mineral ages and pt conditions of late paleozoic high-pressure ...

[American Journal of Science, Vol. 310, November, 2010, P. 916 –950, DOI 10.2475/09.2010.07]

MINERAL AGES AND P-T CONDITIONS OF LATE PALEOZOIC HIGH-PRESSURE ECLOGITE AND PROVENANCE OF ME´LANGE SEDIMENTS FROM ATBASHI IN THE SOUTH TIANSHAN OROGEN OF KYRGYZSTAN ¨ NER***,§, M. CORSINI§§, D. V. ALEXEIEV§§§, E. HEGNER*, R. KLEMD**, A. KRO L. M. IACCHERI*, T. ZACK***, P. DULSKI§§§§, X. XIA†, and B. F. WINDLEY†† ABSTRACT. Ages derived from various isotope systems in high-pressure (HP) rocks of the western Tianshan orogen of NW China have been interpreted as evidence for late Carboniferous and/or Triassic collision of the accretionary margin of the Central Asian Orogenic Belt (CAOB) with the Tarim Craton. In order to elucidate this controversy, we present new P-T data as well as Sm-Nd and 40Ar/39Ar cooling ages for an eclogite sample from Atbashi in the accretionary me´lange of the South Tianshan suture in Kyrgyzstan, some 500 km along strike to the west of the controversial locality in the upper Akeyazhi River Valley in NW China. A clockwise P-T path for the eclogite with peak pressures of 18 to 24 kbar at 520 to 600 °C is consistent with near-isothermal decompression and exhumation in a subduction zone before collision of the CAOB with the Tarim Craton. Geochemical data and an initial ␧Nd value of ⬃ ⴙ9 suggest an N-MORB protolith for the eclogite. The high-pressure mineral assemblage of the eclogite yielded a statistically robust Sm-Nd isochron age of 319 ⴞ 4 Ma (2␴, 5 data points, MSWD ⴝ 0.4) for equilibration and closure of the Sm-Nd system during HP metamorphism. 40Ar/39Ar dating of phengite from the same sample yielded a cooling age of 316 ⴞ 3 Ma (2␴) implying rapid exhumation. Docking of the Tarim Craton with the southern margin of the Middle Tianshan-North Tianshan blocks in Kyrgyzstan during the late Carboniferous is supported by widespread emplacement of A-type granitoids of early Permian age that suggest a setting of consolidated crust. An unmetamorphosed and little deformed molasse-type conglomerate of latest Carboniferous age, overlying the HP rocks, indicates that HP metamorphism, exhumation, and exposure of the HP me´lange occurred from 320 to ⬃300 Ma. The detrital zircon age spectrum of a metagraywacke sample from the accretionary me´lange suggests sources in the Tarim Craton and/or from the Middle and North Tianshan that possibly comprise rifted blocks from Tarim. introduction

Constraining the temporal framework of crustal accretion processes during closure of the Neoproterozoic to Paleozoic Paleoasian Ocean, which culminated in the Central Asian Orogenic Belt (CAOB), has been a target of international scientific investigations in the past 20 years (for example, Coleman, 1989; Sengo¨r and others, 1993; Kro¨ner and others, 2007; Windley and others, 2007). In particular, the timing and nature of metamorphism along suture zones lined with remnants of eclogite-facies ocean-floor and blueschist-facies accretionary-prism lithologies in the Tianshan oro-

* Department fu¨r Geo- und Umweltwissenschaften, Universita¨t Mu¨nchen, Theresienstrasse 41, D-80333 Mu¨nchen, Germany; [email protected] ** GeoZentrum Nordbayern, Universita¨t Erlangen, Schlossgarten 5a, D-91054 Erlangen, Germany *** Institut fu¨r Geowissenschaften, Universita¨t Mainz, Becherweg 21, 55099 Mainz, Germany § Beijing SHRIMP Center, Institute of Geology, Chinese Academy of Geological Sciences, 26 Baiwanzhuang Road, Beijing 100037, China §§ Institut Geosciences AZUR, (CNRS U.M.R. 6526) Site Sophia, Universite´ de Nice-Sophia Antipolis, 250 rue Albert Einstein, 06560 Valbonne-Sophia-Antipolis, France §§§ Geological Institute, Russian Academy of Sciences, Pyzhevskiy 7, 119018 Moscow, Russia §§§§ GeoForschungsZentrum Potsdam, Telegrafenberg, C 326, 14473 Potsdam, Germany † James Lee Building, Department of Earth Sciences, The University of Hong Kong, Pokfulam Road, Hong Kong †† Department of Geology, University of Leicester, University Road, Leicester LE1 7RH, United Kingdom

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E. Hegner & others

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genic belt have been studied in NW China and Kyrgyzstan in order to elucidate the history of docking of the Tarim Craton with the southern accretionary margin of the CAOB (Tagiri and others, 1995; Gao and others, 1999; Klemd and others, 2002; Gao and Klemd, 2003; Zhang and others, 2007; Lu and others, 2008; Simonov and others, 2008). Windley and others (1990) suggested formation of the South Tianshan orogen in China by Paleozoic collision of the Yili-Central Tianshan (considered to be equivalent to the North Tianshan in Kyrgyzstan) and Tarim blocks and reactivation of the collision zone in the Cenozoic. Uplift and unroofing of the Tianshan orogenic belt due to the collision of India with Asia has occurred since the late Oligocene-early Miocene (for example, Hendrix and others, 1994; Yin and others, 1998; Sobel and others, 2006). Several time brackets for collision of the Tarim Craton with the CAOB have been proposed: Late Devonian (Xia and others, 2004), Late Devonian-Early Carboniferous (Allen and others, 1992), early-middle Carboniferous (Coleman, 1989; Gao and others, 1998; Zhou and others, 2001; Gao and Klemd, 2003), Carboniferous to early Permian (Zonenshain and others, 1990; Biske, 1996; Heubeck, 2001; Xiao and others, 2004; Klemd and others, 2005; Charvet and others, 2007; Konopelko and others, 2007; Wang and others, 2008; Li and others, 2008; Gao and others, 2009; Lin and others, 2009), and Early Triassic (Zhang and others, 2007; Xiao and others, 2009a, 2009b). A distinct Triassic collisional event with high-pressure (HP) to ultra-high pressure (UHP) conditions was postulated by Zhang and others (2007) who obtained SHRIMP U-Pb ages of 233 to 226 Ma for metamorphic zircons in the western Tianshan of NW China. The discrepancy between these zircon ages and late Carboniferous Sm-Nd mineral/whole-rock ages of eclogites from the same locality, as well as 40Ar/39Ar and Rb-Sr ages of HP micas (Gao and Klemd, 2003; Klemd and others, 2005) were explained by Zhang and others (2007) as due to isotopic disequilibrium of the Sm-Nd system and inheritance of excess Ar. In order to resolve this controversy, we determined Sm-Nd and 40Ar/39Ar cooling ages, as well as P-T conditions, of an eclogite sample from the Atbashi HP me´lange in southern Kyrgyzstan, ⬃500 km west of the Kekesu eclogite locality in China and part of the same collisional suture. SHRIMP and LA-ICPMS ages of detrital zircons from a phengite paragneiss sample from the Atbashi me´lange were determined to define a maximum age of deposition and provenance of the material. the south tianshan fold-and-thrust belt in kyrgyzstan

The South Tianshan fold-and-thrust belt (hereafter termed South Tianshan) extends from central Uzbekistan, through southern Kyrgyzstan to northwestern China, over a distance of more than 3000 km (fig. 1). The main deformation in the belt was due to convergence and collision of the North Tianshan-Yili-Central Tianshan continent with the Tarim, Alai and Turan (or Karakum-Tadjik) microcontinents in the south during the late Paleozoic (Zonenshain and others, 1990; Biske, 1996; Windley and others, 2007). The older structures were overprinted by younger folds and faults in the late Cenozoic (Biske, 1996) during intracontinental deformation when India collided with Eurasia (Molnar and Tapponnier, 1975). Below we present a model for the geodynamic evolution of the southern accretionary margin of the CAOB and the Tarim Craton as has been inferred from the geologic evidence in southern Kyrgyzstan. Two competing tectonic models are currently under discussion: N-directed subduction of the Tarim Craton and South Tianshan (Windley and others, 1990; Allen and others, 1992; Biske, 1996; Xiao and others, 2004; Zhang and others, 2007; Makarov and others, 2010), and S-directed subduction of the Yili-Central Tianshan in NW China in the late Carboniferous (Charvet and others, 2007; Wang and others, 2008; Lin and others, 2009). Confirmation of any of these models awaits further structural studies in Kyrgyzstan and China.

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E. Hegner & others—Mineral ages and P-T conditions of late Paleozoic high-pressure

Fig. 1. Lithotectonic map of the Tianshan in Kyrgysztan and neighboring countries (modified after Osmonbetov, 1980; Biske and others, 1985). (A) Outline of the western Tianshan fold-and-thrust belt and major faults: (1) Talas Fergana, (2) Nikolaev Line, (3) South Tianshan Suture also referred to as Atbashi-Inylchek Fault in Kyrgyzstan. Localities of HP rocks are indicated by asterisks: At ⫽ Atbashi, Ak ⫽ Akeyazhi in NW China. Note that in NW China the “North Tianshan” of Kyrgyzstan is referred to as “Yili-Central Tianshan”. The inset shows the studied region and tectonic blocks in Eurasia: KZ ⫽ Kazakhstan; T ⫽ Tarim craton; K ⫽ Karakum, IND ⫽ India; SIB ⫽ Siberia; NC ⫽ North China, SC ⫽ South China. (B) Detailed geological map of the Atbashi Ridge with locality of dated eclogite and metagraywacke samples.

In eastern Kyrgyzstan and westernmost China the South Tianshan mainly consists of sedimentary rocks including shales, turbidites, cherts, and carbonates ranging in age from Silurian to Late Carboniferous, and locally to Early Permian; ophiolites, volcanic

eclogite and provenance of me´lange sediments from Atbashi in the South Tianshan

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and metamorphic rocks are subordinate (Biske, 1996). The structure of the South Tianshan in Kyrgyzstan comprises a package of S-vergent thrust sheets (D1 structures; Biske and Seltmann, 2010). These thrusts were overprinted by ENE-trending synforms and antiforms (D2 structures) and by s-shaped plunging folds, left-lateral strike-slip and reverse faults (D3 structures). The D3 reverse faults commonly demonstrate top-to-the-north sense of motion. This contrasts with the structures documented in the Chinese South Tianshan by Charvet and others (2007), Wang and others (2008) and Lin and others (2009) who observed N-vergent D1 thrust sheets. These different structural interpretations will be more fully discussed below. The ages of the deformed rocks in the thrust sheets and those of turbidites and olistostromes that were deposited during deformation indicate that thrusting began in the north of the South Tianshan in the middle Bashkirian (⬃315 Ma) and prograded southward until the early Permian (Biske and others, 1985; Biske, 1996). Within the South Tianshan and lithotectonic units to the north, post-kinematic A-type granites, including Rapakivi-types, yielded SHRIMP zircon ages of 299 to 273 Ma (Wang and others, 2007; Long and others, 2008; Konopelko and others, 2007, 2009; Biske and Seltmann, 2010) and constrain the time of D2 and D3 deformations to the Early Permian. ENE-trending strike-slip faults were produced in the mid-Permian to Triassic (Bazhenov and others, 1999). Metamorphic rocks are exposed along the northern margin of the South Tianshan belt, with the largest outcrops along the Atbashi ridge. The main lithologies are garnet-muscovite schist, garnet-albite-muscovite-chlorite-(⫾amphibole) schist, mica schist and phyllite. Phengite-bearing schist and paragneiss, marble, and mafic metavolcanic rocks are subordinate. HP/UHP eclogites and glaucophane schists form lenses and boudins in both schist and gneiss (Bakirov, 1978; Bakirov and Kotov, 1988). Similar eclogites occur along the 500 km-long South Tianshan suture zone that extends from the Atbashi Ridge in Kyrgyzstan to the Kekesu-Akeyazhi area in the Chinese West Tianshan (fig. 1A). The age of low-grade schists in the Atbashi Ridge is Silurian to Early Devonian based on Tabulata corals in carbonate lenses (Biske and others, 1985). The depositional age of higher-grade metasediments is uncertain. Various authors inferred either a Proterozoic age, based on K-Ar ages of 1100 and 570 Ma (Bakirov and others, 1974; Tursungaziev and Petrov, 2008) or a middle Paleozoic age based on a similar composition and possible gradual transitions between higher-grade and lower-grade metamorphic rocks (Osmonbetov, 1980; Khristov and Mikolaichuk, 1983). Below we report the first detrital zircon ages from a phengite paragneiss, interpreted as metagraywacke, from the Atbashi HP me´lange that place constraints on the maximum depositional age and provenance. The age of HP metamorphism for the Atbashi eclogites was inferred from various geochronological methods as mostly late Paleozoic (see discussion). Geological constraints on the age of HP metamorphism are provided by little deformed, unmetamorphosed, uppermost Carboniferous and Asselian (lowermost Permian, 303 Ma to 295 Ma) fluvial conglomerates, limestones and turbidites, which occur locally along the northern slope of the Atbashi Ridge (Tagiri and others, 1995; Tursungaziev and Petrov, 2008). The presence of eclogite pebbles in the conglomerate (Baslakunov and others, 2007) indicates that these ca. 303 to 295 Ma sediments were deposited after HP metamorphism, exhumation, and accretion of the me´lange. According to Biske and others (1985) and Biske (1996), the allochthonous metamorphic units are structurally underlain by a number of smaller tectonic slices, which consist of jasper and chert alternating with pillow basalt, subordinate black shale, and serpentenite me´lange (fig. 1B) that contains blocks of dunite, pyroxenite, gabbro, metamorphic and siliceous rocks. The presence of graptolites in black shales indicates a Silurian age (Biske and others, 1985). Conodonts constrain the age of a

