Timing of juvenile arc crust formation and evolution in the Sapat Complex (Kohistan–Pakistan)

Chemical Geology 280 (2010) 243–256

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Chemical Geology j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / c h e m g e o

Timing of juvenile arc crust formation and evolution in the Sapat Complex (Kohistan–Pakistan) Pierre Bouilhol a,b,⁎, Urs Schaltegger c, Massimo Chiaradia c, Maria Ovtcharova c, Andreas Stracke a,1, Jean-Pierre Burg a, Hamid Dawood d a

Department of Earth Sciences, ETH Zurich, 8092 Zurich, Switzerland Department of Earth, Atmospheric and Planetary Sciences, MIT, 02139 Cambridge, USA Section des Sciences de la Terre et de l'environnement, Université de Genève, 1205 Genève, Switzerland d Pakistan Museum of Natural History, Garden Avenue-Shakarparian, 44 000 Islamabad, Pakistan b c

a r t i c l e

i n f o

Article history: Received 26 April 2010 Received in revised form 25 October 2010 Accepted 7 November 2010 Available online 12 November 2010 Editor: R.L. Rudnick Keywords: U/Pb on zircons Hf on zircons Arc evolution Crust formation

a b s t r a c t The combination of age determination and geochemical tracers allows understanding the source evolution during magmatism. We studied the Sapat Complex, in the exhumed Cretaceous Kohistan Paleo-Island Arc, to reconstruct the formation of the juvenile lower arc crust and the evolution of the mantle source during arc magmatism. High precision ID-TIMS U/Pb dating on zircon, shows that a protracted period of magmatic accretion formed the Sapat Complex between 105 and 99 Ma. Since continued melt percolation processes that formed the lower crust obscured the original bulk rock Nd–Pb–Sr isotopic composition, we rely on the Hf isotopic composition of zircons of different ages to unravel the source evolution. Nd and Pb bulk isotopic compositions coupled with Hf isotopic composition on zircons allow reconstructing a geodynamical scenario for the Sapat Complex, and the Cretaceous history of the Arc. We suggest that trenchward migration of the hot mantle source at 105 Ma explains the small heterogeneous εHf signal between + 14 and + 16. This heterogeneity vanished within ca. 2 million years, and the εHf of the source evolved from + 16 to + 14 at 99 Ma. Integrated to the Kohistan Cretaceous history, which has a baseline of εHf ≈ 14, these data pinpoint two geodynamical events, with slab retreat and the formation of the Sapat Complex followed by splitting of the Kohistan island arc at 85 Ma. © 2010 Elsevier B.V. All rights reserved.

1. Introduction The petrological and geochemical characteristics of arc magmas are direct consequences of mantle and crustal processes. The predominant factors controlling the influence of the source on arcmelt composition are the intrinsic composition of the mantle, its thermal state and the adiabatic vs flux melting process, together with the slab-derived fluid flux into the sub arc-mantle (e.g. Tatsumi et al., 1983; White and Patchett, 1984; McCulloch and Gamble, 1991; Salters and Hart, 1991; Hawkesworth et al., 1993; Arculus, 1994; Pearce et al., 1995; Elliott et al., 1997; Grove et al., 2002; Grove et al., 2006). Combining bulk rock trace elements and isotopic compositions allows disentangling the integrated contribution of the different components in magmatic arc rocks (e.g. White and Patchett, 1984; Woodhead

⁎ Corresponding author. Massachussets Institute of Technology, Dept EAPS 54-1020, 77 Massachusetts Avenue, 02139 Cambridge MA, USA. E-mail addresses: [email protected] (P. Bouilhol), [email protected] (U. Schaltegger), [email protected] (A. Stracke), [email protected] (J.-P. Burg). 1 Present address: Westfälische Wilhelms Universität Institut für Mineralogie, 48149 Münster, Germany. 0009-2541/$ – see front matter © 2010 Elsevier B.V. All rights reserved. doi:10.1016/j.chemgeo.2010.11.013

