Lithos 212–215 (2015) 45–58

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Paleocene adakitic porphyry in the northern Qiangtang area,north-central Tibet: Evidence for early uplift of the Tibetan Plateau

Hongrui Zhang a,⁎, Tiannan Yang a, Zengqian Hou a, Jingwu Jia a,b, Maode Hu a,b, Jinwei Fan a,Mengning Dai c, Kejun Hou d

a Institute of Geology, Chinese Academy of Geological Sciences, Beijing 100037, Chinab China University of Geosciences, Beijing 100083, Chinac State Key Laboratory of Continental Dynamics, Department of Geology, Northwest University, Xi'an 710069, Chinad Institute of Mineral Resources, Chinese Academy of Geological Sciences, Beijing 100037, China

⁎ Corresponding author.E-mail addresses: zhanghr@yeah.net, hongrui_1982@1

http://dx.doi.org/10.1016/j.lithos.2014.11.0030024-4937/© 2014 Elsevier B.V. All rights reserved.

a b s t r a c t

a r t i c l e i n f o

Article history:Received 10 July 2014Accepted 2 November 2014Available online 13 November 2014

Keywords:AdakitesPaleoceneUpliftCrustal shorteningTibetan Plateau

Uplift of the Tibetan Plateau and related crustal shortening are key issues in understanding collisionalgeodynamics, and magmatic rocks that formed in this compressional setting provide clues to the processesinvolved. Numerous granitic porphyry dikes have been identified in the Angsai area of the northern Qiangtangblock in north-central Tibet. These dikes were emplaced along NW–SE-striking reverse faults that run parallelor sub-parallel to the thrust belt, suggesting that they are syn-collisional. New LA-ICP-MS zircon 206Pb/238Uand magmatic biotite 40Ar/39Ar results demonstrate that the dikes crystallized at ca. 64 Ma, and the bulk geo-chemistry of the dikes shows that they are adakitic, with high Sr/Y and La/Yb ratios, and low Y and Yb contents.Their low MgO, Ni, and Cr contents, along with positive zircon εHf(t) and bulk εNd(t) values, suggest that theadakitic rocks were the result of partial melting of a juvenile thickened lower crust. Our new data indicate theuplift of north-central Tibet which started as early as 64 Ma. Synthesizing available geochronological, geochem-ical, and structural data, it seems that the crustal shortening and uplift of the plateau is a long-lived processconsisting of a rapid early stage, and a slow and more constant later stage. Collision-related magmatic activitypeaked at the time of transition between those two stages.

© 2014 Elsevier B.V. All rights reserved.

1. Introduction

The Tibetan Plateau is the highest area on Earth with an averagealtitude of 5023 m (Fielding et al., 1994), and it plays a significant rolein controlling the global climate (e.g., Dupont-Nivet et al., 2007). Theplateau resulted from the collision of Eurasia and the Indian plate(e.g., Yin and Harrison, 2000). Thus, the timing of the start of plateauuplift is a key to understanding the collisional dynamics as well asCenozoic climate change worldwide; however, the timing remainshotly debated (e.g., Coleman and Hodges, 1995; Molnar et al., 1993;Wang et al., 2014). Numerous geological and geochemical proxieshave been proposed to constrain the timing of uplift. For example,Blisniuk et al. (2001) suggested that the development of normal faultsin central Tibet resulted from the uplift of the plateau. Numerous othersthought that the development of high-K volcanic rocks presents furtherevidence of uplift (Chung et al., 1998; Turner et al., 1993). Recently, re-searchers have started to use adakiticmagmatism as a proxy for plateauuplift (e.g., Wang et al., 2008a). The available data reveal that uplift ofthe Tibetan Plateau was a diachronous tectonic event (e.g., Chung

26.com (H. Zhang).

et al., 1998; Tapponnier et al., 2001), and that it probably started intheQiangtang area of thenorth-central Tibetan Plateau. Thus, determin-ing when the Qiangtang terrane began to be uplifted is vital to under-standing the collisional history of the Tibetan Plateau.

The presence of 46 Ma adakitic rocks in the crust of the Qiangtangarea (zircon U–Pb ages, Wang et al., 2008a) suggests that the crustwas thickened prior to this time. Paleo-altimetry studies have demon-strated that the Qiangtang area reached its current elevation after theEocene (Cyr et al., 2005; Polissar et al., 2009; Rowley and Currie, 2006;Xu et al., 2013). Recent thermochronological data (Dai et al., 2013;Rohrmann et al., 2012) have revealed that the central Tibetan Plateauunderwent rapid cooling and exhumation during the Late Cretaceous–Eocene, which was probably the result of crustal shortening accordingto structural and sedimentary studies (e.g., DeCelles et al., 2007; Kappet al., 2005; Li et al., 2012; Spurlin et al., 2005). Li et al. (2013) reportedLate Cretaceous high-K volcanic rocks in southern Qiangtang, and theyinterpreted them as evidence for volcanism in the context of a thick-ened crust. However, such an interpretation is problematic accordingto Kohn and Parkinson (2002), and it seems that the age of adakiticmagmatism may provide a better constraint on the commencement ofplateau uplift, because such magmatism is commonly interpreted asthe result of partial melting in a high-pressure environment due to

46 H. Zhang et al. / Lithos 212–215 (2015) 45–58

crustal thickening (e. g., Hou et al., 2004). Recently, we identified numer-ous Paleocene adakitic dikes in the northern Qiangtang area, and weperformed a geochronological and geochemical study of these rockswith the aim of constraining the timing and mechanism of uplift ofnorth-central Tibet.

2. Geological background

The Tibetan Plateau consists of four continental blocks separatedby three suture zones (e.g., Yin and Harrison, 2000; Yang et al., 2012;Zhang et al., 2013; Zhu et al., 2013). From north to south, they are theSongpan–Ganzi (SG) fold-and-thrust belt, the northern Qiangtang(NQ) block, the southern Qiangtang (SQ) block, and the Lhasa (LB)block. The SG and NQ are separated by the western Jinshajiang suture(JS), the NQ and SQ blocks by the Longmucuo–Shuanghu suture (LSS),and the SQ and LB blocks by the Bangonghu–Nujiang suture (BNS;Fig. 1).

The NQ block of central Tibet records a complicated tectonic evolu-tion that starts with the closing of two Paleotethyan oceans and subse-quent continent–continent collisions (Yang et al., 2014).