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cherty-volcanic sequence including pillow basalts as Early to Late Devonian, for example earliest Frasnian age (⬃415 to 385 Ma; time scale according to Ogg and others, 2008), whilst conodonts in cherty sequences lacking volcanic rocks indicate Frasnian, Famennian, and Tournaisian (early Carboniferous) ages (Alekseev and others, 2007). The chemical composition of the Devonian pillow basalts suggests mid-ocean ridge and ocean island settings (Biske and Tabuns, 1991; Simonov and others, 2009). Similar REE patterns for the Devonian basalts and Atbashi eclogites were interpreted by Simonov and others (2009) as evidence for a genetic relationship. If the basalts represent the protolith of the eclogite, an early Devonian to earliest Frasnian magmatic age (Alekseev and others, 2007) can be inferred for the latter. The metamorphic rocks and cherty-volcanic strata with ophiolitic fragments are replaced southwards by a series of thrust sheets comprising shallow-marine carbonates, presumably originating from seamounts, as well as cherts and shales, assumed to have been deposited in a deeper marine basin. In the far south they were thrust onto shales and turbidites of the continental margin of the Tarim Craton (Khristov and Mikolaychuk, 1983; Biske, 1996). petrography of eclogite and phengite paragneiss samples

Eclogite sample KG 23 and phengite paragneiss (metagraywacke) sample KG 25 were collected from the accretionary me´lange of the South Tianshan suture on the northern slope of the Atbashi Ridge (fig. 1B). The samples are from the Kembel River valley, south of Aktala village, and ⬃15 km SW of the township of Atbashi. The sample locality is at N41°03⬘55.8⬙ and E75°41⬘29.9⬙ and ⬃500 km SW of the study area of Gao and Klemd (2003) and Zhang and others (2007) in the western Tianshan of NW China (fig. 1A), where these authors reported Late Carboniferous as well as Triassic ages for HP metamorphism (see discussion). Eclogite sample KG 23 contains ⬃30 volume percent of zoned garnet, the cores and mantles of which have inclusions of clinozoisite, rutile, and phengite; inclusions are absent in the outermost garnet rims. Tablet-shaped clinozoisite-paragonite intergrowths also occur as inclusions in the garnet cores, suggesting the former presence of lawsonite. Omphacite is fine-grained, subidiomorphic to xenoblastic and, together with garnet and glaucophane, constitutes the bulk of the matrix. Glaucophane mainly occurs as fine- to medium-grained fibrous aggregates and is commonly replaced by retrograde fibroblastic (fibrous and of equal size) actinolite, albite and chlorite that constitute most of the retrograde matrix minerals. An estimate of the modal composition of peak-metamorphic matrix minerals suggests ⬃30 volume percent fine- to medium-grained omphacite, ⬃15 volume percent phengite and ⬃18 volume percent partly chloritized glaucophane. Retrograde chlorite (⬃5 vol. %), quartz (1-2 vol. %), clinozoisite (⬃1 vol. %), and albite are subordinate. Rutile, which in places is replaced by titanite, as well as pyrite, monazite or xenotime, and apatite are accessories. Quartz pseudomorphs after coesite were not observed. Lepidoblastic phengite, in places intimately intergrown with chlorite, commonly occurs in close proximity to garnet porphyroblasts. P-T conditions for other eclogite samples from the Atbashi HP me´lange were published by Tagiri and others (1995) and Simonov and others (2008). The former authors reported quartz pseudomorphs after coesite and a lack of glaucophane, which is abundant in eclogite sample KG 23. They also recognized prograde and retrograde paths and five stages of metamorphic reactions with peak metamorphic conditions at ⬃660 °C at 25 kbar. Simonov and others (2008) reported similar peak P-T conditions of 23 to 25 kbar but at lower temperatures of 510 to 570 °C, and they showed that the metamorphic processes of eclogite formation involved fluids with 6 to 12 weight percent NaCl.


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The phengite paragneiss sample KG 25, interpreted as metagraywacke, is strongly deformed and foliated. It has a fine- to medium-grained granoblastic texture and a distinct foliation outlined by oriented quartz- and mica-rich layers. The sample contains a textural equilibrium mineral assemblage mainly of quartz (⬃43 vol. %), plagioclase (⬃35 vol. %), chlorite (⬃10 vol. %), phengite (⬃5 vol. %), and clinozoisite (⬃5 vol. %). Accessory minerals are titanite, apatite, actinolite, zircon, and calcite. We also found a single grain of xenoblastic garnet, which was almost completely corroded and replaced by chlorite. previous age determinations of atbashi eclogites 40

Ar/39Ar dating of phengite and glaucophane multi-grain separates from eclogite samples from the Atbashi Ridge yielded cooling ages of 327 to 324 Ma (Simonov and others, 2008). K-Ar ages of 320 to 288 Ma were reported by Udovkina (1985), and Rb-Sr dating of the HP phases in another eclogite sample yielded ⬃270 Ma (Tagiri and others, 1995). Sm-Nd dating of garnet gave an imprecise age of 350 ⫾ 150 Ma (Shatsky and others, 1988). results

Electron Microprobe Analyses of the Eclogite and Phengite Paragneiss Samples Eclogite sample KG 23.—Garnet in this eclogite comprises 58 to 64 mole percent almandine, 4 to 21 mole percent pyrope, 21 to 32 mole percent grossular, and 0.1 to 1 mole percent spessartine. The chemical zoning is prograde with Fe2⫹/(Fe2⫹⫹Mg2⫹) ranging from 0.75 in the rim to 0.93 in the core as determined from detailed line traverses over the garnet grains; grossular and spessartine proportions decrease from core to rim (table A.1, Appendix). The clinopyroxene is omphacite, the jadeite (Jd) content of which ranges from 35 to 48 weight percent (table A.2, Appendix) according to the classification of Morimoto (1988). No chemical zoning was observed in single grains as is evident from detailed line traverses. However, the omphacite composition is rather heterogeneous, implying some re-equilibration during retrograde conditions. Amphibole includes texturally peak-metamorphic primary glaucophane according to the classification of Leake (1978). In places, glaucophane is replaced along rims and fractures by retrograde actinolite and/or actinolitic hornblende (table A.3, Appendix). Other eclogite minerals include unzoned phengite with Si contents of 6.73 to 6.91 p.f.u. (table A.4, Appendix). The pistacite component [Fe3⫹/(Al3⫹⫹Fe3⫹)] ⫻ 100 of the clinozoisite, with Fetot ⫽ Fe3⫹, ranges from 19.50 to 25.81 (table A.5, Appendix). Retrograde plagioclase is albite (An4-8; table A.6, Appendix). Phengite paragneiss sample KG 25.—Colorless phengite is the only white mica in this sample. Large phengite blades display high Si-contents of 6.66 to 6.92 p.f.u. (table A.4, Appendix) similar to the composition of phengite in the eclogite. The pistacite component [Fe3⫹/(Al3⫹⫹Fe3⫹)] ⫻ 100 of the clinozoisite, with Fetot⫽Fe3⫹, ranges from 10.50 to 13.20 (table A.5, Appendix). Plagioclase is albite (An⬍2; table A.6, Appendix). Most chlorites have a ripidolitic composition with Al contents ranging from 4.79 to 4.86 (table A.7, Appendix) and Mg-values of about 0.5. Major and Trace Element Data The major and trace element concentrations of eclogite sample KG 23 and metagraywacke sample KG 25 are listed in table 1, and their trace element data are plotted in figure 2. The eclogite has a chemically evolved basaltic composition as indicated by its moderately high Mg-number of ⬃49 (Mg-number ⫽ molar Mg2⫹/ (Mg2⫹ ⫹ Fe2⫹)) corresponding to a MgO concentration of ⬃7 weight percent, and low Ni and Cr concentrations of 45 and 126 ppm, respectively. A tholeiitic affinity is supported by a high FeO(t)/MgO ratio of ⬃2.1 at ⬃48 weight percent SiO2 (Miyashiro, 1974, table 1) and a high Cr concentration relative to that of Y (Pearce and

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E. Hegner & others—Mineral ages and P-T conditions of late Paleozoic high-pressure Table 1

Chemical composition of eclogite sample KG 23 and metagraywacke sample KG 25 Sample SiO2 TiO 2 Al2O3 Fe2O3(t) MnO MgO CaO Na2O K2 O P2 O 5 Total LOI Sr Ba Hf Y Zr Ta Nb U Th Pb Ga Cr Ni La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

KG 23* 47.73 1.70 13.64 15.98 0.22 6.98 9.54 3.44 0.64 0.09 99.96 1.75 45.5 84.5 2.77 43.9 94.3 0.967 2.88 0.154 0.237 1.67 18 126 45 2.93 9.09 1.61 9.23 3.45 1.30 5.80 1.11 8.01 1.71 5.20 0.755 5.12 0.766

KG 25* 69.75 0.71 12.64 5.99 0.09 3.14 2.80 2.71 1.50 0.71 100.04 2.81 737 298 6.38 20.4 249 1.00 9.73 2.15 9.99 11.0 n.d. n.d. n.d. 27.4 60.1 6.62 25.4 5.07 1.12 4.54 0.659 4.05 0.804 2.30 0.337 2.21 0.343

BCR-1**

Recom.**

324 655 4.97 32.3 188 0.70 11.4 1.79 5.94 14.0 n.d. n.d. n.d. 25.1 53.3 6.79 28.6 6.54 1.92 6.60 1.02 6.52 1.27 3.69 0.51 3.41 0.50

334 682 4.90 36 189 0.79 12.8 1.69 5.90 13.5 23 13 13 25.2 53.7 6.80 28.7 6.62 1.95 6.74 1.07 6.33 1.28 3.62 0.54 3.37 0.50

* Major-element concentrations reported as anhydrous values in wt. %, trace-element concentrations in ␮g/g. ** International rock standard (United States Geological Survey, USGS, Basalt Columbia River); recommended values from GeoReM (http://georem.mpchmainz.gwdg.de).