et al., 1998; Pearce et al., 1999; Woodhead et al., 2001). However, the temporal evolution of both mantle and slab contributions is poorly constrained, and very little is known about how the melt transport mechanism may influence melt chemistry. A juvenile arc crust (i.e. without continental crust being involved) is formed from mantle melting on top of an oceanic lithosphere during intra-oceanic subduction. In that case, the combination of major and trace elements together with isotopic data can map and elucidate the amount of slab-derived material added to mantle melts at the scale of a whole arc (e.g. Pearce et al., 2005; Singer et al., 2007). Combined with geochronology, this approach should allow tracing the evolution of the different components in space and time (e.g. Lee et al., 1995; Ewart et al., 1998; Straub, 2003; Rioux et al., 2007). Yet, the time-scale of magmatic processes and source evolution during the building of the juvenile lower crust of an arc has been rarely documented (e.g. Rioux et al., 2007; Peytcheva et al., 2008) while such investigations have mainly focused on young and active volcanic complexes (e.g. Hawkesworth et al., 2004). In order to unfold magmatic processes and source evolution in intra oceanic-arcs, isotopic tracers should ideally be linked to precise age determination. The Lu–Hf isotopic system is the first choice to understand such an evolution because Hf is highly compatible in

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zircon, which can be dated very precisely with U–Pb. Indeed, owing to typically very low 176Lu⁄177Hf isotopic ratios in zircon, present day Hf isotopic compositions in zircon remain virtually the same as those at the time of melt crystallization. Furthermore, analyzing and comparing zircons that have been assimilated from earlier stages of the same magmatic system (antecrystic) and newly formed (autocrystic) zircon crystals allows the temporal reconstruction of the source evolution within a single sample because the Hf isotopic system is very stable within the zircon crystal structure (Miller et al., 2007). Hence, U/Pb ages and Hf isotope measurements on the same zircon grain are formidable tools for tracking source evolution. In this paper, we intend to unfold the succession of magmatic intrusions within the Sapat Complex, which belongs to the Kohistan Paleo-Island Arc in NW Pakistan (Fig. 1). The Sapat Complex forms a single lower crustal entity mainly composed of metamorphosed sills of mafic rocks either intruded or co-genetic with kilometer-size cumulative mafic and ultramafic magmatic conduits (Bouilhol et al., 2009, 2010). This crustal sequence lies over ultramafic rocks representing the subjacent mantle portion of the complex (Bouilhol et al., 2009). To document the evolution of the mantle source composition during the building of this juvenile crust, high-precision single zircon ages and the corresponding Hf isotopic compositions were analyzed in combination with the major and trace element and Nd–Sr–Pb isotopic compositions of their respective bulk rocks. Moreover, combining these new data with available Hf isotopic compositions in the Kohistan arc leads to a better understanding of the long-lasting processes involved in the building and dynamics of the Arc. The results document a protracted suite of magmatic intrusions forming the lower crust, first creating the host plutonic

sequence followed by the formation of magmatic conduits. The data record the evolution from a refractory mantle during onset of melting to a less radiogenic value similar to the Kohistan mantle source baseline (εHf ≈ 14). We propose that this evolution mirrors the trenchward migration of isotherms, inducing melting of a forearc mantle affected by slab component, at the beginning of slab retreat. 2. Setting The Sapat Complex is situated in the southern boundary of the Kohistan Paleo-Island Arc (Tahirkheli et al., 1979; Bard, 1983) in Northern Pakistan (Fig. 1). This Paleo-Island Arc formed in Mesozoic times above the northward subduction zone of the Tethys (Yoshida et al., 1996; Zaman and Torii, 1999). Exposing a lithospheric section, from mantle rocks (to the south) to volcanics and sediments (to the north) Kohistan offers unique windows into magmatic processes at all structural levels of an island arc. Three lower crustal sections are documented in this paleo-island arc: The Jijal-Patan Complex, the Chilas Complex and the Sapat Complex. The Sapat Complex, object of this study, is a ca. 60 km long and 18 km wide, crescent-shaped unit bounded to the South by the Indus Suture, and separated from the socalled Southern Amphibolites to the North, by the Thoregah Shear Zone (Fig. 1). The complex is comprised of three main lithologies: Mantle rocks at its base (harzburgites-dunites and clinopyroxenites) (Bouilhol et al., 2009), overlain by a series of metaplutonic rocks (amphibolite-greenschist facies meta-gabbros-hornblendites– tonalites) which are intruded/cogenetic with kilometer-sized pyroxenitic bodies representing magmatic conduits (Bouilhol et al., 2010). The dated plagioclase-rich members are found in all of these units,

Fig. 1. The Sapat Complex maps and sample locations. Sample coordinates in Appendix.