Two Permo-Triassic arc-like volcanic belts developed on the NQblock. The earlier one consists of 275–236 Ma andesites, dacites, rhyo-lites, and associated volcaniclastics that resulted from the northwardssubduction of a Paleotethyan ocean along the Longmucuo–Shuanghusuture (Yang et al., 2011; Zhang et al., 2013). The later arc-like volcanicbelt is Late Triassic in age, and it consists of mafic to silicic volcanics andvolcaniclastics that contain magmatic zircons with U–Pb ages of 225 to200Ma (the peak age is 217Ma; e.g., Yang et al., 2012). The Late Triassicvolcanic belt is interpreted to be the result of the southwards subductionof the Songpan–Garze Ocean along the western Jinsha–Garze–Litangsuture (Dewey et al., 1988; Yang et al., 2012). The high- to ultrahigh-pressure Qiangtang metamorphic belt (244–233 Ma) along the LSSshows that the NQ and SQ blocks merged together during the EarlyTriassic (Liu et al., 2011; Pullen et al., 2008; Zhai et al., 2011). The finalamalgamation of the SG, NQ, and SQ blocks was completed in the EarlyJurassic (e.g., Yang et al., 2012), and since then, the Qiangtang area hasbeen in an intra-continental setting.

The Cenozoic Indo-Eurasia collision produced a large amount ofcrustal shortening (Yin and Horrison, 2000; Li et al., 2012; Wang et al.,2014). Three thrust systems are recognized in the Qiangtang region:the Shiquanhe–Gaize–Amdo thrust system (SGA) along the BNS suturezone (Kapp et al., 2005, 2007), the Tanggula thrust system (TTS) alongthe LSS suture zone (Li et al., 2012), and the Fenghuoshan (FTS) andNangqian thrust belt (NTB) in the northern Qiangtang area (Cowardet al., 1988; Dewey et al., 1988; Spurlin et al., 2005). These thrusts result-ed in a total of 200 km crustal shortening. Numerous syn-contractionalbasins developed along with the thrust systems (DeCelles et al., 2007;Horton et al., 2002; Volkmer et al., 2007; Zhu et al., 2006), and the agesof the basin sediments suggest that the thrust systems developed duringthe Late Cretaceous–early Miocene (Kapp et al., 2003, 2005; Li et al.,2012; Spurlin et al., 2005).

Late Cretaceous–Cenozoic magmatic rocks are widespread in theQiangtang area (e.g., Chen et al., 2013; Chung et al., 1998, 2003, 2005;Deng, 1998; Ding et al., 2003; Hou et al., 2003; Li et al., 2013; Rogeret al., 2000; Wang et al., 2008a, 2010; Xia et al., 2011), and they formedduring three magmatic episodes at 79–60, 51–29, and 15–3 Ma, with amagmatic flare-up at ca. 40Ma. The distribution of the Late Cretaceous–Paleocene magmatic rocks in the Qiangtang area is relatively limited;the rocks include alkali basalts (Ding et al., 2003) and high-K calc-alkaline andesites (Li et al., 2013). The latter are intercalated withinLate Cretaceous conglomerates and sandstones; some authors havesuggested that they resulted from the partial melting of a mafic lowercrust (e.g., Li et al., 2013). The Paleocene alkali basalts are Na-richwith high Mg# [100 ∗ molar Mg/(Mg + Fe)] and high ratios of143Nd/144Nd isotopes, while the ratios of 87Sr/86Sr are relatively low;thus, they represent primitive mantle melts (Deng, 1998). Eocene–

early Oligocene extrusive and intrusive rocks are widely distributed inthe central part of the Qiangtang area, and they include high-Mgpotassicand shoshonitic magmatic rocks, and calc-alkaline lavas (Deng et al.,2000; Ding et al., 2003, 2007; Guo et al., 2006; Lai et al., 2003, 2007;Roger et al., 2000; Wang et al., 2008a, 2010; Xia et al., 2011). Sr–Nd–Pbisotopic data suggest that the potassic to ultrapotassic rocks were de-rived from an enriched mantle source (Deng et al., 2001; Guo et al.,2006), but themechanism of partial melting remains a subject of debate.At present, two contendingmodels have been proposed: partial meltingof an enriched upper mantle induced by subducting continental crust(Deng, 1998; Wang et al., 2008a) or partial melting of an enrichedupper mantle induced by delamination of the lithospheric mantle(Chen et al., 2013; Liu et al., 2008). The Miocene to Pliocene rocks arelimited to the northwest Qiangtang area, and they are geochemicallyultra-potassic to potassic. It is interesting that the geochemistry of thesemagmatic rocks is comparable to their equivalents in the Songpan–Ganzi region (Chen et al., 2012; Guo et al., 2006; Turner et al., 1996;Wang et al., 2005; Williams et al., 2004).

In addition to the aforementioned alkali basalts and high-K calc-alkaline andesites, granitic rocks of Paleocene age have been identifiedin various regions of northern Qiangtang (Bureau of Geology andMineral Resources of Qinghai Province, BGMRQP, 1982). Noprecise geo-chronological or high quality geochemical data are available on thesegranitic rocks. In this paper, therefore, we present new in situ zirconU–Pb and magmatic biotite 40Ar/39Ar dating results and bulk geochem-ical (including isotopic) data for these rocks, and we use these data,along with previously published information, to constrain the timingof uplift of the Tibetan Plateau.

3. Field occurrence and petrography

The Angsai dikes investigated in this study are located in the centralportion of the northern Qiangtang block (Fig. 1), a region east of ZaduoCounty, approximately 100 km southwest of Yushu City (Fig. 2). In thisarea there are outcrops of Carboniferous to Cretaceous sedimentary andvolcanic rocks, cut by several NW–SE and NE–SW trending reversefaults. These reverse faults form part of the NTB system (Spurlin et al.,2005). The Angsai granitoids, including several granitic porphyry dikes1000–6000 m long and 10–50 m wide, intrude the Cretaceous clasticrocks (Fig. 3a). They trend NW–SE, sub-parallel to the reverse faults,suggesting that the emplacement of the Angsai dikes was coeval withor slightly later than the development of the large thrust system.

The pinkish-gray granitic porphyries display obvious porphyritictextures (Fig. 3b). Phenocrysts consist of quartz, plagioclase, biotite,andminor K-feldspar, all of which are set in a groundmass ofmicrocrys-talline plagioclase and quartz (Fig. 3c, d); accessory minerals includeapatite and zircon. Very weak late-stage alteration led to the formationof small amounts of kaolinite, chlorite, and zoisite.

To determine the emplacement age and to reveal the geochemicalfeatures of the dikes, we collected several fresh samples for bulk geo-chemical analyses and for separating zircons and biotites for U–Pb andAr–Ar dating. The GPS location of the samples used for zircon U–Pb geo-chronology and biotite thermochronology is 32°47′01.9″N, 95°41′15.7″E, at an altitude of 4617 m.

