Zircon geochronology of the Koraput alkaline complex: Insights from combined geochemical and U–Pb–Hf isotope analyses, and implications for the timing of alkaline magmatism in the Eastern Ghats Belt, India

GR-01421; No of Pages 16 Gondwana Research xxx (2015) xxx–xxx

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Zircon geochronology of the Koraput alkaline complex: Insights from combined geochemical and U–Pb–Hf isotope analyses, and implications for the timing of alkaline magmatism in the Eastern Ghats Belt, India K. Hippe a,b,⁎, A. Möller c, A. von Quadt a, I. Peytcheva a,d, K. Hammerschmidt b a

Institute of Geochemistry and Petrology, ETH Zürich, Clausiusstrasse 25, 8092 Zürich, Switzerland Institute of Geological Sciences, Freie Universität Berlin, Malteserstrasse 74-100, 12249 Berlin, Germany c Department of Geology, The University of Kansas, 1475 Jayhawk Blvd., Lawrence, KS 66045, USA d Geological Institute, Bulgarian Academy of Sciences, Acad. G. Bonchev st. bl. 24, 1113 Sofia, Bulgaria b

a r t i c l e

i n f o

Article history: Received 25 October 2014 Received in revised form 11 February 2015 Accepted 22 February 2015 Available online xxxx Handling Editor: A. Kröner Keywords: Metamorphic zircon U–Pb geochronology Lu–Hf Eastern Ghats Belt Alkaline magmatism

a b s t r a c t Zircon formation and modification during magmatic crystallization and high-grade metamorphism are explored using TIMS and LA-ICP-MS U–Pb geochronology, Lu–Hf isotope chemistry, trace element analysis and textural clues on zircons from the Koraput alkaline intrusion, Eastern Ghats Belt (EGB), India. The zircon host-rock is a granulite-facies nepheline syenite gneiss with an exceptionally low Zr concentration, prohibiting early magmatic Zr saturation. With zircon formation occurring at a late stage of advanced magmatic cooling, significant amounts of Zr were incorporated into biotite, nearly the only other Zr-bearing phase in the nepheline syenite gneisses. Investigated zircons experienced a multi-stage history of magmatic and metamorphic zircon growth with repeated solid-state recrystallization and partial dissolution–precipitation. These processes are recorded by complex patterns of internal zircon structures and a wide range of apparently concordant U–Pb ages between 869 ± 7 Ma and 690 ± 1 Ma. The oldest ages are interpreted to represent the timing of the emplacement of the Koraput alkaline complex, which significantly postdates the intrusion ages of most of the alkaline intrusion in the western EGB. However, Hf model ages of TDM = 1.5 to 1.0 Ga suggest an earlier separation of the nepheline syenite magma from its depleted mantle source, overlapping with the widespread Mesoproterozoic, rift-related alkaline magmatism in the EGB. Zircons yielding ages younger than 860 Ma have most probably experienced partial resetting of their U–Pb ages during repeated and variable recrystallization events. Consistent youngest LA-ICPMS and CA-TIMS U–Pb ages of 700–690 Ma reflect a final pulse of high-grade metamorphism in the Koraput area and underline the recurrence of considerable orogenic activity in the western EGB during the Neoproterozoic. Within the nepheline syenite gneisses this final high-grade metamorphic event caused biotite breakdown, releasing sufficient Zr for local saturation and new subsolidus zircon growth along the biotite grain boundaries. © 2015 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved.

1. Introduction Defining the conditions of zircon crystallization in magmatic and metamorphic environments is one of the main challenges in zircon U– Pb geochronology in order to understand which geological processes are being dated. Particularly in rocks with a high-grade metamorphic overprint, the obtained age information often reflects a complex history of polyphase zircon growth, dissolution and/or recrystallization (e.g., Harley et al., 2007). With the development of high-precision mass spectrometry and small-diameter spot measurements the analysis of single grains or specific domains within single crystals has become ⁎ Corresponding author at: Laboratory of Ion Beam Physics, ETH Zürich, Otto-Stern-Weg 5, 8093 Zürich, Switzerland. Tel.: +41 44 633 06 24; fax: +41 44 633 10 67. E-mail address: [email protected] (K. Hippe).

possible. These techniques allow us to date multiple stages of crystal growth or modification and, thus, to shed light on detailed regional tectonometamorphic histories. In this context, any geologic interpretation of zircon U–Pb data requires a thorough evaluation of zircon textures and morphology as obtained by high-resolution imagery, e.g., cathodoluminescence (CL), electron backscattering (BSE) or secondary electron microscopy (cf. Corfu et al., 2003). Additionally, crucial information about zircon growth conditions and the source of the magmatic host rock can be drawn from trace element distribution as well as the Hf isotope composition (e.g., Watson and Harrison, 1983; Belousova et al., 2002, 2006; Watson et al., 2006; Scherer et al., 2007; Kemp et al., 2009). In this study, we discuss the processes of zircon formation and modification in granulite facies, alkaline rocks from the Indian Eastern Ghats Belt (EGB). This work combines U–Pb zircon dating with analyses of

http://dx.doi.org/10.1016/j.gr.2015.02.021 1342-937X/© 2015 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved.

Please cite this article as: Hippe, K., et al., Zircon geochronology of the Koraput alkaline complex: Insights from combined geochemical and U–Pb– Hf isotope analyses, and implications..., Gondwana Research (2015), http://dx.doi.org/10.1016/j.gr.2015.02.021

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K. Hippe et al. / Gondwana Research xxx (2015) xxx–xxx

zircon chemical composition, especially rare-earth element (REE) abundances and Lu–Hf isotope composition, correlated to zircon textures and zoning patterns. Zircon REE patterns have been suggested to change with different pressure–temperature conditions and therefore might identify specific stages of mineral growth or metamorphism (e.g., Rubatto, 2002; Rubatto and Hermann, 2003). Similarly, for Hf, Y, P, and rarely for U and Th, trace element partitioning related to metamorphic zircon overgrowth and recrystallization has been documented (e.g., Williams et al., 1996; Rubatto et al., 2001; Hoskin and Schaltegger, 2003; Möller et al., 2003). The Lu–Hf system in zircon can be used as a geochemical tracer of magma provenance. Because of the strong fractionation of Lu/Hf in zircon, changes in the initially incorporated 176 Hf/177Hf ratio due to decay of 176Lu are negligible (Hoskin and Black, 2000; Kinny and Maas, 2003; Gerdes and Zeh, 2009). In most instances, Hf isotopes in zircon show closed-system behavior and, thus, effectively preserve the initial 176Hf/177Hf ratio of the igneous host rock even in highly metamorphic environments (Patchett, 1983; Amelin et al., 2000; Hoskin and Black, 2000). Metamorphic zircon overgrowth, however, can exhibit a variation in the 176Hf/177Hf and 176 Lu/177Hf ratio caused, e.g., by co-precipitation with other metamorphic phases and/or alteration by metamorphic fluids (Kinny et al., 1991; Gerdes and Zeh, 2009; Zeh et al., 2010). Timing and conditions of zircon formation will be further evaluated from the whole-rock chemical composition and the use of Ti-in-zircon thermometry, which allow us to estimate Zr saturation in the nepheline syenite host rock. In this regard, the petrographic context of the zircon crystals as documented in thin sections provides crucial constraints on potential Zr sources for metamorphic zircon growth. Finally, we will discuss the integrated geochronological and geochemical data set in terms of the regional geologic evolution and re-evaluate the timing of alkaline magmatism and the duration of orogenic activity (with high-grade metamorphism) along the western EGB.

2. Geological setting 2.1. Regional geology and previous dating The alkaline complex of Koraput is part of the Precambrian Eastern Ghats Belt situated along the east coast of the Indian peninsular (Fig. 1a). Bounded by Archaean cratons along its western margins, the EGB comprises multiply deformed and highly metamorphosed rocks of igneous and supracrustal origin (Mezger and Cosca, 1999). A division of the EGB into geologically coherent entities has been debated by several authors and was summarized and revised by Dobmeier and Raith (2003). Accordingly, the Koraput alkaline complex is located within the Meso- to Neoproterozoic Eastern Ghats Province that is characterized by strongly deformed metasedimentary rocks, granulites and anorthosites, and multi-intrusive granite–charnockite complexes (Dobmeier and Raith, 2003). The western border of the Eastern Ghats province is marked by the prominent Sileru shear zone that forms the contact to the Archean Jeypore Province only few kilometers west of the Koraput alkaline complex (Fig. 1a). The alkaline complex at Koraput is one of several alkaline plutons that are located within the western part of the EGB (Fig. 1a). Most of these occur along shear zones and are associated with the contact between the EGB and the cratonic basement (Ratnakar and Leelanandam, 1989; Leelanandam, 1998). The earliest phase of alkaline magmatism in the Eastern Ghats Province is documented by U–Pb zircon ages of 1500 ± 3/4 Ma (Aftalion et al., 2000) and 1480 ± 17 Ma (Upadhyay et al., 2006a) from the intrusive complex of Khariar (Fig. 1a). Possibly coeval crustal magmatism has been dated at 1464 ± 63 and 1455 ± 80 Ma (Sm–Nd WR; Shaw et al., 1997), followed by a second magmatic pulse dated at 1176 ± 201 Ma (Pb–Pb WR; Paul et al., 1990) and 1159 ± 59/ 30 Ma (U–Pb zircon; Aftalion et al., 1988). Although the emplacement ages of the Khariar nepheline syenites have been challenged by Biswal

Fig. 1. (a) Simplified geological map of the Eastern Ghats Belt (modified after Dobmeier and Raith, 2003) and the location of major alkaline bodies (gray circles); Ra = Rairakhol, Kh = Khariar, Ku = Kunavaram, El = Elchuru, Pu = Purimetla, Up = Uppalapadu (Leelanandam, 1998; Upadhyay, 2008); White rectangle = Chilka Lake domain. (b) Geological sketch map of the Koraput alkaline complex and its country rocks. The star indicates the sample location for KO1.