P. Bouilhol et al. / Chemical Geology 280 (2010) 243–256

thus defining three main structural settings. From bottom to top: (i) patches within metagabbro veins in the mantle rocks (Bouilhol et al., 2009), (ii) zircon-devoid sills interlayered in, and zircon-bearing dykes cross cutting the crustal metagabbros and (iii) dykes in the pyroxenite bodies. The bulk and mineral chemistry of all Sapat lithologies suggests that they originated from the same parental primitive melt (Bouilhol et al., 2010). However, the complex structural settings require chronological constraints and isotopic compositions in order to consolidate petrological considerations that link all these rocks to a common magmatic period and to better constrain the nature of the parental melt. U–Pb age determinations show that all rocks crystallized between 105 and 99 Ma. These dates, coupled to bulk Sr, Nd, Pb isotopic compositions of whole-rocks, link the Sapat Complex to the Kohistan Arc system during its intra-oceanic history and allow reconstructing the evolution of the mantle source during that stage of magmatic activity. To complete our survey, we also dated the tonalitic Thoregah

245

laccolith, which is syn-to post-Thoregah shear zone, the northern boundary of the Sapat Complex (Fig. 1). The age consistently places an lower age limit of 93.58 ± 0.07 Ma for the Sapat Complex formation. 2.1. Dated samples All dated samples re-equilibrated in amphibolite to greenschist facies, but their original magmatic mineralogy has remained. Sample G13 (Fig. 2a) is from a meter-thick, north-dipping (~N070 20NW) leucocratic dyke that crosscuts the tectonic and magmatic foliations of fine-grained metaplutonic rocks. It is coarse-grained and mainly composed of quartz, saussuritized plagioclase with euhedral garnet inclusions and some chlorite. This dyke stops against a 2 m wide, top-to-the-north, ductile normal shear-zone (Fig. 2a). However, several digitations, partly foliated partly isotropic, intrude into the shear zone and the boundary between the dyke and the shear zone shows no fabric deviation. These features indicate that the G13 dyke

Fig. 2. Tonalitic dykes. a) North-dipping dykes feeding a ductile shear zone above the contact with the mantle-derived ultramafic rocks (sample G13: N35°01.206'; E73°44.335'). b) Plagioclase segregate (sample G12) between pegmatitic metagabbro and underlying finer grain metagabbros (N35°02.082'; E73°44.554'). c) Gently north-dipping dyke cutting clinopyroxenites in pipe 1 (P8, N35°01.021'; E73°42.154'). d) Tonalite dyke cross cutting pyroxenites in pipe 2 (P63: N35°04.086'; E73°45.780'). e) Segregate (sample D101: N35°01.751'; E73°45.031) in metagabbro of the crust–mantle transition where melt infiltration led to the formation of amalgamated metagabbros and pyroxenite (Bouilhol et al., 2009); arrows point to assimilated metagabbros pieces. f) View on the contact between the Sapat Complex and the Thoregah Laccolith (taken from N35°11.080'; E73°58.460'). Inset: close up of K08-18.

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was contemporaneous with late movements on the normal shear zone. Sample G12, (Fig. 2b) appears as a segregated product from the pegmatitic metagabbros. It is mainly composed of albitic saussuritized plagioclase and quartz, with little amounts of tremolite, calcite and epidote. Within the biggest pyroxenite body (pipe 1, Fig. 1) hornblende-bearing tonalite veins dipping gently to the North cross cut the magmatic structures and increase in abundance with elevation. One of these veins, sample P8 (Fig. 2c) is made of saussuritized plagioclase with albite rims, magmatic hornblende partly altered to tremolite and Mg-rich chlorite; sphene, Ti-oxides, apatite and zircons are accessories. Hornblende tonalites are cumulates that formed from residual melt fraction produced by the reactional fractional crystallization processes within the pyroxenite pipes (Fig. 1; Bouilhol et al., 2010). Sample P63 (Fig. 2d) is the least altered of the observed hornblende tonalite veins. To maximize the chance of finding zircons, the sample was taken from the most leucocratic part of a tonalite dyke. It is mostly composed of plagioclase and quartz crystallized at the eutectic, hornblende, biotite and magmatic epidote. Sample D101 (Fig. 2e) is found in metagabbro lenses in the crust– mantle transition zone and is interpreted as a products segregated from the host gabbro, which is also enclaved in the sampled vein (Fig. 2e). Sample D101 contains cm-sized magmatic hornblende, saussuritized and partly albitized plagioclase, quartz, ±Mg-rich chlorite. The modal composition of the Thoregah laccolith (Fig. 2f) ranges from granodiorite to tonalite. This laccolith is thus considered to be emplaced through accretion of multiple magma batches of variable modal composition. Sample KO8-18 has a tonalite composition and represents one of the most leucocratic parts of the laccolith (Fig. 2f, inset). The mineral assemblage includes saussuritized plagioclase, quartz, hornblende and biotite, sphene and Ti-oxides. Except for KO8-18, all the dated samples contain a very small amount of zircon. 3. Analytical techniques 3.1. Bulk rock chemistry Samples were crushed with a hydraulic press in a stainless-steel container, and ground to powder with an agate mill. Major element analyses were performed on fused glass-beads prepared from rock powder mixed with Lithium-Tetraborate on a Panalytical Axios wavelength dispersive XRF spectrometer (WDXRF, 2.4 KV) at ETH Zurich. The calibration of the major elements is based on ca. 30 certified international standards with emphasis on igneous and metamorphic rock compositions are presented in the Appendix, together with accuracy and detection limits. Minor and trace elements in whole rocks were analyzed by solution-ICP-MS (except for KO8-18). We used the HF–HClO4–HNO3 digestion procedure described by Ionov et al. (1992) for sample dissolution. The analyses were performed on an Element XR high-resolution ICP-MS at Géosciences Montpellier (France). Concentrations were determined by external calibration for most elements except Nb and Ta, which were calibrated by using Zr and Hf, respectively, as internal standards for the determination of Nb by spark-source mass spectrometry (Jochum et al., 1990). This method avoids memory effects due to the introduction of concentrated Nb–Ta solutions in the instrument. Detection limits obtained by long-term analyses of chemical blanks can be found in Ionov et al. (1992) and Godard et al. (2000). Blanks, and international standards (BEN, UBN and PCC1) have been analyzed between unknowns, (Appendix). Precision and accuracy can be found in Godard et al. (2008) and Godard et al. (2009). The trace elements concentration for sample KO8-18 and Zr and Hf abundances for the other samples, were determined on freshly broken cross-sections of the glass discs at ETH Zurich by laser ablation-inductively coupled plasma-mass spectrom-

