Journal of Petrology | Volume 40 | Number 9 | Pages 1399-1424 | 1999
© Oxford University Press 1999
Post-Collisional Potassic and Ultrapotassic Magmatism in SW Tibet: Geochemical and Sr–Nd–Pb–O Isotopic Constraints for Mantle Source Characteristics and Petrogenesis
1 Institut Für Mineralogie Und Petrographie Innrain 52, A-6020 Innsbruck, Austria
2 Institut Für Geologie Althanstrasse, A-1090 Wien, Austria
Received August 24, 1998; Revised typescript accepted April 12, 1999
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ABSTRACT |
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 TOP ABSTRACT IntroductionGeological SettingAnalytical TechniquesAge of the Post-Collisional...Sample Description and...Geochemical CharacteristicsDiscussionConcluding RemarksReferences |
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Major and trace element, Sr–Nd–Pb–O isotope and mineral chemical data are presented for post-collisional ultrapotassic, silicic potassic and high-K calc-alkaline volcanic rocks from SW Tibet, with 40Ar/39Ar ages in the range 17–25 Ma. The ultrapotassic lavas contain mantle xenocrysts (olivine ± rutile/armalcolite). Their initial 87Sr/86Sr (0.7172–0.7220) and 143Nd/144Nd (0.51190–0.51200) ratios suggest that they originated from lithospheric sources enriched in Rb with low Sm/Nd ratios. Initial Pb isotopic compositions (206Pb/204Pb = 18.41–18.51; 207Pb/204Pb = 15.68–15.72; 208Pb/204Pb = 39.42–39.60) and geochemical features such as high Th/Ta, low Sr/Nd, low Ce/Pb and negative Eu anomalies are consistent with a recycled crustal component. Nd depleted mantle model ages range from 1.3 to 1.9 Ga, whereas Pb model ages record an Archaean event, suggesting that the source had a complex multi-stage evolution. In contrast, the high-K calc-alkaline dacites and rhyolites have less enriched initial Sr (0.7091–0.7097) and Nd (0.51213–0.51225) isotopic compositions. The presence of zircon xenocrysts with a Pb-evaporation age of 471 ± 33 Ma documents the importance of crustal anatexis in their genesis. Processes responsible for the partial melting of metasomatized lithospheric mantle and post-collisional magmatism in the Lhasa block could be a consequence of (1) convective removal of the lower lithosphere or (2) of slab breakoff.
KEY WORDS: lithospheric mantle; Sr–Nd–Pb-isotopes; SW Tibet; ultrapotassic volcanism
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Introduction |
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TOPABSTRACTIntroductionGeological SettingAnalytical TechniquesAge of the Post-Collisional...Sample Description and...Geochemical CharacteristicsDiscussionConcluding RemarksReferences |
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The high plateau of Tibet, the Himalaya and the Karakoram Ranges are the product of the continuing collision of India with the southern margin of Eurasia during the past 50 my (e.g. Klootwijk & Radhakrishnamurty, 1981
). Despite numerous studies of the region, the processes responsible for the formation of the plateau and the surrounding mountain ranges are still controversial. Three hypotheses have been suggested to explain the crustal thickening and altitude of Tibet. In the first (e.g. Argand, 1924; Powell & Conaghan, 1975; Ni & Baranzagi, 1983), virtually the entire Tibetan plateau is underthrust by Indian lithosphere. In the second (e.g. Zhao & Morgan, 1985, 1987), thickening is by inflow of material from India. In the third, the crust of Tibet is interpreted as having thickened by shortening (e.g. Dewey & Burke, 1973). Quantitative analysis of the third model, assuming a vertically average lithosphere rheology, predicts lithospheric thickening and additional uplift of the entire plateau by convective thinning of the lower continental lithosphere (e.g. England & Houseman, 1988, 1989). Molnar et al., (1993) argued that a sudden increase in uplift occurred at 
8 Ma and that the start of post-collisional potassic basaltic volcanism in northern Tibet is a result of rapid heat transfer to the mid-lithosphere caused by removal of the lithospheric root.
Neogene volcanic rocks are also present in the southern part of the Tibetan plateau (e.g. Coulon et al., 1986; Arnaud et al., 1992; Turner et al., 1993, 1996). Compositionally, this volcanism appears to be bimodal, mafic and acidic, with both mantle and crustal sources. In SW Tibet the post-collisional volcanism encompasses calc-alkaline, potassic and ultrapotassic compositions. Conventionally, potassic rocks are defined as those in which K2O exceeds Na2O (wt % or molar). Ultrapotassic igneous rocks were defined by Foley et al., (1987) as having high contents of K2O > 3 wt %, MgO > 3 wt %, and K2O/Na2O > 2 (wt %). Based on chemical characteristics they defined three subgroups, corresponding to different geodynamic settings; Group I (anorogenic lamproites) are products of intracontinental plate magmatic activity; Group II (kamafugites) are associated with continental rifts such as the Uganda segment of the East African Rift; Group III ultrapotassic rocks, which include eruptives from Italy (e.g. Hawkesworth & Vollmer, 1979; Rogers et al., 1985), occur during or after continental collision following ocean-basin closure. The trace element and isotopic characteristics of these K-rich volcanic rocks are often more enriched than those of basalts presumed to be derived from the convecting asthenosphere. The petrogenetic and geodynamic significance of their unusual chemical characteristics remains controversial (e.g. Mitchell & Bergman, 1991; Foley & Peccerillo, 1992).
In central Asia, ultrapotassic Neogene–Recent volcanic rocks have been previously reported from the northern Karakoram (Pognante, 1990) and from NW Tibet (Pearce & Mei, 1988). This paper documents, for the first time, the presence of ultrapotassic rocks in SW Tibet and presents detailed chemical and isotopic data on the timing and composition of these and other volcanic rocks in SW Tibet that post-date the India–Asia collision. It then discusses their origin within the context of models of lithospheric evolution.
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Geological Setting |
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TOPABSTRACTIntroductionGeological SettingAnalytical TechniquesAge of the Post-Collisional...Sample Description and...Geochemical CharacteristicsDiscussionConcluding RemarksReferences |
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The Tibetan plateau is the largest uplifted structure on Earth. It is bounded to the north by the Altyn Tagh fault and the Tarim block and to the south by the Himalayas. Forming in response to the India–Asia collision, it consists of a series of east–west trending crustal blocks that were successively accreted to the southern margin of Eurasia since the Early Palaeozoic. The southernmost of these continental fragments is the Lhasa Block, which separated from Gondwana in the Triassic and collided with Eurasia in the Middle to Late Jurassic, forming the Banggong suture (e.g. Dewey et al., 1988
). Crustal thickening in southern Tibet (e.g. Brandon & Romanowicz, 1986; Molnar, 1988) was followed by E–W extension along N–S trending normal-fault systems (e.g. Armijo et al., 1989). Volumetrically small amounts of dominantly mafic volcanism have occurred within Tibet since the Miocene (e.g. Coulon et al., 1986; Pearce & Mei, 1988; Arnaud et al., 1992; Turner et al., 1993, 1996).
The major geological elements in SW Tibet (Fig. 1) include (1) a Cretaceous to Eocene Andean-type subduction-related magmatic arc (e.g. Debon et al., 1986) consisting of the Transhimalaya Batholith (THB) and the Linzizong volcanic rocks, (2) the Indus–Yarlung suture zone (IYS) and (3) the right-slip Karakoram fault.