4. Analytical methods

Zircon grains were separated using conventional heavy liquid andmagnetic separation techniques at the Hebei Geological Survey,Langfang City, Hebei Province, China. Cathodoluminescence imageswere obtained on a HITACH S-3000N scanning microscope fitted withGatan Chroma at the Beijing SHRIMP Center, Beijing, China, and theywere used to check the internal structures of individual zircon grains,and to select positions for analyses. In situ zircon U–Pb isotopic analyseswere conducted on a Finnigan Neptune multiple collector-inductivelycoupled plasma-mass spectrometer (MC-ICP-MS) equipped with a

Fig. 1. Simplified geological map of the Tibetan Plateau showing major blocks and the temporal–spatial distribution of Cenozoic rocks (modified from Chung et al., 2005 and Yin andHarrison, 2000). Note the Late Cretaceous–early Oligocene magmatic rocks (ages given in black numbers, Ma) and deformation structures (red numbers, Ma) in the Qiangtang area.Sources of age data: Chung et al. (1998), DeCelles et al. (2007), Ding et al. (2003), Horton et al. (2002), Kapp et al. (2005, 2007), Li et al. (2012, 2013), Spurlin et al. (2005), Turneret al. (1996), Wang et al. (2001), and Wang et al. (2008a). AKMS = Anyimaqen–Kunlun–Muztagh suture; JS = Jinshajiang suture; LSS = Longmucuo–Shuanghu suture; BNS =Bangonghu–Nujiang suture; IYS = Indus–Yarlung–Zangbo suture; SGA = Shiquanhe–Gaize–Amdo thrust system; TTS = Tanggula thrust system; FTS = Fenghuoshan thrust system;and NTB = Nangqian thrust belt. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

Fig. 2. Simplified geologic map of the Angsai area showing the locations of samples analyzed during this study.After Yang et al., 2011.

47H. Zhang et al. / Lithos 212–215 (2015) 45–58

Fig. 3. Field observations (a) and photographs (b–d) of the Angsai granitic porphyries. Abbreviations: Qtz—quartz; Kf—K-feldspar; Bt—biotite.

48 H. Zhang et al. / Lithos 212–215 (2015) 45–58

New Wave UP 213 laser-ablation system at the Institute of MineralResources, Chinese Academy of Geological Sciences (CAGS), Beijing,China. These analyses were carried out with a beam diameter of25 μm, a repetition rate of 10 Hz, and an energy of 2.5 J/cm2. ZirconsGJ-1 and Plešovice were used as internal standards during the analyses.The analytical procedures followed those described byHou et al. (2009).The age calculations and concordia diagrams were made using Isoplot/Ex ver. 3.0 (Ludwig, 2003). The results are given in Table 1.

Biotite grains were obtained by magnetic separation and hand pick-ing techniques, and then irradiated, along with the biotite standardZBH-25 (which has an 40Ar/39Ar age of 132.7 ± 1.2 Ma; Fu et al.,1987), in channel H4 of the atomic reactor 49-2 at the China Instituteof Atomic Energy for fast neutron irradiation. The irradiation durationand neutron flux were 24 h and 2.6 × 1013 n/cm2/s, respectively. The40Ar–39Ar analyses were performed using a Helix MC mass spectrome-ter at the Isotope Geology Lab, Institute of Geology, Chinese Academyof Geological Sciences (CAGS). The inter-laboratory correction factorswere as follows: (36Ar/37Ar)Ca = 2.39 × 10−4, (40Ar/39Ar)K = 4.78 ×10−3, and (39Ar/37Ar)Ca = 8.06 × 10−4. The decay constant usedthroughout the calculations was λ = (5.543 ± 0.010) × 10−10 × α−1,as recommended by Steiger and Jäger (1977). For details of the step-heating analytical processes, see Chen et al. (2006). Data were proc-essed using Isoplot/Ex_ver. 3.0 (Ludwig, 2003). Error assignments ofisotope ratios and ages are ±1σ. The results are listed in Table 2.

Major element contents were determined using an X-ray fluores-cence (XRF) instrument, while trace element concentrations were de-termined using an inductively coupled plasma-mass spectrometer(ICP-MS), both at the National Research Center for Geoanalysis, ChineseAcademy of Geological Sciences (CAGS), Beijing, China. Major element

contents have analytical uncertainties usually less than 2%, and traceelement analytical uncertainties are less than 5%–10%. The results aregiven in Table 3.

Bulk Sr–Nd isotopes were analyzed using the ID-TIMS (FinniganMAT Triton TI) at the State Key Laboratory for Mineral DepositsResearch, Nanjing University, China. The chemical separation proce-dures were similar to those described by Pu et al. (2005). The biotite40Ar/39Ar plateau age (64 Ma, see below for details) was used for calcu-lation of the initial Sr and Nd ratios. The εNd(t) values were calculatedbased on the Nd isotopic compositions of 143Nd/144Nd (CHUR) =0.512630 and 147Sm/144Nd (CHUR) = 0.1960 (Bouvier et al., 2008).

The zircon Hf isotope analyses were carried out using a quadrupolemultiple collector-inductively coupled plasma-mass spectrometer(Q-ICP-MS) connected to a single excimer laser-ablation system atthe State Key Laboratory of Continental Dynamics, Northwest University,Xi'an, China. Standards 91500, Mon-1, and GJ-1 were used during theanalyses. All analyses were carried out with a beam diameter of 44 μm,a repetition rate of 8 Hz, and an energy of 2.4 J/cm2. The analytical proce-dures were similar to those described by Yuan et al. (2008).

The measured 176Lu/177Hf ratios and the 176Lu decay constant of1.865 × 10−11 year−1 (reported by Scherer et al., 2001) were usedto calculate initial 176Lu/177Hf ratios. The chondritic values of176Hf/177Hf = 0.0332 and 176Lu/177Hf = 0.282772, reported byBlichert-Toft and Albarède (1997), were adopted for the calculationof εHf(t) values (the parts in 104 deviation of initial Hf isotope ratios be-tween the zircon sample and the chondritic reservoir). The depletedmantle Hf model ages (TDM) were calculated using the measured176Lu/177Hf ratios of zircon, based on the assumption that the depletedmantle reservoir has a linear isotopic growth from 176Hf/177Hf =

Table 1U–Pb data for zircons separated from granitic dikes, central Tibet.

Sample Pb Th U

Th/U

Isotopic ratios Age (Ma)

(ppm) (ppm) (ppm) 207Pb/206Pb 1σ (%) 207Pb/235U 1σ (%) 206Pb/238U 1σ (%) 207Pb/206Pb 1σ (%) 207Pb/235U 1σ (%) 206Pb/238U 1σ (%)

GJ–1std 38.25 9.71 324.60 0.03 0.05989 0.00019 0.80615 0.00967 0.09764 0.00118 598.17 6.63 600.29 5.44 600.56 6.96

GJ–1std 26.61 9.57 312.84 0.03 0.06039 0.00018 0.81245 0.00846 0.09756 0.00102 616.69 11.27 603.83 4.74 600.10 5.99

plesovice 93.10 65.70 459.77 0.14 0.05476 0.00028 0.40537 0.00446 0.05367 0.00054 466.71 11.11 345.54 3.23 337.01 3.29