Please cite this article as: Hippe, K., et al., Zircon geochronology of the Koraput alkaline complex: Insights from combined geochemical and U–Pb– Hf isotope analyses, and implications..., Gondwana Research (2015), http://dx.doi.org/10.1016/j.gr.2015.02.021

K. Hippe et al. / Gondwana Research xxx (2015) xxx–xxx

et al. (2007) who proposed a much younger emplacement age around 511–552 Ma (U–Pb zircon), most EGB alkaline complexes dated so far have provided emplacement ages between 1500 and 1265 Ma (Sarkar and Paul, 1998 and references therein; Aftalion et al., 2000; Upadhyay et al., 2006a,b). Thus, alkaline magmatism has been linked to rift-related continental breakup during the Mesoproterozoic, which was followed by Mesoproterozoic–early Neoproterozoic continent collision with highgrade metamorphism during the Rodinia assembly (Upadhyay, 2008). Pervasive regional deformation and granulite facies metamorphism in the Eastern Ghats Province have been documented at 1100–950 Ma with peak conditions of 750–800 °C and 7–8 kbar (summarized in Dasgupta and Sengupta, 2003). Anorthosite intrusive activity in the northern EGB was dated to two phases, at 933 ± 32 Ma and 792 ± 2 Ma (U–Pb zircon; Krause et al., 2001), contrasting with previous estimates of older, Mesoproterozoic, magmatism (Rb–Sr WR; Sarkar et al., 1981). Prolonged Neoproterozoic orogenic activity with granulite facies metamorphism has been shown for the northern Eastern Ghats Province (Krause et al., 2001; Dobmeier and Simmat, 2002) but also suggested for the Koraput area (Nanda et al., 2008). In contrast to the other alkaline complexes in the EGB, age constraints determined for the Koraput alkaline pluton suggest a significantly later emplacement during the mid-Neoproterozoic, possibly coeval with the anorthosite intrusions in the northern EGB. Rb–Sr whole rock dating of the Koraput nepheline syenite gneiss yielded an age of 856 ± 18 Ma (Sarkar et al., 1989). However, these data have to be judged cautiously with regard to the strong metamorphic overprint of the complex (see below) and a possible disturbance of the Rb–Sr system. Discordant U–Pb data for sphene and 40Ar–39Ar amphibole ages of about 670–570 Ma from the Koraput region were interpreted to reflect a Pan-African amphibolite-facies thermal imprint on the area (Mezger and Cosca, 1999). 2.2. Petrography of the Koraput alkaline complex The Koraput alkaline complex forms a lenticular NE–SW-trending alkaline body enclosed by metasedimentary country rocks consisting mainly of khondalites, with discontinuous layers of quartzofeldspathic and calc-silicate gneisses. The plutonic part comprises a core of gabbroic rocks (Fig. 1b) occurring either as coarse grained, mostly garnet- or olivine-bearing plagioclase–amphibole metagabbros with a distinct magmatic texture, or as fine grained rocks showing variable stages of recrystallization and deformation. This deformation is expressed as a locally well-developed penetrative schistosity and/or as an anastomosing shear zone network. Coarse grained, broadly foliated nepheline syenite gneisses flank the eastern, southeastern and northernmost margins of the complex. Further east, two bands of perthite syenite gneisses border the nepheline syenite gneiss whereas the northwestern margin of the pluton is defined by a continuous band of quartz-bearing monzonite gneiss (Fig. 1b). An extensive network of ultramafic–mafic, syenitic and pegmatitic–granitic dykes truncates both the plutonic complex and the adjacent country rocks. Field observations and thin section examination show that the intrusion was followed by significant hydration of the alkaline rocks under amphibolite facies conditions. A close spatial and temporal relationship between rock hydration and the development of a predominantly NE– SW oriented foliation is particularly well-established in the metagabbroic rocks. The intensity of the foliation varies within the different rock types and is least developed in the nepheline syenite gneisses which mostly preserve their magmatic texture. Locally, rock foliation is cut by roughly NW–SE trending shear zones that represent the latest generation of tectonic overprint. Growth of coronal garnet and orthopyroxene in the metagabbros overprinting foliation and shear zones points at a subsequent syn- to postkinematic dehydration of the alkaline complex and the country rocks at granulite facies conditions. Within the nepheline syenite gneiss, rock dehydration is evidenced by biotite breakdown and the formation of alkali feldspar. A detailed investigation of the stages of

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tectonometamorphic overprint of the Koraput alkaline complex and its country rocks is given in Gupta et al. (2005) and Nanda et al. (2008). According to these authors, the latest amphibolite- to granulite-facies thermal event reached peak conditions of about 600–700 °C at 6–8 kbar (Nanda et al., 2008). 3. Analytical methods Zircons were extracted from a block of nepheline syenite gneiss (KO1) sampled at the location indicated in Fig. 1b. Zircons were concentrated from the b250 μm fraction of crushed and sieved rocks using a Wilfley table, followed by heavy liquids and magnetic separation. Zircon crystals were then handpicked under a binocular microscope to select clear grains free of cracks or inclusions. Zircons were studied by conventional optical microscopy, raster electron microscopy (REM) and cathodoluminescence (CL) imaging. CL images (SEM-CL) were obtained for 21 zircon crystals, mounted in epoxy resin and polished, on a JEOL JXA 8900 RL electron microprobe at the Geochemisches Institut, Göttingen with operating conditions of 15 kV and 12 nA. Color CL images were obtained at the Institut für Geowissenschaften, Universität Potsdam, with a CITL CCL 8200 MkJV cold cathode chamber and control unit run at 370 μA and 10 to 15 kV. The unit is mounted on a Leica DRX LM optical microscope equipped with long working distance micro objectives corrected for the 2 mm thick vacuum chamber window. Images were captured with a Leica DSC DFX fluorescence digital camera (1.3 megapixel) with exposure times of 2–5 s. 3.1. U–Pb geochronology and trace element chemistry U–Pb analysis by isotope dilution thermal ionization mass spectrometry (ID-TIMS) was carried out on air-abraded multi-grain fractions (15 grains) at the Freie Universität Berlin. Dissolution and chemical extraction of U and Pb were performed using an HCl chemistry following Krogh (1973). A mixed 236U–208Pb spike solution was used for each analysis. Pb and U were measured on a Finnigan TRITON mass spectrometer on faraday cups in a static, multi-collector mode supported by a second electron multiplier for measuring mass 204Pb. The measured Pb isotope ratios were corrected for mass discrimination based on NBS 981 analyses with an average factor of 0.07%/amu (n = 6). Mass discrimination for U isotope ratios was 0.58%/amu (n = 6). Corrections for the total procedural Pb blank were on the order of 80–160 pg Pb (n = 4). Corrections for common Pb were based on model of Stacey and Kramers (1975). In order to minimize the effects of Pb loss, additional chemical abrasion TIMS (CA-TIMS) analyses were performed at ETH Zürich. These involved high-temperature annealing followed by a HF and HCl leaching step (Mattinson, 2005). The latter has been shown to be most effective in removing strongly radiation damaged zircon domains that underwent Pb loss during post-crystallization fluid processes. Seven zircon grains from sample KO1 were annealed at 900 °C for approximately 48 h, partially dissolved with concentrated HF at 180 °C for 12–13 h and thoroughly cleaned. Single zircons selected for ID-TIMS measurements were then spiked with the 205Pb–235U spike of ETH Zürich and Pb and U were separated by anion exchange chromatography in 50 μl microcolumns. Isotope analyses were performed on a TritonPlus thermal ionization mass spectrometer (TIMS) equipped with a digital ion counting system of a MasCom multiplier. Lead as well as U (as UO2) isotope ratios were measured sequentially on the electron multiplier. The linearity of the MasCom multiplier was calibrated using the SRM982 and U500 standard solutions. The mass fractionation of Pb was 1.1‰/amu; the total procedural Pb blank was estimated at 1.0 ± 0.25 pg. Common lead in excess of this blank was corrected using the model of Stacey and Kramers (1975). For all TIMS analyses (multi- and single-grain), the uncertainty of the concentration of U and Pb in the spike solution was taken into account and propagated to each individual analysis. The PbMacDAT

Please cite this article as: Hippe, K., et al., Zircon geochronology of the Koraput alkaline complex: Insights from combined geochemical and U–Pb– Hf isotope analyses, and implications..., Gondwana Research (2015), http://dx.doi.org/10.1016/j.gr.2015.02.021

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program was used for age calculation and error propagation (Schmitz and Schoene, 2007). Calculation of concordant ages was done with the Isoplot/Ex v.3 program of Ludwig (2001). Laser ablation inductively coupled plasma mass spectrometry (LAICP-MS) U, Pb and Th isotope measurements and trace element analyses were performed at ETH Zurich using an ELAN 6100 DRC ICP-MS coupled to an ArF 193 nm excimer laser system. Analytical conditions involved a laser pulse frequency of 10 Hz, energy densities of 20 to 25 J/cm2 and pit size of 40 μm. During LA-ICP-MS analyses several isotopes were monitored to detect contaminations by inclusions and/or zoning. After the elimination of outliers, integration windows were then set to avoid possibly contaminated parts of the signal. Normalization and age calculation were performed using the spreadsheet-based data reduction software Glitter (van Achterberg et al., 2001). Instrumental mass discrimination was corrected by using the GJ-1 zircon standard (Jackson et al., 2004). For zircon composition, element abundances were normalized to stoichiometric SiO2 in zircon. Spots for zircon composition and age determination are not overlaying at the same position. However, care was taken to drill the same domain for both spots, if possible. Please note, that the Th/U ratio given in Table 2 was obtained during the U–Pb– Th isotope analyses and calculated from the background corrected measured Th and U counts. LA-ICP-MS analyses of the Zr content in biotite was performed on thin sections under identical analytical conditions described above except for larger pit sizes of up to 100 μm. LA-ICP-MS data for zircon and biotite composition are considered to be accurate within 5% of the measured values. 3.2. Lu–Hf analyses The Nu Plasma MC-ICP-MS was used to measure Lu, Yb and Hf isotope signals. The lenses were set to detect simultaneously the required isotopes (171Yb, 173Yb, 174Hf, 175Lu, 176(Hf + Yb + Lu), 177Hf, 178Hf, 179Hf). In order to monitor the time profile of the measured ratio, a time resolved analytical (TRA) procedure was employed, collecting signal intensities for each isotope within a 60 s integration time. For the Hf LA-ICP-MSMC protocol we used 30 s of baseline measurements prior to a 60 s ablation of the zircon (laser frequency, 5 Hz, spot size = 60 μm). Measurements of the JMC 475 Hf standard solution yielded only minor isobaric interferences by Lu and Yb isotopes on mass 176, making a correction for those negligible for the standard solution. The 176Hf/177Hf ratio obtained over a two month period was 0.282158 ± 12 (2 S.E.). Data accuracy was checked using zircon standard Mud Tank and Monastery, which was previously analyzed by laser ablation and solution MC-ICP-MS techniques. The external reproducibility for the standard zircons Monastery and Mud Tank over several months was 0.282718 ± 16 and 0.282506 ± 19. For zircon samples, isobaric interferences by 176Lu and 176Yb on the 176Hf signal were monitored by the intensity of the 175Lu and 172Yb and/or 171Yb signals and corrected for accordingly. Lutetium and Hf isotope measurements were performed on 15 zircons. From these data, the initial 176Hf/177Hf ratios (176Hf/177Hft) were calculated using the 206Pb/238U ages obtained from the same zircon, preferentially from the same zircon domain. The calculation of the εHft is based on the isotopic composition of the chondritic uniform reservoir (CHUR) as recommended by Bouvier et al. (2008) with a 176Lu/177Hf of 0.0336 ± 1 and a 176Hf/177Hf of 0.282785 ± 11, and the decay constant of λ = 1.865 ± 0.015 × 10−11 (Scherer et al., 2001). Depleted mantle Hf model ages (TDM) were calculated using a 176Lu/177Hf of 0.0384, a 176 Hf/177Hf of 0.28325 for the depleted mantle (Chauvel and BlichertToft, 2001), and a crustal 176Lu/177Hf of 0.015 (Griffin et al., 2002). 4. Results 4.1. Zircon morphology and textures Zircons isolated for geochronological analysis were 50 μm–1 mm long, colorless to light brown or pinkish, with smooth, uncorroded and