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Potassic and ultrapotassic lava flows and pyroclastic rocks (Table 1) were sampled in the volcanic fields south of Xungba and south of Bongba (Fig. 1). The volcanic rocks are underlain by shallow marine and terrigenous sediments of the Cretaceous Takena Formation and by rhyolitic volcanic rocks of the Linzizong Formation. Ten km SSE of Xungba an ultrapotassic lava flow overlies a granite dated at
116 Ma (Table 2). This observation indicates that the northern part of the THB had already been exhumed and eroded by the time of the ultrapotassic volcanism. East of Jarga, ultrapotassic volcanic rocks form an E–W trending dyke (Fig. 1). In addition, post-collisional calc-alkaline dacites and rhyolites (Table 1) have been encountered in the Gegar volcanic field 60 km NNE of Mt Kailas and, redeposited as debris flows 8 km south of Barga, along the shore of Lake Manasarowar (Fig. 1).
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Analytical Techniques |
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TOPABSTRACTIntroductionGeological SettingAnalytical TechniquesAge of the Post-Collisional...Sample Description and...Geochemical CharacteristicsDiscussionConcluding RemarksReferences |
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Major elements, F, Cl, Sc and Ga were determined on powder pellets by X-ray fluorescence (XRF) using standard methods. Other trace elements and rare earth elements (REE) were determined by inductively coupled plasma mass spectrometry (ICP-MS) at the Centre de Recherches Pétrographiques et Géochimiques (Vandoeuvre-les-Nancy, France).
Mineral and glass composition data were obtained with an ARL SEMQ electron microprobe by energy- and/or wavelength-dispersive spectrometry at the University of Innsbruck. The accelerating voltage was 15 kV and sample current 20 nA. Natural and synthetic standards were used for calibration.
Minerals for isotopic analysis were handpicked. Chemical sample digestion and element separation follows the procedure outlined by Thöni & Jagoutz, (1992). Overall blank contributions are 
0.2 ng for Nd and Sm, and 2 ng for Rb and Sr. Nd and Sm concentrations were determined by isotope dilution, using a mixed 147Sm–150Nd spike, or by ICP-MS for some whole-rock samples and run as metals on a Finnigan MAT 262 multicollector mass spectrometer. Nd was ionized using a Re double filament. Within-run isotope fractionation was corrected for 146Nd/144Nd = 0.7219. All errors quoted in Table 8 (see below) correspond to 2
of the scan mean. The 143Nd/144Nd ratio for the La Jolla international standard during the course of this investigation was 0.511846 ± 8 (35 runs). Errors for the 147Sm/144Nd ratio are ±1%, or smaller, based on iterative sample analysis and spike recalibration. The following model parameters were used for the calculation of depleted mantle (DM) ages: 147Sm/144Nd = 0.222, 143Nd/144Nd = fMitchell0.513114 (Michard et al., 1985). A linear evolution of the Nd isotope composition of the DM is assumed throughout geological time; 
Nd values are calculated relative to CHUR. Sr and Rb concentrations were determined using a VG Micromass MM 30 and Ta filaments. Through the course of this study the value for the NBS 987 Sr standard was 0.71011 ± 1. Maximum errors for 87Rb/86Sr ratios are estimated to be ±1%.
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For Pb analysis, 200 mg of rock powder was dissolved using bi-distilled hydrofluoric and perchloric acids in Teflon bombs. Pb was separated from the silicate matrix using standard HCl ion exchange procedures. The Pb isotopic composition was measured on a Finnigan MAT 262 mass spectrometer in static mode using Re single filaments and the silica-gel–phosphoric acid technique. Mass fractionation was corrected off-line using correction factors derived from NBS SRM 981 and 982 standard measurements. Mass fractionation was 0.20%/a.m.u. Total procedural Pb blanks amounted to 1–3 ng and were not corrected for. Measured Pb isotope ratios were corrected for in situ decay using U, Th and Pb concentrations determined by ICP-MS and rock formation ages given by the Ar/Ar ages of the samples in question. Single zircon dating followed modified procedures of the method of Kober, (1987)
. Details of the technique applied have been summmarized by Klötzli, (1997).
For Ar–Ar age determinations the mineral concentrates were irradiated at the 9 MW ASTRA reactor at the Austrian Research Centre Seibersdorf. Ar was released at progressively higher temperatures using a radiofrequency induction furnace and low-blank tantalum capsules, and analysed with a VG-5400 Fisons rare-gas mass spectrometer. Details of the technique have been given by Frimmel & Frank, (1998). Ages were calculated after corrections for mass discrimination and radioactive decay using the formulae given by Dalrymple et al., (1981). J values were determined with internal laboratory standards, calibrated by international standards including muscovite Bern 4M (18.555 ± 0.395 Ma; Burghele, 1987) and amphibole MMhb-1 (520.4 ± 1.7 Ma; Samson & Alexander, 1987). The errors given on the calculated age of an individual step include only the 1 error of the analytical data. The error of the plateau and total gas ages includes an additional error of ±0.4% on the J value, based on standard reproducibility.
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Age of the Post-Collisional Lavas |
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TOPABSTRACTIntroductionGeological SettingAnalytical TechniquesAge of the Post-Collisional...Sample Description and...Geochemical CharacteristicsDiscussionConcluding RemarksReferences |
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As part of this study, 40Ar/39Ar and Rb/Sr age data were obtained on both mineral separates and whole rocks (Table 2). The ultrapotassic and potassic rocks yield ages that fall within the range 18–25 Ma. The Gegar and Manasarowar calc-alkaline lavas and pyroclastic rocks were erupted
16–17 my ago. Phlogopite phenocrysts from ultrapotassic lava flow TE011/93 (south of Xungba) give a reasonably well-defined plateau age of 23.0 ± 0.3 Ma (steps 3–10; 95% of 39Ar released), in good agreement with the Rb–Sr WR (whole-rock) phlogopite date of 22.4 ± 0.3 Ma (Fig. 2a and b. In contrast, the matrix is slightly younger (21.0 ± 0.4 Ma), with a complex Ar–Ar spectrum as a result of devitrification processes of the matrix glass (Fig. 2c). An associated trachytic sample (TE25/93) yielded a plateau age of 22.8 ± 0.2 Ma (steps 5–11; 83% of 39Ar released; Fig. 2d). The Ar–Ar spectrum obtained on phlogopite TE138/98 is complex (Fig. 2e). The first five steps provide only 2% of 39Ar released. They have no age relevance because of their low intensity. A reasonable minimum age of 18.1 ± 0.3 Ma is defined by four steps (39% of 39Ar released). There is no evidence for a major excess Ar component because the Rb–Sr (WR-Phlog) date yielded similar results (Fig. 2f). Phlogopites extracted from ultrapotassic volcanic rocks TE117/93 and TE118/93 east of Jarga gave similar plateau ages of 18.5 ± 0.4 Ma and 18.3 ± 0.4 Ma, respectively (Fig. 2g and h. The younger total gas age of 15.5 ± 0.6 Ma obtained on the matrix of TE117/93 (Table 2) could be explained by Ar loss caused by its very fine grain size. Therefore the total gas ages of 23.2 ± 0.2 Ma and 25.4 ± 0.2 Ma for the glassy matrix of potassic samples TE148/93 and TE150/93, respectively, could also indicate minimum ages. The higher ages in the low-temperature steps (Fig. 3a and b may indicate incorporation of excess Ar or represent 39Ar recoil effects caused by the fine grain size, as commonly observed in fine-grained volcanic rocks.