Q59–10–1 48.84 241.04 262.75 0.92 0.04716 0.00054 0.06349 0.00097 0.00977 0.00011 57.50 25.92 62.50 0.92 62.68 0.68

Q59–10–2 130.83 848.15 453.33 1.87 0.04722 0.00036 0.06392 0.00093 0.00981 0.00012 61.21 18.52 62.91 0.89 62.95 0.76

Q59–10–3 35.78 157.52 83.01 1.90 0.04802 0.00145 0.06422 0.00215 0.00969 0.00014 101.94 67.59 63.20 2.05 62.19 0.86

Q59–10–4 44.59 224.85 116.91 1.92 0.04737 0.00126 0.06336 0.00191 0.00973 0.00015 77.87 –134.24 62.38 1.82 62.43 0.98

Q59–10–5 98.36 81.88 114.75 0.71 0.05622 0.00031 0.59681 0.00809 0.07704 0.00100 461.16 11.11 475.21 5.15 478.46 6.00

Q59–10–6 176.76 517.19 281.61 1.84 0.04858 0.00034 0.14867 0.00213 0.02219 0.00028 127.87 16.67 140.74 1.88 141.48 1.77

Q59–10–7 5.36 41.91 65.78 0.64 0.04801 0.00137 0.06299 0.00198 0.00955 0.00013 98.24 66.66 62.03 1.89 61.27 0.82

Q59–10–8 319.90 267.57 214.08 1.25 0.05697 0.00026 0.59271 0.00704 0.07546 0.00085 500.04 11.11 472.60 4.49 468.96 5.07

Q59–10–9 11.64 55.51 127.33 0.44 0.04913 0.00119 0.06529 0.00164 0.00967 0.00011 153.79 55.55 64.22 1.57 62.02 0.71

Q59–10–10 26.23 118.99 152.48 0.78 0.04803 0.00089 0.06256 0.00136 0.00942 0.00009 101.94 42.59 61.61 1.30 60.43 0.55

GJ–1std 34.07 12.60 381.31 0.03 0.05997 0.00020 0.80646 0.00894 0.09757 0.00106 611.13 7.40 600.47 5.03 600.18 6.23

GJ–1std 31.23 6.67 256.13 0.03 0.06031 0.00023 0.81214 0.00843 0.09763 0.00091 614.52 13.89 603.65 4.72 600.48 5.36

Q59–10–11 23.56 225.34 176.84 1.27 0.04790 0.00094 0.06459 0.00136 0.00980 0.00009 94.54 41.66 63.55 1.30 62.87 0.55

Q59–10–12 14.18 80.74 64.31 1.26 0.04839 0.00136 0.06304 0.00169 0.00953 0.00008 116.76 66.66 62.07 1.61 61.13 0.53

Q59–10–13 60.54 301.76 173.99 1.73 0.04693 0.00076 0.06306 0.00137 0.00973 0.00014 55.65 38.89 62.09 1.31 62.43 0.86

Q59–10–14 6.55 88.72 84.18 1.05 0.04830 0.00153 0.06262 0.00226 0.00944 0.00018 122.31 74.07 61.67 2.16 60.55 1.14

Q59–10–15 15.40 54.97 122.00 0.45 0.04873 0.00092 0.06418 0.00127 0.00960 0.00010 200.08 44.44 63.16 1.21 61.61 0.64

Q59–10–16 12.45 16.28 58.75 0.28 0.04786 0.00152 0.06214 0.00191 0.00951 0.00010 100.09 74.07 61.22 1.83 61.00 0.64

Q59–10–17 24.57 54.69 123.66 0.44 0.04727 0.00094 0.06199 0.00124 0.00955 0.00008 61.21 46.29 61.07 1.18 61.30 0.51

Q59–10–18 11.86 24.87 45.37 0.55 0.04970 0.00206 0.06480 0.00264 0.00951 0.00012 188.97 98.14 63.75 2.52 61.02 0.78

Q59–10–19 26.56 56.42 95.17 0.59 0.04662 0.00101 0.06037 0.00132 0.00944 0.00009 27.88 51.85 59.52 1.27 60.55 0.58

GJ–1std 36.32 11.03 337.19 0.03 0.06018 0.00022 0.80930 0.00843 0.09756 0.00100 609.28 2.78 602.06 4.73 600.11 5.85

GJ–1std 28.55 8.25 300.25 0.03 0.06010 0.00023 0.80930 0.00892 0.09764 0.00101 605.58 9.26 602.06 5.01 600.55 5.92

Plesovice 177.32 117.93 830.32 0.14 0.05341 0.00017 0.39538 0.00354 0.05368 0.00045 346.35 7.41 338.29 2.58 337.08 2.76

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Table 240Ar/39Ar analyses of biotites separated from the granitic porphyries of the Angsai dikes, central Tibet.Q59-10 biotite W = 35.07 mg J = 0.002246.

T(°C)

40Ar/39Ar 36Ar/39Ar 37Ar/39Ar 38Ar/39Ar 40Ar(%)

F 39Ar(×10−14 mol)

39Ar(Cum.)(%)

Age(Ma)

±1σ(Ma)

600 33.2732 0.0606 0.0000 0.0229 46.18 15.3649 0.10 0.55 61.2 3.1700 36.6481 0.0711 0.0000 0.0284 42.67 15.6378 1.32 7.48 62.27 0.69770 19.3714 0.0111 0.0000 0.0170 83.00 16.0788 2.91 22.73 64.00 0.65820 17.3375 0.0035 0.0000 0.0158 94.04 16.3049 3.45 40.80 64.88 0.65860 16.8479 0.0022 0.0000 0.0156 96.11 16.1927 3.24 57.77 64.44 0.65900 16.9076 0.0024 0.0000 0.0155 95.70 16.1813 1.23 64.23 64.40 0.73950 17.1550 0.0052 0.0000 0.0163 90.97 15.6068 0.84 68.66 62.15 0.711000 18.0354 0.0078 0.0000 0.0162 87.16 15.7194 0.48 71.19 62.6 1.11100 17.6252 0.0052 0.0000 0.0157 91.33 16.0971 2.27 83.07 64.07 0.681200 17.0471 0.0028 0.0000 0.0157 95.06 16.2045 3.18 99.74 64.49 0.641300 26.5511 0.0282 0.0000 0.0173 68.57 18.2071 0.05 100.00 72.3 5.7

Total age = 64.1 Ma.F = 40Ar⁎/39Ar, is the ratio of radiogenic Argon40 and Argon39.

50 H. Zhang et al. / Lithos 212–215 (2015) 45–58

0.279718 at 4.55 Ga to 0.283250 at present, with 176Lu/177Hf = 0.0384(Griffin et al., 2000).

5. Results

5.1. Zircon U–Pb and biotite Ar–Ar ages

The results of zirconU–Pb analyses are listed in Table 1. Zircon grainsfrom the granitic porphyry are clear, stubby to elongate prisms(100–600 μm long) with well-defined oscillatory zoning. Nineteenanalyses were conducted on 19 zircon grains. All provided concordantresults. The Th and U contents vary from 16.28 to 848.15 ppm andfrom 45.37 to 453.33 ppm, respectively, with Th/U ratios higher than0.1 (0.28–1.92). Of the 19 analyses, 16 form a tight cluster with aweighted mean 206Pb/238U age of 61.57 ± 0.34 Ma (MSWD = 1.7,Fig. 4). The ages of another three zircons are older, with their206Pb/238U ages being 141.48 ± 1.77, 468.96 ± 5.07, and 478.46 ±6.0 Ma.