Norry, 1979). Its composition is similar to that of a chemically evolved mid-ocean ridge basalt (MORB, for example Klein, 2003), although a K2O concentration of 0.6 weight percent is much higher than typical values of 0.2 weight percent in fresh MORB (Klein, 2003). As potassium is mobile during seafloor alteration and metamorphism, and


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Fig. 2. Normalized trace-element patterns for eclogite sample KG 23 and metagraywacke sample KG 25. For comparison are shown typical patterns of N-MORB (Hofmann, 1988) and island arc basalt (IAB; Sun, 1980). The latter resembles supra-subduction zone ophiolite basalt (Pearce and others, 1984). (A) LREE depletion and overall high REE concentrations in KG 23 indicate a N-type MORB protolith for the eclogite. Metagraywacke sample KG 25 shows typical features of felsic upper crust. (B) Extended REE pattern of eclogite KG 23 showing overprinting by the fluid-mobile elements Ba, U, and Pb. Inset depicts the subduction-related trace-element characteristics in metagraywacke sample KG 25. Normalizing values are from McDonough and Sun (1995).

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Sm-Nd isotope data for eclogite sample KG 23 and metagraywacke sample KG 25

Sample KG 23- whole rock Omphacite Glaucophane Garnet core-mantle Garnet rim KG 25- whole rock

Sm [µg/g] 3.475 2.911 1.265 0.8161 0.9716 5.275

Nd [µg/g] 9.274 8.831 3.803 0.8624 0.8934 27.48

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Sm/144Nd 0.2265 0.1993 0.2012 0.5722 0.6575 0.1160

Nd/144Nd (m.) εNd(t)

143

0.513172 ± 8 0.513117 ± 9 0.513125 ± 9 0.513895 ± 10 0.514082 ± 19 0.512047 ± 12

9.3

-7.4

143 Nd/144Nd normalized to 146Nd/144Nd ⫽ 0.7219. External precision for 143Nd/144Nd is ⬃1.0 ⫻ 10⫺5 (2 SD). Error of 147Sm/144Nd ⬃0.05% (2 SD). The 143Nd/144Nd ratios are relative to 143Nd/144Nd ⫽ 0.511847 ⫾ 8 (2 SD, N ⫽ 10) in the La Jolla Nd standard. m ⫽ measured ratio. εNd(t) ⫽ 320 Ma for KG 23 and 400 Ma for KG 25, yielding tDM ⫽ 1.5 Ga (DePaolo, 1988).

during slab dehydration during subduction, the significance of the high K2O concentration is difficult to assess. It may be due to enrichment of the sample in the subduction channel as suggested by other fluid-mobile trace elements (see below). The eclogite shows a distinct LREE-depletion, a flat HREE pattern (fig. 2A), and overall REE abundances as in N-MORB (for example, Sun, 1980; Hofmann, 1988). Negative Eu- and Sr anomalies (fig. 2B), due to plagioclase fractionation, are in agreement with the evolved chemical composition of the sample. Noteworthy is the high Nb concentration relative to La (fig. 2B) and low La/Nb ratio, which are in contrast to typical arc characteristics, precluding melting of mantle sources that have been affected by fluids and/or melts in arc environments (for example, Pearce, 1982). A high Nb/Th ratio at low Th concentration suggests a mantle source lacking a sedimentary component as commonly observed in settings remote from subductionzones (for example, Plank, 2005). Positive anomalies for the fluid-mobile elements Pb, U, and Ba (for example, Sun, 1980; Kay, 1980; White and Dupre´, 1986), in the presence of a slightly positive Nb anomaly, are best explained by overprinting of the sample by hydrothermal fluids in a subduction channel. Metagraywacke sample KG 25 has a granodioritic to granitic bulk composition (table 1) and exhibits LREE-enrichment and a negative Eu anomaly (fig. 2A). A graywacke protolith derived from a chemically mature upper crust is supported by a negative Nb and a positive Pb anomaly (inset, fig. 2B). Positive anomalies of Zr and Hf (inset, fig. 2B) suggest accumulation of detrital zircons (for example, Taylor and McLennan, 1985). Sm-Nd Isotopes and Mineral Isochron Age of the Eclogite The Sm-Nd isotope compositions of eclogite sample KG 23 and metagraywacke sample KG 25, as well as the HP minerals in the eclogite, are listed in table 2. An isochron for the whole-rock and HP mineral assemblage of the eclogite is shown in figure 3. Data points for omphacite and glaucophane show similar compositions lying on the enriched side of the whole-rock composition. The inclusion-free rim material of the garnets shows the most depleted composition. Material from the core-mantle of the garnets that contain inclusions of clinozoisite, rutile and phengite plot on a tie-line between the inclusion-free garnet rim and whole-rock compositions. This relationship indicates that isotope equilibrium between mineral inclusions and the surrounding garnet and whole-rock was achieved at the time of HP recrystallization of the sample and closure of the Sm-Nd system. A mass balance calculation of the petrographically determined garnet, omphacite, and glaucophane proportions and the Sm and Nd


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Fig. 3. Mineral-whole rock isochron for eclogite sample KG 23. Data point “Grt-c” represents the garnet core-mantle domain with inclusions of clinozoisite, rutile, and phengite. “Grt-r” represents inclusion-free rim material. WR ⫽ whole-rock analysis, Gln ⫽ glaucophane, Omp ⫽ omphacite. Input errors (2 SD) for regression analysis: 147Sm/144Nd ⫽ 0.05 %, 143Nd/144Nd ⫽ 0.002 %, Grt-r ⫽ 0.004 %. Three-point data regression WR, Grt rim, Grt core/mantle gives 320.0 ⫾ 5.6 Ma (2␴, MSWD ⫽ 0.5). Regression analysis was performed with ISOPLOT 3.09 (Ludwig, 2003). For data see table 2.

concentrations in table 2 suggest that large amounts of the LREE are hosted in accessory phases such as rutile, monazite, and/or apatite. A regression analysis of the data for the whole-rock and the high-pressure phases (5 data points) shows excellent alignment of the data points (MSWD ⫽ 0.4) and yields a precise and statistically robust isochron age of 319.2 ⫾ 4.3 Ma (2␴; ISOPLOT 3.09; Ludwig, 2003). The data for whole-rock and the eclogite-facies minerals omphacite and garnet alone yield an indistinguishable age of 319.9 ⫾ 4.6 Ma (2␴, MSWD ⫽ 0.3). The initial εNd value is ⫹9.3, consistent with a long-term LREE-depleted MORB-like source for the protolith of the eclogite. The MSWD value of ⬍1 indicates data scatter well within analytical uncertainty and shows that all analyzed minerals and the whole-rock behaved as isotopically open systems during recrystallization at HP and had attained isotopic equilibrium before closure of the Sm-Nd system against element diffusion during exhumation and cooling of the eclogite. 40

Ar/39Ar Ages of Phengite The results of laser step-heating Ar analyses of two single phengite grains from eclogite sample KG 23 are listed in table 3, and the degassing spectra are shown in figure 4. Age regression of the data of phengite grain KG 23-1 yields a plateau age of 316.5 ⫾ 3 Ma (2␴) based on 88 percent of the released 39Ar (steps 3, 4, 5, and fuse; table 3). The data of the second grain KG 23-2 do not fulfill our criteria given in the Appendix for defining a plateau age, and yield an integrated total gas age of 316.5 ⫾ 2 Ma (2␴), identical to that derived from the degassing plateau of sample KG 23-1.

926


Laser

40

Ar/39Ar analytical data for two phengite grains from eclogite sample KG 23

Temperature Steps Phengite grain 1 1 2 3 4 5 fuse Total gas age Phengite grain 2 1 2 3 4 fuse Total gas age

Atmospheric contam. (%)

39

Ar (%)

37

ArCa /39ArK

40

Ar*/39ArK

Age (Ma ± 1σ)

36.10 22.91 11.79 13.43 10.24 13.00

2.05 9.89 28.40 6.89 39.78 12.99

0.006 0.001 b.d. 0.002 b.d. 0.001

18.07 20.65 19.91 19.57 19.81 19.68

290.9 ± 12 328.9 ± 2 318.2 ± 2 313.1 ± 9 316.6 ± 2 314.6 ± 4 317.5 ± 1.5

4.13 5.31 15.70 10.46 8.68

3.34 5.46 8.11 35.28 47.81

0.003 0.002 0.001 b.d. 1 b.d.

20.853 19.923 20.588 9.743 19.633

331.8 ± 8.1 318.2 ± 5.7 327.9 ± 4.0 315.6 ± 1.6 313.9 ± 1.5 316.5 ± 1.0

40 Ar* ⫽ radiogenic 40Ar; Ca ⫽ produced by Ca-neutron interferences; K ⫽ produced by Kneutron interferences. b.d. ⫽ below detection.

Fig. 4. Laser step-heating Ar degassing spectra for two phengite grains from eclogite sample KG 23. Spectrum of grain KG23-1 is shown in black bars and that of grain KG23-2 in white bars. Age uncertainties are given at the 2␴ confidence level (P ⫽ plateau age; T ⫽ total gas age).


927

Hence we accept 316.5 Ma as the 40Ar/39Ar closure age for cooling of the phengite below 400 ⫾ 50 °C (McDougall and Harrison, 1988). discussion

Pressure-Temperature Conditions of HP Metamorphism The prograde metamorphic evolution of eclogites and blueschists in the western Tianshan of NW China was described by Gao and others (1999) and Klemd and others (2002), who demonstrated a transition from lawsonite-blueschist to eclogitefacies conditions at 530 ⫾ 30 °C and 14 to 21 kbar. For the Atbashi eclogites, Simonov and others (2008) reported a higher maximum pressure range of 23 to 25 kbar at temperatures ranging from 510 to 570 °C. Inclusions of possible coesite pseudomorphs in garnet at Atbashi may indicate a passage through the coesite stability field before peak-temperature conditions were reached (Tagiri and others, 1995). UHP conditions were also reported from NW China by Lu¨ and others (2008) who found inclusions of coesite in eclogitic garnet. The prograde chemical zoning of garnet in eclogite sample KG 23 is interpreted as a result of decreasing pressure and increasing temperature during garnet growth. Thus, the garnet rims crystallized at higher temperature and lower pressure than the cores. This zoning behavior indicates a clockwise P-T path with increasing temperature and associated decreasing pressure. The maximum pressure mineral paragenesis in eclogite sample KG 23 consists of garnet-omphacite-Na-amphibole-phengite-clinozoisite-quartz, ⫾rutile and apatite. An H2Oindependent pressure estimate can be obtained from the reaction albite ⫽ jadeite ⫹ quartz (Holland, 1980, 1983). In the absence of prograde albite, the maximum jadeite content of 48 mole percent indicates minimum pressures of 13.5 and 16 kbar at 500 and 600 °C, respectively. The garnet-omphacite-phengite equilibrium after Waters and Martin (1993) and Waters (1996) gives equilibrium pressures of 18 to 23 kbar at 490 to 590 °C using omphacite with the highest jadeite content, texturally primary phengite (not in contact with garnet and omphacite) with the highest Si-content, and prograde garnet rims. Temperature estimates were obtained with the grt-cpx geothermometers of Ellis and Green (1979) and Krogh-Ravna (2000). The former yields a temperature range of 595 to 655 °C and 605 to 680 °C at 18 and 23 kbar, respectively. The temperatures calculated with the Krogh-Ravna calibration are up to 100 °C lower and vary between 490 and 590 °C. In addition, peak equilibrium conditions in the eclogite are reflected in the paragenesis of garnet ⫹ omphacite ⫹ Na-amphibole ⫹ clinozoisite ⫹ phengite ⫹ rutile, which indicates maximum pressures between 18 and 24 kbar at temperatures ranging from 520 to 600 °C, using the petrogenetic grid for mafic rocks of Evans (1990; fig. 5, this study). These P-T conditions agree with those estimated for other Atbashi eclogites (Simonov and others, 2008) and eclogites from the western Tianshan in NW China (Gao and others, 1999; Klemd and others, 2002). The mineral inclusions in the garnet of eclogite sample KG 23 indicate a prograde transition from lawsonite-blueschist/epidote-blueschist to eclogite-facies conditions, in agreement with the P-T evolution of glaucophane-bearing eclogites in the western Tianshan in NW China (Gao and others, 1999; Klemd and others, 2002). Furthermore, the absence of post-peak retrograde glaucophane and the presence of retrograde albite and associated clinozoisite indicate a transition from eclogite- to albite-epidote amphibolitefacies conditions according to the P-T grid of Evans (1990; fig. 5, this study). The assemblage phengite-chlorite-albite-clinozoisite-quartz in phengite paragneiss sample KG 25 is typical of greenschist-facies metamorphic conditions (for example, Evans, 1990). The phengite analyses in table A.4, Appendix, show Si contents as high as 6.7 to 6.9 p.f.u. as in the phengite of eclogite sample KG 23 and are interpreted to be the result of incomplete re-equilibration during retrograde condi-