etry (LA-ICP-MS) using an Excimer 193 nm laser system with a 60 μm laser beam diameter and 10 Hz repetition rate. The NIST 610 glass was used as an external calibration standard, and Al2O3 and CaO (42Ca) concentrations were used as internal standards and for crosschecking. Further details can be found in Gunther et al. (2001) and Georgiev et al. (2009). Results in Table 1. Rocks were also analyzed for Pb, Sr and Nd isotopic compositions at Geneva University. Between 100 and 150 mg of powdered rock fractions (b70 μm) were dissolved on a hot plate at 140 °C in closed Teflon vials during 7 days in a mixture of 4 ml concentrated HF and 1 ml 15 M HNO3. The sample was then dried on a hot plate, and redissolved in 3 ml of 15 M HNO3 in closed Teflon vials at 140 °C and dried again. Sr, Nd and Pb separation was carried out using cascade columns with Sr-spec, TRU-spec and Ln-spec resins following a method modified from Pin et al. (1994). Pb was further purified with an AG-MP1-M anion exchange resin in hydrobromic medium. Pb, Sr and Nd isotope ratios were measured on a Thermo TRITON mass spectrometer on Faraday cups in static mode. Pb was loaded on Re filaments using the silica gel technique and all samples and standards were measured at a pyrometer controlled temperature of 1220 °C. Pb isotope ratios were corrected for instrumental fractionation by a Table 1 Majors and trace elements compositions of the dated samples. *: Ga and Sc measured by XRF. †: KO8-18 trace elements from ablated glass pills, and Zr Hf from other samples. Sample

P63

D101

G13

P8

G12

K08-18†

wt.% SiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5 LOI Total ppm Ga* Sc* Ti Zn Cu Ni Co Cr V Cs Rb Ba Th U Nb Ta La Ce Pb Pr Sr Nd Sm † Zr † Hf Eu Gd Tb Dy Y Ho Er Ti Tm Yb Lu

79.51 12.13 0.175 0.005 0.05 3.63 3.47 0.02 0.01 0.54 99.59 10.6 3.1 374 4.0 5.0 0.6 2.0 9.4 4.0 0.222 0.219 10.85 0.019 0.004 0.040 0.005 0.209 0.365 1.439 0.037 147 0.177 0.063 0.433 0.051 0.230 0.058 0.010 0.076 0.465 0.017 0.046 374 0.010 0.048 0.008

64.01 20.77 0.303 0.01 0.13 5.61 7.15 0.02 0.01 0.76 98.80 17.9 3.7 66 2.8 17.1 3.9 3.6 5.2 0.0 0.037 0.143 17.95 0.017 0.003 0.018 0.004 0.621 1.398 0.602 0.169 126 0.790 0.167 3.01 0.128 0.646 0.146 0.020 0.128 0.718 0.028 0.077 66 0.013 0.082 0.013