The results of 40Ar/39Ar analyses of the biotites separated fromthe same granitic porphyry sample that provided the zirconsdiscussed above, are given in Table 2 and illustrated in Fig. 5. Forthis study, nine consecutive steps with 92.5% of the total 39Ar re-leased yielded a plateau age of 64.04 ± 0.5 Ma (MSWD = 1.7;Fig. 5a). The isochron age (64.3 ± 1.3 Ma; Fig. 5b) and the inverseisochron age (64.2 ± 1.3 Ma, not shown) are closely comparable tothe plateau age.

5.2. Bulk geochemistry

Six samples were analyzed for bulk geochemistry, and all sixhave similar major element contents (Table 3). They are graniticrocks (Fig. 6a) with high SiO2 (68.73–71.41 wt.%) and K2O (5.53–6.32wt.%), and low TiO2 (0.24–0.27 wt.%) and P2O5 (0.07–0.08 wt.%); theMgO contents are extremely low (0.40–0.52 wt.%) with a narrow rangein Mg# (35.06–42.78) values. These rocks plot in the shoshonitic areaon the K2O–SiO2 diagram (Fig. 6b; Peccerillo and Taylor, 1976)and geochemically they are peraluminous with A/CNK [molecularratio of Al2O3/(CaO + Na2O + K2O)] value higher than 1.0(1.1–1.9).

Primitive-mantle-normalized trace element plots (Fig. 7a) show thatthe granitic porphyries are enriched in the large-ion lithophile elements(LILEs; e.g., Ba, K, and Sr) and depleted in the high field strength ele-ments (HFSEs; e.g., Nb, Ta, P, and Ti). The chondrite-normalized rareearth element (REE) diagrams show an enrichment in light REEs(LREEs) relative to the heavy REEs (HREEs, GdN/YbN = 3.23–4.84),and weak negative Eu anomalies (Eu/Eu* = 0.67–0.85; Fig. 7b).

The Angsai dikes have a relatively narrow range of (87Sr/86Sr)i ratios(0.70642–0.70788) and (143Nd/144Nd)i values (0.512610–0.512697),the latter corresponding to very small ranges of positive εNd(t) values(+1.2 to +2.9) and model ages (TDM, 521–638 Ma; see Table 3). Suchisotopic features are different from those of the Eocene and Mioceneadakitic rocks in western Qiangtang, as shown in the εNd(t) vs.(87Sr/86Sr)i diagram (Fig. 8). On the other hand, the Sr–Nd isotopic char-acteristics of the Angsai dikes are comparable to those of Mioceneadakitic intrusives in southern Tibet (Hou et al., 2004) (Fig. 8).

5.3. Zircon Hf isotopes

Seventeen zircon grains from sample Q59-10 were selected forin situ Hf isotope analyses, including 14 magmatic and 3 xenocrysticzircons (Table 4, Fig. 9). The 176Lu/177Hf ratios of the 14 magmatic zir-cons are very low (0.000287 to 0.001490), and their 176Hf/177Hf ratiosvary from 0.282571 to 0.282764, corresponding to initial 176Hf/177Hfratios of 0.282528 to 0.282948 (64 Ma). The initial εHf(t) values have awide range from −7.27 to +7.59. Single-stage (TDM1) and two-stagemodel ages (TDM2) range from 426 to 1019 Ma and 649 to 1597 Ma,respectively. The 141 Ma xenocrystic zircon has an initial εHf(t) valueof −4.18, while the initial εHf(t) values of the other xenocrysticgrains with ages from 478 to 468 Ma are much smaller (−31.36 to−14.71).

6. Discussion

6.1. Age of magmatism

It is unusual that the U–Pb ages (ca. 61.5 Ma) of zircons from theAngsai granitic porphyry seem to be younger than the biotite40Ar–39Ar weighted mean plateau age (64 Ma), because the closuretemperature of the U–Pb isotopic system in zircon is commonlythought be much higher than that of the K–Ar system in biotite(e.g., Rollinson, 1993). One explanation of this result is that the analyzedbiotites are xenocrysts or that they crystallized earlier than the zircons;however, these possibilities are inconsistent with the pattern of the Ar–Ar plateau (Fig. 5a), and the concordant nature of the weighted meanplateau, isochron, and inverse isochron ages (Fig. 5a and b).We suggest,therefore, that the differences in ages are probably a function of thetechnical limitations of the LA-ICP-MS method. Numerous studies(e.g., Gehrels et al., 2008; Kosler et al., 2013) have pointed out that iso-topic fractionation during laser erosion (i.e., the matrix effect) does notallow a precision better than 2%. For the in situ zirconU–Pb analyses, weused the natural zircon standards GJ-1 and Plešovice as external and in-ternal references, respectively, and they were interspersed with thesamples and measured to monitor the accuracies and reproducibility

Table 3Major, trace element, and Sr–Nd isotopic data for the Angsai dikes, central Tibet.

Sample Q59-1 Q59-2 Q59-6 Q59-7 Q59-8 Q59-9

Major elements (wt.%)SiO2 71.41 68.73 71.36 70.21 70.77 68.79Al2O3 15.91 14.65 14.93 14.83 14.76 14.4Fe2O3 0.62 0.8 0.89 0.95 0.93 0.89FeO 0.59 0.52 0.59 0.49 0.49 0.52CaO 0.56 2.93 0.83 1.57 1.38 2.93MgO 0.47 0.52 0.49 0.47 0.46 0.4K2O 5.53 5.53 6.32 5.94 5.88 5.91Na2O 0.63 0.64 1.76 2.06 2.02 0.74TiO2 0.26 0.24 0.27 0.26 0.26 0.26MnO 0.01 0.01 0.02 0.01 0.01 0.01P2O5 0.08 0.07 0.08 0.08 0.08 0.07H2O+ 3.3 3.34 1.76 1.88 2.26 2.64CO2 0.34 1.64 0.17 0.78 0.26 2.24LOI 3.64 4.99 2.27 2.82 2.6 4.52Total 99.71 99.63 99.81 99.69 99.64 99.44Mg# 42.19 42.78 38.58 38.39 38.20 35.06K2O/Na2O 8.78 8.64 3.59 2.88 2.91 7.99A/CNK 1.98 1.18 1.33 1.17 1.21 1.11