928


319 Ma 20 E

LBS 16 P k b ar

EBS z + Qt Jd 50 Ab

12

8 PA

AEA 316 Ma

GS A

4

300

400 T °C

500

600

Fig. 5. Inferred P-T path of eclogite sample KG 23 (dashed line) and metagraywacke (phengite paragneiss) sample KG 25 (dotted line). P-T grid after Evans (1990). LBS, lawsonite-blueschist-facies; E, eclogite-facies; PA, pumpellyite-actinolite-facies; EBS, epidote-blueschist-facies; GS, greenschist-facies; A, amphibolite-facies; AEA, albite-epidote-amphibolite-facies. Ab ⫽ Jd50 ⫹ Qtz after Holland (1983). For data see tables A.1 to A.7 in the Appendix.

tions after eclogite-facies metamorphism (see discussion of Klemd and others, 1991). This can be concluded because large phengite blades or phengite inclusions in garnet porphyroblasts of otherwise greenschist-facies metasediments may preserve high Si contents, in agreement with eclogite-facies conditions (for example, Klemd and others, 1991). The assumed high-pressure phengite relics in the paragneiss suggest that this sample and the eclogite underwent the same metamorphic development. The preservation of prograde, peak, and retrograde mineral assemblages in the eclogite and phengite paragneiss supports a clockwise P-T path showing nearisothermal decompression from post-eclogite-facies to albite-epidote-amphibolite facies conditions (fig. 5). This suggests that the Atbashi high-pressure rocks formed in an “Alpine-type” regime similar to that suggested for the high-pressure rocks from the Chinese western Tianshan (Klemd and others, 2002). Tectonic Setting of the Eclogite Protolith Eclogite samples from the Atbashi Ridge (Sobolev and others, 1989; Simonov and others, 2008) and the Akeyazhi River in NW China (Gao and Klemd, 2003) show a spectrum of LREE-depleted to enriched patterns that were interpreted as evidence for


929

subducted N-MORB, OIB, and plateau-type basalts. The pattern of immobile REE in eclogite sample KG 23 is similar to that of N-MORB (fig. 2A). The overall REE abundances show that the protolith of the eclogite originated by similar degrees of melting of a depleted upper mantle, as inferred for MORB. An old and highly depleted MORB-like upper mantle source is also reflected in the high initial εNd value of ca. ⫹9 assuming a Devonian-Carboniferous magmatic age for the protolith. A small and slightly positive Nb-anomaly in the eclogite is clear evidence of melting of a mantle source not affected by subduction processes (for example, Sun, 1980; Hofmann, 1988). All these immobile trace element characteristics show that the protolith belongs to the small group of ophiolites with N-MORB characteristics (for example, Metcalf and Shervais, 2008). The high total REE concentrations and lack of a negative Nb-anomaly in the eclogite preclude a supra-subduction zone and arc origin of the protolith as rocks from these settings typically show lower total REE abundances than MORB and exhibit negative Nb-anomalies (see fig. 2A; Sun, 1980; Pearce and others, 1984). An origin of the eclogite protolith in a back-arc spreading center is also unlikely as such samples typically exhibit subduction-related trace element characteristics, albeit of variable magnitudes (for example, Monnier and others, 1995; Bach and others, 1998). Considering the N-MORB affinity of the eclogite protolith, it is surprising to find enrichment of the fluid-mobile elements Ba, U, and Pb in the normalized trace element patterns (fig. 2B). Enrichment of these elements characterizes rocks from subduction-modified mantle. However, these rocks also show a negative Nb-anomaly relative to La. As the negative Nb-anomaly is such a robust criterion for melting of subduction-modified upper mantle, we propose that the Ba, U, and Pb enrichment was caused by hydrothermal fluids in the subduction channel (for example, Gao and others, 2007). Geochronological Evidence for Carboniferous Subduction and Collision in the South Tianshan of Kyrgyzstan The K-Ar, Rb-Sr, and Sm-Nd systems reveal a wide range of cooling ages of 350 to 270 Ma for the HP-LT metamorphism in southern Kyrgyzstan. A Late Carboniferous rather than Triassic HP event, as also constrained by other geological evidence, is supported by the recently published 40Ar/39Ar ages of Simonov and others (2008; see “previous age determinations”). The isochron produced by the whole-rock and the HP mineral assemblage in eclogite sample KG 23 indicates open-system recrystallization and Sm-Nd isotopic closure at about 319 Ma. Isotopic equilibrium among all analyzed phases of the eclogite, including omphacite that commonly shows slow isotope resetting (for example, Mork and Mearns, 1986), may have been facilitated by the small grain size and the presence of pervasive saline fluids (Simonov and others, 2008) as well as temperatures high enough to facilitate element diffusion. The 40Ar/39Ar cooling age of 316 Ma is indistinguishable within error from the Sm-Nd age and consistent with a fast cooling rate for the eclogite. Neglecting the errors of the ages, a high cooling rate of ⬃50 °C/Ma is suggested, assuming temperatures of 600 °C for closure of the Sm-Nd system in garnet and 400 ⫾ 50 °C of the K-Ar system in phengite (McDougall and Harrison, 1988). The geochronological results show that at ⬃319 Ma the basaltic protolith of the eclogite had been subducted to eclogite-facies depth where it recrystallized, and this was followed by rapid isothermal decompression, until it reached neutral buoyancy, probably at mid-crustal levels of 10 to 20 km by ⬃316 Ma (Ernst, 1988, Platt, 1993). Contemporaneously, accretion and subduction continued as can be inferred from deposition of turbidites and ongoing arc-magmatism in the Middle Tianshan. Exhuma-

930


tion and exposure of the HP rocks clearly occurred before ⬃300 Ma when conglomerates were deposited unconformably on the HP me´lange (see the following section). Eclogite samples from the South Tianshan Suture some 500 km to the east near Kekesu in the Akeyazhi River Valley in NW China initially yielded significantly older ages for HP metamorphism and exhumation than in Kyrgyzstan as suggested by 40 Ar/39Ar ages of 344 ⫾ 1 Ma for glaucophane and 331 ⫾ 2 Ma for phengite (Gao and Klemd, 2003) and by an imprecise Sm-Nd errorchron age of 343 ⫾ 44 Ma for eclogite-facies minerals (Gao and Klemd, 2003). These authors interpreted the Sm-Nd and 40Ar/39Ar ages in the context of formation of HP minerals and rock exhumation sometime between 340 and 325 Ma and concluded that peak metamorphism occurred ⬃344 Ma ago. Rb-Sr and 40Ar/39Ar closure ages of peak metamorphic white micas range from 313 to 302 Ma and 323 to 312 Ma, respectively, and shed light on the post-peak metamorphic stages of the complex (Klemd and others, 2005). Ionmicroprobe-dated zircon rims from eclogites from the same locality recently yielded a mean age of ⬃319 Ma that was interpreted as the best estimate for eclogite-facies metamorphism before collision of the Tarim and Yili-Central Tianshan blocks in NW China (Wei and others, 2009; Wen and others, 2009). How these new young ages relate to the previously published older ages remains enigmatic and awaits clarification. In any case, the new zircon ages from the Chinese locality agree well with the Sm-Nd garnet ages presented here from southern Kyrgyzstan. The discussion about the temporal framework of HP/UHP metamorphism in the Tianshan of NW China has received new input after Zhang and others (2007) reported SHRIMP zircon rim ages of 233 to 226 Ma for grains with magmatic cores of ⬃310 to 410 Ma. These authors interpreted the age of the rims as that of the HP metamorphism and Triassic collision of the Tarim craton with the CAOB. However, more recent U-Pb and Sm-Nd age determinations of Zhang and others (2009) also provide evidence for metamorphic events at ⬃309 to 327 Ma and 243 to 263 Ma whose tectonic significance within the evolution of the South Tianshan Suture remains to be determined. We conclude that the apparently straightforward geochronological and geological evidence (see next section) for Late Carboniferous subduction and collision in the South Tianshan of Kyrgyzstan contrasts with the complex metamorphic history in NW China that may only be unraveled with new geological constraints. Geological Evidence for Late Paleozoic Subduction and Collision in the South Tianshan of Kyrgyzstan Geological data for the South Tianshan in Kyrgyzstan indicate a northwarddipping subduction zone, which caused convergence of the Middle and North Tianshan and the Tarim craton in the Late Carboniferous to Early Permian. The inferred subduction polarity is based on: a) prevailing top-to-the-south direction of primary thrusting, b) southward propagation of thrust deformation with time from the late Bashkirian (⬃315-312 Ma, fossil evidence) at Atbashi Ridge to Early Permian (fossil evidence) in the northern Tarim craton (Biske, 1996; Bazhenov and Burtman, 1997; Burtman, 2008), and c) a Late Carboniferous volcanic belt, coeval with thrusting and located north of the thrust belt (Alekseev and others, 2009). Palaeozoic top-to-thesouth thrusting and inferred north-dipping subduction in the Tianshan of southern Kyrgyzstan, however, contrast with tectonic models for the western Tianshan of NW China where top-to-the-north thrusting and south-directed subduction has been inferred (Charvet and others, 2007; Wang and others, 2008; Lin and others, 2009). Clarification of this problem is beyond the scope of this paper and requires more studies, in particular across-border observations. Stratigraphic sequences from Frasnian to lower Bashkirian age in areas north and south of Atbashi from the Middle Tianshan to the Tarim craton do not contain volcanic or tuffaceous sedimentary rocks (Khristov and Mikolaichuk, 1983; Biske,


931

1996; Cook and others, 2002; Alekseev and others, 2007), and this suggests absence of subduction-related magmatism in this part of the belt from ⬃385 to ⬃315 Ma. Paleozoic southward-directed thrusting was accompanied by deposition of deep marine turbidites in the late Bashkirian and Moscovian (311-307 Ma, inferred from fossils) in the South Tianshan Suture (Burtman, 1975; Khristov and Mikolaichuk, 1983; Biske and others, 1985; Biske, 1996; Bazhenov and Burtman 1997; Burtman, 2008) and sub-synchronous volcanic activity in the Naryn area of the Middle Tianshan (Alekseev and others, 2009). These relationships constrain the time of subduction to ⬃315 to 305 Ma. Initiation of a foreland basin along the northern side of the Tarim craton at the end of the Late Carboniferous (Allen and others, 1999) confirms the beginning of collision with the Tarim craton in this part of the orogen. Cessation of volcanism in the Middle Tianshan (Naryn area) in the Kasimovian (⬃303-307 Ma; Alekseev and others, 2009) may also be related to the start of this collision. During the earliest stages of collision in the Gzhelian and Asselian, the north Atbashi area was dominated by sedimentation of slope turbidites with subordinate shallow marine limestones and conglomerates. Both sediment transport directions based on flute casts in turbidites and the composition of pebbles in conglomerates, indicate that the provenance was located in the north, and imply that the Southern Tianshan was not uplifted at that time. The occurrence of pebbles derived from the South Tianshan in the middle Asselian and the appearance of marine basins throughout the Tianshan in Kyrgyzstan at the end of the Asselian, constrain the time of active uplift to after ⬃294 Ma. Uplift, folding, and emplacement of A-type, including Rapakivi-type, granitoids in the early Permian at ⬃300 to 270 Ma in China (Long and others, 2008) and Kyrgyzstan (Konopelko and others, 2007, 2009) indicate a late- to post-collisional environment during that period. Our Sm-Nd age of 319 ⫾ 4 Ma for HP metamorphism of the Atbashi eclogite thus corresponds to a time when subduction of oceanic lithosphere was still ongoing as can be inferred from geological evidence. The HP complex is unconformably overlain by a sequence of conglomerates, sandstones, and limestones (from bottom to top; Tagiri, 1995; Tursungaziev and Petrov, 2008) containing Fusulinida fossils of Gzhelian age (304-299 Ma; A. Neievin, personal communication). The presence of eclogite pebbles in the above conglomerates (Baslakunov and others, 2007) indicates that the HP metamorphic rocks were already exhumed at the end of the Carboniferous. Taking into account the early Permian ages of stitching A-type granitoids in the South Tianshan and the Tarim craton (for example Solomovich, 2007) it is unlikely that ocean closure and plate collision occurred in southern Kyrgyzstan during the Early Triassic. Detrital Zircon Ages and Provenance of the Me´lange Metagraywacke Detrital zircons in phengite paragneiss sample KG-25, interpreted as a metagraywacke, were dated in order to constrain the maximum depositional age and provenance of the sedimentary material in the Atbashi HP me´lange. We report 206 Pb/238U ages for grains ⬍1000 Ma and 207Pb/206Pb ages for grains ⬎1000 Ma (for explanation see the Appendix). The whole-rock sample has a mean Nd crustal residence age of 1.5 Ga and an initial εNd value of ⫺7.4 for a depositional age of ⬃400 Ma (table 2). The zircon grains are colorless to light brown and have shapes ranging from near-euhedral to well-rounded (figs. 6A and 6B). Mechanical rounding is particularly evident through pitted surfaces seen at high magnification under a binocular microscope, whereas no multifaceted grains due to metamorphic growth were observed. These features suggest variable mechanical reworking during short (near-idiomorphic zircons) and long (well-rounded zircons) sediment transport and imply that proximal as well as distal sources contributed to sediment formation.

