 1503.83 to 4563.76 ppm, respectively, and these produced Th/U ratios that range from 0.001 to 0.222, whereas the 206Pb/238U ages vary from 350 Ma to 631 Ma (Table 2).
Five of the 20 samples that were analysed yielded 206Pb/238U ages that ranged from 341 to 585 Ma, and these highlight an obvious deviation from the concordant line. These samples are likely xenocrysts because of their high Th/U ratios of 0.005 to 0.1. The other 15 samples produced a weighted average age of 360 ± 5.2 Ma, and this probably represents the age of emplacement of the Tiemulete pegmatite (Figure 4c).

4.2. Zircon Hf Isotope Compositions

4.2.1. Jiamanhaba Pegmatite

Ten zircon grains from the Jiamanhaba pegmatite were analysed for Lu/Hf isotope compositions, and these samples yielded comparable 176Hf/177Hf ratios that vary from 0.2826946 to 0.2827535. The εHf(t) values that were obtained based on the U-Pb age data range from 2.87 to 4.94, whereas the TDM2 ages vary from 969 to 1100 Ma (Table 3).

4.2.2. Amulagong Pegmatite

Seventeen zircon grains from the Amulagong pegmatite that were analysed for Lu/Hf isotope compositions also produced comparable 176Hf/177Hf ratios of 0.282705 to 0.2827494. The εHf(t) values that were obtained based on the U-Pb data range from 5.83 to 7.39, whereas the TDM2 ages vary from 900 to 999 Ma (Table 3).

4.2.3. Tiemulete Pegmatite

Regarding the Tiemulete pegmatite, 10 zircon grains that were separated and analysed for Lu/Hf isotope compositions produced uniform 176Hf/177Hf ratios of 0.2826354–0.2827263. The εHf(t) values that were obtained based on the U-Pb data range from 2.95 to 6.17, whereas the TDM2 ages vary from 967 to 1171 Ma (Table 3).

4.3. Zircon O Isotopic Compositions

Ten zircons samples that were separated from each of the of the Jiamanhaba, Amulagong and Tiemulete pegmatites were analysed for O isotope compositions and the data are presented in Table 4. The δ18O data for the Jiamanhaba, Amulagong and Tiemulete samples correspondingly vary from 6.05 to 7.32 ‰ (mean = 6.73‰), 9.55 to 14.27‰ (mean = 12.00‰), and 11.47 to 15.86‰ (mean = 12.61‰).

5. Discussion

5.1. Zircon U-Pb Geochronology

The U-Pb data for samples that were used to calculate Concordia ages are presented in Table 2. However, because data for the main zircon populations from the Jiamanhaba and Tiemulete pegmatites failed to provide corresponding concordant ages, the lower intersection age that was obtained using a T-W diagram was considered the age of emplacement of the associated magma. According to zircon U-Pb dating data, the Jiamanhaba pegmatites were formed in the Early Permian, whereas the other two pegmatites were emplaced in the Late Devonian. Analysis data produced wide Th/U ratios of 0.04 to 1.37 and from 0.001 to 0.222 for samples from the Jiamanhaba and Tiemulete pegmatites, respectively. According to the values of Th/U, several zircons that were obtained from the Jiamanhaba and Tiemulete pegmatites are metamictised, but these are of a magmatic origin. Conversely, data generated from zircons that were collected from the Amulagong pegmatite show a high concordance, and thus, the weighted average age of 373.0 ± 7.8 Ma that was produced from these data is considered representative of the age of emplacement. The metamictisation of zircon grains is suggested to decrease radiogenic Pb content. However, Peucat et al. [102] and Davis and Paces [103] suggested that in shallow deposits, because of the low temperature and lack of fluid activity, the loss of radiogenic Pb from metamictised zircons is negligible. Therefore, if a loss of radiogenic Pb occurred in several zircon grains that are used for dating because of metamictisation, the U-Pb ages obtained can deviate from the concordant line. The deduction of effects of metamorphism on the generated data can thus provide representative ages of emplacement for the pegmatites. According to previous studies on the Chinese Altay, metamict zircon grains can also provide reliable U-Pb age data for the pegmatites [21,22,23,49,50,104]. Therefore, the Jiamanhaba, Amulagong, and Tiemulete pegmatites are assigned average weighted ages of 260.6 ± 2.5, 373.0 ± 7.8, and 360 ± 5.2, respectively.
Zircon U-Pb data can also be used to evaluate the associated tectonic settings and petrogenesis of these Early Permian and Late Devonian pegmatites. According to a previous study on tectonism in the Chinese Altay, the Permian and Devonian were peak periods of magmatism. In the study area, granites that are coeval with pegmatites are present in outcrops, and these facilitate investigations on the coupling between tectonism and the formation of the pegmatites studied.

