Zircon U-Pb and Hf isotope constraints on the Mesozoic tectonics and crustal evolution

Zircon U –Pb and Hf isotopic constraints from eastern Transhimalayan batholiths on the precollisional magmatic and tectonic evolution in southern TibetHan-Yi Chiu a ,Sun-Lin Chung a ,⁎,Fu-Yuan Wu b ,Dunyi Liu c ,Yu-Hsuan Liang a ,I-Jhen Lin a ,Yoshiyuki Iizuka d ,Lie-Wen Xie b ,Yanbin Wang c ,Mei-Fei Chu aaDepartment of Geosciences,National Taiwan University,Taipei 106,TaiwanbInstitute of Geology and Geophysics,Chinese Academy of Sciences,Beijing 100029,China cInstitute of Geology,Chinese Academy of Geological Sciences,Beijing 100037,China dInstitute of Earth Sciences,Academia Sinica,Taipei 115,Taiwana b s t r a c ta r t i c l e i n f o Article history:Received 1May 2008Received in revised form 15January 2009Accepted 23February 2009Available online 9March 2009Keywords:ZirconU –Pb and Hf isotope Tethyan subductionTranshimalayan batholiths TibetTectonic evolutionBefore the Indian collision with Asia,northward subduction of the Neo-Tethyan oceanic lithosphere produced an Andean-type convergent margin in South Asia characterized by arc magmatism starting from the early Jurassic and lasting until the Eocene.The magmatic arc is now represented by widespread intrusive bodies or the so-called Transhimalayan batholiths in the Lhasa terrane of southern Tibet that have been divided into two main magmatic suites,i.e.,the northern plutonic belt and the southern Gangdese Batholith.Their temporal distribution,concerning how exactly the magmatic suites correlate eastwards and then southeastwards around the eastern Himalayan syntaxis,however,remains poorly constrained.Here we report the first combined zircon U –Pb and Hf isotopic study of the Transhimalayan batholiths from the eastern part of the Lhasa terrane (∼95–97.5°E and ∼28.5–30°N),SE Tibet.Zircon U –Pb dating results of 24granitoids indicate that the rocks were emplaced principally in the Early Cretaceous (∼133–110Ma)and subordinately in the Paleocene (∼66–57Ma),long after an older granite intrusion in the earliest Jurassic (∼198Ma).The zircon εHf (T )values range from +5to −20,yielding Hf crustal model ages (T DM C )between 0.8and 2.4Ga that peak at ∼1.7Ga,suggesting a major episode of crustal growth in the Proterozoic and a predominantly,or even exclusively,crustal source for the granitoid petrogenesis.These zircon U –Pb and Hf isotopic constraints,together with whole-rock geochemical characteristics,allow us to correlate them to S-type granitoids in the northern plutonic belt,rather than to the I-type Gangdese Batholith from the central part of the southern Lhasa terrane.In contrast to the Gangdese magmatism that formed with a signi ficant juvenile mantle input related to the Neo-Tethyan subduction,these S-type granitoids show geochemical similarities to the North American Cordilleran Interior batholiths that did not form directly from a subduction setting.Thus,we attribute the petrogenesis of the widespread Early Cretaceous granitoids in the northern belt and eastern Himalayan batholiths to a postcollisional regime due to the Late Jurassic –Early Cretaceous continental collision between the Lhasa and Qiangtang terranes.Under this framework,the S-type magmatism was generated in response to collision-induced crustal thickening,which may have capabilities of causing crustal anatexis by itself in the region.However,continued interplay with the Neo-Tethyan subduction zone processes such as back-arc extension could also have played a crucial,and long-lasting,role in the magma generation.