932


Fig. 6. Cathodoluminescence images of a representative selection of detrital zircons in metagraywacke sample KG 25. (A) Grains analyzed on SHRIMP; circles define analyzed spots and ages refer to table A.8. (B) Grains analyzed by LA-ICPMS; circles define analyzed spots and ages refer to table A.9. Note lack of metamorphic zircon rims.

Under cathodoluminescence, most zircons show oscillatory zoning but, surprisingly, no metamorphic low-U overgrowths were recognized (fig. 6). A few grains are highly metamict and were not analyzed. Eight detrital zircon grains were analyzed on the Beijing SHRIMP II (table A.8, Appendix). All analyses are concordant (fig. 7A, table A.8, Appendix) and vary widely in age from 830⫾11 Ma (206Pb/238U age) to 2527⫾4 Ma (207Pb/206Pb age). Twenty-four further detrital grains were analyzed by LA-ICPMS in Mainz, and these results also scatter widely between an even wider range of 427 and 2774 Ma (fig. 7B; table A.9, Appendix). We note that the number of 32 zircon ages for the paragneiss sample is far too low for a detailed provenance analysis and only provides a sketch of the age spectrum of the source. The four youngest, near-idiomorphic grains have 206Pb/238U ages of 427 ⫾ 7,429 ⫾ 7, 433 ⫾ 9, and 445 ⫾ 7 Ma (table A.9, Appendix). Such ages are typical of granitoids in the North Tianshan in Kyrgyzstan (Jenchuraeva, 2001; Kiselev and Maksumova, 2001), and these grains therefore could have been derived from those granitoids. A more proximal source may be an island arc complex in the South Tianshan where a calc-alkaline diorite has a U-Pb zircon age of 436 ⫾ 3 Ma (Kro¨ner and others, 2009). The graywacke therefore has a likely depositional age ⬍427 Ma and most probably is a high-grade equivalent of the low-grade Silurian and Devonian shale and sandstone units, which have similar compositions and are widespread in the Atbashi Ridge. The age spectrum of the older zircons in the phengite paragneiss sample is similar to that reported from the Tarim Craton (Lu and others, 2008, fig. 8) and the North and Middle Tianshan (Kiselev and Maksumova, 2001; Kro¨ner and others, 2009). In


933

Fig. 7. Concordia diagrams showing (A) SHRIMP detrital zircon analyses from metagraywacke sample KG 25. For data see table A.8, Appendix. (B) LA-ICPMS detrital zircon analyses from metagraywacke sample KG 25. For data see table A.9, Appendix.

particular, ages in the Grenvillian range of ⬃1300 to 900 Ma are ubiquitously absent in Siberia and northern Gondwana but are abundant in the Tarim craton. Ages ranging from 850 to 750 Ma are known in the Tarim craton, and the North and the Middle

934


Fig. 8. Cumulative probability plots for detrital zircon ages for metagraywacke sample KG 25 and magmatic zircon ages for the Tarim craton (Lu and others, 2008). Plots were constructed with ISOPLOT version 3.09 (Ludwig, 2003).

Tianshan, which raises the question whether the North and Middle Tianshan may have been a part of the Tarim craton in Neoproterozoic times. The presence of Archean igneous rocks has been documented in the Tarim craton (Lu and others, 2008, fig. 8), but is questioned in the North Tianshan (Kro¨ner and others, 2009). However, abundant Archean and Paleoproterozoic detrital zircons were found in Neoprotero-


935

zoic sandstones of the Talass Ridge (Khudoley and Semiletkin, 2008) and show that Archean rocks probably comprise parts of the basement of the North and/or Middle Tianshan. Paleoproterozoic and Archean zircon ages are also known in the Anrakhai Mountains of southern Kazakhstan (Kro¨ner and others, 2007). Tectonic reconstructions for the middle Paleozoic South Tianshan basin of Kyrgyzstan (Khristov and Mikolaychuk, 1983; Biske, 1996) suggest that clastic sediments on the northern side of the South Tianshan could have been deposited on the continental slope of the Kazakhstan continent, whereas the Tarim craton was located on the opposite side of an oceanic basin of unknown width, making it less plausible that sediment was derived from this continental block. Accepting these observations we suggest that the provenance of the Atbashi protolith clastic sediments was probably to the north of the ocean basin in the North and Middle Tianshan. acknowledgments

A. Bakirov, K. Sakiev, S. Mikolaichuk are thanked for sharing their geological expertise during the International Excursion and Workshop “Tectonic evolution and crustal structure of the Tien Shan Belt”, Bishkek, June 2009. The manuscript was reviewed and improved by G. Ernst, J. Gao, and S. Wilde. This study was supported by Deutsche Forschungsgemeinschaft (grant KR 590/90-1 to AK), the Russian Foundation for Basic Research (grant 09-05-91331-NNIO-a` to DVA), the Beijing SHRIMP Center, the Germany-Hong Kong Research Scheme of DAAD, and the Hong Kong Research Council. M. G. Barth helped with LA-ICP-MS analyses and D. Jacob, University of Mainz, kindly provided access to the laboratory. Appendix: analytical methods and data tables Electron Microprobe Analysis

Electron microprobe studies were carried out on a SX-50 CAMECA microprobe at the Department of Mineralogy at Wu¨rzburg University. Operating conditions were 15 kV acceleration voltage, 15 nA beam current, and counting times of 20 to 30s. The beam diameter was set at 1 ␮m for all phases except for micas that were analyzed with a 3 to 5 ␮m beam diameter. Natural and synthetic minerals were used for standardization. The raw data were corrected with the ZAF procedure of the PAP software provided by CAMECA. The amphibole formulae normalization and estimates of the ferric/ferrous iron ratio followed the procedure of Robinson and others (1982). The amphibole nomenclature is based on the classification of Leake (1978). The Fe3⫹ content of pyroxene was inferred with the method of Vieten and Hamm (1978). Mineral abbreviations follow the recommendations of Kretz (1983). Major and Trace Element Analysis

The major element concentrations of the eclogite and phengite paragneiss samples were determined by XRF (Magic Pro PANalytical) on lithium-metaborate disks at Munich University. Concentrations of REE and other trace elements were determined by ICP-MS at the GeoForschungsZentrum, Potsdam. Details of the analytical methods and element concentration data showing the precision and accuracy of the method are reported in Dulski (2001) and given in table 1. Sm and Nd Isotope Analysis

The eclogite sample weighing ca. 6 kg was crushed and the ⬍0.5 mm heavy mineral fraction separated with a Wilfley shaking table for separation of zircons. A portion of the heavy mineral fraction was cleaned with water and acetone in an ultrasonic bath and sieved for 355 to 125 and 125 to 63 micron size fractions. A Frantz magnetic separator was employed to enrich garnet and omphacite in the size fractions. Separates of glaucophane, omphacite, and garnet were handpicked under a binocular microscope. Rims of the garnets (almost 100 % pure), as optically identified by their clear, inclusion-free appearance, were obtained by handpicking the 125 to 63 micron size-fraction. Core/mantle material of garnet (⬃80 % enriched) was picked from the 355 to 125 micron fraction. Clean omphacite and glaucophane grains were picked from the same size fraction. The garnet-leaching method is a streamlined version of that reported in Anczkiewicz and Thirlwall (2003). For the leaching procedures and other wet chemical methods ultra-pure water and quartz-distilled

936

E. Hegner & others—Mineral ages and P-T conditions of late Paleozoic high-pressure Table A.1

Representative electron microprobe data for garnet in eclogite sample KG 23* Mineral Position Rim SiO2 38.06 TiO 2 0.05 Al2O3 21.50 Cr2O3 0.03 Fe2O3 1.08 FeO 26.53 MnO 0.17 MgO 5.27 CaO 7.39 Total 100.08 Normalized to 24 oxygen Si 2.975 Ti 0.003 Al 1.980 Cr 0.002 Fe3+ 0.064 Fe2+ 1.734 Mn 0.011 Mg 0.614 Ca 0.619 Total 8.000 Fe/(Fe+Mg) 0.75 Xalm 0.58 Xprp 0.21 Xgrs 0.21 Xsps --

Rim 38.06 0.05 21.37 b.d. 0.54 28.42 0.43 3.85 7.68 100.39 2.991 0.003 1.980 -0.032 1.868 0.029 0.451 0.646 8.000 0.81 0.62 0.15 0.22 0.01

Garnet Core 37.08 0.45 20.64 b.d. 0.40 24.93 4.26 1.16 10.49 99.40

Core 37.15 0.28 20.48 0.04 0.58 24.86 3.59 1.07 11.13 99.18

Rim 37.18 0.16 20.91 0.01 1.01 28.45 0.40 1.95 9.59 99.66

Rim 36.90 0.15 20.85 0.06 1.22 28.19 0.33 2.10 9.37 99.16

2.978 0.030 1.960 0.001 0.020 1.680 0.290 0.140 0.900 8.000 0.92 0.56 0.05 0.30 0.10

2.992 0.017 1.944 0.003 0.035 1.674 0.245 0.129 0.960 8.000 0.93 0.56 0.04 0.32 0.08

2.974 0.010 1.971 0.001 0.061 1.903 0.027 0.233 0.822 8.000 0.89 0.64 0.08 0.28 0.01

2.965 0.009 1.974 0.004 0.074 1.894 0.022 0.251 0.807 8.000 0.89 0.64 0.08 0.27 0.01

* Data for analytical traverses are available from R.K. b.d. ⫽ below detection. acids were used. The handpicked garnet separates were washed in an ultrasonic bath and ground to a fine powder in a boron carbide mortar. About 200 mg of this garnet powder was leached in concentrated HNO3 at ⬃140 °C for 30 minutes. The sample solution was diluted with water, transferred to a centrifuge tip, and centrifuged. The leachate was discarded. The residue was washed in water using a Vortex vibrator and subsequently centrifuged for a short time; the leachate was again discarded. This step was repeated three times. The sample residue was then transferred to a PFA beaker and leached in ⬃3 ml of analytical-grade concentrated H2SO4 at 140 °C for one day. The H2SO4-leachate was discarded, and the residue was washed five times in ultrapure water and centrifuged as described above. The residue was then transferred in a few drops of ultrapure water to a pre-weighed PTFE-high-pressure dissolution bomb. The sample weight was determined after drying the residue. The garnet and whole-rock powders were decomposed in a mixture of 2 ml concentrated HF and ⬃50 ␮l HClO4 in steel-lined PTFE-bombs at ⬃180 °C over a period of one week. After drying the sample solution, the fluorides/perchlorates were converted in the same bomb at ⬃180 °C for one day. Then followed treatment of the fluoride/perchlorate residue with 6N HCl in PFA vessels until the sample cake was completely dissolved. An aliquot of the garnet sample was spiked for determination of Sm and Nd concentrations. With the Sm and Nd concentrations at hand, the remaining garnet sample was optimally spiked, and a spike-sample isotopic equilibrium was achieved by heating the sample solution overnight in a closed PFA-vessel. After evaporation of the garnet and whole-rock solutions, the samples were dissolved in 2 ml of 2.5 N HCl for chromatography.