crack-free crystal faces. Subhedral to anhedral grains are most common while typical magmatic euhedral bipyramidal zircons are rare. Few zircons show simple evolved prism shapes but most crystals have either rounded or multi-facetted terminations. Rounded grains are commonly composed of multiple distinct crystal faces resembling the so-called ‘soccerball zircons’ (Schaltegger et al., 1999). Some zircons have striated surfaces and protrusions on the grain edges. Looking at the entire population of zircons separated from the nepheline syenite gneiss, the majority of crystals is anhedral. Many zircons that were not picked for U–Pb dating are intergrown with each other or suggest interstitial zircon growth. The common observation of biotite–zircon intergrowth in thin sections supports this idea and will be discussed in detail below. SEM-CL imaging of mounted zircon grains reveals blurred and patchy internal structures (Fig. 2a). Most zircons are dominated by featureless domains (Fig. 2a: ru2, lp3, pl1) and planar and/or lobate growth banding (Fig. 2a: kp4, lp1, lp5, pl3) which might correspond to the striated surfaces. In some grains, planar structures define a poorly evolved sector zoning (Fig. 2a: ru4) and a zigzag shaped ‘fir-tree zoning’ (Fig. 2a: kp1) with an eccentric position of the crystal growth center. Typical oscillatory growth zonation in prismatic shape is mainly absent. A few grains exhibit low-CL domains surrounding a high-CL central part, which could be interpreted as a relict core with significant overgrowth (Fig. 2a: kp3, pl6, ru1). This was also found with optical CL-imaging (Fig. 2b: 3). Altogether, internal structures of nearly all investigated zircons provide evidence for multiple episodes of crystal growth and/or modification. Zircon recrystallization is indicated from homogeneous crystal domains and distinct reaction fronts (Fig. 2a: pl1, pl2, pl4, kp4; Hoskin and Black, 2000; Corfu et al., 2003 and references therein). Optical CL-imaging reveals two distinctive zircon types. One type is represented by subhedral, rounded crystals of low CL intensity and blurred internal structures that can be surrounded by a thin rim of yellowish high CL intensity (Fig. 2b: 1, 2). The second type is found as mostly smaller, irregularly shaped crystals of typically on one-sided higher (yellowish) CL intensity, but equally undefined internal structure (Fig. 2b: 4, 5). 4.2. Pb geochronology and the chemical composition of zircon U–Pb data obtained by ID-TIMS multi-grain analyses of 10 zircon fractions and by CA-TIMS analyses of 7 single grains are listed in Table 1 and illustrated in Fig. 3a. Results of 43 U–Pb–Th analyses by LA-ICP-MS measurements on 20 zircon grains are summarized in Table 2 and Figs. 2a and 3b. Single and mean ages are given with 2σ uncertainty. Three CATIMS U–Pb analyses give concordant ages of 754 ± 1 Ma, 699 ± 1 Ma, and 690 ± 1 Ma (Fig. 3a), whereas the other results are discordant with apparent 206Pb/238U ages between 701 ± 11 Ma and 756 ± 1 Ma. The multi-grain ID-TIMS data are all discordant. A discordia based on the IDTIMS dates and the four oldest CA-TIMS dates yields an upper intercept age of 762 ± 20 Ma, which agrees with the oldest concordant CA-TIMS age of 754 ± 1 Ma. The lower intercept does not provide a geological meaningful age. The three youngest concordant or nearly concordant CA-TIMS ages are interpreted to record separate events and were therefore not included into the calculation of the discordia. Processes that could be responsible for the observed discordance are non-zero age Pb loss from metamict domains, new overgrowths leading to mixed ages, or partial recrystallization with or without complete Pb expulsion in the recrystallized domains (e.g., Mezger and Krogstad, 1997; McFarlane et al., 2005; Kooijman et al., 2011). With regard to the internal zircon textures described above, we consider mixing of different age components and zircon recrystallization as the most likely causes of discordance. Results from LA-ICP-MS analyses support this view. LA-ICP-MS data are concordant or nearly concordant, almost continuously covering a 206 Pb/238U age range of 701 ± 15 Ma to 877 ± 20 Ma. Thus, the youngest ages obtained by TIMS and LA-ICP-MS agree reasonably well, while the older spot ages are not recovered by any of the TIMS analyses.

Please cite this article as: Hippe, K., et al., Zircon geochronology of the Koraput alkaline complex: Insights from combined geochemical and U–Pb– Hf isotope analyses, and implications..., Gondwana Research (2015), http://dx.doi.org/10.1016/j.gr.2015.02.021

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Fig. 2. (a) SEM cathodoluminescence (CL) images of all zircons analyzed by LA-ICP-MS for U–Pb–Th (Table 2), Lu–Hf isotopic composition (Table 4) and trace element composition (Table 3). (b) Optical microscope color CL images of zircon obtained within thin section. Images 1, 2, 4, and 5 show thin overgrowths of yellowish high CL intensity on predominantly featureless zircons. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

Presentation of the data in a cumulative probability plot (Fig. 3b) leads us to propose three distinct 206Pb/238U age populations of 869 ± 11 Ma (n = 3), 801 ± 9 Ma (n = 26), and 728 ± 11 Ma (n = 14).

From the three analyses defining the oldest composite age, only one (ru1) was obtained from a distinct high-CL domain that can be interpreted as an inherited core (Fig. 2a). The other two ages were

Please cite this article as: Hippe, K., et al., Zircon geochronology of the Koraput alkaline complex: Insights from combined geochemical and U–Pb– Hf isotope analyses, and implications..., Gondwana Research (2015), http://dx.doi.org/10.1016/j.gr.2015.02.021

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Zircon/fraction IDa

Weight (mg)b

U (ppm)

Pb (ppm)

Isotope ratiosc 206

204

Pb/

Pb

Age (Ma)d 208

Pb/

206

Pb

206

238

Pb/

U

2 σ (%)

207

Pb/

235

U

2 σ (%)

207

Pb/

206

Pb

2 σ (%)

206

Pb/238U

Rho 207

Pb/235U

207

Pb/206Pb

CA-TIMS (single-grain) KO1-1 0.0093 KO1-2 0.0080 KO1-3 0.0138 KO1-5 0.0061 KO1-6 0.0064 KO1-7 0.0020 KO1-8 0.0068

178 146 103 170 278 847 159

22 22 14 21 34 115 26

2850 421 4831 6274 4363 373 219

0.1569 0.1788 0.1694 0.1487 0.1670 0.2346 0.1563

0.1167 0.1240 0.1244 0.1204 0.1129 0.1145 0.1204

0.04 0.11 0.11 0.11 0.18 0.09 0.11

1.0201 1.1038 1.1095 1.0642 0.9726 0.9945 1.0698

0.08 0.28 0.13 0.13 0.22 0.39 0.51

0.0634 0.0645 0.0647 0.0641 0.0625 0.0629 0.0644

0.07 0.24 0.06 0.06 0.12 0.36 0.47

711.6 753.8 755.7 732.9 689.4 699.1 732.9

(0.3) (0.8) (0.9) (0.8) (1.3) (0.7) (0.8)

714.0 755.2 757.9 735.9 689.8 701.0 738.7

(0.6) (2.1) (1.0) (0.9) (1.5) (2.8) (3.7)

721.4 759.1 764.5 745.2 691.2 707.3 756.2

(0.5) (1.8) (0.5) (0.5) (0.8) (7.6) (3.5)

0.5 0.5 0.9 0.9 0.8 0.5 0.4

ID-TIMS (multi-grain) PB 1231 PB 1234 PB 1256 PB 1257 PB 1258 PB 1259 PB 1260 PB 1261 PB 1232 PB 1233

– – – – – – – – – –

– – – – – – – – – –

510 149 116 418 775 1109 594 1894 5193 956

0.1557 0.1729 0.1704 0.1628 0.1701 0.1766 0.1671 0.1534 0.1698 0.1612

0.1160 0.1168 0.1169 0.1159 0.1159 0.1148 0.1201 0.1179 0.1180 0.1183

1.57 1.57 1.59 1.65 1.56 1.57 1.56 1.56 1.56 1.56

1.0338 1.0529 1.0365 1.0348 1.0242 1.0116 1.0688 1.0456 1.0446 1.0501

1.68 2.26 2.74 1.69 1.59 1.57 1.57 1.56 1.56 1.60

0.0646 0.0654 0.0643 0.0648 0.0641 0.0639 0.0645 0.0643 0.0642 0.0644

0.60 1.53 2.10 0.40 0.29 0.13 0.20 0.12 0.05 0.31

707.4 712.0 712.4 706.7 706.7 700.6 731.1 718.6 719.1 721.0

(11.1) (11.2) (11.3) (11.6) (11.0) (11.0) (11.4) (11.2) (11.2) (11.3)

720.8 730.3 722.2 721.3 716.0 709.7 738.1 726.7 726.2 728.9

(12.1) (16.5) (19.8) (12.2) (11.4) (11.1) (11.6) (11.4) (11.3) (11.7)

762.7 786.9 752.7 767.1 745.5 738.4 759.5 751.6 748.4 753.3

(4.6) (12.0) (15.8) (3.1) (2.1) (1.0) (1.6) (0.9) (0.4) (2.3)

0.9 0.7 0.6 1.0 1.0 1.0 1.0 1.0 1.0 1.0

a b c d

– – – – – – – – – –

Single zircons (or zircon fragments) annealed and chemically abraded after Mattinson (2005); zircons for multi-grain analyses were air-abraded. Nominal fraction weights measured after chemical abrasion; no weights taken for multi-grain analyses. Pb isotope ratios corrected for fractionation, spike, and common Pb. Errors are 2-sigma.