5.2. Source Characteristics

Owing to the significantly higher concentration of Hf relative to its radioactive parent Lu, the Hf isotope composition of zircons remains essentially constant after crystallisation. In addition, the closure temperature of 900 °C for the zircon Hf isotope system is similar to that of the zircon U-Pb system. Thus, the Hf isotope composition of zircon can accurately record the compositions of parent magmas of pegmatites [105]. Furthermore, the narrow ranges of εHf(t) values for the samples studied indicate that the associated Hf isotope systems were closed. The εHf(t) ranges of 2.87–4.94, 5.83–7.39, and 2.95–6.17 for the Jiamanhaba, Amulagong, and Tiemulete samples produced corresponding TDM2 age ranges of 969–1100, 900–999, and 967–1171 Ma, respectively, and these values fall between the depleted mantle and CHUR evolution lines. These results indicate that the precursor magmas of the three pegmatites originated from heterogeneous sources. The εHf(t) values of Triassic pegmatites including the Keketuohai No.3, Kelumute No. 112, Kaluan and Asikaerte range correspondingly from 1.25 to 2.39, 0.03 to 2.35, 0.65 to 2.50, and −1.50 to 0.77, and these are associated with model ages of 1102–1174, 1112–1225, and 1090–1213 Ma, respectively [49,50]. Overall, the εHf(t) values for the Triassic pegmatites range from −1.50 to +2.50, and the model ages vary from 1090 to 1225 Ma, which highlight low Hf isotope compositions and high model ages. The εHf(t) values for the Xiaokalasu, Dakalasu, Yelaman, and Jiamanhaba pegmatites range from 1.6 to 4.2, 1.6 to 4.3, 2.4 to 3.3, and 2.87 to 4.94, respectively, and the corresponding model ages involve ranges of 997–1293, 989–1064, and 1050–1110 Ma. Relatedly, Permian pegmatites yield εHf(t) values of 1.6–4.94 and two-stage model ages of 989–1293 Ma, whereas Devonian pegmatites produced εHf(t) values of 2.87–7.39 and two-stage model ages that range from 900 to 1171 Ma. Apparently, the εHf(t) values of the Permian and Devonian pegmatites are significantly higher than those of the Triassic pegmatites. The ranges of εHf(t) values for pegmatites that were formed in these three periods reflect differences in the composition of their sources. These results suggest that the magma of the Permian pegmatites involved a higher input of mantle materials relative to that of the Triassic pegmatites. Alternatively, the parent magma of the Permian pegmatites originated from the partial melting of sedimentary or granitic rocks that are associated with high εHf(t) values. Geochemical data on pegmatites in the Chinese Altay from previous studies show that these rocks were derived from the partial melting of Palaeozoic granites and sedimentary/metasedimentary rocks [21,49,50,106]. Therefore, we compiled εHf(t) data for Palaeozoic granites and sedimentary rocks from the Chinese Altay. Data for the Jiamanhaba–Amulagong–Tiemulete pegmatites fall within the ranges of 17 to 9.8 and 17 to 26 that were obtained for the Palaeozoic granites/gneiss and sedimentary rocks, respectively (Figure 5). These results demonstrate that excluding the mixing of magma, the remelting of pre-existing Palaeozoic granites or sedimentary rocks was adequate to generate the initial magma from which the studied pegmatites were formed. However, the origin of the Jiamanhaba–Amulagong–Tiemulete pegmatites cannot be definitively demonstrated in the present study.
Zircon is a very stable mineral, and the diffusion rate of its oxygen isotope is low; thus, its initial oxygen isotope composition is preserved even after advanced metamorphism [107,108,109,110]. Therefore, the oxygen isotope composition of zircon can be used as a tracer of sources of ore-forming materials [110,111,112,113,114,115]. The zircon O isotope values of the Jiamanhaba–Amulagong–Tiemulete pegmatites range from 6.05 to 15.86 ‰ (Table 4, Figure 6), which are higher than those of mantle-derived zircons (5.3 ± 0.3‰). These results indicate that mantle or juvenile materials were involved in the initial magmas from which the Jiamanhaba–Amulagong–Tiemulete pegmatites were formed. Among major rock-forming minerals in these pegmatites, quartz and feldspar exhibit oxygen isotope enrichments. Pegmatites formed through the partial melting of granites should involve quartz and feldspar that show oxygen isotope enrichment relative to the granites [116,117]. According to data shown in Figure 6, the O isotope values of Palaeozoic granites are clearly higher than those of the Jiamanhaba pegmatites and comparable to those of the Amulagong and Tiemulete pegmatites, which demonstrate that these pegmatites were not derived from the partial melting of Palaeozoic granites. According to data that are displayed in Figure 5, the Habahe Group and Kulumuti Formation are part of Unit 2 (Central Altay Terrane). The dating data show that ages of samples from the Habahe Group mostly range from Cambrian to Ordovician, whereas for the Habahe/Kulumuti Formation, these vary from Neoproterozoic to Silurian [35,37,38,118,119]. Unit 3 (Qiongkuer Terrance) comprises the Early Devonian Kangbutiebao and Middle Devonian Altay formations as well as the Cambrian–Ordovician Habahe Group [35,38,84,120,121,122,123,124,125,126,127,128,129,130]. Therefore, the Permian Jiamanhaba pegmatite was most likely derived from the partial melting of sedimentary/metasedimentary rocks of the Kangbutiebao Formation in the northwest of the Qiongkuer Terrance, and minor mantle-derived or juvenile materials were added to the original magma. In contrast, the Devonian Amulagong and Tiemulete pegmatites were probably derived from the partial melting of sedimentary/metasedimentary rocks of the Habahe/Kulumuti Formation in the southeast of the Qiongkuer Terrance, and the original magmas were contaminated by higher amounts of mantle-derived or juvenile materials than that of the Permian pegmatites.
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Figure 5. The age vs. εHf(t) diagrams of the Jiamanhaba, Amulagong and Tiemulete pegmatites compared with Paleozoic granites and sedimentary rocks in the Chinese Altay. (The data are from Zhang et al. [21]; Long et al. [35,37]; Sun et al. [38,40]; Lv et al. [50]; Cai et al. [62]; Zhang et al. [73]; Tong et al. [77]; Liu et al. [78]; Yu et al. [85]; Zhao et al. [131]; He et al. [132]; Shen et al. [133]; and Zhang et al. [134]).
Figure 5. The age vs. εHf(t) diagrams of the Jiamanhaba, Amulagong and Tiemulete pegmatites compared with Paleozoic granites and sedimentary rocks in the Chinese Altay. (The data are from Zhang et al. [21]; Long et al. [35,37]; Sun et al. [38,40]; Lv et al. [50]; Cai et al. [62]; Zhang et al. [73]; Tong et al. [77]; Liu et al. [78]; Yu et al. [85]; Zhao et al. [131]; He et al. [132]; Shen et al. [133]; and Zhang et al. [134]).
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Figure 6. The age vs. δ18O diagrams of the Jiamanhaba, Amulagong and Tiemulete pegmatites, compared with Paleozoic granites in the Chinese Altay. (The data are from Wang et al. [64]; Kong et al. [135]; Zhang et al. [136]).
Figure 6. The age vs. δ18O diagrams of the Jiamanhaba, Amulagong and Tiemulete pegmatites, compared with Paleozoic granites in the Chinese Altay. (The data are from Wang et al. [64]; Kong et al. [135]; Zhang et al. [136]).
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5.3. Petrogenesis