937

Table A.2

Representative electron microprobe data for clinopyroxene in eclogite sample KG 23* Mineral SiO2 55.10 TiO 2 0.02 Al2O3 8.70 FeO 8.09 MnO 0.01 MgO 7.22 CaO 11.95 Na2O 7.84 K2O 0.01 Total 98.94 Normalized to 6 oxygen Si 1.987 Ti 0.001 Al (IV) 0.013 Al (VI) 0.357 Fe3+ 0.203 Fe2+ 0.041 Mn -Mg 0.388 Ca 0.462 Na 0.548 K -Total 4.000 WEF 44.83 Jd 35.19 Ae 19.98

55.25 0.01 8.86 7.95 0.03 7.13 11.95 7.93 0.01 99.13

Clinopyroxene 55.69 55.76 0.01 0.02 8.94 9.77 7.31 4.78 0.16 b.d. 7.76 8.64 12.01 13.19 7.91 7.45 b.d. b.d. 99.64 99.61

1.987 -0.013 0.363 0.202 0.037 0.001 0.383 0.461 0.553 -4.000 44.35 35.77 19.89

1.984 -0.016 0.360 0.201 0.017 0.005 0.412 0.458 0.547 -4.000 44.95 35.33 19.72

55.36 0.02 9.53 7.53 0.02 7.07 11.62 8.18 b.d. 99.31

1.978 0.001 0.022 0.386 -0.142 -0.457 0.501 0.513 -4.000 52.77 48.23 0.00

1.982 0.001 0.018 0.383 0.200 0.025 -0.377 0.446 0.567 -4.000 42.78 37.59 19.63

55.24 0.01 9.57 7.72 0.10 6.87 11.62 7.81 0.02 99.06 1.992 -0.008 0.399 0.154 0.078 0.003 0.369 0.449 0.546 0.001 4.000 45.18 39.53 15.29

* Data for analytical traverses are available from R.K. b.d. ⫽ below detection.

The glaucophane and omphacite mineral separates were washed in acetone and leached in 6N HCl for 1 hour at ⬃120 °C. The mineral grains were washed three times in ultrapure water, dried, weighed in a PFA beaker, and spiked with a 149Sm150Nd tracer solution. The grains were then decomposed in a mixture of HF-HClO4 at ⬃100 °C over a period of one week. The sample solution was dried down and the sample cake was dissolved in HCl for column chemistry. Chromatographic element separation followed the procedure described in Hegner and others (1995). The LREE of the samples were separated on quartz columns with a 5 ml resin bed of AG 50W-X12 of 200 to 400 mesh. Nd and Sm were separated on a quartz column filled with 1.7 ml Teflon powder coated with di-ethylhexylphosphoric acid (HDEHP). Nd and Sm were loaded on Re-filaments using diluted phosphoric acid and measured as metals in a double filament configuration. Total procedural blanks of ⬍200 pg for Nd and Sm are not significant for the analyzed concentration levels. Sm and Nd isotope abundance ratios were determined on an upgraded MAT 261 mass spectrometer. Sm isotope ratios were determined in static data collection mode. The 143Nd/144Nd abundances were measured with a dynamic triple collector routine in order to minimize cup effects (for example, Thirlwall, 1991), and monitoring for interfering 144Sm. The 143Nd/144Nd ratios were normalized to 146Nd/144Nd ⫽ 0.7219 and those of Sm to 147Sm/152Sm ⫽ 0.56081. During this study the La Jolla Nd reference material yielded 143Nd/144Nd ⫽ 0.511847 ⫾ 0.000008 (2 SD of population, N ⫽ 10). Analyses of the Nd reference material JNdi yielded ⫽ 0.712109 ⫾ 0.000003 (2 SD, N ⫽ 4). The long-term external precision for 143 Nd/144Nd is ⬃1 ⫻ 10⫺5 (2 SD). Analysis of the 147Sm/144Nd ratio in BCR-1 gave 0.1381 ⫾ 0.0001 (2 SD, N ⫽ 10). All uncertainties reported in this study are given at the 95 percent confidence interval.

938

E. Hegner & others—Mineral ages and P-T conditions of late Paleozoic high-pressure Table A.3A

Representative electron microprobe data for actinolite in eclogite sample KG 23* Mineral Position SiO2 TiO 2 Al2O3 Cr2O3 FeO MnO MgO CaO Na2O K2O Total Si Al (IV) Sum T Al (VI) Ti Fe3+ Cr Mg Fe2+ Mn Sum C Fe2+ Mn Ca Na Sum B Na K Sum A Total

Rim 53.32 0.03 3.07 0.03 16.72 0.18 12.71 11.09 1.07 0.11 95.59 7.513 0.487 8.000 0.053 0.003 0.544 0.004 2.829 1.544 0.023 5.000 --1.774 0.226 2.000 0.084 0.021 0.105 15.105

Core 50.81 0.05 4.46 0.01 16.54 0.16 12.41 9.93 2.10 0.16 96.67 7.457 0.543 8.000 0.228 0.006 0.552 0.001 2.715 1.478 0.02 5.000 --1.56 0.439 2.000 0.159 0.030 0.189 15.189

Core 53.82 0.06 4.12 b.d. 9.63 0.14 16.58 9.79 2.11 0.10 96.33 7.638 0.362 8.000 0.326 0.006 0.447 -3.508 0.695 0.017 5.000 --1.49 0.511 2.000 0.069 0.018 0.087 15.087

Actinolite Rim 53.60 0.07 4.00 b.d. 10.17 0.11 16.46 9.64 2.12 0.11 96.26 7.615 0.385 8.000 0.285 0.007 0.546 -3.486 0.662 0.013 5.000 --1.47 0.533 2.000 0.051 0.020 0.071 15.071

Rim 53.49 0.09 3.86 0.05 11.32 0.16 15.54 9.94 1.92 0.08 96.46 7.649 0.351 8.000 0.299 0.010 0.438 0.006 3.313 0.916 0.019 5.000 --1.52 0.477 2.000 0.055 0.011 0.066 15.066

Core 54.47 0.06 3.72 b.d. 10.85 0.16 16.37 9.88 2.09 0.11 97.70 7.654 0.346 8.000 0.270 0.006 0.498 -3.429 0.777 0.019 5.000 --1.49 0.512 2.000 0.057 0.020 0.077 15.077

Core 54.39 0.02 3.53 0.02 10.03 0.15 16.68 10.04 2.02 0.09 96.96 7.689 0.311 8.000 0.277 0.002 0.417 0.002 3.515 0.769 0.018 5.000 --1.52 0.479 2.000 0.074 0.016 0.091 15.091


40

Ar/39Ar Age Determination

Argon isotope analyses of single phengite grains were performed at Nice University, France. Phengite grains of 0.5 to 1 mm size were handpicked under a binocular microscope, packed in aluminium foil, and irradiated for 30 hours in the nuclear reactor at McMaster University in Hamilton (Canada), along with the MMHb-1 hornblende neutron flux monitor dated at 520.4 ⫾ 1.7 Ma (2␴). The total neutron flux density during irradiation was 9.0 ⫻ 1018 neutrons/cm2. The estimated error on the corresponding 40Ar*/39ArK ratio is 0.2 percent (1␴) in the volume where the samples were set. Isotope analyses were performed by step-heating the mineral grains with a 50 W CO2 Synrad 48-5 continuous laser probe. The laser beam was larger than the analyzed grain to facilitate homogeneous heating. Measurement of isotopic ratios was performed with a VG-3600 mass spectrometer equipped with a Daly detector system. The decay constants used for age calculations are those of Steiger and Ja¨ger (1977). Blanks were measured routinely each first or third run and subtracted from the subsequently measured gas fractions of the samples. The mass-discrimination was monitored by regularly analysing the air pipette volume.


939

Table A.3B

Representative electron microprobe data for glaucophane in eclogite sample KG 23 Mineral Position Rim SiO2 57.10 TiO 2 0.00 Al2O3 10.69 Cr2O3 0.01 FeO 10.03 MnO 0.05 MgO 10.76 CaO 1.32 Na2O 6.84 K2O 0.03 Total 96.81 Normalized to 23 oxygen Si 7.883 Al (IV) 0.117 Sum T 8.000 Al (VI) 1.621 Ti 0.000 Fe3+ 0.269 Cr 0.001 Mg 2.214 Fe2+ 0.889 Mn 0.006 Sum C 5.000 Fe2+ 0.000 Mn 0.000 Ca 0.195 Na 1.805 Sum B 2.000 Na 0.026 K 0.005 Sum A 0.031 Total 15.031

Core 57.51 0.00 10.33 0.00 9.17 0.11 11.75 1.68 6.78 0.02 97.35

Glaucophane Core Core 57.35 56.75 0.07 0.06 10.05 10.70 0.04 0.02 9.51 9.58 0.05 0.04 11.98 11.46 1.99 1.51 6.50 6.53 0.03 0.02 97.56 96.67

Rim 57.11 0.03 10.62 0.04 10.30 0.06 10.83 1.58 6.55 0.06 97.18

Rim 57.44 0.05 10.73 0.01 9.38 0.07 11.15 1.34 6.84 0.03 97.04

7.874 0.126 8.000 1.539 0.000 0.291 0.000 2.398 0.759 0.013 5.000 0.000 0.000 0.246 1.754 2.000 0.046 0.003 0.050 15.066

7.838 0.162 8.000 1.455 0.007 0.378 0.004 2.441 0.709 0.006 5.000 0.000 0.000 0.291 1.709 2.000 0.014 0.005 0.019 15.019

7.868 0.132 8.000 1.591 0.003 0.269 0.004 2.224 0.908 0.000 5.000 0.010 0.007 0.233 1.748 2.000 0.000 0.011 0.011 15.011

7.889 0.111 8.000 1.625 0.005 0.253 0.001 2.283 0.825 0.008 5.000 0.000 0.000 0.197 1.803 2.000 0.019 0.005 0.024 15.024

7.822 0.178 8.000 1.559 0.006 0.345 0.002 2.355 0.732 0.000 5.000 0.027 0.005 0.223 1.745 2.000 0.000 0.004 0.004 15.004

The criteria used to define a plateau age are as follows: (1) the plateau age should be based on 70 percent or more of the total 39Ar released; (2) it should comprise three or more successive step-heating fractions; (3) the integrated age of the plateau should agree with each apparent age of the plateau fractions within its 2␴ error. The uncertainty on the 40Ar/39Ar ratio of the monitor is included in the calculation of the integrated plateau age uncertainties, but the error on the age of the monitor is not included in the age calculations. The plateau and integrated ages and their uncertainties are listed in table 3. The Ar-degassing spectra are presented in figure 4. SHRIMP and LA-ICPMS Zircon Dating

A heavy mineral fraction from metagraywacke sample KG 25 was prepared using conventional magnetic and heavy mineral separation methods. The zircons were handpicked and mounted in epoxy resin together with chips of the Perth Consortium zircon standard CZ3. Cathodoluminescence imaging of the sectioned zircons and isotopic analyses were performed in the Beijing SHRIMP Center using a Hitachi S-3000N