75.80 13.07 0.930 0.04 0.37 2.94 4.72 0.02 0.02 0.68 98.69 13.4 5 532 10.3 117.5 2.8 3.8 3.3 8.3 0.065 0.101 15.88 0.019 0.004 0.290 0.021 0.892 2.083 1.173 0.276 58 1.33 0.300 46.9 1.345 0.509 0.283 0.052 0.395 2.95 0.095 0.343 532 0.063 0.536 0.101

62.64 20.11 1.32 0.06 1.45 4.13 9.06 0.05 0.02 0.69 99.63 12.4 3.9 605 40.4 3.41 7.68 4.09 5.95 3.06 2.39 2.07 72.23 0.010 0.004 0.308 0.018 0.551 1.636 1.269 0.293 199 1.91 0.815 3.42 0.209 0.765 1.39 0.257 1.83 9.74 0.391 1.09 605 0.162 1.04 0.174

63.63 19.30 1.15 0.02 0.69 3.66 9.66 0.04 0.01 1.13 99.39 2.1 0.1 619 23.3 22.4 19.1 4.3 3.9 12.5 0.090 0.210 9.65 0.004 0.002 0.062 0.003 0.329 0.730 0.087 0.097 41 0.494 0.120 2.07 0.021 0.550 0.163 0.028 0.193 1.14 0.042 0.120 619 0.019 0.127 0.021

62.21 16.27 6.01 0.12 2.20 6.40 3.26 0.79 0.17 1.87 99.94 18.6 13.3 3432 76.2 13.5 13.8 16.3 9.1 110.7 0.738 17.41 251.6 1.47 0.262 2.26 0.085 10.7 26.4 2.661 3.470 417 15.48 3.39 142.9 4.08 0.973 3.03 0.510 2.79 16.16 0.594 1.92 3432 0.249 1.70 0.261

P. Bouilhol et al. / Chemical Geology 280 (2010) 243–256

247

Table 2 Isotopic compositions of the different lithologies of the Sapat Complex. Sample description and location in the Appendix. 2σ

εNd (i)

87

2σ

206

0.000008 0.000006

5.59 7.88

0.703376 0.703210

0.000004 0.000002

Mantle Ol-clinopyroxenite D64 0.513118 0.000009 D18 0.513108 0.000018 LD5 0.513061 0.000025

7.04 4.45 4.82

0.703804 0.704239 0.703746

Crustal metagabbros G36 0.513137 G49 0.513104 G59 0.513204 G11 0.513071 G54 0.513025

0.000004 0.000004 0.000010 0.000009 0.000006

7.18 6.45 6.89 7.55 6.15

Interlayered tonalites G14 0.513040 G24 0.513008 G32 0.513050

0.000008 0.000013 0.000007

Pipes-Pyroxenites P7 – P9 0.513072 P18 0.513082 P42 0.513049 P31 0.512954 P51 0.51 Dated samples P63 0.512923 D101 0.512892 G13 0.512930 P8 0.513047 G12 0.512879