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The biotite of dacite TE194/93 (Fig. 3c) gives a well-defined age plateau of 16.7 ± 0.2 Ma (steps 2–18; 96% of 39Ar released), in good agreement with the plagioclase total gas age of 17.0 ± 2.0 Ma (Fig. 3d) and the date of 17.0 ± 0.3 Ma obtained by the Rb–Sr method for the biotite-WR isochron (Fig. 3e). Rhyolite TE192/93 gives ages in the same range (Fig. 3e and f).
The biotite of rhyolite TE47/93 gives an age plateau of 17.7 ± 0.9 Ma (Fig. 3h). It is characterized by Ar loss in the low-temperature steps as a result of exsolution of rutile and opaque oxide. The plagioclase of this sample is slightly discordant, with a total gas age of 16.2 ± 1.1 Ma (Fig. 3g). In contrast, zircons separated from this sample and analysed by the single zircon method (Klötzli, 1997) yielded a Pb–Pb age of 471 ± 33 Ma (Table 3), clearly indicating their xenocrystic nature.
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40Ar/39Ar data on other post-collisional volcanic rocks from the southern part of the Tibetan plateau are limited. Coulon et al., (1986)
reported 40Ar/39Ar ages in the range of 10.1–15.8 Ma for mineral separates from a rhyodacitic lava and three ignimbritic tuffs from the Maquiang area east of Lhasa. These are the youngest ages known so far from the Lhasa block. Arnaud, (1992) reported K–Ar ages in the range of 16–20 Ma for intermediate to silicic volcanic rocks from Shiquanhe in SW Tibet. In contrast, no age older than 6.4 Ma has been reported for volcanic rocks from NW Tibet (e.g. Arnaud et al., 1992; Turner et al., 1996), whereas Turner et al., (1996) reported ages in the range of 1–18 Ma for volcanic rocks from north–central Tibet. |
Sample Description and Petrography |
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TOPABSTRACTIntroductionGeological SettingAnalytical TechniquesAge of the Post-Collisional...Sample Description and...Geochemical CharacteristicsDiscussionConcluding RemarksReferences |
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Geochemical data (Tables 6 and 7; see below) and Fig. 4 indicate that the analysed Late Oligocene–Early Miocene eruptives from SW Tibet are potassium rich, and may be divided in three groups: ultrapotassic, silicic potassic and silicic high-K calc-alkaline. The rocks are generally fresh. Small crustal xenoliths (granitic, pegmatitic and gabbroic rocks from the underlying THB) are sometimes present in the ultrapotassic and potassic volcanic rocks.
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Ultrapotassic rocks (UPV)
Table 1 summarizes the textural and mineralogical characteristics of the investigated samples. In TE011/93 and TE125/93, phenocrysts of phlogopite and clinopyroxene, and microphenocrysts of apatite and titano-magnetite are set in a pale brown glassy matrix. In other samples, the groundmass consists of microcrystalline sanidine, clinopyroxene, titanian-phlogopite, olivine, apatite, titano-magnetite and varying amounts of glass. Phlogopite, olivine, clinopyroxene and, rarely, orthopyroxene are phenocryst phases. Plagioclase is conspicuously absent. In addition, some samples contain anhedral olivine macrocrysts with undulose extinction, kink banding and orthopyroxene or phlogopite reaction rims. These are interpreted as mantle xenocrysts. In two samples, the xenocryst population also includes corroded macrocrysts of rutile mantled by armalcolite or intergrown with magnesian ilmenite.
Low-Ti phlogopite phenocrysts have high mg-numbers (84–95), and high Ni (1600–5100 ppm) and BaO (up to 3.1 wt %) contents. They are mantled by thin rims (<100 µm) with a composition similar to that of the groundmass phlogopite. These are Ti rich (5–9 wt % TiO2 ) with mg-numbers of 72–81 (Table 4). The analysed phlogopite phenocryst cores overlap the compositions of phlogopites from lamproites of SE Spain and Leucite Hills, Wyoming (e.g. Venturelli et al., 1984a; Mitchell & Bergman, 1991; Contini et al., 1993). Olivine occurs as unzoned, deformed xenocrysts (Fo91–93), characterized by high NiO (0.50–0.79 wt %) and low CaO (<0.1 wt %) contents. In contrast, olivine phenocrysts are less magnesian and zoned (cores Fo85–91; rims Fo76–85) with somewhat lower NiO (0.06–0.49 wt %) and higher CaO (0.07–0.33 wt %) contents. Groundmass grains (Fo73–82) contain slightly more CaO (0.07–0.43 wt %) and less NiO (0.02–0.29 wt %). Orthopyroxene forms reaction coronas mantling olivine xenocrysts (En89–92Fs6–9Wo1–2) and rare subhedral to euhedral phenocrysts (En82–88Fs9–16Wo2–3) or microlites (En71–84Fs13–25Wo3–4). Subcalcic clinopyroxene (Wo41En54Fs5 to Wo44En47Fs8) with low TiO2/Al2O3 ratios is present as phenocrysts and in the groundmass of all samples. Green cores with more Fe-rich compositions (Fs16–19) and/or concentric, oscillatory zoning patterns reflecting varying Al, Mg and Fe contents are occasionally observed. Sanidine is the only feldspar in these lavas and is restricted to the groundmass. It has a compositional range of Or60–75Ab18–30An2–5 with minor Fe2O3 (<0.9 wt %), TiO2 (<0.6 wt %) and BaO (0.2–3 wt %) substitutions. Titaniferous magnetite is commonly the only Fe–Ti oxide phase present in the groundmass, although Ti–Al–Mg–Cr spinels (3–38 wt % Cr2O3) have been observed in several samples (Table 5). Irregular grains of blue pleonaste occur as inclusions in phlogopite in samples TE137/93 and TE138/93 (Table 5). Similar inclusions have been reported in lamproites from West Kimberley, Australia (Jaques & Foley, 1985), Leucite Hills and SE Spain (Venturelli et al., 1984a, 1988; Wagner & Velde, 1987). Mg-ilmenite (3.5–6.7 wt % MgO) is rare and commonly only found as bleb-like inclusions in phlogopite. F-apatite is an important accessory phase.
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It is unlikely that the few single crystals of rutile mantled by armalcolite and the sigmoidal rutile–magnesian (9–11 wt % MgO) ilmenite intergrowths (Table 5) are high-pressure phenocrysts, as the coexisting liquids should be very TiO2 rich (Dickinson & Hess, 1982
). Their textures and compositions, however, are similar to those of rutile–ilmenite–armalcolite grains known from metasomatic upper-mantle xenoliths (e.g. Pasteris, 1980; Mitchell, 1986; Haggerty, 1987), from kimberlites (e.g. Haggerty, 1983) and from the Smoky Butte, Moon Canyon and Spanish lamproites (Velde, 1975; Wagner & Velde, 1986; Mitchell et al., 1987; Venturelli et al., 1988; Contini et al., 1993). The Tibetan examples are similar in composition to armalcolite in the Smoky Butte lamproites.
Potassic volcanic rocks (PVR)
The potassic volcanic rocks are characterized byphenocrysts of phlogopite ± orthopyroxene (En76–84Fs15–20Wo1–4) ± diopsidic clinopyroxene (En39–50Fs6–14Wo41–49) and a glass-rich groundmass (Table 1). Sanidine, minor apatite and Ti-magnetite are restricted to the groundmass. Phlogopite phenocrysts are commonly zoned and oxidized, with mg-numbers (74–87) and TiO2 contents (3.7–6.8 wt %) usually decreasing from core to rim. Ba is generally low (BaO > 0.4 wt %) and F highly variable (0.1–3.4 wt %). The ubiquitous groundmass-glass is rich in K2O (7–10.3 wt %).