51.17 0.28 25.04 b.d. 2.24 b.d. 4.53 0.13 b.d. 0.32 10.85 94.55

6.792 6.909 0.027 0.028 4.260 3.985 --0.269 0.253 --0.751 0.912 0.008 0.007 --0.134 0.083 1.751 1.870 13.993 14.046

50.52 0.27 26.89 b.d. 2.39 b.d. 3.74 0.16 b.d. 0.51 10.21 94.69

KG 23 49.93 0.19 27.33 0.06 2.62 0.01 3.48 0.20 0.02 0.51 10.19 94.55

49.68 0.24 26.94 0.04 2.40 b.d. 3.52 0.09 0.22 0.45 10.22 93.81

6.799 6.739 6.751 0.023 0.020 0.024 4.229 4.347 4.315 -0.007 0.005 0.254 0.296 0.273 0.006 0.001 -0.774 0.700 0.714 0.009 0.011 0.005 -0.003 0.033 0.157 0.133 0.119 1.784 1.754 1.772 14.035 14.009 14.010

50.51 0.23 26.66 b.d. 2.26 0.06 3.86 0.17 b.d. 0.60 10.39 94.73

Phengite 50.62 51.82 50.73 0.13 0.17 0.22 25.69 25.33 25.74 0.06 0.06 0.04 4.00 3.64 3.78 0.09 0.04 0.08 3.44 3.48 3.44 0.21 0.38 0.31 0.01 b.d. 0.01 0.22 0.22 0.14 10.55 10.63 10.45 95.01 95.76 94.94 Normalized to 22oxygen 6.853 6.945 6.862 0.014 0.017 0.022 4.099 4.002 4.104 0.006 0.007 0.004 0.453 0.408 0.428 0.010 0.004 0.009 0.694 0.694 0.694 0.011 0.020 0.017 0.001 -0.001 0.057 0.056 0.038 1.822 1.817 1.803 14.021 13.970 13.982


49.79 50.74 SiO2 TiO 2 0.26 0.31 27.16 26.62 Al2O3 Cr2O3 0.08 0.05 FeO 2.58 2.38 MnO 0.03 b.d. MgO 3.52 3.92 BaO 0.22 0.19 CaO b.d. b.d. Na2O 0.50 0.56 K2 O 10.36 10.35 Total 94.48 95.13 Normalized to 24 oxygen Si 6.734 6.802 Ti 0.026 0.032 Al 4.329 4.205 Cr 0.008 0.005 Fe2+ 0.292 0.267 Mn 0.003 -Mg 0.710 0.783 Ba 0.012 0.010 Ca --Na 0.131 0.146 K 1.787 1.770 Total 14.030 14.020

Mineral Sample

Table A.4

51.30 0.14 25.61 0.06 3.41 0.07 3.49 0.23 0.03 0.31 10.44 95.1

51.78 0.16 26.10 0.05 3.35 0.09 3.26 0.28 0.03 0.25 10.32 95.68

KG 25 50.45 0.19 28.20 0.11 2.51 0.02 3.40 0.31 0.02 0.50 9.96 95.67

50.40 0.21 26.75 0.06 3.10 b.d. 3.32 0.26 b.d. 0.28 10.30 94.67

49.99 0.18 28.16 0.06 2.53 0.04 3.10 0.28 0.08 0.52 9.94 94.87

6.936 6.909 6.916 6.713 6.805 6.710 0.019 0.014 0.016 0.019 0.021 0.018 4.010 4.065 4.108 4.423 4.256 4.455 0.004 0.007 0.005 0.012 0.006 0.006 0.385 0.384 0.375 0.279 0.349 0.284 0.012 0.007 0.010 0.002 -0.004 0.732 0.701 0.649 0.674 0.667 0.620 0.021 0.012 0.015 0.016 0.014 0.015 0.003 0.005 0.005 0.003 -0.011 0.056 0.082 0.066 0.129 0.073 0.135 1.779 1.793 1.758 1.690 1.774 1.702 13.956 13.979 13.923 13.961 13.966 13.960

51.47 0.18 25.25 0.04 3.41 0.10 3.64 0.40 0.02 0.22 10.35 95.08

Representative electron microprobe data for phengite in eclogite sample KG 23 and metagraywacke sample KG 25*

50.54 0.24 26.48 0.07 3.10 0.02 3.32 0.35 0.01 0.34 10.45 94.91 6.919 6.820 0.023 0.024 4.089 4.212 -0.008 0.359 0.350 0.002 0.002 0.660 0.668 0.015 0.018 0.002 0.002 0.094 0.088 1.796 1.799 13.959 13.990

51.63 0.23 25.89 b.d. 3.20 0.02 3.30 0.28 0.02 0.36 10.51 95.43

940 E. Hegner & others—Mineral ages and P-T conditions of late Paleozoic high-pressure

Table A.5

36.95 0.20 24.93 0.03 11.37 0.39 b.d. b.d. 23.51 0.09 0.01 97.48 5.892 0.024 4.686 0.004 1.364 0.053 --4.018 0.029 0.002 16.072 22.55

37.34 0.14 23.45 0.04 13.26 0.09 b.d. b.d. 23.30 b.d. 0.04 97.65

5.968 0.016 4.416 0.006 1.595 0.012 --3.989 -0.008 16.011 26.53

5.946 0.017 4.847 0.005 1.168 0.025 --4.010 0.019 0.007 16.041 19.42

KG 23 37.56 0.14 25.98 0.04 9.80 0.19 b.d. b.d. 23.64 0.06 0.03 97.45 5.944 0.031 4.804 0.002 1.215 0.022 --3.993 -0.004 16.016 20.19

37.47 0.26 25.69 0.01 10.18 0.17 b.d. b.d. 23.49 0.00 0.02 97.27


Mineral Sample 37.57 SiO2 TiO 2 0.13 23.61 Al2O3 0.05 Cr2O3 12.87 Fe2O3 MnO 0.10 MgO b.d. BaO b.d. CaO 23.46 Na2O 0.04 K2O b.d. Total 97.64 Normalized to 25 oxygen Si 5.986 Ti 0.016 Al 4.433 Cr 0.007 Fe3+ 1.542 Mn 0.013 Mg -Ba -Ca 4.005 Na 0.013 -K Total 16.014 Fe3+/(Fe3++Al)x100 25.81 5.944 0.005 4.863 0.003 1.178 0.010 --4.020 0.003 0.006 16.034 19.50

37.73 0.04 26.19 0.02 9.94 0.07 b.d. b.d 23.82 0.01 0.03 97.52

K G 25 38.36 38.56 38.39 38.80 0.10 0.15 0.28 0.16 27.98 28.24 28.63 28.96 0.10 0.04 0.05 0.04 6.56 6.20 6.00 5.33 0.18 0.16 0.13 0.15 b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. 23.75 23.05 23.37 24.10 b.d. 0.01 0.06 0.04 0.02 b.d. 0.01 0.02 97.03 96.42 96.91 97.60 Normalized to 25 oxygen 6.020 6.060 6.001 6.020 0.020 0.020 0.040 0.020 5.160 5.220 5.280 5.300 0.020 0.003 0.003 0.002 0.780 0.740 0.700 0.620 0.020 0.020 0.020 0.020 --------3.980 3.880 3.920 0.000 -0.002 0.020 0.020 0.001 0.002 -0.001 16.000 15.940 15.980 16.000 13.02 12.30 11.81 10.52

Clinozoisite

6.100 0.020 5.220 0.004 0.700 0.020 -0.002 3.840 0.040 0.001 15.940 11.80

38.77 0.14 28.17 0.06 5.90 0.20 b.d. 0.06 22.77 0.11 0.01 96.20

Representative electron microprobe data for clinozoisite in eclogite sample KG 23 and metagraywacke sample KG 25*

6.000 0.020 5.280 0.003 0.740 0.020 --3.900 --15.960 12.35

38.56 0.18 28.75 0.05 6.35 0.13 b.d. b.d. 23.33 0.00 0.00 97.35

eclogite and provenance of me´lange sediments from Atbashi in the South Tianshan 941

942

E. Hegner & others—Mineral ages and P-T conditions of late Paleozoic high-pressure Table A.6

Representative electron microprobe data for albite in eclogite sample KG 23 and metagraywacke sample KG 25* Mineral Sample KG 23 SiO2 67.54 Al2O3 19.41 FeO 0.19 MgO b.d. BaO 0.05 CaO 0.16 Na2O 11.57 K2O 0.03 Total 98.92 Normalized to 8 oxygen Si 2.990 Al 1.010 Fe3+ 0.010 Mg -Ba -Ca 0.010 Na 0.990 K -Total 5.010 xAn 1.37 xAb 98.38 xOr 0.25

Albite 68.08 19.30 0.05 0.04 b.d. 0.08 11.70 0.03 99.28

67.77 19.31 0.01 0.11 0.09 0.11 11.70 0.06 99.16

KG 25 68.06 19.41 0.09 0.06 0.05 0.24 11.58 0.04 99.53

67.78 19.38 0.03 1.02 0.01 0.18 11.37 0.06 99.84

68.22 19.28 0.10 0.14 0.02 0.03 11.49 0.07 99.35

2.995 1.001 0.002 0.003 -0.004 0.998 0.002 5.003 0.35 99.46 0.18

2.989 1.004 -0.007 0.001 0.005 1.000 0.004 5.011 0.53 99.11 0.36

2.989 1.004 0.003 0.004 0.001 0.011 0.986 0.002 5.001 1.15 98.61 0.24

2.969 1.001 0.001 0.067 -0.009 0.966 0.003 5.015 0.87 98.78 0.34

2.997 0.998 0.004 0.009 -0.002 0.979 0.004 4.993 0.16 99.47 0.38


scanning electron microscope (accelerating voltage 10 kV, beam current 109 mA) and the SHRIMP II high-resolution ion-microprobe (De Laeter and Kennedy, 1998). For analytical details of SHRIMP II such as instrumental conditions, data reduction procedures and error assessment see Kro¨ner and others (2003) and references cited therein. For data collection, six scans through the critical mass range were made. Primary beam intensity was about 4.0 nA, and a Ko¨hler aperture of 100 ␮m diameter was used, giving a slightly elliptical spot size of about 30 ␮m. Sensitivity was about 20 cps/ppm/nA Pb on the standard CZ3. Analyses of samples and standards were alternated to allow assessment of Pb⫹/U⫹ discrimination. The 1␴ error of the 206 Pb/238U ratio during analysis of all standard zircons during this study was 1.3 percent. Common-Pb corrections were applied using the 204Pb-correction method. Because of very low counts on 204Pb it was assumed that common lead is surface-related (Kinny, 1986) and therefore the isotopic composition of Broken Hill lead was used for correction. The analytical data are presented in table A.8, Appendix. Errors of individual analyses are given at the 1␴ level and are based on counting statistics and include the uncertainty of the U/Pb age of the standard (Nelson, 1997). Errors for pooled analyses are reported at the 2␴ or 95 percent confidence interval. The data are graphically presented on the conventional concordia diagram of figure 7A. Additional detrital zircons from phengite paragneiss KG 25 were analyzed by LA-ICPMS at Mainz University using a New Wave 213 nm laser coupled with an Agilent 7500ce quadrupole ICP-MS (Zack and others, 2009, 2010 submitted1). Prior to analysis the sample mount was cleaned by polishing the

1 Zack, T., Stockli, D., Luvizotto, G. L., Barth, M. G., Belousova, E., Wolfe, M. R., and Hinton, R. W., 2010, In-situ U/Pb rutile dating by LA-ICP-MS: correction and prospects for geological application: Submitted to Contributions to Mineralogy and Petrology.