Pegmatites are petrogenetically distinguished based on their relationships to granites because these are considered late evolution segregation and crystallisation products of granitic magmas. An example is the Li-rich pegmatite in the Songpan–Ganzi Orogenic Belt [137,138,139]. However, based on geochemical and field studies, an anatexis model was advanced recently for the petrogenesis of pegmatites in the Chinese Altay. Lv et al. [22], for example, suggested that the Baicheng pegmatites originated from the anatexis of metasedimentary rocks in the Qinghe area. In addition, Zhang et al. [21] reported that several Triassic pegmatites originated from partial melting of the Halong granite and sedimentary rocks of the Kulumuti Group. According to previous studies and zircon Hf-O isotope compositions of the Jiamanhaba–Amulagong–Tiemulete pegmatites, it is likely that the pegmatites studied were derived from the partial melting of mature sedimentary rocks, and the original melts were contaminated with mantle-derived or juvenile materials. The differences in δ18O values between these pegmatites and the Palaeozoic granites indicate the lack of a direct genetic relationship between the former and the latter. Based on the discrepancy of δ18O value between these pegmatites and Paleozoic granites, we conclude that the pegmatites do not have the direct relationship with Paleozoic granites. This is because the earth crust’s average content of beryllium ranges from 2.4 to 4 ppm [140], which is far lower than the beryl saturated crystallisation threshold value of higher than 35 ppm [141,142]. Even if the sedimentary rocks were partially melted in small proportion, it is far from reaching the threshold value of beryl saturated crystallisation. In addition, the Jiamanhaba–Amulagong–Tiemulete pegmatites have a low fraction degree, which also cannot enrich the beryllium content to reach the threshold value of beryl saturated crystallisation. According to the zircon U-Pb dating data, the Jiamanhaba pegmatite was formed in the early Permian, and the others (Amulagong and Tiemulete) were formed in the late Devonian, which show the strong crust–mantle interaction in the two periods. In addition, the available data show two important peaks in Devonian (ca. 420~380 Ma) and Permian (ca. 290~250 Ma), which counterpart to the Amulagong pegmatite (ca. 373 Ma), Tiemulete pegmatite (ca. 360 Ma) and Jiamanhaba (ca. 261 Ma) pegmatite [40,48,72,77,79,82,88,143,144,145,146,147]. According to classification diagrams (Figure 7), the Permian granites are dominantly I- and S-types, whereas the Devonian granites are mostly the A-type. The A-type granite is commonly associated with anorogenic or post-orogenic settings. Plutons of the Permian granites are approximately elliptical and are restricted by fault zones. Based on geochemical characteristics that were reported in previous studies, the Permian granites were formed in a post-orogenic setting [77,78,87,88,89,132,134,148,149,150]. In addition, εHf(t) values of the Permian pegmatites are positive, whereas the coeval granites show both positive and negative values, and these data highlight mantle–crust interactions during the Permian in the Chinese Altay. Discriminant diagrams, including the Nb vs. Y and Rb vs. Y + Nb [151], show almost all Devonian granite samples in the VAG or VAG-SYN-COLG region (Figure 8). These results suggest that the granites are associated with the volcanic arc setting. This is consistent with the previous studies, which show that these Devonian granites were attributed a subduction setting [38,61,62,73,79,85,152]. The positive εHf(t) values for the Middle Devonian to Late Carboniferous granites shown in Figure 9 reveal that mantle-derived magmas associated with the slab window were added to initial magmas of the granites and pegmatites in the Devonian. The late evolution of a granitic magma that was rich in incompatible elements and volatiles elevated the rare metal contents of the pegmatites, which was crucial for the formation of deposits. Mantle-derived magma underplating of the lower crust provided heat and materials, which caused the partial melting of mature sedimentary rocks. These felsic magmas then moved upward and formed pegmatites after crystallisation. The incompatible elements and volatiles associated with coeval granites promote mobility and low solidus temperatures during the late evolution of such magmas. These components can migrate over long distances to elevate the rare metal contents of other fluids, which subsequently form ores [153].
In previous studies, a clear genetic relationship between granites and pegmatites in the Chinese Altay was not established. The genetic relationship between the Alar granite and Keketuohai No. 3 pegmatite, for example, remains a debate, whereas the Jideke and Halong granites have been disproved as parent granites of the Kelumute No. 112 pegmatite and Kaluan–Azubai–Jiamukai pegmatites, respectively. Lv et al. [22,23] therefore proposed an anatexis model for the genesis of these pegmatites. Therefore, even though pegmatites in the Chinese Altay are associated with granitic magmas, they exhibit no direct genetic relationship to coeval granites.