K. Hippe et al. / Gondwana Research xxx (2015) xxx–xxx

Please cite this article as: Hippe, K., et al., Zircon geochronology of the Koraput alkaline complex: Insights from combined geochemical and U–Pb– Hf isotope analyses, and implications..., Gondwana Research (2015), http://dx.doi.org/10.1016/j.gr.2015.02.021

Table 1 Results from U–Pb analysis from zircon by CA- and ID-TIMS.

K. Hippe et al. / Gondwana Research xxx (2015) xxx–xxx

Fig. 3. Concordia diagrams showing results for TIMS and LA-ICP-MS U–Pb analyses of zircons from sample KO1 from the Koraput nepheline syenite gneiss. (a) Data obtained by TIMS analyses. Light gray ellipses are results of multi-grain ID-TIMS analyses (n = 10); dark gray ellipses show single-grain CA-TIMS analyses (n = 6). The discordia is based on all ID-TIMS data and the four oldest CA-TIMS results (see text for discussion). The lower intercept of the discordia is indistinguishable from zero and geologically meaningless. (b) Data from LA-ICP-MS analyses with pooled results representing weighted mean 206 Pb/238U ages. The inset shows a cumulative probability plot of all 43 laser ICP-MS analyses.

gained from homogeneous, recrystallized domains in the zircon center (pl2) and at a grain margin (ru4), respectively. There is no consistent trend in the age distribution with regard to zircon morphology or internal textures (Fig. 2). Sector zoned domains yield ages from 842 ± 14 Ma to 731 ± 16 Ma (Fig. 2a: kp2 and ru4); ages from featureless domains range from 866 ± 19 Ma to 709 ± 17 Ma (Fig. 2a: ru4 and ru2). Marginal domains and overgrowth structures, which are considered to represent late crystal growth, yield young ages around 720–700 Ma (Fig. 2a: kp1, pl6) as well as significantly older ages around 810–800 Ma (Fig. 2a: kp3, pl2). Notable age variations also occur within single zircon grains even if apparently identical

7

domains have been analyzed. For example, homogenous and presumably recrystallized domains of grains lp2, kp5, ru2, and pl4 show intracrystalline age differences of up to 60 Ma suggesting incomplete recrystallization with partial resetting of the U–Pb system by redistribution of radiogenic Pb. Generally, most zircons that have been analyzed with two or three laser spots reveal an older central part or ‘core’ (Fig. 2a: kp4, kp6, pl2, pl4, pl6, ru1) and younger overgrowths and/or recrystallized domains. One exception is grain ru4 with an apparently younger sector zoned central part of 799 ± 15 Ma and an older homogeneous, marginal domain of 866 ± 19 Ma. The analysis of the sector zoned domain at the margin of the grain yields a significantly younger age of 731 ± 16 Ma. This pattern reflects how variable different zircon domains are affected by recrystallization and Pb redistribution. All zircons yield medium U (245–518 ppm) and Th (106–249 ppm) contents with Th–U ratios of 0.42–0.63 from LA-ICP-MS trace element analyses (Table 3). The Th–U ratios calculated from the measured Th and U counts during U–Th–Pb isotope analysis give a slightly wider range of 0.33–0.69 (Table 2). In both cases, the Th–U ratio is not correlated with age (Fig. 4a) or texture of the analyzed zircon or zircon domain, respectively. Similarly, no clear correlation is found between the age and the concentrations of U, Th or other trace elements (e.g., Y, Yb, Hf). The zircons have a medium HfO2 content of 1.0–2.2 wt.%, and low concentrations of P (58.6–95.7 ppm), Y (51.8–364.6 ppm) and Nb (0.4–1.6 ppm); their total REE content is low, ranging at 48–354 ppm. The chondrite normalized REE patterns (Fig. 4b) indicate a similar chemical behavior of all zircons with a strong enrichment in HREE (LuN/GdN = 25–58) and a pronounced positive Ce anomaly (Ce/ Ce* = 3–58). There is a tendency towards higher HREE concentrations in zircon cores and higher LREE concentrations in zircon domains identified as overgrowths. This trend is supported by the distribution of Y showing the highest concentrations in 4 of 5 analyzed zircon cores and the lowest concentrations in the recrystallized domains (Fig. 4c). However, as for the Th–U ratio there is no systematic change in the Y or REE content with age. Most zircons exhibit no or only a weak negative Eu anomaly (Eu/ Eu* = 0.47–0.95) implying that zircon did not compete with feldspar for Eu during crystallization. A slightly positive Eu anomaly (Eu/ Eu* = 1.19–2.31) is recorded in few, mostly marginal zircon domains (Fig. 2a: grains ru2, ru4, pl1, pl6-spot at the cone end). Interestingly, those analyses that yield a positive Eu anomaly exhibit a much lower positive Ce anomaly (Ce/Ce* ≤ 10) than analyses with no or a negative Eu anomaly (Ce/Ce* = 14–58) (Fig. 4d). Except for grain pl6 there is, however, no discernible pattern in the relationship of those domains of positive Eu and small positive Ce anomalies and their associated U– Pb ages (see Discussion section). 4.3. Zircon Lu–Hf isotopic systematics Results of the Lu–Hf isotope measurements are given in Table 4. Initial 176Hf/177Hf ratios range from 0.282419 to 0.282603, with εHft values of +4.6 to +10.9 indicating formation from a slightly depleted mantle source. The 176Lu/177Hf ratios scatter moderately from 0.00015 to 0.00037, with the exception of zircon sample pl6, which yielded an anomalously high 176Lu/177Hf ratio of 0.00176. Yb/Hf ratios scatter widely from 0.00010 to 0.01460. The Lu–Hf isotope data do not form any distinct clusters and show no correlation with the analyzed zircon domain. Overgrowths and recrystallized domains have εHft values of 4.8–10.9; core domains show identical εHft values of 4.6–10.6 (Fig. 5a). Equally, both domains yielded the same range of 176Hf/177Hft and 176Lu/177Hf ratios (Fig. 5b, c), except for the very high 176Lu/177Hf in the core domain of sample pl6. There is no obvious correlation between the Lu–Hf data and their associated 206Pb/238U ages. This might be due to the fact that the laser spots for U–Pb and Lu–Hf analyses did not overlap, but partially sampled different domains of the zircon grain making it difficult to compare both

Please cite this article as: Hippe, K., et al., Zircon geochronology of the Koraput alkaline complex: Insights from combined geochemical and U–Pb– Hf isotope analyses, and implications..., Gondwana Research (2015), http://dx.doi.org/10.1016/j.gr.2015.02.021

8

K. Hippe et al. / Gondwana Research xxx (2015) xxx–xxx

Table 2 Results of U–Pb–Th analyses from zircon by LA-ICP-MS. Analysis no.

Zircon ID

Th/U

au10a04 au10a05 au10a06 au10a07 au10a08 au10a09 au10a11 au10a12 au10a13 au10a14 au10a15 au10a16 au10a17 au10b04a au10b05 au10b06 au10b07 au10b08 au10b11 au10b12 au10b13 au10b14 au10b15 au10b16 au10b17 11mr07a03 11mr07a04 11mr07a05 11mr07a06 11mr07a07 11mr07a08 11mr07a09 11mr07a10 11mr07a13 11mr07a14 11mr07a15 11mr07a16 11mr07a18 11mr07a19 11mr07a20 11mr07a21 11mr07a22 11mr07a23

lp1 lp2 lp3 lp4 lp5 ru3 ru1 ru1 ru2 ru4 ru4 pl2 pl4 pl6 pl6 pl1 pl3 pl4 kp3 kp4 kp5 kp6 kp6 kp1 kp2 kp5 kp3 kp1 kp4 lp1 lp2 lp3 lp4 pl2 pl3 pl3 pl4 pl6 pl6 ru1 ru2 ru3 ru4

0.48 0.45 0.40 0.55 0.49 0.40 0.45 0.52 0.37 0.53 0.69 0.37 0.48 0.50 0.44 0.60 0.45 0.50 0.45 0.41 0.48 0.48 0.42 0.46 0.46 0.46 0.41 0.41 0.34 0.49 0.42 0.39 0.52 0.40 0.50 0.43 0.53 0.41 0.46 0.45 0.33 0.40 0.64

206

a b

Age (Ma)b

Isotope ratios Pb/238U

0.1352 0.1254 0.1235 0.1273 0.1313 0.1299 0.1457 0.1324 0.1162 0.1438 0.1320 0.1299 0.1350 0.1386 0.1160 0.1365 0.1385 0.1308 0.1376 0.1337 0.1190 0.1354 0.1161 0.1223 0.1395 0.1271 0.1338 0.1179 0.1196 0.1290 0.1362 0.1286 0.1280 0.1434 0.1176 0.1327 0.1275 0.1278 0.1148 0.1241 0.1226 0.1307 0.1202

2σ (%)

207

2.0 2.1 2.4 2.1 2.1 2.4 2.4 2.2 2.5 2.3 2.0 2.3 2.4 2.2 2.0 2.3 2.5 2.4 2.2 2.1 2.2 2.2 2.3 2.1 2.2 2.3 2.3 2.2 2.3 2.2 2.3 2.3 2.2 2.2 2.2 2.4 2.4 2.3 2.2 2.3 2.3 2.4 2.3

1.2454 1.1733 1.1895 1.2066 1.2735 1.1395 1.4044 1.2309 1.1088 1.4518 1.2164 1.0995 1.2411 1.4038 1.0406 1.2580 1.1657 1.2623 1.2620 1.2541 1.1345 1.2585 1.0109 1.1317 1.2462 1.1968 1.2582 1.0369 1.0496 1.1517 1.2470 1.1418 1.1279 1.2901 0.9720 1.1215 1.1757 1.1631 1.0082 1.1731 1.0866 1.1939 1.0347

Pb/235U

2σ (%)