943

Table A.7

Representative electron microprobe data for chlorite in eclogite sample KG 23 and metagraywacke sample KG 25* Mineral Sample KG 23 SiO2 25.92 24.98 25.48 TiO 2 0.05 0.05 b.d. Al2O3 19.32 18.58 19.72 Cr2O3 0.01 0.05 0.01 Fe2O3 0.00 0.43 0.50 FeO 28.68 29.32 27.50 MnO 0.24 0.29 0.20 MgO 12.94 12.12 13.97 CaO 0.08 0.12 0.08 Na2O b.d. 0.06 0.02 K2 O 0.09 0.03 0.02 Total 87.33 86.02 87.50 Normalized to 28 oxygen Si 5.610 5.529 5.465 Ti 0.009 0.008 -Al 4.930 4.846 4.986 Cr 0.002 0.009 0.002 Fe3+ -0.071 0.080 Fe2+ 5.192 5.428 4.932 Mn 0.043 0.053 0.037 Mg 4.176 3.997 4.466 Ca 0.018 0.029 0.018 Na -0.024 0.008 K 0.019 0.006 0.004 Total 20.000 20.000 20.000

Chlorite KG 25 25.83 0.04 19.28 b.d. 0.71 25.41 0.28 15.46 0.01 0.02 0.02 87.07

25.43 0.03 19.26 b.d. 0.61 26.73 0.19 14.35 0.03 0.02 b.d. 86.64

25.53 0.01 18.93 0.01 1.46 25.24 0.41 15.16 b.d. 0.04 0.02 86.81

25.63 0.04 19.19 0.02 0.88 26.16 0.34 14.85 b.d. b.d. 0.02 87.13

5.494 0.004 4.904 -0.100 4.830 0.034 4.619 0.006 0.008 0.001 20.000

5.485 0.001 4.792 0.001 0.236 4.535 0.075 4.855 -0.016 0.005 20.000

5.495 5.511 0.007 0.007 4.850 4.848 0.003 0.001 0.143 0.115 4.690 4.534 0.062 0.051 4.745 4.917 -0.003 -0.009 0.004 0.005 20.000 20.000

25.88 0.01 19.42 0.01 0.90 25.86 0.35 15.19 0.02 0.07 0.01 87.69

25.65 0.04 19.38 0.01 1.23 25.42 0.27 15.36 0.04 0.01 0.03 87.42

5.496 0.001 4.861 0.002 0.144 4.593 0.062 4.808 0.004 0.028 0.002 20.000

5.462 0.006 4.864 0.002 0.197 4.527 0.049 4.875 0.009 0.002 0.006 20.000


surface with ␥-alumina powder to remove any carbon coating, then placed in an ultrasonic bath for 5 minutes with ultra-clean water, and finally dried with an ethanol-soaked paper tissue. Further purification of the zircon surface was achieved by preablating the analysis spots with 5 laser shots with a beam diameter of 40 ␮m. Each analysis consisted of a background measurement of 40 seconds, followed by 30 seconds of sample analysis. The U-Pb data were collected by ablating zircons with a laser beam diameter of 30 ␮m, a beam energy density of ca. 3.5 J/cm2, and a repetition rate of 10 Hz. The aerosol produced during ablation was transported to the ICP-MS in a mixed Ar-He carrier gas at a flow rate of 1.3 l/minute. Isotopes were measured in time-resolved mode. Dwell times for each isotope of individual mass scans are 10 ms for 232 Th and 238U, 30 ms for 201Hg, 204Hg⫹Pb and 206Pb, and 50 ms for 207Pb and 208Pb. Th and U concentrations, 206Pb/204Pb ratios, as well as 207Pb/235U, 206Pb/238U and 208Pb/232Th ages were calculated off-line from time-resolved raw counts. The zircon standard PL (Plesˇovice; Sla´ma and others, 2008) was used as a primary standard to correct for laser-induced as well as ICP-induced mass fractionation by integrating the same time segments for each sample and standard zircon (Jackson and others, 2004). The accuracy of 207 Pb/235U, 206Pb/238U and 208Pb/232Th ages is currently ⬃2.0 percent (2␴), based on analysis of the zircon standards GJ-1, which was also used to calculate U and Th concentrations (Jackson and others, 2004), 91500 (Wiedenbeck and others, 1995), and Mud Tank (Black and Gulson, 1978). The analytical data are presented in table A.9, Appendix and plotted in figure 7B. Precise dating of young zircons by ion-microprobe and ICPMS is best achieved by using 206Pb/238Uages (see Black and Jagodzinski, 2003, for explanation) because the proportions of these two isotopes have changed by easily measurable amounts over most of that time, whereas the 207Pb/206Pb isotopic system is better suited to the dating of zircons older than about 1000 Ma, because of the short half-life of 235 U producing 207Pb (Black and others, 2003). We follow this practice in quoting ages in this paper.

Spot U Th Th/U # [ppm] [ppm] 1 115 338 2.94 2 958 371 0.40 592 325 0.55 3 4 211 102 0.48 5 160 61 0.38 6 356 298 0.84 46 0.12 7 388 8 182 15 0.08

5028 20913 13885 5087 3051 6780 11533 4650

Pb/204Pb

206

Pb/206Pb ± 1σ 0.7888 ± 47 0.1125 ± 6 0.1970 ± 15 0.1406 ± 33 0.1211 ± 58 0.2530 ± 30 0.0450 ± 14 0.0300 ± 52

208

Pb/206Pb ± 1σ 0.1667 ± 15 0.1669 ± 4 0.0940 ± 6 0.0733 ± 14 0.0664 ± 24 0.0740 ± 11 0.0829 ± 8 0.0708 ± 23

207

Pb/238U ± 1σ 0.4814 ± 68 0.4776 ± 63 0.2625 ± 35 0.1699 ± 23 0.1374 ± 19 0.1731 ± 23 0.2169 ± 29 0.1608 ± 22

206

Pb/235U ± 1σ 11.06 ± 20 10.99 ± 15 3.404 ± 53 1.717 ± 43 1.257 ± 51 1.766 ± 37 2.480 ± 44 1.568 ± 58

207

Pb/238U age ± 1σ 2533 ± 30 2517 ± 27 1503 ± 18 1012 ± 13 830 ± 11 1029 ± 13 1265 ± 15 961 ± 12

206

Pb/235U age ± 1σ 2528 ± 16 2522 ± 13 1505 ± 12 1015 ± 16 826 ± 23 1033 ± 14 1266 ± 13 958 ± 23

207

Pb/206Pb % of conage ± 1σ cordance 2525 ± 15 100 2527 ± 4 99.6 1509 ± 13 99.6 1022 ± 38 99.0 818 ± 77 101 1042 ± 30 98.8 1268 ± 20 99.8 950 ± 67 101

207

SHRIMP II analytical data for spot analyses of detrital zircons from metagraywacke sample KG 25, Atbashi me´lange, Kyrgystan

Table A.8

944 E. Hegner & others—Mineral ages and P-T conditions of late Paleozoic high-pressure

>1900 >2500 >20800 >18500 >16400 4500 >13700 >8300 3900 5064 8260 >9800 >9600 1040 7600 2010 >4962 2800 >7200 >33300 3800 13700 >3450 >1360

Pb/204Pb

206

Pb/232Th ±1σ 0.0223 ± 6 0.0360 ± 22 0.0736 ± 21 0.0912 ± 31 0.1081 ± 33 0.0571 ± 16 0.1425 ± 91 0.0501 ± 13 0.0220 ± 7 0.1377 ± 49 0.114 ± 12 0.0525 ± 14 0.0691 ± 18 0.0230 ± 7 0.0843 ± 23 0.1575 ± 66 0.0476 ± 31 0.0481 ± 12 0.0270 ± 7 0.0698 ± 18 0.0242 ± 8 0.1259 ± 46 0.0405 ± 12 0.0220 ± 6

208

Pb/235U ±1σ 0.558 ± 17 1.120 ± 35 3.211 ± 67 4.90 ± 14 9.77 ± 22 2.084 ± 43 13.94 ± 29 1.621 ± 38 0.569 ± 13 8.73 ± 17 10.11 ± 38 2.155 ± 54 2.947 ± 58 0.638 ± 18 3.514 ± 84 10.00 ± 28 1.728 ± 47 1.575 ± 37 0.714 ± 16 3.978 ± 75 0.688 ± 18 10.16 ± 25 1.173 ± 29 0.552 ± 15

207

Pb/238U ±1σ 0.0694 ± 14 0.1217 ± 29 0.2487 ± 42 0.3155 ± 83 0.3937 ± 62 0.1869 ± 27 0.5219 ± 87 0.1644 ± 33 0.0714 ± 12 0.3943 ± 60 0.426 ± 14 0.1756 ± 27 0.2301 ± 33 0.0689 ± 12 0.2593 ± 52 0.4503 ± 92 0.1650 ± 36 0.1606 ± 30 0.0868 ± 13 0.2461 ± 33 0.0759 ± 15 0.4476 ± 94 0.1150 ± 20 0.0685 ± 11

206

Pb/206Pb ±1σ 0.0583 ± 16 0.0667 ± 19 0.0936 ± 15 0.1127 ± 30 0.1800 ± 33 0.0809 ± 12 0.1937 ± 19 0.0715 ± 14 0.0577 ± 9 0.1606 ± 17 0.1720 ± 64 0.0890 ± 18 0.0929 ± 11 0.0671 ± 16 0.0983 ± 13 0.1612 ± 44 0.0760 ± 15 0.0711 ± 11 0.0596 ± 9 0.1172 ± 13 0.0657 ± 13 0.1646 ± 36 0.0740 ± 15 0.0584 ± 14

207

Pb/232Th age ± 1σ 445 ± 12 715 ± 43 1435 ± 39 1764 ± 58 2075 ± 61 1122 ± 30 2693 ± 162 988 ± 25 439 ± 13 2607 ± 87 2189 ± 227 1035 ± 27 1350 ± 34 460 ± 14 1635 ± 44 2957 ± 115 940 ± 59 950 ± 24 538 ± 14 1364 ± 34 484 ± 16 2397 ± 82 802 ± 24 440 ± 12

208

* Discordant samples with irregular signal behavior during ablation; excluded from further discussion.

Spot U Th Th/U # [ppm] [ppm] 9 0.82 181 148 10 148 57 0.39 11 526 381 0.72 12 439 155 0.35 13* 275 159 0.58 14 157 60 0.38 15 11 0.07 161 16 0.58 291 168 17 530 390 0.74 18* 512 66 0.13 19* 608 73 0.12 20* 384 112 0.29 21 278 163 0.59 223 201 22* 0.90 270 123 23 0.46 24 338 87 0.26 25 186 164 0.88 26 206 170 0.83 27 502 198 0.39 28* 867 614 0.71 29* 324 189 0.58 296 433 30* 1.47 31 189 100 0.53 32 132 96 0.72

Pb/238U age ± 1σ 433 ± 9 741 ± 17 1432 ± 22 1768 ± 41 2140 ± 29 1105 ± 15 2707 ± 37 981 ± 18 445 ± 7 2143 ± 28 2289 ± 62 1043 ± 15 1335 ± 17 429 ± 7 1486 ± 27 2396 ± 41 984 ± 20 960 ± 17 537 ± 8 1418 ± 17 472 ± 9 2384 ± 42 701 ± 12 427 ± 7

206

LA-ICPMS data for spot analyses of detrital zircons from metagraywacke sample KG 25

Table A.9

Pb/235U age ± 1σ 450 ± 11 763 ± 17 1460 ± 16 1803 ± 25 2414 ± 21 1144 ± 14 2745 ± 20 978 ± 15 457 ± 9 2311 ± 18 2445 ± 36 1167 ± 17 1394 ± 15 501 ± 11 1530 ± 19 2435 ± 26 1019 ± 17 960 ± 15 547 ± 9 1630 ± 15 531 ± 11 2449 ± 23 788 ± 14 446 ± 10

207

Pb/206Pb age ± 1σ 542 ± 61 830 ± 61 1501 ± 30 1844 ± 49 2653 ± 31 1218 ± 30 2774 ± 16 972 ± 41 519 ± 35 2462 ± 18 2577 ± 64 1405 ± 39 1486 ± 23 841 ± 50 1591 ± 25 2468 ± 47 1094 ± 41 960 ± 32 589 ± 32 1914 ± 20 796 ± 43 2503 ± 38 1041 ± 41 546 ± 52

207

eclogite and provenance of me´lange sediments from Atbashi in the South Tianshan 945

946

E. Hegner & others—Mineral ages and P-T conditions of late Paleozoic high-pressure References

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