5.4. Implications for Magmatism in the Chinese Altay

The evolution of the Central Asian Orogenic Belt (CAOB) from the Precambrian to Late Palaeozoic remains controversial. According to previous studies, in the Early Palaeozoic, the tectonic setting in this belt was an island arc [31,154] or a passive continental margin [155] or a Precambrian microcontinent [156]. However, based on zircon dating data of granites [35,37,38,40] and petrological and geochemical data for gneissic rocks [38], the CAOB was partitioned into different tectonic units, including the active continental margin, accretionary wedge complex, fore-arc and back-arc basins, and island arcs [33,38,61,62,72,157]. According to the Wilson Cycle of plate movements, the contraction of an ocean basin begins with subduction of the oceanic crust. In the CAOB, subduction and accretionary orogeny have been experienced since the Neoproterozoic. The Late Palaeozoic collision of the Siberian and Kazakh plates marked the closure of the Irtysh–Zaisan Ocean. Owing to the closure of the ocean basin, the Irtysh Suture Zone then formed through a land-to-land collision, and subsequently, the CAOB was characterised as a post-collision extensional tectonic environment. However, the time of closure of the Irtysh–Zaisan Ocean remains under debate. According to some studies, this occurred during the Early–Middle Devonian [80,158], whereas other studies suggest that this occurred during the Late Devonian–Early Carboniferous [159,160,161], and closure in the Permian was advanced in others [162,163,164,165].
According to the zircon U-Pb dating data for pegmatites in the Chinese Altay, pegmatites in the Qinghe ore deposits area, such as the Amulagong and Tiemulete, were formed during the Devonian [23]. Others including the Xiaokalasu, Dakalasu, Yelaman, Jiamanhaba, Keketuohai No. 3 pegmatite, Kelumute No. 112, Kaluan and Asikaerte are of Permian and Triassic ages [21,44,45,46,47,48,49,50,51]. Previous studies have demonstrated that pegmatite formation commonly occurs in a post-collisional or another extensional setting [3,5,166,167,168,169,170]. Therefore, published geochronological and Hf-O isotopic data for felsic igneous rocks in the Chinese Altay were also examined (Figure 6 and Figure 9). Starting from the Neoproterozoic, the continuous subduction of the oceanic crust created frequent magmatic activities, and these attained a climax at ca. 400 Ma. A significant change in the Hf isotope composition occurred at ca. 400 Ma because before this time, εHf(t) values of zircons were both positive and negative, whereas after this time, all data from zircons analysed are positive. This implies that compositions of initial magmas significantly changed within a short period because of the input of juvenile or depleted mantle materials [33,40,62,171,172]. In addition, the high Hf isotope TDM2 ages (900–1171 Ma) and high δ18O values indicate the presence of abundant ancient sedimentary material in source magmas. In fact, previous studies suggest that the upwelling of mantle materials and heat triggered the partial melting of sedimentary rocks in the middle–lower crust to produce initial magmas of the granites and pegmatites (Amulagong and Tiemulete) [23,38,40,62]. Interestingly, Figure 9 shows obvious changes in εHf(t) and δ18O values at ca. 300 Ma. In addition, these dramatic changes in εHf(t) and δ18O values indicate a change in the tectonic setting. According to previous studies, syn-subduction magmatism in the Chinese Altay ended at ca. 313 Ma [72,83]. Furthermore, according to the recent studies, the chronology and Hf isotopic composition of detrital zircons located on the north (The Chinese Altay) and south (West Junggar and East Junggar) sides of the Irtysh suture zone not only reveal that the Irtysh suture zone is the tectonic boundary between the Chinese Altai and East Junggar but also that the Irtysh–Zaisan Ocean is closed at 323 Ma [173,174,175]. Furthermore, Zircon chronology data from structures in different phase of deformation have constrained the timing of deformation, magmatism and metamorphism events of the Irtysh suture [77,86,119,176]. Therefore, the closure of the Irtysh–Zaisan Ocean likely occurred between Late Carboniferous and Early Permian. Starting from Early Permian to Triassic, pegmatites and A-type granites were increasingly formed in Altay, which involved a post-orogenic extensional tectonic setting. The most common magmatic activity that has been reported for a post-collision tectonic setting is delamination of the lithospheric mantle beneath a thickened crust [177,178,179]. The upwelling of asthenospheric mantle materials can trigger partial melting of the lower crust and, thus, trigger a largescale magmatism [180,181,182,183,184,185]. The changes in εHf(t) and δ18O values in the study area can be attributed to magmatic source differences. The higher TDM2 ages, εHf(t), and δ18O values for Permian relative to Triassic pegmatites are consistent with an older magmatic source for the former.
Therefore, a multi-stage magmatism model is proposed for the study area from the Early Palaeozoic to Mesozoic (Figure 10), and this is summarised subsequently. (1) The Irtysh–Zaisan Ocean started declining in the Neoproterozoic, and this involved subduction of the oceanic crust on both sides of the ocean ridge into the continental crust of Siberia and Kazakhstan, respectively. However, until the slab window was formed, no clear interaction occurred between the crust and lithospheric mantle. (2) Following the formation of the slab window because of the oceanic ridge subduction, abundant materials were upwelled from mantle. The upwelling elevated the temperature field of the lower crust and altered the composition of the original magma. Partial melting of the lower crust and mixing with mantle materials produced initial magmas of the granites and pegmatites that involved positive εHf(t) values. (3) Starting from Late Carboniferous to Early Permian, the CAOB was transformed from a syn-collision orogenic to a post-collision setting. Delamination of the lithosphere promoted the upwelling of mantle materials under the extensional setting [64]. (4) Upon the onset of asthenospheric upwelling, the magmatism probably originated as deep in the crust as near the top of the Moho surface.

6. Conclusions

In the present study, the Jiamanhaba, Amulagong and Tiemulete pegmatites were characterised using zircon U-Pb geochronology and Hf and O isotopes, and the main findings can be summarised as follows:
  • The Jiamanhaba, Amulagong and Tiemulete pegmatites were emplaced correspondingly at approximately 261, 373, and 360 Ma. These results indicated that the Jiamanhaba pegmatites were of Permian age, whereas those of the Amulagong and Tiemulete were Devonian.
  • The εHf(t) values of 2.87–7.39, two-stage model ages of 900–1171 Ma and δ18O values of 9.55‰–15.86‰ that were obtained from samples of the Amulagong and Tiemulete pegmatites suggested that these were derived from anatexis of mature sedimentary rocks deep in the crust. In contrast, the Jiamanhaba pegmatite samples produced εHf(t) values of 2.87–4.94 and δ18O values of 6.05‰–7.32‰, which indicate the addition of minor amounts of mantle/juvenile materials to the original magma.
  • Using a combination of available geochronological and Hf-O isotopic data of felsic igneous, sedimentary, and metamorphic rocks that occur in the Chinese Altay, the formation of the Jiamanhaba pegmatite was assigned to a post-collision tectonic setting, where intensive mantle–crustal interactions occurred in the magma source region. In contrast, the Amulagong and Tiemulete pegmatites were formed in a syn-collision tectonic setting, and juvenile or mantle materials that were produced through the slab window were incorporated in the initial magma.
  • This paper puts forward an independent petrogenesis model of pegmatite in Chinese Altay, that is, the genesis of deep melting, which is different from the previous hypothesis that the pegmatite in Chinese Altay consists of fractionation products of granites during the late fractionation stage and provides reliable evidence for the deep melting of pegmatite. The Hf-O isotopic data of pegmatite and granites and sedimentary rocks in this region reveal the tectonic evolution of the region from Devonian to Permian, and they put forward the closing time of the Irtysh–Zaisan Ocean at about 300 Ma in the Late Carboniferous. The Hf-O isotope data of pegmatite provide a new perspective for revealing the tectonic setting of pegmatite formation, which in turn confines the petrogenesis of pegmatite. The coupling of the two relations provides a reliable guarantee for the study of petrogenesis of pegmatites in different periods in the Chinese Altay.