Trace elements (ppm)Ba 800 791 818 869 840 905Rb 191 184 224 208 202 210Sr 364 329 456 542 507 386Y 4.24 4.31 4.63 4.95 4.89 4.73Zr 119 108 144 142 138 144Nb 3.51 3.37 4.57 4.82 4.73 4.69Th 12.2 10.7 16.3 16.9 16.8 15.7Pb 12.6 12.5 11.7 11.5 11.4 11.6Ga 20.2 18.2 19.9 19.8 19.4 19.2Zn 100 32.6 24.7 24 23.8 30.7Cu 15.4 3.47 6.41 4.09 16.5 4.65Ni 5.31 7.8 2.79 1.82 2.77 2.63V 21.7 18.9 21.2 21.2 20.7 21.7Cr 3.56 11.1 1.95 1.74 2.95 1.33Hf 3.5 3.17 4.26 4.08 4.18 4.07Cs 8.79 7.2 9.62 12.4 8.51 12.2Sc 2.6 2.39 2.26 2.4 2.37 2.01Ta 0.38 0.34 0.5 0.5 0.56 0.52Co 3 3.65 3.75 2.6 2.79 2.72Be 1.43 1.4 1.68 1.81 1.62 1.71U 1.98 2.04 2.37 2.19 1.87 2.75La 24.9 21.4 34.9 34.9 36.8 32.7Ce 44.5 38.3 62.8 60.1 66.1 58.1Pr 4.13 3.82 6.15 6.2 6.59 5.82Nd 14.9 13.5 21.1 21.5 22.5 20.1Sm 2.25 2.29 3.04 3.22 3.46 3.06Eu 0.43 0.54 0.73 0.79 0.83 0.72Gd 1.56 1.67 2.19 2.21 2.28 2.14Tb 0.21 0.2 0.24 0.24 0.27 0.23Dy 0.9 0.91 1.06 1.23 1.16 1.05Ho 0.15 0.15 0.17 0.2 0.17 0.18Er 0.44 0.42 0.48 0.51 0.49 0.47Tm 0.06 0.06 0.06 0.06 0.06 0.06Yb 0.39 0.33 0.42 0.43 0.38 0.4Lu 0.06 0.05 0.06 0.06 0.06 0.06Sr/Y 85.85 76.33 98.49 109.49 103.68 81.61Eu/Eu* 0.67 0.81 0.83 0.86 0.85 0.82(La/Yb)N 43.04 43.72 56.02 54.72 65.29 55.12(Gd/Yb)N 3.23 4.08 4.21 4.15 4.84 4.32

Sr–Nd isotope compositions(87Sr/86Sr)m 0.70788 0.70681 0.70644 0.706422SE 13 12 12 11(143Nd/144Nd)m 0.512735 0.512670 0.512656 0.5126492SE 18 8 7 10(87Sr/86Sr)i 0.70648 0.70550 0.70542 0.70536(143Nd/144Nd)i 0.512697 0.512633 0.512618 0.512610εNd(t) 2.9 1.7 1.4 1.2TDM/Ma 521 583 616 638

51H. Zhang et al. / Lithos 212–215 (2015) 45–58

of our analyses.While theweightedmean 206Pb/238U age of the internalstandard Plešovice (337.1 ± 2.1 Ma) is the same as the TIMS result(337.1 ± 0.4 Ma; Sláma et al., 2008), the weighted mean 206Pb/238U

age of GJ-1 (600.3 ± 2.4 Ma) seems to be ca. 1.5% younger than itsTIMS result (608.5 ± 0.4 Ma; Jackson et al., 2004). Thus, the weightedmean 206Pb/238U ages of the zircons may be slightly younger than the

Fig. 4. (a) Cathodoluminescence (CL) images showing the textures, spots, and respective ages of zircon U–Pb dating. (b) and (c) Zircon U–Pb concordia diagrams for the Angsai granitic dikes.

52 H. Zhang et al. / Lithos 212–215 (2015) 45–58

real age. Taking into account the fact that the Angsai granitic porphyrydikes were emplaced at a shallow level in the crust, we prefer to thinkthat the biotite 40Ar–39Ar weighted mean plateau age (64 Ma) betterrepresents the time of emplacement.

Fig. 5. 40Ar/39Ar age spectra (a) and isochron plots

6.2. Petrogenesis

The granitic porphyries exhibit very low HREE contents (Yb =0.33–0.43 ppm, Y = 4.24–4.95 ppm) but high Sr contents (329–542

(b) for biotites from the Angsai granitic dikes.

Fig. 6. (a) Na2O + K2O vs. SiO2 diagram for the Angsai dikes. (b) K2O vs. SiO2 diagram for the Angsai dikes. Data sources: Songpan–Ganzi adakitic rocks (18–15 Ma), Wang et al. (2005);Qiangtang adakitic rocks (46–38Ma), Wang et al. (2008a); Qiangtang alkali basalts (65–45 Ma), Ding et al. (2003); xenoliths from Cenozoic volcanic rocks in the Qiangtang area, Hackeret al. (2005) and Lai et al. (2011).

53H. Zhang et al. / Lithos 212–215 (2015) 45–58

ppm), giving rise to high La/Yb and Sr/Y ratios (63.85–96.84 and76.33–109.49, respectively). Thus, these rocks are geochemicallyadakitic (Defant and Drummond, 1990), as shown on the Sr/Y–Y and(La/Yb)N–YbN diagrams (Fig. 10a and b). The high Sr and very low Yand Yb contents suggest that garnet was involved as a remnant phaseat a stage of the evolutionary history of the adakitic magma, accordingto the results of high-temperature and high-pressure experiments(e.g., Rapp et al., 2002). Accordingly, several petrogenetic models havebeen proposed for the formation of adakitic rocks: (1) assimilation

Fig. 7. Primitive-mantle-normalized trace element diagrams (a) and chondrite-normalizedREE patterns (b) for the Angsai adakitic rocks from the Qiangtang area. Normalizing valuesare from Sun andMcDonough (1989) and Boynton (1984). The data for the Eocene adakiticrocks are from Wang et al. (2008a) and for the Miocene adakitic rocks from Wang et al.(2005).

and fractional crystallization (AFC) of a mantle-derived basaltic magma(Castillo et al., 1999; Macpherson et al., 2006; Richards and Kerrich,2007); (2) partial melting of a young and hot subducted oceanic crustwith subsequent interaction with the overlying mantle wedge (Defantand Drummond, 1990; Martin et al., 2005; Rapp et al., 1999; Stern andKilian, 1996); and (3) partial melting of a thickened or delaminatedlower crust (Atherton and Petford, 1993; Chung et al., 2003; Hou et al.,2004; Rapp et al., 2002, 2003; Wang et al., 2005; Xiao and Clemens,2007; Xu et al., 2002).

The AFC of basaltic magma seems to be impossible for the Paleocenegranitic porphyries of northernQiangtang. The bulk Sr–Ndand zirconHfisotopic data clearly demonstrate that the degree of crustal contamina-tion of the Qiangtang porphyries is limited, as shown by the positivebulk εNd(t) and zircon εHf(t) values. Furthermore, there is no correlationbetween 167Lu/177Hf and εHf(t) (not shown). On the La/Yb vs. La dia-gram (Fig. 10d), the La/Yb ratio increases with La content. If AFC had

Fig. 8. Initial 87Sr/86Sr vs. εNd(t) diagram for the Angsai dikes. Data sources: Songpan–Ganzi adakitic rocks (18–15 Ma), Wang et al. (2005); Qiangtang adakitic rocks(46–38 Ma), Chen et al. (2013) and Wang et al. (2008a); Qiangtang alkali basalts(65–45 Ma), Ding et al. (2003); regional Permian volcanic rocks, Zhang et al. (2013);xenoliths from Cenozoic volcanic rocks in the Qiangtang area, Deng et al. (1998), Hackeret al. (2005), and Lai et al. (2011).