207

7.9 8.6 11.7 8.7 8.4 11.7 11.5 9.6 11.8 10.4 7.9 10.2 11.6 8.5 6.8 9.8 11.5 10.8 9.2 7.8 9.0 9.0 9.8 8.2 9.3 6.3 6.6 5.9 6.6 5.9 6.3 6.4 5.8 6.1 5.9 7.7 7.5 6.3 5.8 6.5 7.0 7.3 5.9

0.0653 0.0663 0.0643 0.0661 0.0674 0.0615 0.0651 0.0645 0.0636 0.0695 0.0639 0.0610 0.0655 0.0705 0.0622 0.0656 0.0607 0.0673 0.0648 0.0661 0.0657 0.0659 0.0631 0.0633 0.0634 0.0648 0.0656 0.0633 0.0644 0.0645 0.0651 0.0651 0.0643 0.0649 0.0614 0.0626 0.0664 0.0661 0.0627 0.0672 0.0627 0.0675 0.0636

Pb/206Pb

2σ (%)

207

4.7 5.2 7.0 5.1 4.8 7.1 6.5 5.4 7.0 5.5 4.0 5.7 6.4 4.9 4.1 5.8 7.0 6.4 5.3 4.4 5.2 5.0 5.9 4.5 5.0 4.3 4.3 4.1 4.5 4.0 4.2 4.4 4.0 4.1 4.2 5.1 5.0 4.4 4.1 4.4 4.8 4.9 4.2

784 817 750 809 851 657 777 758 728 913 740 639 790 943 682 793 630 846 767 811 795 803 713 718 721 767 794 717 755 757 777 779 752 771 653 696 818 811 698 844 700 853 729

Pb/206Pb (97) (107) (145) (104) (98) (148) (133) (113) (144) (111) (84) (121) (131) (99) (87) (119) (147) (129) (109) (90) (108) (104) (122) (94) (105) (88) (90) (86) (94) (84) (88) (91) (83) (86) (88) (107) (103) (90) (87) (91) (101) (100) (88)

206

Pb/238U

817 762 751 772 795 787 877 802 709 866 799 787 816 837 707 825 836 793 831 809 725 819 708 744 842 771 809 718 728 782 823 780 776 864 717 803 773 775 701 754 746 792 731

(16) (15) (17) (16) (16) (18) (20) (17) (17) (19) (15) (17) (19) (17) (14) (18) (20) (18) (17) (16) (15) (17) (16) (15) (18) (16) (17) (15) (16) (16) (17) (17) (16) (18) (15) (18) (17) (17) (15) (16) (16) (18) (16)

207

Pb/235U

821 788 796 804 834 772 891 815 758 911 808 753 819 891 724 827 785 829 829 825 770 827 709 769 822 799 827 722 729 778 822 773 767 841 690 764 789 783 708 788 747 798 721

(45) (47) (65) (48) (48) (63) (68) (54) (63) (62) (44) (54) (65) (50) (35) (56) (63) (61) (52) (44) (49) (51) (50) (44) (52) (35) (37) (30) (34) (32) (36) (35) (31) (35) (29) (41) (41) (35) (29) (36) (37) (40) (30)

Rho 0.7 0.6 0.5 0.6 0.7 0.6 0.6 0.6 0.5 0.6 0.7 0.6 0.6 0.7 0.8 0.6 0.6 0.6 0.7 0.7 0.6 0.7 0.6 0.7 0.7 0.9 0.9 1.0 0.9 1.0 1.0 1.0 1.0 1.0 1.0 0.9 0.8 1.0 1.0 0.9 0.9 0.9 1.0

Outlier; not included in Fig. 3b. Errors are 2-sigma.

data sets. Hafnium model ages have a narrow range from TDM = 1.0 to TDM = 1.5 Ga and mark the timing when the magma source of the syenite melt was extracted from the depleted mantle reservoir. Model ages are about 200–700 Ma older than the associated 206Pb/238U ages. There are different reasons that can account for these older Hf model ages, e.g., contamination of the juvenile magma with isotopically older country rocks during intrusion, complete resetting of the U–Pb system during metamorphic overprint not affecting the Lu–Hf system (Nebel et al., 2007), time difference between the crust formation and the intrusion event (Hawkesworth and Kemp, 2006). With regard to the observed overall uniformity in the Lu–Hf isotope system there is, however, no evidence for a contamination of the host magma.

from the nepheline syenite gneiss in the Koraput alkaline complex are characteristic of zircons from high-grade metamorphic rocks (e.g., Kröner et al., 1991; Schaltegger et al., 1999; Bingen et al., 2001). The majority of studies on metamorphic zircon formation describe composite crystals with inherited, distinct magmatic cores surrounded by variably wide metamorphic overgrowths (e.g., Vavra et al., 1996; Hoskin and Black, 2000; Möller et al., 2002; Gerdes and Zeh, 2009). Most of the zircons from the Koraput nepheline syenite gneiss do not provide such a clear distinction, but rather have metamorphic features (homogeneity, sector zoning, blurred structures) defining the entire zircon crystal. In the following, we will discuss the probable zircon growth conditions of the magmatic formation as well as the metamorphic growth and modification in detail.

5. Discussion 5.1. Zr saturation and magmatic zircon growth conditions Crystal morphology and internal structures of separated zircons from the Koraput nepheline syenite gneiss show evidence for metamorphic growth and pervasive recrystallization. The rarity of typical magmatic zircon attributes such as oscillatory zoning or bipyramidal prismatic shape is an important observation for the interpretation of the analytical data. Instead of magmatic features, the predominantly rounded and multi-faceted or simple prismatic shapes of the zircons

Alkaline rocks are generally thought to be quite enriched in Zr (Mielke, 1979; Watson, 1979), and Zr concentrations of several 100 up to 1000 ppm have been reported for nepheline syenites (Jones and Larsen, 1985; Andersen and Sørensen, 1993; Eby and Sclar, 1993; Eby et al., 1998). However, whole rock chemical analyses from the Koraput nepheline syenite gneisses (Table S1 in the Supplementary material)

Please cite this article as: Hippe, K., et al., Zircon geochronology of the Koraput alkaline complex: Insights from combined geochemical and U–Pb– Hf isotope analyses, and implications..., Gondwana Research (2015), http://dx.doi.org/10.1016/j.gr.2015.02.021

Analysis no.

au10c03

au10c04

au10c05

au10c06

au10c07

au10c08

au10c09

au10c10

au10c12

au10c13

au10c14

au10c15

au10c16

au10c17

au10c18

Zircon ID

lp2

lp1

lp5

lp4

ru2

ru3

ru1

ru4

pl2

pl1

pl6

pl6

kp2

kp1

kp6

Domain Ti P Nb Y Th U Th/U Hf

Core b2.72 66 0.66 190 161 311 0.52 11559

Core

Recryst. 5.61 86 0.61 169 112 248 0.45 11919

Recryst. 6.14 84 0.66 146 118 252 0.47 8463

Overgrowth 4.72 77 0.94 133 111 266 0.42 11096

Core

Overgrowth 6.22 96 0.83 155 104 243 0.43 10603

Recryst. 8.14 59 0.49 52 140 312 0.45 11434

Recryst. 5.05 62 0.60 201 101 235 0.43 9507

Recryst. 6.91 78 0.70 166 135 259 0.52 12006

Overgrowth 6.18 63 1.64 234 239 498 0.48 19121

Core b4.77 62 1.27 365 124 269 0.46 10053

Recryst. 5.69 92 0.94 175 139 302 0.46 12120

Core 6.98 70 0.76 276 205 397 0.52 9888

Recryst. 6.33 73 0.92 107 122 279 0.44 11745

Rare earth elements La 0.005 Ce 2.16 Pr 0.016 Nd 0.13 Sm 0.41 Eu 0.36 Gd 3.4 Tb 1.19 Dy 16.3 Ho 6.1 Er 33 Tm 7.8 Yb 85 Lu 17.2 Σ REE 173 Ce/*Ce 58.32 Eu/*Eu 0.92

3.66 73 0.43 267 135 284 0.48 10840 b0.019 2.13 0.035 0.47 1.11 0.41 6.4 1.97 23.1 8.8 46 9.8 101 21.0 223 19.96 0.47

b0.020 1.78 0.005 0.25 0.55 0.33 3.9 1.20 15.4 5.8 28 5.8 61 12.4 137 42.99 0.69

0.007 2.69 b0.026 0.13 0.60 0.37 2.4 0.86 12.5 4.6 24 5.5 59 12.1 15 48.22 0.95

0.240 4.53 0.193 0.89 0.60 0.64 3.2 0.89 12.3 4.2 23 4.9 54 11.2 120 5.09 1.42

5.39 84 1.31 355 161 256 0.63 11149

0.081 4.19 0.067 0.82 1.18 0.62 6.2 2.36 30.6 10.8 60 14.0 152 31.0 313 13.77 0.70

b0.022 2.99 0.016 0.12 0.46 0.28 2.7 0.83 12.9 4.7 25 5.9 62 12.6 131 38.52 0.77

0.019 0.75 b0.017 0.17 0.17 0.35 1.3 0.42 5.2 1.6 9.0 1.9 23 4.0 48 10.12 2.31

0.030 2.48 0.043 0.28 0.45 0.39 3.5 1.02 14.5 6.2 35 8.1 96 19.8 188 16.73 0.95

0.276 3.16 0.172 1.18 1.42 0.89 3.7 1.27 15.0 5.2 28 6.0 64 13.1 143 3.50 1.19

0.436 5.53 0.350 1.69 1.36 1.53 5.3 1.85 21.6 7.6 39 9.7 107 21.8 225 3.42 1.74

b0.034 4.33 0.049 0.55 0.92 0.50 5.4 2.16 28.0 11.3 64 16.3 182 38.8 354 25.64 0.68

0.035 2.24 0.027 0.24 0.45 0.28 3.6 1.24 16.1 5.5 30 6.8 66 14.2 147 17.61 0.66

0.005 3.50 0.057 0.91 1.24 0.78 5.8 2.03 23.9 8.8 43 10.1 103 20.2 223 50.19 0.89

0.000 1.80 0.012 b0.193 0.30 0.23 2.4 0.64 8.4 3.1 16 3.7 40 8.8 86 0.00 0.81

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Table 3 Representative LA-ICP-MS trace element analyses of zircon (in ppm). Please note that the classification of the different zircon domains is not unambiguous and should be used with caution (see Fig. 2).