Author Contributions

Conceptualization, M.W. and X.Z.; methodology, M.W.; software, M.W.; validation, M.W. and X.Z.; formal analysis, M.W.; investigation, M.W.; resources, M.W.; data curation, M.W.; writing—original draft preparation, M.W.; writing—review and editing, M.W.; visualization, M.W.; supervision, X.Z.; project administration, X.Z.; funding acquisition, X.Z. All authors have read and agreed to the published version of the manuscript.

Funding

This research was funded by Regional Science Fund Project of National Natural Science Foundation of China 42062004 And The APC was funded by 42062004.

Data Availability Statement

Data will be made available on request.

Conflicts of Interest

The authors declare no conflict of interest.

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Figure 2. Photographs of the field outcrops and hand specimens of the Jiamanhaba pegmatite, Amulagong pegmatite, Tiemulete pegmatite: (a) contact relationship between the graphic assemblage and saccharoidal albite assemblage of Jiamanhaba pegmatite; (b) mineral assemblage of the nested quartz–muscovite zone of the Jiamanhaba pegmatite; (c) mineral assemblage of the saccharoidal albite zone of the Jiamanhaba pegmatite. (d) contact relationship between the saccharoidal albite assemblage and massive microcline assemblage of Amulagong pegmatite; (e) mineral assemblage of the graphic zone of the Amulagong pegmatite; (f) mineral assemblage of the saccharoidal albite zone of the Amulagong pegmatite. (g) contact relationship between the Muscovite ore and the wall rocks of the Tiemulete pegmatite; (h) mineral assemblage of the graphic zone of the Tiemulete pegmatite. (i) mineral assemblage of the saccharoidal albite zone of the Tiemulete pegmatite. Abbreviations: Ab, albite; Brl, beryl; Kfs, K-feldspar; Ms, muscovite; Qtz, quartz; Scl, schorl; Mc, microcline; Grt, garnet.
Figure 2. Photographs of the field outcrops and hand specimens of the Jiamanhaba pegmatite, Amulagong pegmatite, Tiemulete pegmatite: (a) contact relationship between the graphic assemblage and saccharoidal albite assemblage of Jiamanhaba pegmatite; (b) mineral assemblage of the nested quartz–muscovite zone of the Jiamanhaba pegmatite; (c) mineral assemblage of the saccharoidal albite zone of the Jiamanhaba pegmatite. (d) contact relationship between the saccharoidal albite assemblage and massive microcline assemblage of Amulagong pegmatite; (e) mineral assemblage of the graphic zone of the Amulagong pegmatite; (f) mineral assemblage of the saccharoidal albite zone of the Amulagong pegmatite. (g) contact relationship between the Muscovite ore and the wall rocks of the Tiemulete pegmatite; (h) mineral assemblage of the graphic zone of the Tiemulete pegmatite. (i) mineral assemblage of the saccharoidal albite zone of the Tiemulete pegmatite. Abbreviations: Ab, albite; Brl, beryl; Kfs, K-feldspar; Ms, muscovite; Qtz, quartz; Scl, schorl; Mc, microcline; Grt, garnet.
Minerals 13 01127 g002
Minerals 13 01127 g003
Figure 3. Zircon cathodoluminescence images of the Jiamanhaba, Amulagong and Tiemulete pegmatites. (a) Cathodoluminescence images of zircon grains from Jimanhaba pegmatite and their ablation locations; (b) Cathodoluminescence images of zircon grains from Amulagong pegmatite and their ablation locations; (c) Cathodoluminescence images of zircon grains from Tiemulete pegmatite and their ablation locations.
Figure 3. Zircon cathodoluminescence images of the Jiamanhaba, Amulagong and Tiemulete pegmatites. (a) Cathodoluminescence images of zircon grains from Jimanhaba pegmatite and their ablation locations; (b) Cathodoluminescence images of zircon grains from Amulagong pegmatite and their ablation locations; (c) Cathodoluminescence images of zircon grains from Tiemulete pegmatite and their ablation locations.
Minerals 13 01127 g003
Minerals 13 01127 g004
Figure 4. U–Pb concordia diagrams of the Jiamanhaba, Amulagong and Tiemulete pegmatites. (a) U–Pb concordia diagrams of Jiamanhaba pegmatite; (b) U–Pb concordia diagrams of Amulagong pegmatite; (c) U–Pb concordia diagrams of Tiemulete pegmatite.
Figure 4. U–Pb concordia diagrams of the Jiamanhaba, Amulagong and Tiemulete pegmatites. (a) U–Pb concordia diagrams of Jiamanhaba pegmatite; (b) U–Pb concordia diagrams of Amulagong pegmatite; (c) U–Pb concordia diagrams of Tiemulete pegmatite.
Minerals 13 01127 g004
Minerals 13 01127 g007
Figure 7. Chemical classification diagrams: (a) K2O + Na2O versus 10000 Ga/Al; (b) Nb versus 10000 Ga/Al; (c) Y versus 10000 Ga/Al; and (d) K2O/MgO versus 10000 Ga/Al (The data are from Sun et al. [38]; Cai et al. [61]; Zhang et al. [73]; Tong et al. [77]; Liu et al. [78]; Tong et al. [79]; Yu et al. [85]; Wang et al. [87]; Zhou et al. [88]; Gao et al. [89]; He et al. [132]; Zhang et al. [134]; Tong et al. [148]; Zhang et al. [149]; Lin et al. [150]; and Tong et al. [152]).