Calc-alkaline dacites and rhyolites (CAV)
The dacites and rhyolites commonly have phenocrysts of quartz, biotite (mg-number 57–63; 4.5–4.9 wt % TiO2; no F or Cl), plagioclase, often with complex zoning (An58–33Ab41–63Or1–4), ± magnesio-hornblende in a brown, glassy matrix (Table 1). In addition, minor ilmenite, Ti-magnetite, zircon ± allanite are present. The zircons are clear with a brownish tinge and have similar morphological features: pr {100} = {110} and py {101} = {211}. In the typological distribution diagram of Pupin, (1980) the percentage of crystals decreases in the following order: S13, S8, S3.
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Geochemical Characteristics |
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TOPABSTRACTIntroductionGeological SettingAnalytical TechniquesAge of the Post-Collisional...Sample Description and...Geochemical CharacteristicsDiscussionConcluding RemarksReferences |
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Major and compatible trace elements
The UPV (Table 6) are characterized by elevated SiO2 contents (53.5–57.4 wt %), high mg-numbers (65–75), high Ni (87–358 ppm), high Cr (295–528 ppm), moderate Al2O3, and uniformly low CaO (4–5.5 wt %) and TiO2 (0.86–1.37 wt %) contents. The compatible trace elements Ni, Co, Cr, V and Sc correlate positively with MgO or mg-number values and negatively with SiO2, although trends are not well defined (e.g. Fig. 5). The major element compositions of these rocks (Fig. 4c) are transitional between Groups I (anorogenic lamproites) and III (orogenic ultrapotassic series) according to the classification of Foley et al., (1987)
. The ultrapotassic samples have a clear affinity with the lamproites from SE Spain (Contini et al., 1993, and references therein; Nixon et al., 1984; Venturelli et al., 1984a). However, as they contain >12 wt % Al2O3 and are neither peralkaline [molar (Na2O + K2O)/Al2O3 ratios are 0.75–0.90], nor perpotassic (molar K2O/Al2O3 = 0.42–0.56), they are not bona fide members of the lamproite clan (Mitchell & Bergman, 1991).
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Fe)] for ultrapotassic (filled squares) and potassic (open squares) volcanic rocks from SW Tibet. It should be noted that the high mg-number ultrapotassic rocks are enriched in compatible as well as incompatible elements relative to the silicic potassic volcanic rocks.Although of similar K2O content, the potassic rocks differ from the ultrapotassic samples in being richer in SiO2, Al2O3 and Na2O, and poorer in MgO, FeO and CaO (Table 7). K2O/Na2O (wt %) ratios vary between 1.3 and 3.5. Compositionally similar rocks have been reported from the Shiquanhe area in SW Tibet (Fig. 4a and b by Arnaud et al., (1992)
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The metaluminous to slightly peraluminous SiO2-rich samples from Gegar and Manasarowar plot within the medium- to high-K calc-alkaline field (Fig. 4a) defined by Pecerillo & Taylor, (1976). Silica contents range from 64.8 to 72.9 wt %, MgO from 0.48 to 1.5 wt %, and total alkalis from 6.2 to 7.5 wt % (Table 7). Mafic compositions have not been encountered in the study area. In the Lhasa block, however, post-collisional high-K calc-alkaline rhyolites have also been erupted in association with basaltic andesites in the Maquiang area west of Lhasa (Coulon et al., 1986; Turner et al., 1996).
Incompatible trace elements
Although trends are scattered, incompatible element abundances in the UPV generally decrease with fractionation, as revealed by the positive correlations between Ba, Sr, Nd and mg-number values (Fig. 5). Chondrite-normalized REE patterns are shown in Fig. 6a. All UPV samples are considerably enriched in light REE (LREE) (Ce: 229–352 times) relative to chondrites, whereas heavy REE (HREE) are less enriched (Yb: 6.3–10 times chondrites), resulting in CeN/YbN ratios that range from 25 to 55. All samples have similar patterns, which tend to flatten out in both the LREE and HREE and exhibit pronounced negative Eu anomalies (mean Eu/Eu* = 0.63). Similar REE patterns have been previously reported for lamproites from SE Spain (e.g. Nixon et al., 1984; Contini et al., 1993), for ultrapotassic lavas from NW Italy (Venturelli et al., 1984b) and from central Italy (e.g. Hawkesworth & Vollmer, 1979; Rogers et al., 1985).
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The large ion lithophile elements (LILE), particularly Rb (437–753 ppm), Ba (1863–3488 ppm) and Th (114–186 ppm), are significantly enriched relative to the high field strength elements (HFSE) and HREE. The UPV mantle-normalized abundance curves (Fig. 7a) are characterized by negative Ba, Ta, Nb, Sr and Ti anomalies, which occur despite their high abundances. They are remarkably similar to those of the lamproites from SE Spain (Venturelli et al., 1984a
; Contini et al., 1993). They are also broadly similar to the post-collisional lavas from northern Tibet (Turner et al., 1996). However, negative Ba anomalies are more pronounced, and Rb and Th are far more enriched in the UPV relative to the lavas from northern Tibet (Fig. 7b).
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The normalized REE (Fig. 6b) and incompatible element abundance curves (Fig. 7b) of the potassic lavas are similar to those of the ultrapotassic rocks and of the Shiquanhe rhyolites (Turner et al., 1996
), suggesting a genetic relationship. CeN/YbN ratios range from 35 to 56, and mean Eu/Eu* ratios are 0.67. Absolute REE and incompatible element abundances, however, are generally lower in the potassic lavas despite their more evolved (higher SiO2 and lower MgO) compositions (e.g. Fig. 5).
Figures 6c and 7c show that the CAV are characterized by selective enrichments of Rb, Ba, Sr, K, Th and LREE, and low abundances of Ta, Nb, P, Ti and HREE. Compared with the ultrapotassic and potassic rocks, they have lower LREE/HREE ratios and much lower abundances of incompatible elements, with the exception of Sr. The chondrite-normalized REE patterns (Fig. 6c) are straight near-parallel trends characterized by moderate LREE enrichment relative to chondrite (CeN/YbN = 16–22). Their trace element characteristics are similar to those of the 16–10 Ma lavas described from the Maquiang area near Lhasa by Coulon et al., (1986) and Turner et al., (1996).
Volatile trace elements
In the UPV, fluorine, hosted in phlogopite and apatite, varies from 3800 to 6100 ppm. These values are similar to those previously reported for lamproites (Mitchell & Bergman, 1991). Chlorine abundances are variable (230–1100, mean 508 ppm) and higher than in lamproites (Mitchell & Bergman, 1991). Sulphur contents are low (68–183, mean 123 ppm).
Sr–Nd isotopic composition
The Sr and Nd isotopic compositions of analysed samples are given in Table 8. The UPV have highly radiogenic initial 87Sr/86Sr (0.717–0.722) and unradiogenic 143Nd/144Nd ratios corresponding to Nd values between -11.9 and -14.3. As illustrated in Fig. 8a, the lavas plot in the enriched quadrant of a conventional Sr–Nd isotope diagram, where they overlap the fields defined by the lamproites from SE Spain (Nelson et al., 1986) and the leucite lamproites from the Kimberley region of Western Australia (McCulloch et al., 1983). Inspection of Fig. 8a also shows that the 16–20 Ma rhyolitic lavas from Shiquanhe (Turner et al., 1996) have similar isotopic characteristics. The Nd model ages of the UPV relative to depleted mantle (Table 8) range from 1.3 to 1.9 Ga (mean 1.7 Ga). Because the Sm/Nd of a melt generally is lower than that of its source, these TDM model ages represent minimum ages of enrichment.