Table 4Hf isotopic data for zircons separated from the Angsai dikes, central Tibet.

Sample Age 176Yb/177Hf 176Lu/177Hf 176Hf/177Hf fLu/Hf 176Hf/177Hfi 2σ εHf(0) εHf(t) TDM1 TDM2

Q59-10-01 62.68 0.033661 0.001140 0.282606 −0.97 0.282604 0.000022 −5.9 −4.56 919 1425Q59-10-02 62.95 0.018066 0.000544 0.282711 −0.98 0.282711 0.000037 −2.1 −0.78 757 1185Q59-10-03 62.19 0.017092 0.000597 0.282826 −0.98 0.282825 0.000027 1.9 3.25 597 927Q59-10-04 62.43 0.007854 0.000287 0.282885 −0.99 0.282885 0.000027 4.0 5.36 510 792Q59-10-05 478.46 0.016564 0.000720 0.282065 −0.98 0.282058 0.000032 −25.0 −14.71 1659 2379Q59-10-06 141.48 0.042665 0.001683 0.282571 −0.95 0.282566 0.000026 −7.1 −4.18 982 1460Q59-10-07 61.27 0.009376 0.000367 0.282775 −0.99 0.282774 0.000022 0.1 1.43 665 1043Q59-10-08 468.96 0.059641 0.002149 0.281613 −0.94 0.281594 0.000030 −41.0 −31.36 2365 3407Q59-10-09 62.02 0.024097 0.000867 0.282529 −0.97 0.282528 0.000024 −8.6 −7.27 1019 1597Q59-10-10 60.43 0.020475 0.000758 0.282744 −0.98 0.282744 0.000031 −1.0 0.32 715 1113Q59-10-11 62.87 0.022841 0.000834 0.282638 −0.97 0.282637 0.000028 −4.7 −3.39 865 1351Q59-10-12 61.13 0.012056 0.000390 0.282901 −0.99 0.282901 0.000023 4.6 5.90 489 756Q59-10-13 62.43 0.018609 0.000626 0.282949 −0.98 0.282948 0.000026 6.2 7.59 426 649Q59-10-14 60.55 0.020310 0.000724 0.282825 −0.98 0.282824 0.000025 1.9 3.17 601 931Q59-10-15 61.61 0.016107 0.000583 0.282814 −0.98 0.282813 0.000025 1.5 2.81 614 954Q59-10-16 61.00 0.012971 0.000501 0.282840 −0.98 0.282840 0.000025 2.4 3.74 576 895Q59-10-17 61.30 0.022293 0.000855 0.282831 −0.97 0.282830 0.000025 2.1 3.38 595 918

54 H. Zhang et al. / Lithos 212–215 (2015) 45–58

been important, there should be a strong correlation between the ratiosof some trace elements (e.g., La/Y, Dy/Yb, Sr/Y, and La/Yb) and SiO2 con-tents, but this is not the case for the Angsai samples (Fig. 10c).

Partial melting of a subducted, young, hot oceanic crust can also beruled out, given the regional geology and the geochemistry of theAngsaiadakitic rocks. For example, there has been no suggestion in the litera-ture that Paleocene oceanic crust exists in the northern Tibet region.On the contrary, the region is thought to have been in an intra-continental setting since the Early Jurassic (e.g., Yang et al., 2012). Theclosest subduction belt, between the Lhasa and southern Qiangtangblocks (part of the Eurasian continent), formed during the Late Juras-sic–Early Cretaceous (Yin and Harrison, 2000; Kapp et al., 2005, 2007;Zhu et al., 2009), and is obviously much older than the Angsai adakiticmagmatism. In the Paleocene, an ocean-slab subduction event tookplace along the southern margin of the Lhasa block, more than500 km from the northern Qiangtang block where the Angsai adakiticrocks developed. Taking Cenozoic crustal shortening of ca. 370 kminto account (60 km in the Lhasa block, 250 km in the SGA belt, and60 km by the FTS and NTB; cf. Yin and Harrison, 2000), the distance be-tween the Angsai adakitic porphyries and the Cenozoic subduction zonemust have been greater than 800 km. Studies of modern subductionzones (e.g., Stern, 2002) have demonstrated that the distance betweenan arc volcanic belt and the trench should not be greater than 166 ±60 km, while the magmatic arc itself has an average width of 97 km.Thus, it is unreasonable to propose any relationship between the sub-duction of the Neotethyan oceanic crust and the adakitic rocks in the

Fig. 9. εHf(t) vs. 206Pb/238U ages for zircons from the Angsai adakitic dikes, central Tibet.

northern Qiangtang area. Furthermore, adakites that have been derivedfrom the melting of a subducted young, hot, oceanic slab commonlyhave relatively high Na2O contents, low K2O and Th contents, and verylow K2O/Na2O ratios (Defant and Drummond, 1990; Martin et al.,2005; Moyen, 2009). In contrast, our geochemical data show that theAngsai adakitic samples have high K2O (5.53–6.32 wt.%) and Th(10.7–16.9 ppm) contents, and high K2O/Na2O ratios.

Very high SiO2 (68.73–71.41 wt.%) and high K2O (5.53–6.32 wt.%)contents, and low MgO (0.40–0.52 wt.%), Cr (1.33–11.1 ppm), and Ni(1.82–7.80 ppm) contents (Fig. 10e) suggest that the Angsai adakiticrocks resulted from the partial melting of crustal rocks (Fig. 10f). Thepositive bulk εNd(t) and zircon εHf(t) values further indicate that thecrustal rocks that underwent partial melting were juvenile. Such aninterpretation was proposed to explain the geochemistry of similarCenozoic adakitic rocks from southern Tibet, thought to have been de-rived from a garnet-bearing lower crust (e.g., Chung et al., 2003; Guoet al., 2007; Hou et al., 2004).

Numerous lower-crustal xenoliths are hosted by the Eocene–Miocene lavas in the northern Qiangtang area (Hacker et al., 2000,2005; Ding et al., 2007; Lai et al., 2011). The xenoliths are mafic, inter-mediate, or felsic with a wide range of SiO2 (44.63–75.93 wt.%) andK2O (1.58–8.03 wt.%) contents (Fig. 6b). The xenoliths include eclogite,high-pressure granulites, and glimmerite. The predominant mineralsare biotite, clinopyroxene, orthopyroxene, sanidine, quartz, plagioclase,and more importantly, garnet. The host rocks are Eocene to Miocenelavas that were derived from the partial melting of mantle rocks ratherthan crustal rocks (e.g., Deng et al., 2001; Guo et al., 2006), and thelower-crustal xenoliths cannot possibly represent the remnant phasesof the Eocene to Miocene partial melting. In other words, during theEocene–Miocene, the lower crust of the northern Qiangtang area wasgarnet-bearing, and it is reasonable to suggest that the garnet-bearingxenoliths represent the remnant phase of the partial melting of lowercrust under high-pressure conditions which gave rise to the Angsaiadakitic porphyries.