©2009Elsevier B.V.All rights reserved.1.IntroductionIn southern Tibet,the elongated belt of granitoids that occurs in the Lhasa terrane north of the Yarlu –Tsangpo suture (Fig.1)has been termed the “Transhimalayan batholiths ”(Searle et al.,1987;for review and references therein)and widely regarded as a major component of the Andean-type continental margin along South Asia resulting from northward subduction of the Neo-Tethyan oceanic lithosphere before India started colliding with Asia (e.g.,Allègre et al.,1984;Debon et al.,1986;Searle et al.,1987;Harris et al.,1988;Yin and Harrison,2000;Chung et al.,2005;Kapp et al.,2005a;Mo et al.,2005;Wen et al.,2008a ).Knowledge about the precollisional history in this region,however,remains limited because previous studies mostly concen-trated on issues of postcollisional geology in particular those related to uplift of the Tibetan plateaus and surrounding mountain ranges.As part of a systematic study of the precollisional Transhimalayan magmatism (Chu et al.,2006;Lee et al.,2007,2009-this volume;Liang et al.,2008;Wen et al.,2008a,2008b ),this paper reports for the first time a combined in-situ analysis of zircon U –Pb ages and Lu –Hf isotopes for the eastern Transhimalayan batholiths,emplaced in the easternmost part of the Lhasa terrane,the southeastern Tibetan plateau (Fig.1).The results signi ficantly improve our understanding of the age distribution and geochemical characteristics of the eastern Transhimalayan batholiths,Tectonophysics 477(2009)3–19⁎Corresponding author.E-mail address:sunlin@.tw (S.-L.Chung).0040-1951/$–see front matter ©2009Elsevier B.V.All rights reserved.doi:10.1016/j.tecto.2009.02.034Contents lists available at ScienceDirectTectonophysicsj o u r n a l h o m e p a g e :w ww.e l s ev i e r.c o m /l o c a t e /t e c towhich,together with our comparative studies from the central part of the Lhasa terrane,shed new insights into not only the magmatic but also tectonic evolution in southern Tibet before the India –Asia collision.2.Background and samplesThe Lhasa terrane is widely believed to have dispersed from Gondwana during the Permian or Triassic (Allègre et al.,1984;Chang et al.,1986),then drifted northward and finally collided with the Qiangtang terrane during the Late Jurassic –Early Cretaceous (Kapp et al.,2005a ).The terrane is bordered by the Bangong –Nujiang suture in north and by the Yarlu –Tsangpo suture in south (Fig.1),with the latter resulting from closure of the Neo-Tethys ocean owing to the continental collision between India and Asia (Yin and Harrison,2000).The Transhimalayan batholiths,distributing for a distance of ∼2500km in the Lhasa terrane,have been divided into two principal magmatic suites:(1)a southern Gangdese belt represented by the gigantic Gangdese Batholith that consists dominantly of Late Cretaceous to Eocene granotoids with I-type geochemical composition (Debon et al.,1986;Searle et al.,1987;Wen et al.,2008a ),and (2)a northern plutonic belt of peraluminous or S-type granitoids such as those exposed in the Nyainqentanghla Range that were emplaced largely in the Early Cretaceous (Xu et al.,1985;Harris et al.,1990;Kapp et al.,2005b ).It is generally thought that the Gangdese Batholith,and associated Linzizong volcanic successions (Fig.1),are the magmatic products by northward subduction of the Neo-Tethyan oceanic lithosphere under South Asia.However,the petrogenesis of the northern plutonic belt remains an issue of debates and different models have been proposed,including:(1)crustal anatexis during and/or postdating the con-tinental collision between the Lhasa and Qiangtang terranes (Xu et al.,1985;Pearce and Mei,1988),(2)high-temperature crustal melting due to upwelling of the asthenosphere after the Lhasa –Qiangtang collision (Harris et al.,1990),and (3)a low-angle,or “shallow,”subduction of the Neo-Tethyan oceanic lithosphere (Coulon et al.,1986;Ding et al.,2003;Zhang et al.,2004;Kapp et al.,2005a;Chu et al.,2006).Besides,how these magmatic belts correlate to the east,and then southeastwards around the eastern Himalayan syntaxis,the answer is uncertain because there have been few studies on the eastern Transhimalayan batholiths east of 95°E since the pioneering works carried out in the late seventies (Tu et al.,1982).Therefore,we conducted field excursions along the main roads