Sample

143

Nd/144Nd

Mantle metagabbros LD10 0.513011 D104 0.513150

Sr/86Sr

Pb/204Pb

2σ

207

18.382 18.497

0.013 0.007

0.000019 0.000012 0.000076

18.590 18.668 18.439

0.704350 0.704222 0.704195 0.703485 0.704445

0.000003 0.000002 0.000004 0.000004 0.000003

7.56 5.94 7.30

0.704275 0.704326 0.704306

– 0.000023 0.000070 0.000006 0.000008 0.000008

– 5.59 5.81 5.84 5.16 6.26

0.000042 0.000010 0.000004 0.000016 0.000008

5.35 5.84 6.48 7.21 5.34

Pb/204Pb

2σ

208

15.636 15.638

0.011 0.006

0.006 0.018 0.014

15.633 15.614 15.592

18.471 18.654 18.618 18.362 18.697

0.031 0.012 0.015 0.007 0.007

0.000008 0.000006 0.000008

18.462 18.450 18.738

0.705646 0.705341 0.705591 0.704885 0.704389 0.704548

0.000012 0.000010 0.000026 0.000010 0.000008 0.000014

0.705032 0.704161 0.704336 0.704426 0.703744

0.000006 0.000020 0.000008 0.000004 0.000006

Pb/204Pb

2σ

206/204i

207/204i

208/204i

38.291 38.544

0.026 0.014

18.327 18.485

15.633 15.638

38.272 38.532

0.005 0.015 0.011

38.731 38.805 38.452

0.012 0.038 0.028

18.582 18.662 18.434

15.633 15.614 15.592

38.708 38.801 38.443

15.636 15.622 15.607 15.564 15.633

0.025 0.011 0.012 0.006 0.006

38.606 38.483 38.743 38.400 38.806

0.064 0.025 0.032 0.014 0.014

18.348 18.627 18.579 18.358 18.685

15.630 15.621 15.605 15.563 15.632

38.587 38.461 38.723 38.396 38.801

0.008 0.013 0.002

15.605 15.629 15.627

0.007 0.011 0.002

38.530 38.561 38.960

0.016 0.028 0.005

18.412 18.433 18.735

15.602 15.629 15.627

38.499 38.541 38.958

18.489 18.565 18.480 18.487 18.528 –

0.015 0.009 0.007 0.012 0.006 –

15.596 15.629 15.623 15.636 15.626 –

0.012 0.007 0.006 0.010 0.005 –

38.511 38.679 38.557 38.611 38.683 –

0.031 0.016 0.015 0.024 0.012 –

18.443 18.547 18.467 18.472 18.516 –

15.593 15.628 15.623 15.635 15.626 –

38.484 38.659 38.550 38.596 38.678 –

18.652 18.676 18.476 18.537 18.361

0.003 0.009 0.003 0.002 0.012

15.615 15.631 15.615 15.630 15.597

0.002 0.007 0.002 0.001 0.010

38.834 38.856 38.626 38.727 38.459

0.006 0.019 0.005 0.003 0.024

18.650 18.671 18.473 18.534 18.332

15.614 15.630 15.615 15.630 15.596

38.829 38.847 38.620 38.725 38.445

factor of 0.07% per amu based on more than 90 measurements of the SRM981 standard and using the standard values of Todt et al. (1996). External reproducibility (2σ) of the standard ratios are 0.05% for 206 Pb/204Pb, 0.08% for 207Pb/204Pb and 0.10% for 208Pb/204Pb. Sr was loaded on single Re filaments with a Ta oxide solution and measured at a pyrometer-controlled temperature of 1490 °C. 87Sr/86Sr values were internally corrected for fractionation using a 88Sr/86Sr value of 8.375209. Raw values were further corrected for external fractionation by a value of +0.03‰, determined by repeated measurements of the SRM987 standard (87Sr/86Sr=0.710250). External reproducibility (1σ) of the SRM987 standard is b7 ppm. Nd was loaded on double filaments with 1 M HNO3. 143Nd/144Nd values were internally corrected for fractionation using a 146Nd/144Nd value of 0.7219 and the 144Sm interference on 144Nd was monitored on the mass 147Sm and corrected by

using a 144Sm/147Sm value of 0.206700. External reproducibility (1σ) of the JNdi-1 standard (143Nd/144Nd=0.512115±7 Tanaka et al., 2000) is 4 ppm. Nd, Sr and Pb initial values are calculated for 100 Ma using the Sm, Nd, Rb, Sr, U, Th, and Pb concentrations obtained from ICPMS measurements. Results in Table 2. 3.2. U/Pb and Hf on zircons Rock samples have been crushed using a disk mill and sieved to retain the fraction below 250 μm. Heavy minerals were then separated using a Wilfley table, methylene iodide and a Frantz magnetic separator. The most euhedral and clear zircons without inclusions and cracks were hand-picked under a binocular microscope from the non-magnetic fraction at N1.5 A. Zircons were rare in the

Fig. 3. Cathodoluminescence image representative of dated grains for sample G13, P63 and K08-18.

248 Table 3 Zircon Pb and Hf isotopic compositions. a) calculated on the basis of radiogenic 208Pb/206Pb ratios, assuming concordance, b) corrected for fractionation and spike, c) corrected for fractionation, spike, blank and common lead (Stacey & Kramers, 1975) d) corrected for Th/U disequilibrium accepting a Th/U = 4 for the source. No

weigh

Concentration

(mg)

U (ppm)

Th/Ua Pb com. (pg)

Atomic ratios rad/com

206/204b

Ages

corr.

207/235c

error 2σ (%)

206/238cd

error 2σ (%)

207/206cd

error 2σ (%)

206/238d

207/235

207/206d

coeff.

176Hf/177Hf (meas.)