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In sample TE11/93, the nearly identical Nd isotopic compositions in coexisting phlogopite and diopside phenocrysts suggest isotopic equilibration with the whole rock. In sample TE137/93 (Table 8) phlogopite is in isotopic equilibrium with the host rock, whereas the
Nd value of clinopyroxene indicates disequilibrium. The Nd isotopic composition (Nd = -16.6) of the kinked mantle olivine xenocrysts is even more unradiogenic than that of the host rock, indicating an ancient enrichment event (TDM = 3.3 Ga).
Most of the analysed PVR lavas (Table 8, Fig. 8a) have similar 143Nd/144Nd ratios (Nd = -13.1 to -14.5) to the ultrapotassic samples and the Shiquanhe rhyolites (Turner et al., 1966), but even more extreme initial 87Sr/86Sr ratios (0.718–0.737). The Nd model ages relative to depleted mantle range from 1.3 to 1.4 Ga.
The 17 Ma silicic CAV rocks have distinctly lower 87Sr/86Sr (0.7091–0.7097) and higher 143Nd/144Nd (Nd = -7.1 to -9.5) isotopic compositions relative to the ultrapotassic and potassic samples (Table 8, Fig. 8b). However, their isotopic compositions are not quite as primitive as those reported for the 10–16 Ma mafic and silicic lavas from Maquiang in the southern part of the plateau (Turner et al., 1996). Their depleted mantle Nd model ages are 1.1–1.3 Ga.
Pb isotopic composition
The Pb isotopic compositions of UPV and PVR are similar (Table 9). Initial 206Pb/204Pb ratios are restricted, lying in the range of 18.41–18.76. 207Pb/204Pb (15.67–15.76) and 208Pb/204Pb (39.14–39.75) ratios are unusually radiogenic and plot well above the Northern Hemisphere Reference Line (NHRL; Hart, 1984) in conventional Pb isotope diagrams (Fig. 9), implying that they must have had a complex multi-stage evolution. The uniform 206Pb/204Pb leads to a vertical array on a plot of 206Pb/204Pb vs 207Pb/204Pb (Fig. 9a), similar to that reported for the lamproites from SE Spain (Nelson et al., 1986) and for other post-collisional, shoshonitic volcanic rocks from Tibet (Turner et al., 1996). As already pointed out by Nelson et al., (1986) such a correlation is unlikely to have resulted directly from the closed decay of U, as the range in 207Pb/204Pb requires the existence of long-term variation in U/Pb, which should have resulted in a large variation in 206Pb/204Pb. Secondary isochrons drawn from individual Pb data points to the intersection between the geochron and the NHRL (Silver et al., 1988) give Pb model ages ranging from 3.6 to 3.8 Ga.
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O isotopic composition
The UPV and PVR are characterized by oxygen isotope whole-rock data (Table 8) in the range 9.4–11.0%° (relative to SMOW), showing that the lavas are enriched in 18O relative to typical mantle-derived rocks (5–8%°; Kyser, 1986
). They are, however, within the range of the Leucite Hills lamproites and Roman province type (RPT) lavas (Mitchell & Bergman, 1991). Xenocrystal mantle olivine and clinopyroxene separates from sample TE137/93 are within the range reported for direct or indirect samples of the mantle. None of these minerals, however, plots within the field of unmetasomatized mantle because of their strongly crustal Nd isotopic signature. Phlogopite phenocrysts have somewhat elevated 
18O (7.9%°). High 18O (8.8%°) of phlogopite phenocrysts in Leucite Hills lamproites has been interpreted as a primary feature (Kuehner, 1980). |
Discussion |
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TOPABSTRACTIntroductionGeological SettingAnalytical TechniquesAge of the Post-Collisional...Sample Description and...Geochemical CharacteristicsDiscussionConcluding RemarksReferences |
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Ultrapotassic volcanic rocks
Simple models explaining the formation of ultrapotassic rocks do not exist. Their genesis is controlled by the evolution and mineral assemblage of the mantle source, the type and degree of partial melting, and the subsequent fractionation and contamination history of the melts. The anomalous geochemical signature of the UPV echoes that of the Spanish lamproites and suggests that the ultrapotassic samples are derived from, or contain a contribution from, a source enriched in LREE and Rb relative to the Bulk Earth.
The UPV have high mg-numbers and high Ni contents, and contain mantle xenocrysts. The most primitive samples have the highest compatible as well as incompatible element abundances (Fig. 5). Sr (621–1042 ppm), Nd (109–200 ppm) and other incompatible element contents are significantly higher than those in average upper crust (Taylor & McLennan, 1985) or typical upper-crustal melts (Pearce et al., 1984). The fact that they were erupted through continental crust raises the possibility that crustal contamination produced some of the isotopic and incompatible trace element signatures. However, assimilation of crustal rocks would be difficult to resolve, as it would only dilute LILE and LREE contents of the UPV. The crustal xenoliths observed in some of these rocks and the basement THB granitoids have much lower 87Sr/86Sr and higher 143Nd/144Nd ratios relative to the lavas. Mixing between these crustal xenoliths or THB basement rocks and a mid-ocean ridge basalt (MORB) or ocean-island basalt (OIB)-like parental magma cannot create the UPV (Fig. 8b). In addition, Sr and Nd isotope ratios do not show any significant correlation with the degree of fractionation. This suggests that crustal assimilation and fractional crystallization (AFC) processes (DePaolo, 1981) were not important in the development of the extreme compositions of the ultrapotassic lavas. Therefore, the trace element and isotopic variations of the UPV should reflect those of their mantle sources.
Current models for the formation of highly potassic melts such as lamproites and group II kimberlites invoke the presence of a veined mantle source (e.g. Foley, 1992a, 1992b). The vein assemblages are thought to consist of both hydrous and anhydrous phases enriched in incompatible elements, such as phlogopite, K-amphibole and apatite.
Isotopic constraints
The fact that the Sr, Nd and Pb isotope compositions of the investigated samples (Figs 8a and 9) fall far outside the ranges for oceanic basalts argues against exclusively asthenospheric or mantle plume sources. Their high 87Sr/86Sr and low 143Nd/144Nd isotope signatures require sources with a time-integrated history of enrichment in LREE and Rb relative to the HREE and Sr. One possible model is to ascribe these geochemical characteristics to a metasomatized subcontinental lithospheric mantle source that was isolated for >1.3 Ga (minimum TDM) before the partial melting event that led to the formation of the ultrapotassic magmas. In principle, the observed isotopic compositions of the lavas might also have resulted from the mixing of Sr, Nd and Pb from a variety of sources (e.g. McCulloch et al., 1983; Nelson et al., 1986) where potential end-members could include asthenospheric and highly enriched lithospheric sources. However, the lack of convincing evidence such as correlations between 207Pb/204Pb and Ce/Pb or between 143Nd/144Nd and Ta/Nd or Ba/Nb ratios argues against this possibility.