in the area from Bomi to Basu and Chayu (Shama),between ∼95–97.5°E and ∼30–28.5°N,SE Tibet (Fig.1),through which 80+granitoids of the eastern Transhimalayan batholiths were recovered for detailed petrochemical analyses (e.g.,Lin et al.,2007).The granitoids are generally fresh looking and composed mainly of median-grained monzogranites,granodiorites and granites that contain rather sim-ple mineral constituents including K-feldspar (∼40%),plagioclase (∼30%),quartz (∼20%),biotite (∼5–10%)and accessary minerals such as hornblende,muscovite,apatite and opaque phases.24among the samples were subjected to in-situ zircon U –Pb and Hf isotopic determination by this study.For the convenience,we divide the northwest –southeast-trending Transhimalayan batholiths inthisFig.1.Simpli fied geological map showing the sample localities and distribution of major intrusive bodies in the eastern Lhasa terrane,SE Tibet (Pan et al.,2004).New zircon U –Pb age results conducted by this paper and those reported by Booth et al.(2004)are given in the parentheses and hexagons,respectively (numbers in Ma).BNS —Bangong –Nujiang suture;YTS —Yarlu –Tsangpo suture;STDS —South Tibet Detachment System.4H.-Y.Chiu et al./Tectonophysics 477(2009)3–19region into three groups,namely,from north to south,the Basu –Ranwu,Bomi –Ranwu and Chayu –Shama batholiths (Fig.1).3.Analytical methodsZircons were separated from ∼3kg samples using conventional heavy-liquid and magnetic separation techniques.Cathodolumines-cence (CL)images (Fig.2)were taken at the Institute of Earth Sciences,Academia Sinica,Taipei for examining the internal structures of individual zircon grains and selecting suitable positions for U –Pb and Lu –Hf isotope determinations.3.1.Zircon U –Pb geochronologyZircon U –Pb isotopic analyses were performed using two analytical methods,the sensitive high-resolution ion microprobe (SHRIMP)and the laser ablation-inductively coupled plasma –mass spectrometry (LA-ICPMS).These include 10age results that were obtained using the SHRIMP II equipped at the Beijing SHRIMP Center,Institute of Geology,Chinese Academy of Geological Sciences and reported in Liang et al.(2008)for tracing the source provenances of the Irrawaddy detrital zircons,and 14new age data that were obtained by the LA-ICPMS recently set up at the Department of Geosciences,National Taiwan University.The analytical procedures for the SHRIMP method were same as those reported in Wen et al.(2008a).We describe below the LA-ICPMS techniques,using an Agilent 7500s quadrupole ICPMS and a New Wave UP213laser ablation system housed at NTU.The operating conditions and data acquisition parameters are summarized in Table 1.The laser ablation was performed with a helium carrier gas that can substantially reduce the deposition of ablated material onto the sample surface and greatly improve transport ef ficiency,and thus increase the signal intensities,as compared to “conventional ”ablationusing argon as the carrier gas (Eggins et al.,1998;Günther and Heinrich,1999;Jackson et al.,2004).During the experiments,about 1min was spent for measuring gas blank and the results indicate sensitivities of less than 1000counts per second (cps)for all measured isotopes.Calibration was performed using the GJ-1zircon standard (provided by the Australian Research Council National Key Centre for Geochemical Evolution and Metallogeny of Continents,atMacquarieFig.2.Cathodoluminescence (CL)images of representative zircons separate from the studied samples.Circles indicate the positions of LA-ICPMS U –Pb age analyses,sizing 30–50µ-MC-ICPMS Hf isotope analyses were performed on the same positions with a slightly larger,∼50–65µm,laser beam size.Table 1LA-ICPMS operating conditions and data acquisition parameters.Agilent 7500s ICPMS Forward power 1500W Gas flow rate:Plasma (Ar)15L/min Auxiliary (Ar)1L/minCarrier (He)0.65–0.7L/min Make-up (Ar)1.3–1.4L/min Oxide ion (ThO +/Th +)≦0.3%New Wave UP213laser ablation system Wavelength 213nm UV Repetition rate 4and 5Hz Spot size 30–50μm Energy8–14J/cm 2Data acquisitionData acquisition protocol Time-resolved analysis (TRA)Scanning modePeak hopping,1point per peakDwell time per isotope:206Pb 15ms 207Pb 30ms 208Pb 30ms 232Th 10ms 238U15msData acquisition time130s (60s gas blank and 70s ablation)5H.