Figure 7. Chemical classification diagrams: (a) K2O + Na2O versus 10000 Ga/Al; (b) Nb versus 10000 Ga/Al; (c) Y versus 10000 Ga/Al; and (d) K2O/MgO versus 10000 Ga/Al (The data are from Sun et al. [38]; Cai et al. [61]; Zhang et al. [73]; Tong et al. [77]; Liu et al. [78]; Tong et al. [79]; Yu et al. [85]; Wang et al. [87]; Zhou et al. [88]; Gao et al. [89]; He et al. [132]; Zhang et al. [134]; Tong et al. [148]; Zhang et al. [149]; Lin et al. [150]; and Tong et al. [152]).
Minerals 13 01127 g007
Minerals 13 01127 g008
Figure 8. Discriminant diagrams for Devonian granites (a) Nb vs. Y diagram, and (b) Rb vs. Y + Nb diagram. (The data are collected from Cai et al. [61]; Zhang et al. [73]; Yu et al. [85]; Tong et al. [152]).
Figure 8. Discriminant diagrams for Devonian granites (a) Nb vs. Y diagram, and (b) Rb vs. Y + Nb diagram. (The data are collected from Cai et al. [61]; Zhang et al. [73]; Yu et al. [85]; Tong et al. [152]).
Minerals 13 01127 g008
Minerals 13 01127 g009
Figure 9. εHf(t) values of granites and pegmatite in four stages in the Chinese Altay (the data are collected from Zhang et al. [21]; Lv et al. [22,23,50]; Sun et al. [38,40]; Ma et al. [49]; Cai et al. [62]; Wang et al. [64]; Zhang et al. [73]; Tong et al. [77]; Liu et al. [78]; Yu et al. [85]; Zhao et al. [131]; He et al. [132]; Shen et al. [133]; and Zhang et al. [134]).
Figure 9. εHf(t) values of granites and pegmatite in four stages in the Chinese Altay (the data are collected from Zhang et al. [21]; Lv et al. [22,23,50]; Sun et al. [38,40]; Ma et al. [49]; Cai et al. [62]; Wang et al. [64]; Zhang et al. [73]; Tong et al. [77]; Liu et al. [78]; Yu et al. [85]; Zhao et al. [131]; He et al. [132]; Shen et al. [133]; and Zhang et al. [134]).
Minerals 13 01127 g009
Minerals 13 01127 g010
Figure 10. Multi-stage magmatism in the Chinese Altay from Devonian to Early Permian.
Figure 10. Multi-stage magmatism in the Chinese Altay from Devonian to Early Permian.
Minerals 13 01127 g010
Table 2. The zircon U–Pb dating results of the Jiamanhaba pegmatite, the Amulagong pegmatite, and the Tiemulete pegmatite, determined by LA-ICP-MS.
Table 2. The zircon U–Pb dating results of the Jiamanhaba pegmatite, the Amulagong pegmatite, and the Tiemulete pegmatite, determined by LA-ICP-MS.
SamplePbThU207Pb/206Pb206Pb/238U238U/232Th206Pb/238UConcordanceTh/U
ppmppmppmRatio1 SigmaRatio1 SigmaRatioAge (Ma)1 Sigma
Jiamanhaba 01183.99 346.06 4348.64 0.057361 0.001638 0.042048 0.000735 12.825747 266 4.55 89%0.080
Jiamanhaba 0263.78 121.82 1552.95 0.052802 0.001812 0.040358 0.000557 13.239468 255 3.45 97%0.078
Jiamanhaba 03116.29 173.67 2616.00 0.062280 0.001717 0.042801 0.000698 15.654055 270 4.32 83%0.066
Jiamanhaba 0493.60 267.74 1999.70 0.068614 0.002253 0.044488 0.000866 7.574216 281 5.34 75%0.134
Jiamanhaba 05109.54 188.73 2197.48 0.080826 0.004052 0.045201 0.000879 12.989826 285 5.42 65%0.086
Jiamanhaba 06112.22 1851.54 2221.95 0.095521 0.002978 0.043001 0.000630 2.312056 271 3.90 48%0.833
Jiamanhaba 07102.10 114.38 2499.22 0.051724 0.001511 0.041081 0.000686 22.546406 260 4.25 99%0.046
Jiamanhaba 08156.34 631.29 3001.38 0.128373 0.003875 0.039502 0.000562 4.944176 250 3.49 26%0.210
Jiamanhaba 09127.71 1356.84 2788.61 0.081997 0.002592 0.041318 0.000862 2.109520 261 5.34 60%0.487
Jiamanhaba 10124.62 197.33 2724.45 0.065753 0.002145 0.043838 0.001020 14.085445 277 6.30 78%0.072
Jiamanhaba 11170.92 369.89 3724.09 0.071823 0.002019 0.042634 0.000757 10.290779 269 4.68 72%0.099
Jiamanhaba 12128.46 715.75 3212.64 0.126546 0.007357 0.032398 0.000987 5.282133 206 6.16 29%0.223
Jiamanhaba 13200.14 333.14 4261.21 0.078899 0.001743 0.043126 0.000639 13.280314 272 3.95 64%0.078
Jiamanhaba 14110.68 269.43 2258.75 0.088961 0.004326 0.043104 0.000804 8.872965 272 4.97 52%0.119
Jiamanhaba 15170.81 918.64 3848.93 0.084320 0.001868 0.040264 0.000740 5.836801 254 4.59 58%0.239
Jiamanhaba 16122.76 217.49 2864.89 0.058676 0.001283 0.041919 0.000550 13.835594 265 3.41 88%0.076
Jiamanhaba 17179.89 517.26 3929.58 0.070328 0.001989 0.043096 0.000679 9.016384 272 4.20 72%0.132
Jiamanhaba 18123.40 520.78 2920.75 0.065649 0.001453 0.039962 0.000747 5.902105 253 4.63 78%0.178
Jiamanhaba 19184.86 564.75 3334.16 0.163997 0.008366 0.037974 0.000658 6.026068 240 4.09 8%0.169
Jiamanhaba 20122.68 766.15 2529.16 0.091700 0.002380 0.041449 0.000684 3.807381 262 4.24 52%0.303
Jiamanhaba 21122.33 365.77 2737.11 0.065215 0.001683 0.041792 0.000563 7.643289 264 3.49 79%0.134
Jiamanhaba 22230.68 538.50 5199.49 0.060367 0.001270 0.043736 0.000639 9.864293 276 3.95 86%0.104