Of particular interest are the Pb isotopic compositions. The relatively elevated 207Pb/204Pb ratios suggest involvement of an old radiogenic component that was probably present in the source region as it can be identified in all compositions. The steep trend on the 207Pb/204Pb vs 206Pb/204Pb diagram and the shift towards the composition of pelagic sediments (Fig. 9a) suggests a component of recycled continent-derived material in the source of these volcanic rocks. The restricted range in 206Pb/204Pb, however, that is also seen in the lamproites from SE Spain (Nelson et al., 1986) is a problem, as it does not favour typical sediments.
Evidence for subducted sediment
The elevated 207Pb/204Pb ratios and the negative correlation between 
7/4 (the displacement to high 207Pb/204Pb above the NHRL) and Nd isotope composition in the ultrapotassic lavas from SW Tibet are consistent with contamination of their mantle source by a component resembling oceanic sediments. In addition, the Ce/Pb ratios are within the range 2–5 and distinctly lower than in oceanic (MORB and OIB) basalts (Hofmann et al., 1986), but typical of a subducted sediment component (Ben Othman et al., 1989). A sedimentary signature is also reflected in the low Sr/Nd (4.0–7.2), low Nb/La (0.21–0.52), elevated Ce/Sr (0.2–0.4), high Th/Ta (40–129), high Th/La (1–2.3) ratios and in the negative Eu anomalies that are unrelated to plagioclase fractionation. The 18O values for the UPV differ from the mantle values, implying the involvement of surface processes. They could be explained by a model proposing seawater alteration before subduction (e.g. Jacob et al., 1994). However, these 18O values need not be primary, as oxygen isotopes are very sensitive to weathering and only values obtained from rocks with H2O < 0.3 wt % should be considered representative (Kyser et al., 1982).
The presence of negative Ta–Nb–Ti anomalies in the mantle-normalized hygromagmatophile element abundance patterns (Fig. 7a) could indicate the involvement of a subduction-related component. An alternative explanation of the relative Ta, Nb and Ti depletions requires the presence of residual titanates in the mantle source during partial melting (Foley & Wheller, 1990). The rutile xenocrysts in UPV samples TE137/93 and TE138/93 could indeed indicate the stability of titanate minerals in the source. The UPV are characterized by low Ti/Y (250 ± 68), high Rb/Ba and negative Nd values. This combination is a feature of continental material subducted back into the upper mantle and suggests a major contribution from sediment-contaminated lithospheric mantle in the generation of these magmas. Affinities to lamproites are highlighted on the La/Nb vs Ba/Nb diagram (Fig. 10). The UPV plot on a trend defined by potassic melts that are derived from subcontinental lithospheric sources, such as lamproites and micaceous kimberlites (Mitchell, 1986; Mitchell & Bergman, 1991). This trend cannot be caused by any model involving partial melting from a four-phase peridotite (MORB or OIB) source followed by crustal contamination, but instead requires a distinct mantle source.
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Constraints on source region mineralogy
More direct evidence for involvement of deep-seated materials that may have contributed to the UPV magmas is provided by the presence of rutile and kinked olivine xenocrysts, which may have resulted from the disaggregation of ancient, metasomatically modified lithospheric mantle. The olivine has unexpectedly high Nd contents (0.64 ppm) that support mantle metasomatism, an
Nd of -17 at 18 Ma and a Nd model age of 3.3 Ga relative to depleted mantle. Interestingly, this Archaean enrichment event in the southern part of the Tibetan subcontinental mantle lithosphere might also account for the Archaean Pb model ages.
As pointed out by Hawkesworth et al., (1985), most basalts, including MORB, OIB and non-orogenic lamproites, have Rb/Ba ratios between 0.03 and 0.1 and define a positive trend in Fig. 11, suggesting that processes responsible for these magmas do not fractionate Rb/Ba significantly. Primitive mantle and garnet peridotite xenoliths also plot on this trend. The ultrapotassic and potassic lavas from SW Tibet are clearly displaced to higher Rb/Sr (0.5–1.0) and to higher Rb/Ba (0.15–0.51), revealing an enrichment of Rb relative to Ba and Sr, similar to that seen in the lamproites from SE Spain (Fig. 11). High Rb/Sr, high Sr basalts are relatively rare, and tend to be characterized by low Ti (e.g. Ti/K = 0.10–0.16). The cause of these characteristics is still debated, as they cannot be adequately explained by mica–liquid fractionation or small-degree melting models. High Rb/Sr, high Rb/Ba, high Sr and low Ti/K ratios are also a feature of phlogopite peridotite xenoliths (Fig. 11). Such xenoliths are clearly H2O rich relative to normal mantle and may be products of mantle metasomatism induced by H2O-rich fluids (e.g. Erlank et al., 1987). In contrast, the mafic lavas from the northern part of the plateau (Turner et al., 1996) are characterized by much lower Rb/Ba (0.05–0.11) and Rb/Sr (0.06–0.14) and higher Ti/K (0.19–0.49) ratios, reflecting a different style of enrichment, possibly by migration of small-volume melts. According to Hawkesworth et al., (1990), the negative correlation between U/Pb and K/Nb ratios in lamproites and other mantle-derived magmas indicates that U/Pb fractionation is strongly influenced by a potassic phase. In the UPV, the high K/Nb (1241–2278) and low U/Pb (0.2–0.5) is associated with high Rb/Sr, suggesting stabilization of phlogopite in the source region. In addition, Venturelli et al., (1984a) have argued that the fairly high Ni/MgO ratios of the UPV (30 ± 6) relative to MORB (17 ± 3) and OIB (14) could indicate disequilibrium partial melting of a phlogopite-bearing source.
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Additional petrological characteristics of the UPV mantle source can be qualitatively illustrated using La/Yb vs Yb and La/Yb vs Dy/Yb diagrams (Fig. 12) that distinguish between melting in the spinel and garnet stability fields (Thirlwall et al., 1994
). In the UPV, La/Yb ranges from 37 to 90. This range is uncorrelated with the mg-number, demonstrating that partial melting and regional source variations exerted the dominant control on the trace element variability within the lavas. There is little change in La/Yb ratios and Dy/Yb ratios remain almost constant during melting in the spinel stability field, whereas melting in the garnet stability field produces large changes in Dy/Yb and La/Yb ratios. Partial melting and mixing curves shown in Fig. 12 indicate that variable degrees of partial melting of a hypothetical LREE-enriched mantle source cannot generate the La/Yb–Dy/Yb systematics of the UPV either in the garnet or on the spinel stability fields. However, mixing of very small melt fractions from garnet-facies mantle with distinctly larger melt fractions from spinel-facies mantle could account for the observed REE data. Although the model is dependent on the nature of the mantle source, the overall shape of the trajectories for non-modal batch melts from spinel- and garnet-facies mantle changes little with changing source composition, suggesting that the ultrapotassic lavas from SW Tibet require melt components from both the garnet and spinel stability fields. Garnet in the source region has also been suggested for the northern Tibetan lavas (Turner et al., 1996) and for the Pleistocene trachyandesites from NW Tibet (Cooper et al., 1997).
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Partial melting conditions
Currently, sanidine–phlogopite lamproites are interpreted to be derived by partial melting of phlogopite- bearing metasomatic veins within a harzburgitic lithospheric mantle source (e.g. Mitchell & Bergman, 1991
). Foley, (1992b) has reviewed the aspects of melt generation from a veined mantle source, introducing his vein-plus-wallrock melting model. The composition of melts derived from veins containing hydrous potassic phases is controlled by the degree of partial melting and the relative contributions of vein and wall-rock material. However, the exact mineralogy and the degree of partial melting of such a source cannot be unambiguously estimated from geochemical parameters. Quantitative modelling of partial melting processes of such a source is difficult because variations in mineralogy, modal and chemical compositions, and trace element distribution coefficient data are not well constrained. A mechanism involving melting of phlogopite veins in a harzburgitic substrate, however, could explain the fact that the UPV have very low CaO/Al2O3, ranging from 0.28 to 0.39, and low CaO (4.9–5.5 wt %) and Sc contents (14–20 ppm).