-Y.Chiu et al./Tectonophysics 477(2009)3–19University,Sydney),well established for a precise207Pb/206Pb age and an intercept age using isotope-dilution thermal ionization mass spectrometry(ID-TIMS)at608.5±0.4Ma(2σ)and608.5±1.5Ma (2σ),respectively(Jackson et al.,2004).The Harvard reference zircon 91500and Australian Mud Tank Carbonatite zircon(MT)were used as secondary standards for data quality control.All U–Th–Pb isotope ratios were calculated using the GLITTER4.0(GEMOC)software,and common lead was corrected using the common lead correction function proposed by Anderson(2002).The weighted mean U–Pb ages and concordia plots were carried out using Isoplot v.3.0(Ludwig, 2003).Our LA-ICPMS results for the secondary zircon standards91500and MT,obtained during April2006and October2007,are shown in the concordia diagrams(Fig.3).The intercept,206Pb/238U and207Pb/206Pb ages of zircon91500are1061.8±2.1Ma(2σ),1062.1±1.7Ma(2σ) and1066.7±2.1Ma(2σ),respectively(Fig.3a–c),which are in perfect accordance with the consensus results reported by Wiendenbeck et al. (1995)using ID-TIMS method that gave207Pb/206Pb age at1065.4±0.6Ma(2σ)and206Pb/238U age at1062.4±0.8Ma(2σ).The intercept,206Pb/238U and207Pb/206Pb ages of zircon MT are736.8±5.2Ma(2σ), 734.0±2.4Ma(2σ)and739.2±8.1Ma(2σ),respectively(Fig.3d–f), also well match the U–Pb concordia intercept age of732±5Ma reported by Black and Gulson(1978)using ID-TIMS method. Additionally,before starting routine U–Pb dating analysis,we established three“in-house”secondary zircons,which are igneous zircons separated from three Gangdese I-type granites,samples ST129A,ST143A and ST147A,whose206Pb/238U ages have been measured using the SHRIMP method that gave94.1±2.4Ma(2σ), 84.8±1.6Ma(2σ)and50.6±0.7Ma(2σ),respectively(Wen et al., 2008a).Our LA-ICPMS results of206Pb/238U ages of these three in-house zircon standards are96.4±0.6Ma(2σ),86.2±0.6Ma(2σ)and 49.9±0.2Ma(2σ),respectively(Fig.4),which are all in good consistency with the SHRIMP results.3.2.Zircon Lu–Hf isotopic measurementsIn-situ Lu–Hf isotopic measurements were subsequently per-formed by the LA-MC-ICPMS method using a Thermo FinniganNeptune Fig.3.U–Pb concordia diagrams of the zircon standards,(a)91500and(d)MT,analyzed by the LA-ICPMS at NTU.The long-term results and weighted averages of206Pb/238U and 207Pb/206Pb ages are also shown.6H.-Y.Chiu et al./Tectonophysics477(2009)3–19multicollectors-ICPMS and a Geolas CQ 193nm laser ablation system housed at the Institute of Geology and Geophysics,Chinese Academy of Sciences,Beijing.Detailed descriptions for the analytical techniques can be found in Wu et al.(2006).The Lu –Hf isotopes were measured on the dated spots of individual zircons to minimize zoning effect but the laser ablation size is ∼50–65µm,slightly larger than that of preexisting pits,sizing 30–50mm,made by the U –Pb dating.4.Analytical results 4.1.Zircon U –Pb agesThe zircons are mostly euhedral and reveal long to short prismatic forms (Fig.2),with average crystal lengths of ∼150–300µm and length-to-width ratios from 2:1to 3:1.Most zircons are transparent,colorless to pale brown and show oscillatory zoning indicative of mag-matic growth (Hoskin and Schaltegger,2003).Zircons with rounded or oval shape and complex internal textures are rare.Thus,the interpretation of the zircon U –Pb isotopic data (see below)is simple and the obtained ages are interpreted as representing the crystal-lization time of the zircons dated and thus the emplacement timing of the host rocks.The U –Pb age results are plotted in terms of concordia diagrams in Fig.5and summarized together with the coordinates of sample localities and rock types in Table 2.These include the 206Pb/238U ages of 14samples,based on a total of 215individual zircon analyses,obtained by the LA-ICPMS method and those of 10remaining samples,on 90zircon grains,reported in Liang et al.