± 2σ

176Hf/177Hf (T)

εHf (T)

0.283142 0.283129

8 8

0.283141 0.283128

15.4 15.0

6 10 5 5 6 9

0.40 0.57 0.30 0.33 0.20 0.34

1.38 2.09 1.22 1.19 1.50 1.88

0.00 0.13 0.08 0.01 0.12 0.14

0.2 0.3 0.2 0.2 0.7 0.5

33 38 36 32 69 54

0.13398 0.11436 0.11566 0.11578 0.10671 0.10784

13.70 8.40 8.96 12.78 4.53 4.49

0.01547 0.01598 0.01574 0.01544 0.01611 0.01624

1.01 0.61 0.76 0.91 0.37 0.38

0.06283 0.05192 0.05331 0.05440 0.04805 0.04817

12.91 7.90 8.40 12.05 4.32 4.27

98.94 102.18 100.65 98.75 103.00 103.83

127.66 109.94 111.13 111.24 102.95 103.99

702.48 281.91 342.02 387.63 101.71 107.62

0.80 0.81 0.75 0.81 0.61 0.61

17 8 22 6 6

0.29 0.15 0.39 0.12 0.17

0.75 0.69 0.82 0.72 0.93

0.65 0.94 0.92 1.21 1.22

1.7 0.6 1.5 1.2 0.6

123 52 107 85 50

0.11446 0.09690 0.10541 0.09545 0.10672

4.68 16.80 5.99 8.99 12.13

0.01560 0.01546 0.01566 0.01546 0.01613

0.28 0.53 0.28 0.36 0.40

0.05320 0.04545 0.04881 0.04477 0.04798

4.55 16.45 5.85 8.78 11.93

99.81 98.92 100.19 98.91 103.16

110.04 93.91 101.76 92.56 102.96

337.35 0.67 138.74

0.49

0.283125

51

0.283124

14.8

98.34

0.52 0.60 0.51

0.283133 0.283111 0.283146

14 4 13

0.283132 0.283110 0.283145

15.1 14.3 15.6

24 9 16 11 16 8

0.77 0.46 0.73 0.52 0.44 0.14

0.49 0.90 0.55 0.84 0.58 0.53

1.66 2.59 3.14 1.95 2.42 0.53

1.8 0.8 1.6 0.9 0.3 0.6

103 50 75 58 32 60

0.11186 0.11224 0.11302 0.12131 0.12234 0.11653

3.84 11.39 5.05 7.70 45.24 13.82

0.01672 0.01678 0.01667 0.01646 0.01610 0.01625

0.31 0.64 0.39 0.53 1.98 0.60

0.04852 0.04853 0.04917 0.05345 0.05510 0.05202

3.66 10.86 4.81 7.29 43.65 13.37

106.90 107.24 106.59 105.24 102.98 103.89

107.67 108.01 108.73 116.26 117.19 111.93

124.74 124.97 155.77 348.15 416.26 286.47

0.60 0.84 0.64 0.79 0.81 0.75

0.283119

11

0.283118

14.6

0.283150

10

0.283149

15.7

14 42 15

0.30 1.89 1.32

1.52 2.74 2.09

0.22 0.21 0.25

0.4 0.2 0.1

45 34 26

0.10900 0.10834 0.10875

15.89 15.49 34.22

0.01640 0.01621 0.01651

1.11 1.14 2.31

0.04822 0.04849 0.04778

14.89 14.48 32.15

104.84 103.63 105.53

105.05 104.45 104.82

110.07 123.29 88.38

0.91 0.89 0.90

0.283132 0.283163

17 14

0.283131 0.283162

15.1 16.2

30 64

0.47 1.08

0.63 0.62

0.24 0.41

1.5 4.7

115 306

0.11106 0.10888

4.48 1.55

0.01645 0.01639

0.34 0.13

0.04898 0.04819

4.20 1.45

105.15 104.77

106.96 104.94

146.86 108.64

0.84 0.79

0.283155

19

0.283154

15.9

69 128 108 94 79

1.092 1.996 1.886 1.422 1.241

0.43 0.21 0.29 0.34 0.36

22.4 8.6 4.6 30 15.2

1393 582 311 1899 971

0.09829 0.09712 0.09708 0.09797 0.09882

0.41 0.74 0.46 0.33 0.63

0.01484 0.01462 0.01462 0.01478 0.01485

0.08 0.11 0.12 0.07 0.10

0.04802 0.04817 0.04814 0.04807 0.04826

0.37 0.69 0.42 0.29 0.59

94.99 93.57 93.59 94.60 95.03

95.20 94.11 94.08 94.90 95.69

100.45 107.78 106.31 102.56 112.19

0.60 0.49 0.48 0.58 0.48

0.283095

4

0.283094

13.2

1.7 2.6 16.4 1.1 0.6

P. Bouilhol et al. / Chemical Geology 280 (2010) 243–256

G13 1 0.0032 2 0.0040 3 0.0038 4 0.0034 5 0.0112 6 0.0073 P63 1 0.0044 2 0.0027 3 0.0032 4 0.0072 5 0.0044 D101 1 0.0016 2 0.0029 3 0.0018 4 0.0029 5 0.0004 6 0.0025 G12 1 0.0028 2 0.0020 3 0.0010 P8 1 0.0020 2 0.0027 KO8-18 1 0.0370 2 0.0125 3 0.0484 4 0.0244 5 0.0082