Liquidus and near-liquidus studies on various lamproite compositions [see the review by Edgar & Vukadinovic, (1992)] indicate that the Tibetan ultrapotassic magmas could have been derived by partial melting of a phlogopite-bearing mantle source. Partial melts of phlogopite harzburgite can be represented by the peritectic melting points Ol + Opx + Phl + Lq in the system kalsilite–forsterite–quartz (Ka–Fo–Qz) at different pressures under water-saturated conditions (Foley, 1993). The SiO2-rich UPV recalculated into the Ks–Fo–Qz subsystem plot near the peritectic point at 2 GPa, or near the enstatite–forsterite–sanidine eutectic at 1.85–1.95 GPa (Wendlandt & Eggler, 1980), corresponding to shallow depths of origin of 60 km. This and the substantial melt contributions from spinel-facies mantle suggest that the sources of the UPV are located in the lithosphere.
The formation of an armalcolite rim upon the rutile macrocrysts may be a growth feature and a consequence of pressure decreasing below 1.6 GPa (Arima & Edgar, 1983) during their transport in the magma. Only two coexisting ilmenite–Ti-magnetite pairs were observed. They record oxygen fugacities below the QFM (quartz–fayalite–magnetite) buffer (log fO2 -15.5 to -16.7) and low temperatures (730–800°C). As both oxide phases are homogeneous they could record late magmatic temperatures.
Potassic volcanic rocks
The silica-rich potassic lithologies have lower incompatible element concentrations than any of the ultrapotassic samples. In addition, their Sr isotopic composition is significantly higher relative to the related UPV rocks. Therefore, derivation from the UPV through fractional crystallization alone cannot explain their trace element and isotopic signatures. A model in which regional isotopic heterogeneities occur within the subcontinental mantle beneath SW Tibet and/or of contamination by more evolved crustal material seems inevitable. The evolved Sr isotopic characteristics, however, imply that the crustal component must have had an even higher radiogenic Sr ratio. This rules out contamination by material from the underlying Transhimalaya Batholith that is also sometimes observed as xenoliths in these and the ultrapotassic lavas. As Table 8 and Fig. 8b show, the 87Sr/86Sr ratios of these crustal rocks are far too low.
High-K calc-alkaline dacites and rhyolites
The dacitic and rhyolitic rocks occur in the same tectonic setting as the UPV and PVR, but they are mineralogically and chemically distinct. Using the Zr solubility data of Watson & Harrison, (1983), temperature estimates of 740–780°C are obtained. The abundance and early crystallization of the hydrous phases biotite and amphibole indicate that water was an important magmatic component. Their high-K calc-alkaline signature argues against an origin by crustal anatexis alone and demands addition of heat that was probably provided by mantle-derived melts intruded into the crust. However, in the absence of pertinent basaltic magmas their origin is difficult to assess.
Rhyolitic rocks are also known from the Ulugh Muztagh area in the northern part of the Tibetan plateau. They have been interpreted by McKenna & Walker, (1990) to represent partial melts of pelitic source rocks within the thickened Tibetan crust. They differ from the CAV by being distinctly peraluminous, by having large negative Eu anomalies, significantly higher Rb/Ba and higher initial 87Sr/86Sr ratios. On the other hand, similarities in trace element patterns with the Maquiang basalts and rhyolites (Figs 6c and 7c) suggest a relationship with these extrusive rocks from the southern part of the plateau. The Maquiang basalts are clearly not primary magmas and were erupted through continental crust, yet their trace element, Sr (0.7049) and Nd (0.5127) isotope compositions (Fig. 8b) and the high Sr/Nd ratios (Coulon et al., 1986; Turner et al., 1996) could indicate an asthenospheric input, possibly from a subduction-affected wedge, but data are insufficient to reach real conclusions.
Model Nd ages of the Maquiang lavas and the CAV are significantly older than the age of extrusion and so suggest some crustal input. The fact that the CAV contain abundant zircon xenocrysts with a Pb-evaporation age of 471 ± 33 Ma (Table 3) confirms that their origin must be sought in context with crustal anatexis. In addition, the zircon xenocrysts document the existence of Palaeozoic basement material in this part of the Lhasa block. They record an event coincident with the widespread Ordovician felsic magmatism in the High Himalaya Crystalline, which in turn suggests a common evolution of the Lhasa block and India during the Early Palaeozoic. The nature of the crustal component, however, remains ambiguous. As shown in Fig. 8b, simple bulk mixing with the 531 Ma Amdo orthogneiss, which is the only analysed basement sample of the Lhasa block (Harris et al., 1988), or the Ordovician granitoids cannot generate the isotopic variation seen in the post-collisional calc-alkaline volcanic rocks from southern Tibet. However, additional input from a lower-crustal source with 87Sr/86Sr 0.7100 and 143Nd/144Nd 0.5115 (Fig. 8b) could account for some of the observed Sr–Nd isotopic variations in the Maquiang lavas and the CAV. AFC trajectories (not shown) for the initial 87Sr/86Sr vs Nd plot (Fig. 8b) do not deviate markedly from the bulk mixing curves for F > 0.8, r = 0.2–0.6, DNd = 0.1 and DSr = 0.1 and 1.0. Quantification of the amounts of crustal contamination, however, is impossible because of the uncertainties regarding the trace element and isotopic compositions of the primary melts and crustal end-member(s), the variable bulk distribution coefficients during assimilation and fractional crystallization, and the variable ratios (r) of the rate of assimilation to the rate of fractional crystallization.
Tectonic implications
The collision of India and Asia at about 50–55 Ma marked a significant shift in the tectonic regime of southern Tibet. One of the consequences of collision has been the thickening of the crust beneath the Tibetan plateau to 60–80 km (Brandon & Romanowicz, 1986) and the increase in its surface elevation. The change at the end of the Eocene, however, did not immediately lead to the raising of the plateau. Earlier than 25 my ago, very little of the convergence was taken up by lifting the surface of the Earth (Yin et al., 1994). The processes of subduction of the relatively thin crust of the northern margin of India beneath the accretionary complex and of tectonic escape (e.g. Tapponnier et al., 1982; Peltzer & Tapponnier, 1988) seem to have accommodated the continued northward progression of India. Further convergence has been accommodated by folding and major north-dipping fault systems in southern Tibet [Gangdese thrust system (GTS)], and within the Himalaya [Main Central Thrust (MCT) and Main Boundary Thrust (MBT)]. Thermochronology constrains initiation of motion on the GTS to 27 Ma (Yin et al., 1994) and suggests that the MCT was first active in the earliest Miocene (22–24 Ma; e.g. Hubbard & Harrison, 1989), possibly coeval with movement on the North Himalayan (normal) Fault (NHF; e.g. Yin, 1993). The active tectonics of the plateau are characterized by east–west extension and normal faulting (e.g. Armijo et al., 1989). These are thought to be a consequence of plateau uplift and seem to have begun at 8 Ma (Harrison et al., 1992; Molnar et al., 1993). The mechanisms by which the plateau had attained an elevation sufficient to trigger extension are not well understood.