(2008)using the SHRMIP method.Given that in zircon U –Pb isotopic analysis,precise age measurements using 207Pb/235U and 207Pb/206Pb ratios are feasible usually only for Precambrian zircons,due essentially to the fact that 235U comprises less than 1%of natural U and thus relatively little 207Pb can be produced in the Phanerozoic (cf.Ireland and Williams,2003),the weighted mean of pooled 206Pb/238U ages are taken to represent the crystallization ages of the dated samples.As shown in Fig.5,all the mean ages are given at 95%con fidence level,i.e.,2s analytical uncertainties;coupled with their mean square of weighted deviates (MSWD).All U –Pb and Lu –Hf isotopic data are listed in Table 3.In the Basu –Ranwu batholiths (Fig.5a –h),eight samples dated indicate that the rocks were emplaced mainly in the Early Cretaceous (∼130–115Ma)except a granite (sample ET223A)that shows an apparently older emplacement age of ∼198Ma (197.7±1.5Ma;Fig.5b)in the Jurassic.In the Bomi –Ranwu batholiths,there are ten samples being dated and eight of them exhibit Early Cretaceous U –Pb ages clustering around ∼125–109Ma.Two remaining samples,i.e.,ET122E and ET014C (Table 2),show younger ages in the latest Cretaceous (∼66Ma)and Paleocene (∼59Ma),respectively.Six samples from the Chayu –Shama batholiths were dated and the results show two age populations (Table 2).These are a group of four granite samples from areas around the Chayu township (Fig.1)that yielded Early Cretaceous ages (∼133Ma)and another group of two samples from areas close to Shama that show younger ages in the Paleocene (∼60Ma).4.2.Zircon Hf isotopic ratiosA total of 305dated zircon grains from the 24samples were analyzed for 176Hf/177Hf isotopic ratios.The results are listed in Table 3and e Hf data are plotted in Fig.6,8and 9.The εHf (T )value,i.e.,parts in 104deviation of initial 176Hf/177Hf isotopic ratios between the sample and the chondritic uniform reservoir,and T DM C ,the zircon Hf isotope “crustal ”model age based on a depleted-mantle source and an assumption that the protolith of the zircon's host magma has the average continental crustal 176Lu/177Hf ratio of 0.015,were calculated after Grif fin et al.(2002)by utilizing the 176Lu –176Hf decay constants adopted in Scherer et al.(2001).Our conclusions would not be signi ficantly affected if the alternative decay constants proposed by more recent studies were used.In the Basu –Ranwu batholiths (Fig.6a –h),magmatic zircons exhibit exclusively negative εHf (T )values ranging from −1.9to −18.1except an “outlier ”grain from sample ET234A that shows positive and the highest value of +0.5(Fig.6d).The zircon εHf (T )values are uniform within each sample,generally b 4−εunit variations,i.e.,about the same as the external errors of the LA-MC-ICPMS analyses (Wu et al.,2006),except the Jurassic sample ET223A that displays a larger variation and the lowest εHf (T )value of −18.1(Fig.6b).In the Bomi –Ranwu batholiths (Fig.6i –r),zircon εHf (T )values show a slightly wider range from +5.0to −15.1.In addition,there are three samples,i.e.,ET122E (Fig.6k),ET120B (Fig.6m)and ET117B (Fig.6r),containing zircons that show positive εHf (T )values.Zircons from the Chayu –Shama batholiths,have exclusively negative εHf (T )values ranging from −4.4to −20.0(Fig.6s –x).Occasionally,Fig. 4.U –Pb concordia diagrams of three in-house Gangdese zircon standards,(a)ST129A,(b)ST143A and (c)ST147A,analyzed by the LA-ICPMS at NTU.The SHRIMP results are also listed for comparison.7H.-Y.Chiu et al./Tectonophysics 477(2009)3–198H.