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K. Hippe et al. / Gondwana Research xxx (2015) xxx–xxx

Fig. 4. (a) Th/U versus U–Pb age diagram showing a narrow range of Th/U for all analyzed zircon domains independent from the determined age. Data obtained by LA-ICP-MS U–Pb–Th isotope analyses; ages are given with 2σ error bars. (b) Chondrite normalized REE patterns for the Koraput nepheline syenite gneiss zircons; normalization after Sun and McDonough (1989). Zircon core domains tend to have the highest HREE concentrations, while overgrowths show the highest LREE concentrations. (c) Y concentration (in ppm) versus Th/U diagram with highest Y concentrations in zircon core domains. (d) Diagram showing the relation between the measured Eu and Ce anomalies. Analyses with Eu/Eu* N 1.0 (positive Eu anomaly) and low positive Ce/Ce* are clearly distinguishable from those with Eu/Eu* b 1.0 and a high positive Ce/Ce*. Core analyses all belong to the latter group. Please note that the assigned classification of the different CL-textural zircon domains is not unambiguous and should be used with caution (see Fig. 2).

yielded on average only 136 ppm Zr (n = 13), with 106 ppm Zr in the nepheline syenite gneiss sample KO1 hosting the zircons analyzed in this study. We argue that such low Zr concentrations restricted primary zircon crystallization during early magmatic cooling. Watson (1979) showed that the saturation level required for zircon crystallization strongly depends on the molar proportions of (Na2O + K2O) / Al2O3

with rather little sensitivity to the SiO2 content. Based on this tenet, Watson and Harrison (1983) presented a Zr saturation model as a function of temperature and melt composition using the cation ratio M = (Na + K + 2 Ca) / (Al · Si) as a parameter for melt basicity. They found that the Zr saturation level increases with increasing melt alkalinity and rising temperature. In Fig. 6, we have applied the model of

Table 4 Lu–Hf isotope analyses from zircons of the Koraput complex. Analyses from the zircon center/core domain are marked with an asterisk. Zircon ID

Lu/Hf

Yb/Hf

176

pl3⁎

0.00230 0.00266 0.01266 0.00128 0.00177 0.00110 0.00160 0.00116 0.00199 0.00111 0.00174 0.00120 0.00179 0.00106 0.00184

0.01444 0.01460 0.00754 0.00783 0.01087 0.00659 0.00963 0.00579 0.01030 0.00614 0.01050 0.00713 0.00974 0.00593 0.00010

0.00032 0.00037 0.00176 0.00018 0.00025 0.00015 0.00022 0.00016 0.00028 0.00015 0.00024 0.00017 0.00025 0.00015 0.00026

pl4 pl6⁎ pl1 pl2 kp3⁎ kp4 kp2⁎ kp1⁎ lp2⁎ lp1⁎ lp4 ru1 ru2 ru3

Lu/177Hf

176

Hf/177Hf

0.282487 0.282425 0.282585 0.282571 0.282475 0.282431 0.282492 0.282451 0.282483 0.282532 0.282425 0.282605 0.282514 0.282546 0.282594

±2σ

176

Hf/177Hfta

0.000028 0.000034 0.000030 0.000034 0.000050 0.000032 0.000036 0.000036 0.000056 0.000032 0.000042 0.000040 0.000024 0.000042 0.000030

0.282482 0.282419 0.282557 0.282568 0.282471 0.282429 0.282489 0.282448 0.282479 0.282530 0.282421 0.282603 0.282510 0.282544 0.282590

εHft

±2σ

Age (Ma)b

±2σ

TDM (Ga)c

7.2 4.8 10.6 10.8 8.2 6.0 5.8 6.9 5.8 9.4 4.6 10.9 9.9 8.1 10.7

1.0 1.2 1.1 1.2 1.8 1.1 1.3 1.3 2.0 1.1 1.5 1.4 0.9 1.5 1.1

803 793 837 825 864 831 728 842 744 823 782 776 877 746 787

18 18 17 18 18 17 16 18 15 17 16 16 20 16 18

1.3 1.5 1.4 1.1 1.3 1.4 1.3 1.3 1.3 1.1 1.4 1.0 1.2 1.2 1.0

a Initial 176Hf/177Hft and εHft calculated using the 206Pb/238U age determined by LA-ICP-MS dating, and the CHUR parameters 176Lu/177Hf = 0.0336, and 176Hf/177Hf = 0.282785 (Bouvier et al., 2008). b 206 Pb/238U age. c Two-stage Hf model ages (TDM) in Ga calculated using a 176Lu/177Hf of 0.0384, a 176Hf/177Hf of 0.28325 for the depleted mantle (Chauvel and Blichert-Toft, 2001), and a crustal 176Lu/177Hf of 0.015 (Griffin et al., 2002).

Please cite this article as: Hippe, K., et al., Zircon geochronology of the Koraput alkaline complex: Insights from combined geochemical and U–Pb– Hf isotope analyses, and implications..., Gondwana Research (2015), http://dx.doi.org/10.1016/j.gr.2015.02.021

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Fig. 6. Zircon saturation temperatures Tzir sat for the Koraput nepheline syenite gneiss as predicted from the whole rock cation ratio M = (Na + K + 2 · Ca) / (Al · Si) and the Zr concentration. Zr saturation curves are based on the zircon solubility model of Watson and Harrison (1983). For comparison, zircon crystallization temperatures Tzir Ti calculated from the Ti-in-zircon thermometer (Watson et al., 2006) are given in the inset.

These results are consistent with zircon crystallization temperatures of 640–720 °C determined by Ti-in-zircon thermometry in the analyzed grains. The Ti-in-zircon thermometer is based on the temperature dependent uptake of Ti into zircon and is expressed by the equation log(Tizircon) = (6.01 ± 0.03) − 5080 ± 30/T(K), with Tizircon (in ppm) being the measured Ti content in zircon (Watson et al., 2006). Based on these considerations and the observed zircon morphologies and textures, we propose that the majority of zircons formed during late stages of advanced magmatic cooling and (to a minor extent) during subsequent high-grade metamorphism. A locally restricted zircon saturation and nucleation is indicated by the occurrence of zircon clusters and intergrown zircon crystals (Fig. 7a, b, e). The presence of internal sector zoning, which is thought to result from low lattice diffusion and slow growth from a fluid-enriched boundary layer (Watson and Liang, 1995; Schaltegger et al., 1999), suggests fluid-enhanced crystal growth. However, ‘fir-tree zoning’, the commonly eccentric position of the crystal growth center as well as the occurrence of rounded growth banding indicates a rather limited fluid activity causing fluctuating growth rates and repeated growth inhibition (Vavra et al., 1996, 1999). Fig. 5. Results from LA-ICP-MS U–Pb and Lu–Hf isotope analyses. The εHft values (a), 176 Hf/177Hft ratios (b), and 176Lu/177Hf ratios (c) are plotted versus the associated 206 Pb/238U ages. All errors are 2σ; filled black symbols mark analyses from zircon core domains, open symbols are rim or recrystallized domains (see Fig. 2). 176Lu/177Hf results for sample pl6 are not shown in diagram due to very high value (c). Because laser spots for Lu–Hf and U–Pb analyses do not overlap, comparing both data is difficult. However, diagrams show overall moderate scatter in the Lu–Hf distribution with no correlation between the Lu–Hf data and the zircon domain or the 206Pb/238U age.

Watson and Harrison (1983) to identify Zr saturation temperatures in the nepheline syenite gneisses (M = 1.9–2.6). Temperature curves for 650 °C to 800 °C are based on the model equation lnDzircon/melt = Zr (−3.80 − [0.85 ⋅ (M − 1)]) + 12, 900/T(K), where lnDzircon/melt repreZr sents the concentration ratio of Zr in stoichiometric zircon to that in the melt (Watson and Harrison, 1983). The experimental range of Watson and Harrison (1983) did not include M-values below 2.0 or temperatures below 750 °C. However, an overall good agreement between the model predictions and the experimental data was observed for the low temperature experiments (Watson and Harrison, 1983). Thus, the extrapolation of the temperature curves up to M = 2.6 and down to 650 °C is considered to be a valid approach to gain an estimate on the Zr saturation temperature in the Koraput nepheline syenite gneisses. As illustrated in Fig. 6, most data scatter between the 650 °C and 750 °C saturation curves, strongly suggesting that magmatic zircon did not crystallize during early magma cooling due to undersaturation.

5.2. Metamorphic zircon growth Zircon internal structures as observed from CL-imaging allow us to differentiate at least three different textural events (see Fig. 2a: e.g., ru1, pl6, pl4, kp4). Primary zircon growth was followed by two or more events of metamorphic overprint of which zircon recrystallization is the dominant feature affecting a larger volume of the imaged grains. Characteristic reaction fronts reaching into rather featureless, presumably recrystallized, domains or weakly zoned cores imply that partial recrystallization may have occurred more than once (Fig. 2a: kp4, pl1, pl3, ru1, lp5). The final stage of zircon growth is represented by lobate overgrowths interpreted as evidence for zircon dissolution and reprecipitation (Fig. 2a: pl4, kp2, kp4). In some cases, this final zircon formation can be associated with growth along grain-boundaries that changes the prior grain shape and is characterized by protruding overgrowths (Fig. 2a: kp1, lp4; 2B: e, f) that may produce typical “snail-foot” shape (Fig. 2a: kp1). Thin sections display a frequent paragenesis of zircon with (partially decomposed) biotite, where zircon crystals grow along the biotite grain boundary or in between two biotite grains (Fig. 7). In few cases, zircons are similarly associated with ilmenite (Fig. 7b, f). Similar parageneses of zircon with other minerals have been reported in the literature and interpreted as metamorphic zircon formation after the breakdown of different Zr-bearing phases. For example, Davidson and van Breemen (1988) documented coronal zircon grown around baddeleyite in high-

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Fig. 7. Photomicrographs from three thin sections (section KO1-b1: pictures a–d, section KO1-2: picture e, and section KO1-b3: picture f) of the nepheline syenite gneiss sample KO1 illustrating the textural relationship between zircon and biotite (and ilmenite). White circles on biotite show the location of LA-ICP-MS analysis spots with the measured Zr concentration given in ppm; circle diameters vary as they represent the actual spot size. Photographs were taken under cross-polarized light, except the insets in image (b), which are optical CL images.