Jiamanhaba 2392.71 304.48 1987.76 0.071074 0.001739 0.042786 0.000523 6.661006 270 3.24 72%0.153
Jiamanhaba 24123.06 152.18 2937.72 0.051092 0.001126 0.041639 0.000539 19.728021 263 3.34 99%0.052
Jiamanhaba 25132.48 323.06 3776.90 0.089738 0.001970 0.030039 0.000360 11.893359 191 2.25 48%0.086
Jiamanhaba 26129.15 110.43 3129.59 0.050805 0.001134 0.040989 0.000522 29.070767 259 3.23 99%0.035
Jiamanhaba 277.64 106.64 77.86 0.053176 0.003443 0.067581 0.001249 0.750263 422 7.55 96%1.370
Jiamanhaba 28116.67 146.08 2760.23 0.056010 0.001246 0.041123 0.000469 19.506242 260 2.91 92%0.053
Jiamanhaba 2977.23 120.85 1868.87 0.055988 0.001390 0.040142 0.000451 15.833978 254 2.80 92%0.065
Jiamanhaba 30283.45 271.94 6619.34 0.060757 0.001734 0.041453 0.000509 24.754284 262 3.15 84%0.041
Amulagong 0115.83 0.83 1497.69 0.054734 0.001571 0.058838 0.001381 2911.375377 369 8.41 98%0.001
Amulagong 0295.80 4.07 1534.46 0.056518 0.001645 0.058977 0.001903 612.062711 369 11.58 95%0.003
Amulagong 032.74 3.71 2304.46 0.054474 0.001666 0.060976 0.001447 642.321830 382 8.80 99%0.002
Amulagong 040.10 2.99 1777.93 0.053565 0.001941 0.061402 0.001472 628.623270 384 8.94 99%0.002
Amulagong 050.01 6.17 2559.35 0.054624 0.001675 0.060100 0.001405 424.756847 376 8.55 98%0.002
Amulagong 06416.81 7.33 2855.01 0.056667 0.001788 0.058720 0.001484 418.729516 368 9.04 95%0.003
Amulagong 071.41 6.32 2306.08 0.055047 0.001822 0.058851 0.001395 375.576182 369 8.50 97%0.003
Amulagong 0827.34 1.16 1159.36 0.055431 0.002187 0.059972 0.001409 1067.137138 375 8.57 97%0.001
Amulagong 0920.29 4.43 2782.25 0.053819 0.001593 0.060174 0.001506 644.785354 377 9.16 99%0.002
Amulagong 106.44 1.72 1317.29 0.054472 0.002047 0.059648 0.001419 807.019061 373 8.64 98%0.001
Amulagong 111.44 10.46 3837.11 0.056838 0.001774 0.060027 0.001546 374.449433 376 9.40 95%0.003
Amulagong 1214.86 4.51 1374.01 0.057896 0.001842 0.060309 0.001492 315.800502 378 9.07 94%0.003
Amulagong 130.47 1.32 1338.74 0.055119 0.001875 0.060582 0.001422 1082.378851 379 8.65 98%0.001
Amulagong 140.31 3.31 1120.51 0.056670 0.001687 0.068982 0.001576 352.629848 430 9.51 97%0.003
Amulagong 1511.04 4.11 2089.54 0.058551 0.002129 0.062841 0.001495 543.342589 393 9.07 93%0.002
Tiemulete 01740.00 696.37 3135.78 0.425251 0.013040 0.095003 0.001494 4.974462 585 8.80 −7%0.222
Tiemulete 02670.93 319.98 2773.61 0.465163 0.011346 0.094877 0.001727 8.799213 584 10.17 −10%0.115
Tiemulete 03488.70 100.55 4255.11 0.220619 0.008916 0.069567 0.000926 50.724626 434 5.59 9%0.024
Tiemulete 04127.64 21.48 2195.46 0.059284 0.001709 0.057870 0.000822 102.790913 363 5.01 91%0.010
Tiemulete 05205.14 37.64 3137.53 0.079958 0.004096 0.059860 0.000930 86.099086 375 5.66 67%0.012
Tiemulete 06197.92 25.33 3346.64 0.060191 0.001824 0.057921 0.000765 131.858292 363 4.67 90%0.008
Tiemulete 07386.66 462.48 2493.92 0.315203 0.006811 0.075770 0.001113 5.654171 471 6.67 −4%0.185
Tiemulete 081284.41 572.01 4300.98 0.540468 0.011528 0.102901 0.001842 7.829960 631 10.77 −11%0.133
Tiemulete 09730.01 144.22 4273.16 0.346816 0.008355 0.079125 0.001447 32.160684 491 8.65 −7%0.034
Tiemulete 10482.21 309.95 3812.63 0.250423 0.006939 0.070590 0.001051 13.259689 440 6.33 3%0.081
Tiemulete 11167.87 7.15 1767.59 0.148054 0.005202 0.079313 0.001958 256.452435 492 11.70 33%0.004
Tiemulete 12223.69 41.03 4563.76 0.112827 0.003811 0.069254 0.001795 199.341776 432 10.82 46%0.009
Tiemulete 139.32 6.95 2717.30 0.057503 0.001856 0.056426 0.001676 418.718268 354 10.23 93%0.003
Tiemulete 14482.54 57.66 3505.16 0.173041 0.005323 0.066851 0.001657 61.916343 417 10.01 20%0.016
Tiemulete 157.29 6.31 2838.11 0.055940 0.001825 0.055738 0.001400 458.074782 350 8.55 95%0.002
Tiemulete 1644.49 2.58 1503.83 0.126596 0.005312 0.063587 0.001844 615.414580 397 11.18 36%0.002
Tiemulete 17621.47 11.25 2240.30 0.228988 0.007939 0.073218 0.001650 207.634680 456 9.91 8%0.005
Tiemulete 180.03 3.90 4478.65 0.054815 0.001724 0.056220 0.001328 1215.092676 353 8.11 97%0.001
Tiemulete 19669.41 51.03 2395.59 0.219201 0.012017 0.076159 0.002587 56.493580 473 15.50 10%0.021
Tiemulete 2092.97 15.21 4232.11 0.111963 0.004121 0.062123 0.001507 295.115984 389 9.15 44%0.004
Table 3. Hf isotopic compositions of the Jiamanhaba pegmatite, the Amulagong pegmatite, and the Tiemulete pegmatite, determined by LA-MC-ICPMS.
Table 3. Hf isotopic compositions of the Jiamanhaba pegmatite, the Amulagong pegmatite, and the Tiemulete pegmatite, determined by LA-MC-ICPMS.
Sample176Yb/177Hf1 Sigma176Lu/177Hf1 Sigma176Hf/177Hf1 SigmaAgeεHf(t)1 SigmaTDM1TDM2
Jiamanhaba 010.021535 0.000140 0.000693 0.00000320.28269460.000012612.87 0.4 783 1100
Jiamanhaba 020.032983 0.000180 0.001146 0.00000520.28272390.00000852613.83 0.3 751 1039