-Y.Chiu et al./Tectonophysics477(2009)3–19Fig.5.Concordia diagrams of zircon U–Pb results of the eastern Himalayan batholiths.Note that the“mean”values of each sample are206Pb/238U ages in Ma.there are grains of inherited zircon detected in some samples,which are listed in Table 3but not plotted in Fig.6.5.Discussion5.1.Age distribution of the eastern Transhimalayan batholithsOur results,combined with five zircon U –Pb ages reported by Booth et al.(2004)from the Bomi –Ranwu batholiths (Fig.1)were plotted as an age histogram (Fig.7a)that illustrates a major intrusive stage in the Early Cretaceous (∼133–110Ma)and a subordinate stage in the Paleocene (∼66–57Ma).This,however,does not imply the absence of igneous activity during Late Cretaceous time because the current dataset is far from being complete for making a conclusion on the age distribution.In fact,Liang et al.(2004)reported a detrital zircon study along the Yarlu –Tsangpo river system showing that zircons aged ∼80–90Ma are abundant in riverbank sediments recovered from the Parlung river flowing around the Bomi township.An inference is hence that magmatic rocks of Late Cretaceous ages,although not sampled so far from the surface exposures,are likely present in the region.Collectively,the age distribution pattern of the eastern Himalayan batholiths resembles in general to that of the northern plutonic belt (Fig.7b)but differs from that of the Gangdese Batholith (Fig.7c).The latter,emplaced in the southern margin of the central Lhasa terrane,shows a predominant stage of magmatism,or a magmatic “flare-up ”event at ∼50Ma,in early Eocene time (Lee et al.,2007,2009-this volume;Wen et al.,2008a ).In addition,this study identi fies the first outcrop of Jurassic igneous rocks from the eastern Lhasa terrane,i.e.,a granite body aged ∼198Ma in the Basu –Ranwu batholiths.Regardless of the scarcity,Jurassic magmatic records that are preserved/exposed sporadically in the central part (∼84–94°E)of the Lhasa terrane have been reported in numerous recent studies using the zircon U –Pb method.These include:(1)Liang et al.(2004)that first reported the existence of abundant Jurassic inherited zircons (aged ∼194–169Ma)in a Cretaceous tuffaceous sandstone from the Gangdese belt;(2)Kapp et al.(2005b)reporting abundant Jurassic inherited zircons (∼204–181Ma)in S-type granitoids from the Nyainqentanglha Range;(3)Chu et al.(2006)reporting Jurassic inherited zircons (∼203–193Ma)in two S-type granites from the northern plutonic belt,together with identifying an ∼188Ma granite body in the Gangdese Batholith;(4)Zhang et al.(2007a,b)reporting an ∼178Ma intrusion in the Gangdese Batholith and two Jurassic S-type granites (∼205and 202Ma)in the southern margin of the northern plutonic belt;(5)Leier et al.(2007)reporting Jurassic and older detrital zircons (∼180–220Ma)in the Tekena Formation;(6)Zhu et al.(2008)reporting a Jurassic dacite (∼174Ma)in the Yeba Formation,and (7)Wu et al.(2009)reporting abundant Jurassic detrital zircons (∼190–150Ma)from the Late Cretaceous Xigaze fore-arc basin.5.2.Zircon Hf isotopic characteristics of Transhimalayan batholiths U –Pb and Hf isotope data of magmatic zircons from the eastern Himalayan batholiths are plotted in Fig.8,from which several important features are observed:(1)they are dominated by negative εHf (T )values,(2)they show a wide spectrum of εHf (T )values ranging from +5to −18,i.e.,showing an ∼23units of εHf (T )variation,(3)zircons in two Cretaceous samples from the Bomi –Ranwu batho-liths have the highest and positive εHf (T )values,in contrast to those from the Paleocene Chayu –Shama batholiths that show the lowest and negative εHf (T )values,and (4)overall speaking,the Hf isotope compositions of the Chayu –Shama batholiths appear to have become more radiogenic through time,a tendency that is unseen or absent in the Basu –Ranwu and Bomi –Ranwu batholiths.Furthermore,also shown in Fig.8,the overall distribution pattern of the eastern Himalayan batholiths emplaced in the eastern Lhasa terrane differs markedly from that of the Gangdese Batholith from the southern part of the