Convective instability of a thickened mantle boundary layer has been proposed as a mechanism that could lead to such an event (Houseman et al., 1981). Modelling lithospheric evolution as a viscous continuum suggests convective thinning of the thermal boundary layer in a relatively brief period of time. The change in potential energy in the remaining lithosphere would then lead to a sudden uplift of the plateau (e.g. England & Houseman, 1988, 1989; Houseman & England, 1993), triggering in turn its current extensional deformation. In addition, it has been argued that the post-collisional high-K basaltic volcanism in Tibet is a thermal consequence of lithospheric thinning (Molnar et al., 1993; Turner et al., 1993, 1996).
The eruption of the post-collisional ultrapotassic magmas in SW Tibet (18–25 Ma) clearly post-dates continental collision. This delay would not be expected if fluid release accompanying oceanic or continental subduction were the trigger. The volcanism pre-dates regional extension and normal faulting, precluding decompression melting of vein assemblages stored in the mantle as a cause for this magmatism. If the ultrapotassic volcanism was caused by a temperature increase and melting in the mantle part of the continental lithosphere, then associated convective thinning of the subcontinental lithospheric mantle is a necessary prerequisite (Turner et al., 1992). This, in turn, should have dramatic consequences, in particular, a change in thermal structure and, most importantly, surface uplift (e.g. Houseman et al., 1981). Therefore, the age data of this volcanic event can be interpreted in two ways: (1) Uplift of the southern Tibetan plateau occurred at the generally accepted time around 8 Ma (e.g. Molnar et al., 1993; Turner et al., 1996) and the thinning between 25 and 18 Ma was not accompanied by uplift. In this case, however, a process must be invoked that would allow negligible surface uplift despite lithospheric thinning, such as the possibility that the region of investigation was not in isostatic equilibrium and that the surface elevation was dynamically, rather than isostatically supported. (2) Uplift of the southern Tibetan plateau occurred much earlier than previously assumed, i.e. around 25–18 Ma. Thermochronological studies and the fact that the drainage from southern Tibet across the Himalaya is antecedent to elevated topography do indeed suggest that significant uplift preceded the rapid unroofing of the Transhimalaya plutonic belt at 21 Ma (Harrison et al., 1992). This could indicate operation of a mechanism that allows for delayed instability of the thickened thermal boundary layer as proposed by Lenardic & Kaula, (1995). Their model predicts that the mantle lithosphere at the margin of a thickened region could be thermally eroded until the base of the crust comes into contact with asthenospheric mantle. This would create a peripheral weak zone that in turn allows thickened mantle lithosphere of the interior plateau region to become mechanically detached producing rapid uplift. In this scenario, post-collisional magmatism should start in the southern part of the Tibetan plateau with partial melting within enriched lithospheric mantle. Although this solution is appealing, it is not the only plausible mechanism of inducing partial melting in the mantle lithosphere of the Lhasa block.
An alternative possibility is predicted by the slab breakoff model of Davies & von Blanckenburg, (1995). As the subducted oceanic plate breaks off, hot asthenosphere will rise and impinge upon the mantle lithosphere of the overriding plate. The resulting thermal perturbation should induce melting of domains within the mantle lithosphere that have been metasomatized and hydrated during the preceding oceanic subduction (possibly involving ancient sediments) and/or during earlier events. Depending on the age and enrichment history of the lithosphere a variety of mantle melts could be produced, ranging from ultrapotassic for small-degree melting of enriched and veined domains up to calc-alkaline for higher-degree melting of more fertile or hydrated peridotite layers. These basalts would rise into the crust, where they could induce anatexis, resulting in felsic magmatism. In this case, however, different processes must be invoked to explain the post-collisional magmatism in the northern parts of the Tibetan plateau, such as lithospheric delamination or intracontinental southward subduction of the Precambrian Tarim basin lithosphere (Pearce & Mei, 1988; Arnaud et al., 1992).
The high-K calc-alkaline volcanic rocks that occur in the same tectonic setting as the UPV are mineralogically and chemically distinct. Their Sr–Nd systematics and trace element signatures are within the range shown by undoubted subduction-related suites, such as Transhimalaya Batholith and associated volcanic rocks (Fig. 8b). In the Maquiang area, evolved dacites and rhyolites that are similar to the CAV may be related to more mafic rocks by AFC processes, suggesting that the enrichment history of their mantle source is different from that of the UPV and must be sought in context with the subduction of India.
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Concluding Remarks |
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TOPABSTRACTIntroductionGeological SettingAnalytical TechniquesAge of the Post-Collisional...Sample Description and...Geochemical CharacteristicsDiscussionConcluding RemarksReferences |
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The post-collisional lavas in Tibet formed as a part of the complex geodynamic evolution of this area. The eruption of the mantle-derived ultrapotassic lavas in SW Tibet pre-dates the onset of mafic volcanism in northern Tibet. The ultrapotassic magmas originated from old lithospheric sources enriched in Rb with low Sm/Nd ratios. Their isotopic and trace element systematics probably reflect processes far preceding the India–Asia collision. The Nd model ages of the olivine xenocrysts may even date an Archaean metasomatic event and thus indicate preservation of ancient mantle lithosphere beneath the Lhasa block. However, certain features of the trace element and isotope systematics require a more complex petrogenetic model involving contamination and veining of the source by a component derived from subducted sediments. The post-collisional mafic lavas from northern Tibet also represent lithospheric mantle melts (Turner et al., 1993
, 1996), but differ from the UPV by their less evolved isotopic compositions and by their distinctly lower Rb/Ba, Rb/Sr and higher Ti/K ratios, suggesting different mantle source characteristics for the different domains of the plateau. Given the complexities of geological processes in assembling Tibet, it is to be expected that metasomatism and melt–fluid infiltration of its subcontinental lithosphere will not necessarily be a uniform, single-stage process. The dilemma is that currently available data cannot provide unequivocal evidence concerning the character and timing of these enrichment event(s). The uprise of the ultrapotassic magmas from metasomatized lithospheric sources some 30 my after continental collision may signal heating of the subcontinental lithospheric mantle either by convective thinning or by another process that allows exposure of the former mechanical boundary layer mantle to hot asthenosphere. In any case, mantle dynamics beneath Tibet are not yet well constrained and our data suggest that more complex physical models may be required.
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Acknowledgements |
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Samples were collected during the 1993 field trip financed by the Fonds zur Förderung der wissenschaftlichen Forschung (FWF), Project P9420-GEO and jointly co-ordinated in co-operation with the Chengdu College of Geology and the Bureau of Geology and Mineral Resources of Xizang (China). We especially thank Jianming Liu (Chinese Centre of Mineral Resources, Beijing) and Xiang Zhou (Bureau of Geology and Mineral Resources, Lhasa) for their support in the field. We thank V. Mair for assistance with fieldwork in Tibet, S. Hoernes for providing the oxygen isotope analyses, and M. Jelenc and M. Thöni for assistance with Sr and Nd isotope analyses. S. F. Foley, J. Konzett, J. A. Pearce, N. W. Rogers, I. Ryabchikov, K. Stüwe and P. Ulmers provided helpful comments on an earlier version of this manuscript. We thank S. Turner and an anonymous reviewer for their critical reviews. Marjorie Wilson's editorial assistance also improved this manuscript.
* Corresponding author.
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TOPABSTRACTIntroductionGeological SettingAnalytical TechniquesAge of the Post-Collisional...Sample Description and...Geochemical CharacteristicsDiscussionConcluding RemarksReferences |
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