6.3. Implications

While the temperature in the crust is easily disrupted by thermalevents due to lithospheric activity, changes in pressure in the crust areexclusively related to changes in crustal thickness. The discussionsabove have demonstrated that partial melting of the lower crustunder high-pressure conditions took place in the northern Qiangtangarea during the Paleocene, and that this led to adakitic magmatism(this study) and garnet-bearing remnants (the lower crustal xenoliths;e.g., Lai et al., 2011). This implies that the crust of the northern

Fig. 10. Diagrams of Sr/Y vs. Y (a), (La/Yb)N vs. YbN (b), Sr/Y vs. SiO2 (c), La/Yb vs. La diagram (d), Ni vs. SiO2 (e), and MgO vs. SiO2 (f) for the Angsai adakitic dikes.

55H. Zhang et al. / Lithos 212–215 (2015) 45–58

Qiangtang area had been thickened at ca. 64 Ma. The crustal thicknesscan be estimated from the light/heavy REE ratios (e.g., Chung et al.,2009; Kay and Kay, 2002; Mantle and Collins, 2008). The La/Yb ratiosof the Angsai porphyries are 63.85–96.84, corresponding to a N55 kmthickness of the paleo-crust (Fig. 11). Airy isostasy theory predicts thata thickened crust would lead to a higher elevation at the surface. Thus,the uplift of central Tibet probably started as early as 64 Ma, and thisresult is consistent with the thermochronological data that revealed arapid cooling and exhumation at 85–45 Ma for the central TibetanPlateau (Fig. 12a; Rohrmann et al., 2012).

Numerous structural studies have suggested that the early Paleocenethickened crust in northern Qiangtang resulted from crustal shortening.The crustal shortening and coeval sedimentation in central Tibet (includ-ing northern Qiangtang) started during the Late Cretaceous (Dai et al.,2012; DeCelles et al., 2007; Horton et al., 2002; Kapp et al., 2003, 2005,2007; Li et al., 2012; Spurlin et al., 2005; Volkmer et al., 2007; Wanget al., 2008b). The Qiangtang area underwent regional compression dur-ing the Late Cretaceous–Eocene, resulting in three large thrust systems(i.e., TTS, FTS, and NTB), and the Angsai dikes were emplaced along

NW–SE-striking reverse faults at ca. 64 Ma. This temporal and spatial re-lationship suggests that the dikes were emplaced in a setting of regionalcompression.

Just like the long-lived regional compression, at least three stagesof adakitic magmatism have been identified in the Qiangtang area:the Paleocene Angsai granitic porphyries (this study); the Eoceneperaluminous and metaluminous lavas in the western Qiangtangarea, 600 km west of the Angsai dikes; and the Miocene K-rich lavasin the Hohxil area, 800 km northwest of the Angsai dikes. The Eocenelavas were mainly erupted at 46–38 Ma, and they have a wide rangeof SiO2 (55–73 wt.%) and K2O (2.46–4.96 wt.%) contents, and relativelysmall K2O/Na2O ratios (0.63–1.80) (Chen et al., 2013; Wang et al.,2008a); they have been divided into peraluminous (smaller Mg# valuesof 20–50) and metaluminous (higher Mg# values of ca. 46–69) groupsof lava (Wang et al., 2008a). The Miocene (18–15 Ma) K-rich lavas aredistributed sporadically within the western Qiangtang area, but arewidespread in the Songpan–Ganzi region. They have a narrow rangeof SiO2 (61–67wt.%) and K2O (4.02–4.68wt.%) contents, and also a nar-row range of K2O/Na2O values (1.10–1.27) (Wang et al., 2005).

Fig. 11. La/Yb vs. age for magmatic rocks in the Qiangtang area. Data sources: magmatismin an intra-continental setting, Ding et al. (2003), Li et al. (2013), Spurlin et al. (2005),Wanget al. (2005, 2008a, and 2010), and this study;magmatism related to slab subduction, Yanget al. (2011) and Zhang et al. (2010, 2011, and 2013). The crustal thickness correlation isinferred after Chung et al. (2009).

56 H. Zhang et al. / Lithos 212–215 (2015) 45–58

The structural and geochemical studies mentioned above suggesta long-lived crustal shortening and thickening history for northernTibet, starting at least at around 64 Ma (this study) and continuingto the present time. To graphically display the crustal shortening

Fig. 12.Age distribution of Cenozoicmagmatic rocks and cooling curves from theQiangtangarea. Ages are from Chen et al. (2013 and references therein), Ding et al. (2003), Guo et al.(2006 and references therein), Hou et al. (2003 and references therein), Li et al. (2013),Roger et al. (2000), Spurlin et al. (2005), Wang et al. (2008a, 2010), and this study. Coolingcurves (bold lines in 12a) are from Dai et al. (2013) and Rohrmann et al. (2012).

and thickening history of northern Tibet, the available geo/thermochronological and geochemical data for the late Mesozoic toCenozoic magmatic rocks of the Qiangtang area are summarized inFig. 11. The La/Yb ratios are used as an index for the thickness ofthe paleo-crust (Kay and Kay, 2002; Mantle and Collins, 2008). Thedata (Fig. 12a) show that the thickness of the crust increased rapidlyduring the period from ca. 80 to 45 Ma. After 45 Ma, the thickness ofthe crust showed little change. It is interesting that the age distributionofmagmatic rocks related to crustal shortening displays amajor peak ataround 40 Ma (Fig. 12b). This means that the long-lived continent–continent collision had a peak of collision-related magmatism immedi-ately after the period of rapid crustal shortening. Further study is requiredto reveal the geodynamic implications of this phenomenon.

7. Conclusions

(1) The Angsai granitic porphyry dikes in the northern Qiangtangblock, north-central Tibet, have the geochemical characteristicsof adakitic rocks, as indicated by high Sr and low Y and Yb con-tents, and high Sr/Y and La/Yb ratios.

(2) Zircon U–Pb and biotite 40Ar–39Ar dating demonstrates that theadakitic magmas crystallized in the early Paleocene (64 Ma).

(3) The early Paleocene adakitic magmas were probably the result ofthe partial melting of a thickened juvenile lower crust, which sug-gests that the uplift of northern Tibet started as early as 64 Ma.

(4) The processes of crustal shortening and uplift of the TibetanPlateau seem to have been long-lived, with a rapid early stageand a relatively slow but continuous later stage. Collisional mag-matic activity peaked at the timeof transition between those twostages.

Acknowledgments

This study was supported financially by the State Key ResearchDevelopment Program of China (973, No. 2015CB452601), the NSFC(programs 41472067 and 41320104004), the International GeologicalCorrelation Program (IGCP/SIDA-600), and the Geological Survey ofChina (1212011220908).

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