central Lhasa terrane (Chu,2006;Chu et al.,2006,in press;Zhang et al.,2007a;Wen et al.,2008a ).Hence,in contrast to the Gangdese Batholith that requires involvement of a depleted or juvenile mantle as the major source component in the magma generation,af filiated closely with the Neo-Tethyan subduction zoneTable 2Summary of sample localities and zircon U –Pb ages of the eastern Himalayan batholiths,SE Tibet.SampleLongitude (°E)Latitude (°N)Elevation (m)Rock type Age (Ma)±2σBasu –Ranwu batholiths ET224A 96.682029.90013843granite 130.00.9ET223A 96.686929.86903870granite197.7 1.5ET225A 96.689330.00583770granodiorite 124.00.8ET234A 96.709629.76494140granite 114.80.7ET234C 96.709629.76494140enclave115.70.6ET226A 96.725230.03213668granodiorite 125.30.7ET227A 96.786830.03833462granite 122.70.7ET227C 96.786830.03833462andesite127.60.6Bomi –Ranwu batholiths ET014C a 95.070030.11002176granite58.7 3.1ET103B a 95.384529.95422636granite (deformed)118.6 2.0ET122E 95.697129.76493745granite66.1 1.1ET125B a 95.716329.75654072granite (deformed)125.1 1.5ET120B a 96.020929.74173021granite 109.0 1.1ET104A a 96.604429.50753773granite 114.8 1.7ET219A 96.851529.39224201granite 125.00.6ET218A 96.868329.38554297granite 122.50.8ET106A a 96.869029.38554207granite 122.5 1.5ET117B a97.134329.32133848granite 116.9 1.8Chayu –Shama batholiths ET113B a 97.085228.56161644granite 59.1 3.2ET207A 97.086528.56391701granite56.50.4ET115E a97.249628.59911896granite (deformed)133.1 1.3ET210A 97.366128.61912166granite (deformed)130.40.7ET215A 97.403428.98492884granite 132.90.8ET116A a97.469828.67242319granite 132.60.9aSHRIMP results from Liang et al.(2008).9H.-Y.Chiu et al./Tectonophysics 477(2009)3–19Table3U–Pb and Hf isotope data of igneous zircons from eastern Himalayan batholiths.Spot Th/U206Pb/238U±1σ207Pb/235U±1σError corr.206Pb/238Uage(Ma±1σ)176Hf/177Hf±1σ176Lu/177Hf±1σεHf(T)±1σT DM T DMCET224A wt.mean=130.0±0.9Ma(2σ)ET224A-010.81300.02030.00020.13570.00200.6840129.0 1.00.28244360.0011084−8.910.2011471752 ET224A-020.87720.02050.00020.13990.00250.5663131.0 1.00.28240260.0007781−10.290.2311941841 ET224A-030.84750.02060.00020.13900.00230.6082131.0 1.00.28240770.00107011−10.130.2611961830 ET224A-04 1.51520.02010.00020.13830.00240.6087128.0 1.00.28240470.00134213−10.340.2512101841 ET224A-070.71940.02000.00020.13560.00220.6076128.0 1.00.28240160.0008064−10.400.2211971845 ET224A-080.38610.03590.00040.26600.00320.8034227.0 2.00.28237860.0023061−9.300.2212791849 ET224A-090.78130.02020.00020.14260.00320.4844129.0 1.00.282444100.0006035−8.820.3611301746 ET224A-11 1.49250.02060.00020.14380.00280.5460131.0 1.00.28242370.0007087−9.530.2611621792 ET224A-12 1.12360.02030.00020.13180.00310.4752130.0 1.00.28241370.0009214−9.930.2511831817 ET224A-140.94340.02060.00020.13920.00390.4135131.0 2.00.28239980.0007041−10.380.2911961846 ET224A-180.72460.02050.00020.13740.00230.6208131.0 1.00.28238770.0008233−10.800.2612161873 ET224A-200.37170.07160.00070.56560.00620.8736446.0 4.00.28240280.0014007−3.670.2912131659 ET224A-23 1.26580.02070.00020.14320.00170.8004132.0 1.00.28243580.0013404−9.130.2711651768 ET224A-240.56180.10810.0011 1.03610.01130.8887661.0 6.00.28242480.0008172 1.910.2811651470ET223A wt.mean=197.7±1.5Ma(2σ)ET223A-010.03320.03090.00030.21890.00220.9609196.0 2.00.28242750.0017393−8.130.1911891752 ET223A-020.04760.03170.00030.21970.00240.8950201.0 2.00.28246690.00220418−6.690.3011471664 ET223A-040.15900.05570.00060.42510.00480.8744349.0 3.00.28217780.00111516−13.620.2815192214 ET223A-050.00780.03160.00030.21770.00260.8204201.0 2.00.282141100.00156714−18.110.3515892385 ET223A-070.03940.03180.00030.22260.00330.6834202.0 2.00.282468130.00157123−6.510.4511251654 ET223A-100.05270.05330.00060.40560.00540.8273335.0 4.00.282488100.0020216−3.130.3511101540 ET223A-110.46300.03120.00030.22200.00530.4197198.0 2.00.28243080.0021789−8.050.2811991748 ET223A-130.03880.03090.00030.22360.00250.8501196.0 2.00.282537130.0020548−4.270.4510411508 