Jiamanhaba 030.029732 0.000180 0.000808 0.00000390.2827410.00000992614.49 0.4 720 997
Jiamanhaba 040.024594 0.000250 0.000634 0.00000510.28275070.00000832614.87 0.3 703 974
Jiamanhaba 050.027514 0.000240 0.000781 0.0000050.28275350.00000872614.94 0.3 702 969
Jiamanhaba 060.072113 0.000320 0.002193 0.00000790.28270710.00000732613.05 0.3 797 1089
Jiamanhaba 070.021297 0.000210 0.000629 0.00000570.28272680.0000092614.02 0.3 737 1027
Jiamanhaba 080.039408 0.000063 0.001061 0.00000190.28274750.00000842614.68 0.3 716 985
Jiamanhaba 090.040564 0.000210 0.001077 0.00000440.28272980.0000082614.05 0.3 741 1025
Jiamanhaba 100.024028 0.000110 0.000704 0.00000330.28272980.00000782614.11 0.3 734 1021
Amulagong 010.003888 0.000012 0.000093 0.000000240.28274540.0000113737.25 0.4 701 909
Amulagong 020.001193 0.000004 0.000031 0.000000110.282720.00000923736.36 0.3 735 965
Amulagong 030.003669 0.000032 0.000088 0.000000920.28271150.00000873736.05 0.3 748 985
Amulagong 040.001810 0.000005 0.000046 0.000000120.28271050.00000923736.02 0.3 748 987
Amulagong 050.003921 0.000045 0.000103 0.00000120.28270690.000013735.88 0.4 754 996
Amulagong 060.007036 0.000083 0.000170 0.000002 0.28273980.0000353737.03 1.2 710 923
Amulagong 070.004397 0.000024 0.000103 0.000000660.28271080.00000883736.02 0.3 749 987
Amulagong 080.002171 0.000036 0.000052 0.000000870.2827050.00000993735.83 0.4 756 999
Amulagong 090.001989 0.000004 0.000046 0.000000170.28272660.00000923736.59 0.3 726 950
Amulagong 100.003800 0.000016 0.000094 0.000000390.28274940.0000153737.39 0.5 695 900
Tiemulete 010.031169 0.000310 0.000625 0.00000560.28269410.00000943605.02 0.3 782 1040
Tiemulete 020.028038 0.000100 0.000610 0.00000230.28264480.00000943603.28 0.3 851 1151
Tiemulete 030.028071 0.000160 0.000588 0.00000420.28263540.00000973602.95 0.3 864 1171
Tiemulete 040.047595 0.000160 0.001020 0.00000570.28269990.0000113605.13 0.4 782 1033
Tiemulete 050.037695 0.000270 0.000772 0.00000590.28269190.00000843604.90 0.3 789 1048
Tiemulete 060.026417 0.000160 0.000580 0.00000280.28272630.00000853606.17 0.3 737 967
Tiemulete 070.033801 0.000320 0.000745 0.00000650.28272180.00000963605.97 0.3 746 980
Tiemulete 080.069240 0.000460 0.001552 0.0000130.28270490.00000793605.18 0.3 787 1030
Tiemulete 090.026999 0.000910 0.000668 0.000018 0.28271150.0000143605.62 0.5 759 1002
Tiemulete 100.009072 0.000210 0.000191 0.000005 0.28269610.0000133605.19 0.5 771 1029
Table 4. O isotope compositions of Jiamanhaba, Amulagong and Tiemulete pegmatites, determined by sims.
Table 4. O isotope compositions of Jiamanhaba, Amulagong and Tiemulete pegmatites, determined by sims.
Sample18O/16OSE (%)δ18Ocorrect2SE (‰)
Jiamanhaba 010.00201880.0156.770.31
Jiamanhaba 020.00201930.0127.020.23
Jiamanhaba 030.00201740.0086.090.15
Jiamanhaba 040.00201990.0127.320.24
Jiamanhaba 050.00201730.0196.050.38
Jiamanhaba 060.00201960.0147.160.28
Jiamanhaba 070.00201810.0166.420.31
Jiamanhaba 080.00201860.0146.680.29
Jiamanhaba 090.00201830.0166.510.31
Jiamanhaba 100.00201980.0157.290.3
Amulagong 010.00202980.01712.240.33
Amulagong 020.0020270.01610.890.32
Amulagong 030.00203240.01413.580.29
Amulagong 040.00202570.01610.220.32
Amulagong 050.00203170.01313.230.27
Amulagong 060.00202790.01811.310.37
Amulagong 070.00203380.01414.270.27
Amulagong 080.00203310.01513.90.31
Amulagong 090.00202680.01410.790.29
Amulagong 100.00202430.0169.550.33
Tiemulete 010.00202960.02312.150.46
Tiemulete 020.00202860.01611.660.32
Tiemulete 030.00203170.02213.210.43
Tiemulete 040.00203510.01514.920.29
Tiemulete 050.00202830.01211.540.24
Tiemulete 060.0020290.01811.850.37
Tiemulete 070.00202820.01611.470.32
Tiemulete 080.0020370.01815.860.36
Tiemulete 090.00202880.01811.790.35
Tiemulete 100.00202860.01911.690.37
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Wang, M.; Zhang, X. Petrogenesis of Devonian and Permian Pegmatites in the Chinese Altay: Insights into the Closure of the Irtysh–Zaisan Ocean. Minerals 2023, 13, 1127. https://doi.org/10.3390/min13091127

AMA Style

Wang M, Zhang X. Petrogenesis of Devonian and Permian Pegmatites in the Chinese Altay: Insights into the Closure of the Irtysh–Zaisan Ocean. Minerals. 2023; 13(9):1127. https://doi.org/10.3390/min13091127

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Wang, Mengtao, and Xin Zhang. 2023. "Petrogenesis of Devonian and Permian Pegmatites in the Chinese Altay: Insights into the Closure of the Irtysh–Zaisan Ocean" Minerals 13, no. 9: 1127. https://doi.org/10.3390/min13091127

APA Style

Wang, M., & Zhang, X. (2023). Petrogenesis of Devonian and Permian Pegmatites in the Chinese Altay: Insights into the Closure of the Irtysh–Zaisan Ocean. Minerals, 13(9), 1127. https://doi.org/10.3390/min13091127

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