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grade metagabbros, while Degeling et al. (2001) and Tomkins et al. (2005) found metamorphic zircon formed by the breakdown of Zrbearing garnet in metapelitic rocks. Bingen et al. (2001) observed zircon coronas and “hat”-shaped zircons in contact with ilmenite suggesting a conversion of pre-existing baddeleyite to zircon due to Zr release from ilmenite at high Si-activity levels during metamorphism. Accordingly, trace element analyses on potentially Zr-bearing phases in felsic granulites and alkaline rocks yielded the highest Zr abundances in garnet (up to ~ 60 ppm), amphibole (up to ~ 190 ppm), rutile (up to ~ 250 ppm), and titanite (up to 2100 ppm), whereas lower concentrations of Zr (~20–50 ppm) were also measured in pyroxene, magnetite, and ilmenite (Green, 1994; Bea et al., 2006). For biotite, these studies report Zr concentrations well below 5 ppm. However, Deer et al. (2003) presented a summary of data with slightly higher Zr abundances of 14.9 ppm in phlogopites and of 17.6 ppm in Mg–Fe-biotites. Biotite decomposition as a probable Zr source for metamorphic zircon overgrowth was proposed by Vavra et al. (1996, 1999) for metasediments from the Ivrea Zone, but unfortunately, no data was given on the Zr concentration for those biotites. To explore the particular zircon–biotite paragenesis found in the Koraput nepheline syenite gneiss, we used LA-ICP-MS analysis on thin sections and determined the Zr content within biotite grains both in the vicinity and far away from zircon grains. Measured concentrations range from 1.2 ppm to 30.1 ppm (n = 20; Fig. 7; Table S2 in the Supplementary material) and are thus higher than reported elsewhere. Two factors might have contributed to the comparatively high Zr concentrations: 1) late-stage magmatic zircon formation, i.e. Zr was not removed by zircon crystallization during early magmatic cooling, but was still largely present at the time of biotite crystallization, and 2) the absence of other major rock forming minerals that can incorporate Zr into their crystal lattice. The nepheline syenite gneisses of the Koraput alkaline complex contain only accessory amounts of amphibole and ilmenite and none of the other above mentioned potentially Zr-bearing minerals. Consequently, biotite could have taken up Zr, substituting two octahedral cations Fe2+ and/or Mg2+ by one Zr4+. As biotite breakdown during the granulite-facies overprint of the Koraput alkaline complex is well documented in thin sections, we propose that Zr was then released from the biotite lattice inducing local Zr saturation and metamorphic zircon growth. Small amounts of metamorphic fluids released simultaneously from biotite during dehydration could have facilitated Zr mobility. The spatial distribution of Zr in the analyzed biotites strongly supports the idea of biotite as Zr source for metamorphic zircon growth (Fig. 7). In most cases, the lowest Zr concentrations are detected in biotites adjacent to zircon crystals. Moreover, the Zr content within a single biotite or a group of adjoining biotites tends to decrease significantly in proximity to zircon (Fig. 7a, b, e) supporting the idea of Zr diffusion out of biotite and Zr uptake into newly grown zircon. Two exceptions are given in Fig. 7c and d. Here, biotites exhibit comparatively high Zr concentrations even close to intergrown zircons. These zircons have a simple prismatic morphology unlike the strikingly irregular shapes of most other zircons observed in the thin sections. They are therefore most probably of magmatic origin, which is in line with the observation that the zircon in Fig. 7c is dark in optical CL, while the clearly metamorphic zircons (anhedral shape, growing along grain boundaries) of Fig. 7b have a high CL-intensity. In addition to biotite, three ilmenite crystals, one of them adjacent to zircon, have been analyzed for their Zr content. Although the detected counts did not exceed the analytical background level, textural relations between zircon and ilmenite (Fig. 7b, f) point at ilmenite as an additional, though minor, Zr source during metamorphic zircon growth. Although biotite-connected metamorphic zircons occur frequently in thin sections, they are not equally represented in the zircon populations separated for U–Pb dating or chemical composition analysis. Due to their anhedral, partially interstitial morphology, these zircons were probably destroyed and this effectively excluded early on in the mineral separation process.

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5.3. Geochemical implications The overall homogeneous geochemical properties of the zircons from the Koraput nepheline syenite gneiss strongly contrast with the diversity in textural patterns and the wide range of concordant U–Pb ages, which both argue for a complex zircon growth and alteration history. In cases of multiple-stage zircon growth and modification and/or extensive recrystallization, numerous studies have reported significant variability in the trace element composition and concentration between zircons or zircon domains of different age and growth mechanisms. For example, differences were typically found in the Th–U ratios as well as the distribution of REE and other trace elements between magmatic zircon cores and metamorphic overgrowths or between pristine and recrystallized zircon domains (e.g., Williams et al., 1996; Rubatto et al., 2001; Hoskin and Schaltegger, 2003; Möller et al., 2003; Rubatto and Hermann, 2003). Trace element data from the zircons studied here, however, show little variability. The measured Th/U (0.3–0.7) is in the typical range of magmatic zircons and is equally distributed throughout all crystal domains as well as independent from the associated U–Pb ages. The same observations are made for the Lu–Hf data. Although the moderate data scatter is too large to be explained by variations in zircon crystallization ages, and thus requires an additional source of slightly different Hf isotopic composition (probably metamorphic fluids), no trend in the Lu–Hf distribution with zircon domain or age is apparent. Similarly, the REE pattern shows an overall homogeneous distribution with only slightly higher HREE (and Y) concentrations from zircon center/core domains. These results suggest only very limited exchange (due to growth of competing accessory phases such as xenotime or monazite) or open-system (e.g. melt-production and -loss) behavior during metamorphic zircon recrystallization. We find no evidence for trace element modification during recrystallization, implying an overall closed system through time. This was also observed in zircons with polymetamorphic overgrowths documented by Möller et al. (2002, 2003) where only the second, ultra-high temperature, metamorphic event produced significant changes in the Th/U. The analyses of metamorphic overgrowths in the zircons studied here show little variation in their trace element distribution. The only exception is a variation in the Eu- and Ce-anomalies. The general absence of a pronounced negative Eu-anomaly, which typically results from co-precipitation with feldspar, supports the interpretation of zircon growth during advanced fractionation. The occurrence of zircons (or zircon domains) with a positive Eu/Eu* and a Ce/Ce* ≤ 10 in contrast to those with no or a negative Eu/Eu* and a Ce/Ce* = 14–58 might reflect slight changes in the geochemical composition of newly grown, metamorphic zircon domains (pos. Eu/Eu*) compared to primary, magmatic zircon (no/neg. Eu/Eu*). A lower positive Ce anomaly is indicative for crystal growth under less oxidizing conditions while the occurrence of a positive Eu anomaly might result from a release of the Eu-rich albite component from plagioclase during high-T metamorphism. This interpretation is illustrated best in zircon sample pl6, in which two distinct domains have been analyzed for trace elements (Fig. 2a). While the older zircon core (837 ± 17 Ma) has a weakly negative Eu-anomaly (Eu/Eu* = 0.68) and a large Ce-anomaly (Ce/Ce* = 26), the younger metamorphic overgrowth (701 ± 15 Ma and 707 ± 14 Ma) shows a positive Eu anomaly (Eu/Eu* = 1.74) and a smaller Ce anomaly (Ce/ Ce* = 3). The Th–U ratios from both domains are, however, identical with 0.46 and 0.48. Because CL image interpretation for other zircon grains is ambiguous, the correlation between Eu- and Ce-anomalies and the associated U–Pb ages is not entirely clear. 5.4. Geochronology of the Koraput alkaline complex The U–Pb zircon data obtained from the nepheline syenite gneisses are interpreted to be strongly affected by multiple and pervasive metamorphic overprints of the Koraput alkaline complex. The data set is complex with limited clustering of data to allow a clear differentiation

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between the timing of alkaline magmatism and one or more metamorphic events. Instead, results obtained by LA-ICP-MS show concordant (or near concordant) ages spreading over about 180 Ma (Fig. 3b). Of these, a single zircon core provided the oldest 206Pb/238U age of 877 ± 20 Ma, consistent with two other 206Pb/238U ages of 866–864 Ma. These overlap with the published Rb–Sr whole rock age of 856 ± 18 Ma of Sarkar et al. (1989).We interpret the mean age of the three oldest LAICP-MS 206Pb/238U ages (869 ± 11 Ma) to reflect the time of pluton emplacement. Following this interpretation, the intrusion of the Koraput alkaline complex during the Mid-Neoproterozoic clearly postdates the major phases of alkaline magmatism proposed previously at 1.5–1.2 Ga along the western EGB (Aftalion et al., 1988; Paul et al., 1990; Shaw et al., 1997; Aftalion et al., 2000; Upadhyay et al., 2006a). The age of about 870 Ma for the Koraput alkaline complex nepheline syenite correlates better, but is still distinctly younger than ages of ca. 990–980 Ma (U–Pb zircon and perrierite; Grew and Manton, 1986) and ca. 960–990 Ma (U–Pb zircon and monazite; Paul et al., 1990) which have been associated with granulite facies metamorphism and charnockite intrusions in the northern and southeastern Eastern Ghats Province as well as zircon U–Pb ages of 980–900 Ma dating high-grade metamorphism in the central EGB (Bose et al., 2011). A magmatic age of 869 ± 11 Ma for the Koraput complex is also younger than the upper intercept zircon U–Pb age for ferrodiorites at Bolangir of 933 ± 32 Ma in the Northwestern EGB interpreted as magmatic (Krause et al., 2001), but older than the 792 ± 2 Ma concordant zircon results from a ferrodiorite dating the intrusion of the Chilka Lake complex in the Northeastern EGB (Krause et al., 2001). The titanite age of 935 ± 25 Ma from the northern EGB was interpreted to date metamorphism at the time of magmatic emplacement at Bolangir (Mezger and Cosca, 1999). However, monazite U–Th–Pb ages from the eastern EGB suggest regional high-grade metamorphism at ca. 900 Ma (Simmat and Raith, 2008) and correlate with the proposed emplacement age of the Koraput alkaline complex. In contrast, Hf model ages of TDM = 1.5–1.0 Ga are significantly older and coincide with the Mid-Proterozoic EGB alkaline magmatic events. These ages are in line with the supposed recurring ascent of mantle-derived magma and the production of new crust throughout the Mesoproterozoic (e.g., Sarkar and Paul, 1998). Hence, the intrusion of the Koraput alkaline body and the crystallization of zircons from the nepheline syenite magma did not closely follow the crust-forming event recorded by the Lu–Hf isotope signature. Evidence for a metamorphic overprint of the Koraput pluton is reflected in the U–Pb TIMS data, which give an oldest concordant CATIMS age of 754 ± 1 Ma. The discrepancy between the oldest TIMS and LA-ICP-MS ages can be explained by the small volume of old domains within the entire zircon population as well as within single zircon grains. The few old domains are therefore only detected by spot measurements but may have been leached away during chemical abrasion of the CA-TIMS analyses and cause the discordance of the conventional TIMS data. Due to the pronounced internal complexity of the Koraput zircons, even single-grain TIMS results likely represent a mixture of different zircon domains and, thus, different ages. However, even the good spatial resolution provided by the LA-ICP-MS does not allow us to distinguish the timing of specific metamorphic events from the large pool of concordant U–Pb ages. Because zircon recrystallization caused variable rejuvenation by partial resetting of the U–Pb ages, many apparent U–Pb ages likely represent a composite age of growth and recrystallization events. Some of these recrystallization events may have mobilized radiogenic Pb, but did not purge it from the grains, as documented by McFarlane et al. (2005) in zircons from high-grade gneisses from Canada. With a relative short time span between initial crystallization and recrystallization or overgrowth formation, mixed analyses may appear concordant within analytical uncertainty, but are in fact short mixing lines that are near-parallel to the Concordia curve. Without further information, we therefore cannot assign a distinct geological significance to the intermediate age group of zircon LA-ICP-MS data at ca. 800 Ma, despite the excellent correlation with the emplacement age