ET223A-140.87720.09850.00100.80530.00970.8389605.0 6.00.28228590.0008173−4.230.3213591815 ET223A-150.02860.03070.00030.21460.00230.8662195.0 2.00.28234870.0019994−10.960.2513111931 ET223A-160.05800.03130.00030.22420.00240.9212199.0 2.00.28241390.0015325−8.530.3112021780 ET223A-170.02730.03080.00030.21260.00230.9118196.0 2.00.282509100.00339910−5.430.3411221581 ET223A-180.02280.03080.00030.21020.00220.9537195.0 2.00.28240570.0029455−9.090.2612621812 ET223A-190.18180.03080.00030.22290.00250.8866195.0 2.00.282283110.0019946−13.290.4014052077 ET223A-200.01920.03120.00030.21610.00230.9310198.0 2.00.282498120.00275359−5.710.4311191600 ET223A-210.27250.03420.00030.23710.00250.9182217.0 2.00.28248490.0029717−5.860.3111461624 ET223A-220.05000.03130.00030.21460.00240.9054198.0 2.00.282462100.0026686−6.970.3411691680ET225A wt.mean=124.0±0.8Ma(2σ)ET225A-020.80000.01890.00020.13360.00290.4948121.0 1.00.28249170.0007551−7.330.2510691646 ET225A-030.84030.01930.00020.13490.00470.3607123.0 2.00.28248370.0007504−7.590.2510811664 ET225A-050.87720.01950.00020.12310.00460.3326125.0 2.00.28246170.0005621−8.300.2411051710 ET225A-070.90910.01940.00020.13050.00410.3751124.0 1.00.28246980.0007933−8.070.2711011695 ET225A-08 1.01010.01900.00020.13180.00470.3557121.0 2.00.28246180.0006481−8.410.2811081714 ET225A-090.68490.01930.00020.12870.00210.6244123.0 1.00.28244880.0009853−8.850.2711361723 ET225A-100.95240.01920.00020.14140.00210.6785123.0 1.00.28245780.0011991−8.530.2811301723 ET225A-110.72460.01930.00020.13350.00230.5906123.0 1.00.28247870.0009997−7.790.2610951676 ET225A-120.98040.01940.00020.13540.00400.4034124.0 1.00.28247970.0007174−7.690.2410851671 ET225A-130.91740.01990.00020.13380.00330.4468127.0 1.00.28245960.0007861−8.340.2311141714 ET225A-140.90910.01960.00020.12850.00400.3745125.0 1.00.28249980.0007671−6.970.2710581625 ET225A-150.76920.02010.00020.13900.00410.4039128.0 2.00.28247680.0005703−7.730.2810851676 ET225A-16 1.02040.01970.00020.13800.00390.4113126.0 1.00.28248980.0007823−7.330.2710731649 ET225A-170.93460.01950.00020.12770.00410.3699124.0 1.00.28246180.0006751−8.320.2711081711 ET225A-180.98040.01920.00020.13150.00480.3456123.0 2.00.28248180.0006062−7.640.2710791666 ET225A-190.88500.01950.00030.13670.00520.3393124.0 2.00.28247680.0006211−7.810.3010861678 ET225A-200.71940.01940.00020.13000.00240.5980124.0 1.00.28248870.0009076−7.420.2510781653 ET225A-21 1.28210.01940.00020.13870.00420.4068124.0 2.00.28247570.0009921−7.850.2510981681ET234A wt.mean=114.8±0.7Ma(2σ)ET234A-010.70920.01790.00020.12080.00170.6793114.0 1.00.28262270.00124112−2.890.268971358 ET234A-020.78740.01790.00020.11980.00150.7387114.0 1.00.28259570.0016313−3.870.249451420 ET234A-030.69440.01780.00020.12840.00170.7119114.0 1.00.28261970.0013026−3.000.239031365 ET234A-040.78130.01810.00020.12610.00200.6362116.0 1.00.28265080.0011818−1.860.278571294 ET234A-060.67570.01790.00020.12300.00160.7293114.0 1.00.28262270.0016344−2.940.269081361 ET234A-070.63290.01820.00020.12260.00180.6796117.0 1.00.28261860.0011975−2.970.209021365 ET234A-080.66670.01790.00020.11970.00160.7119114.0 1.00.28261180.00224716−3.360.279381388 ET234A-090.72460.01770.00020.12040.00160.7351113.0 1.00.28262670.00193932−2.840.259091354 ET234A-100.88500.01830.00020.12140.00160.7615117.0 1.00.28260780.0014925−3.400.279261393 ET234A-120.75190.01790.00020.11890.00510.2716114.0 1.00.282720150.001727110.540.517681140 ET234A-160.65790.01780.00020.11990.00170.7279113.0 1.00.28261880.00172914−3.110.289161371 ET234A-180.78130.01820.00020.12190.00150.8102116.0 1.00.28261060.0020695−3.350.239361389 ET234A-190.89290.01830.00020.12440.00180.6777117.0 1.00.28264170.0013589−2.180.268741315 ET234A-200.74630.01800.00020.12250.00180.6688115.0 1.00.28261280.0019834−3.280.289301383 ET234A-220.75190.01790.00020.12130.00160.7499114.0 1.00.28262680.00209718−2.810.279121353 ET234A-230.67110.01800.00020.12800.00240.5609115.0 1.00.28262160.00084310−2.890.218901359 10H.-Y.Chiu et al./Tectonophysics477(2009)3–19。