of the Chilka Lake anorthosite at 792 ± 2 Ma (Krause et al., 2001) and a close time relationship with the Chilka lake leucogranites at 743–762 Ma (Dobmeier and Simmat, 2002). Further evidence for metamorphic activity at ca. 800 Ma was found by zircon U–Pb dating of metapelites in the eastern-central part (Shaw et al., 1997) as well as in the northern and southern part of the Eastern Ghats Province (Upadhyay et al., 2009). Nevertheless, there is a reasonably good agreement between the youngest LA-ICP-MS and TIMS ages suggesting that a final metamorphic event occurred around 700–690 Ma. As petrographic evidence identifies granulite facies metamorphism as the final metamorphic overprint, we propose that this event can be associated with the youngest U–Pb ages, best defined by a concordant CA-TIMS age at 690 ± 1 Ma. This is in excellent agreement with the chemical U–Th–Pb monazite ages of 690–660 Ma, obtained by Dobmeier and Simmat (2002) on Chilka Lake complex leucosomes and interpreted to define a granulite facies metamorphic event producing a strong foliation. In summary, we interpret the data provided here to indicate that alkaline magmatism along the western border of the EGB was not restricted to Mesoproterozoic activity. Consistent with results from the northern EGB (e.g., Krause et al., 2001), the emplacement of the Koraput alkaline complex suggests prolonged or recurring Neoproterozoic orogenic activity followed by deformation and extensive granulite-facies metamorphism throughout the Mid-Neoproterozoic. No direct evidence was found for late Pan-African events at ca. 500 Ma as reported elsewhere in the EGB (e.g., Mezger and Cosca, 1999; Dobmeier and Simmat, 2002; Upadhyay et al., 2006a). 6. Conclusions Zircons from the Koraput alkaline complex record a multiple-stage history of zircon growth and repeated modification. Our analyses show that magmatic zircon formation occurred exceptionally late during magmatic cooling due to low Zr concentrations, requiring relatively low temperatures for Zr saturation. The complex pattern of internal zircon structures and the wide range of U–Pb ages are interpreted to result from recurring zircon recrystallization, partial dissolution–precipitation, followed by overgrowths and new subsolidus growth triggered by biotite breakdown during metamorphism (see below). The overall uniform trace element distribution implies a homogeneous source and indicates an effective decoupling between the U–Pb ages and most trace elements. Calculated Hf model ages range in the Mesoproterozoic and overlap with ages of most alkaline bodies along the western EGB (e.g., Aftalion et al., 1988; Sarkar and Paul, 1998; Aftalion et al., 2000; Upadhyay et al., 2006a,b) as well as with estimates for the age of earliest high-grade metamorphism in the EGB (see review in Dasgupta and Sengupta, 2003). This is interpreted as evidence for a juvenile source of the nepheline syenites within the Koraput alkaline complex. Extensive and probably repeated recrystallization caused incomplete resetting of the U–Pb system leading to widely spread apparently concordant U–Pb ages between 869 ± 11 Ma and 690 ± 1 Ma, the former interpreted as the magmatic crystallization age, the latter as the best estimate for final high-grade metamorphism. One consequence of late zircon crystallization and the (near) absence of major Zr-bearing minerals in the nepheline syenite gneisses was the incorporation of significant amounts of Zr (up to ~ 30 ppm) into the crystal lattice of biotite. The breakdown of biotite during high-grade metamorphism then released sufficient Zr to induce local Zr saturation along the fluid-enriched biotite grain boundaries, which led to substantial metamorphic zircon growth. With regard to regional geochronology, U–Pb ages suggest an emplacement age for the Koraput alkaline complex at around 870 Ma. Thus, the intrusion of the Koraput pluton cannot be linked to the widespread, rift-related alkaline magmatism generating a large number of alkaline plutons in the western EGB during the Mesoproterozoic. The data presented here further imply that the final pulse of high-grade

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thermal activity occurred at ca. 700–690 Ma. Together with numerous lines of evidence for concurrent orogenic events in the northern EGB (see e.g. review in Dobmeier and Raith, 2003), these results suggest considerable Neoproterozoic orogenic activity and either prolonged or repeated episodes of high-grade metamorphic conditions not only throughout the northern EGB, but also along the western EGB. Acknowledgments This study would not have been carried out without the encouragement and enthusiasm of Christoph Dobmeier who initiated this project and provided substantial support in the field and in the discussion of the data. Thanks to Saibal Gupta for his assistance in the field and his introduction into the regional geology of the Koraput area, to Monika Feth for the lab assistance at the FU Berlin, and to Svetoslav Georgiev for the assistance with the LA-ICP-MS analyses at ETHZ. KH wishes to express special thanks to Bhibuti Bhusan Roy for his invaluable help during fieldwork and to Rainer Wieler for his patience and support throughout the writing process of this manuscript. Furthermore, we thank Klaus Mezger for his constructive review. Appendix A. Supplementary data Supplementary data to this article can be found online at http://dx. doi.org/10.1016/j.gr.2015.02.021. References Aftalion, M., Bowes, D.R., Dash, B., Dempster, T.J., 1988. Late Proterozoic charnockites in Orissa, India: U–Pb and Rb–Sr isotopic study. Journal of Geology 96, 663–676. Aftalion, M., Bowes, D.R., Dash, B., Fallick, A.E., 2000. Late Pan-African thermal history in the Eastern Ghats terrane, India, from U–Pb and K–Ar isotopic study of the MidProterozoic Khariar alkali syenite, Orissa. Geological Survey of India Special Publication 57, 25–33. Amelin, Y., Lee, D.-C., Halliday, A.N., 2000. Early-middle Archaean crustal evolution deduced from Lu–Hf and U–Pb isotopic studies of single zircon crystals. Geochimica et Cosmochimica Acta 64, 4205–4225. Andersen, T., Sørensen, H., 1993. Crystallization and metasomatism of nepheline syenite xenoliths in quartz-bearing intrusive rocks in the Permian Oslo rift, SE Norway. Norsk Geologisk Tidsskrift 73, 250–266. Bea, F., Montero, P., Ortega, M., 2006. A LA-ICP-MS evaluation of Zr reservoirs in common crustal rocks: implications for Zr and Hf geochemistry, and zircon-forming processes. Canadian Minerals 44, 693–714. Belousova, E.A., Griffin, W.L., O'Reilly, S.Y., Fisher, N.I., 2002. Igneous zircon: trace element composition as an indicator of source rock type. Contributions to Mineralogy and Petrology 143, 602–622. Belousova, E.A., Griffin, W.L., O'Reilly, S.Y., 2006. Zircon crystal morphology, trace element signatures and Hf isotope composition as a tool for petrogenic modelling: examples from eastern Australian granitoids. Journal of Petrology 47, 329–353. Bingen, B., Austrheim, H., Whitehouse, M., 2001. Ilmenite as a source for zirconium during high-grade metamorphism? Textural evidence from the Caledonides of Western Norway and implications for zircon geochronology. Journal of Petrology 42, 355–375. Biswal, T.K., De Waele, B., Ahuja, H., 2007. Timing and dynamics of the juxtaposition of the Eastern Ghats Mobile Belt against the Bhandara Craton, India: a structural and zircon U–Pb SHRIMP study of the fold-thrust belt and associated nepheline syenite plutons. Tectonics 26. http://dx.doi.org/10.1029/2006TC002005. Bose, S., Dunkley, D.J., Dasgupta, S., Das, K., Arima, M., 2011. India–Antarctica–Australia– Laurentia connection in the Paleoproterozoic–Mesoproterozoic revisited: evidence from new zircon U–Pb and monazite chemical age data from the Eastern Ghats Belt, India. Geological Society of America Bulletin 123, 2031–2049. Bouvier, A., Vervoort, J.D., Patchett, P.J., 2008. The Lu–Hf and Sm–Nd isotopic composition of CHUR: constraints from unequilibrated chondrites and implications for the bulk composition of terrestrial planets. Earth and Planetary Science Letters 273, 48–57. Chauvel, C., Blichert-Toft, J., 2001. A hafnium isotope and trace element perspective on melting of the depleted mantle. Earth and Planetary Science Letters 190, 137–151. Corfu, F., Hanchar, J.M., Hoskin, P.W.O., Kinny, P., 2003. Atlas of zircon textures. In: Hanchar, J.M., Hoskin, P.W.O. (Eds.), Zircon. Reviews in Mineralogy and Geochemistry 53, pp. 469–500. Dasgupta, S., Sengupta, P., 2003. Indo-Antarctic correlation: a perspective from the Eastern Ghats Belt, India. In: Yoshida, M., Windley, B.F., Dasgupta, S. (Eds.), Proterozoic East Gondwana: Supercontinent Assembly and Breakup. Geological Society, London, Special Publications 206, pp. 131–143. Davidson, A., van Breemen, O., 1988. Baddeleyite-zircon relationships in coronitic metagabbro, Grenville Province, Ontario: implications for geochronology. Contributions to Mineralogy and Petrology 100, 291–299. Deer, W.A., Howie, R.A., Zussman, J., 2003. 2nd edn. Rock-forming Minerals, Sheet Silicates: Micas vol. 3A. Geol Soc, London.

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Please cite this article as: Hippe, K., et al., Zircon geochronology of the Koraput alkaline complex: Insights from combined geochemical and U–Pb– Hf isotope analyses, and implications..., Gondwana Research (2015), http://dx.doi.org/10.1016/j.gr.2015.02.021