Pb (ppm)

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dated rocks, and only a tenth of them were suitable for isotopic analysis. Needing all selected zircons, they have not been imaged by cathodoluminescence prior to analyses. Nevertheless, Fig. 3 contains a representative subset of cathodoluminescence images from dated grains, showing a weak magmatic zoning for sample G13 and P63, whereas a strong magmatic sector zoning is observed in KO8-18. U–Pb ages on zircon have been obtained at the Department of Mineralogy of the University of Geneva on a Thermo Triton Thermal Ionization Mass Spectrometer. The analytical techniques of U–Pb dating closely follow those outlined in Schaltegger et al. (2008). Prior to analysis, the zircons were treated by annealing-leaching (chemical abrasion), minimizing the effects of post-crystallization lead loss (Mattinson, 2005). The samples were partly analyzed using a 205Pb– 235 U tracer (ETH-55: samples G13, P63, D101), partly the EARTHTIME 205 Pb–233U–235U tracer (ET535: samples G12, P8, K08-18). Results are given in Table 3 and blank compositions are given in the Appendix, as well as the blank amount used to correct the data. Whenever the amount of common Pb equalled the amount of blank Pb, no correction was applied, whereas the model of Stacey and Kramers (1975) was used to correct the common Pb contribution whenever the common Pb exceeded the blank contribution. Concordia plots were drawn using ISOPLOT (Ludwig, 2003). Mean ages are given at the 95% confidence level. The Hf fraction was isolated using EichromTM Ln-spec resin, and measured in static mode on a NuPlasmaTM multi-collector ICP-MS at ETH Zürich using a CetacTM Aridus sample introduction system. The 176 Lu/177Hf ratios of the analyzed zircons were not determined but, instead, 176Hf/177Hf ratios were corrected for a value of 176Lu/177Hf in zircon of 0.005 (which has been used in previous publications); the correction stayed within limits of analytical precision of the measured 176Hf/177Hf ratios in all cases and are therefore insignificant. Initial 176Hf/177Hf values are given together with the zircon ages in Table 3. The Hf isotopic ratios were corrected for mass fractionation using a 179Hf/177Hf value of 0.7325 and normalized to 176Hf/177Hf of 0.282160 of the JMC-475 standard (BlichertToft et al., 1997). Errors of the measured 176Hf/177Hf ratios are given as individual 2σ errors. εHf values were calculated with (176Hf/177Hf CHUR(0) = 0.282772; BlichertToft and Albarede, 1997). 4. Results 4.1. Bulk rocks 4.1.1. Mantle and crustal rocks Samples from the mantle and crustal lithologies have been analyzed for bulk rock isotopic compositions (Table 2; more detailed sample description and composition in (Bouilhol et al., 2009, 2010). Overall, isotopic compositions of mantle clinopyroxenite, crustal pyroxenite, gabbros, and tonalites are scattered and no specific distinction between different lithologies is obvious. The initial isotopic composition of the whole complex ranges from 0.512737 to 0.512913 for 143Nd/144Nd(i), from 0.703200 to 0.705563 for 87Sr/86Sr(i), from 38.272 to 38.139 for 208Pb/204Pb(i), from 18.327 to 18.885 for 206Pb/ 204 Pb(i), and from 15.563 to 15.700 for 207Pb/204Pb(i). 4.1.2. Dated samples Samples G12, G13, P63 and D101 show variable silica content ranging from 63 to 79 SiO2 wt.% (Table 1), together with low MgO (0.05–1.45 wt.%), TiO2 (0.04–0.10 wt.%), and MnO (b0.06 wt.%) contents. They have high Al2O3 (12.13–20.11 wt.%), Na2O (0.38– 9.66 wt.%) and CaO contents ranging from 2.95 to 5.6 wt.%. Hence these rocks are plagioclase-dominated consistent with their segregated aspects. P8 and KO8-18 are plagioclase-poorer and contain more FeO (1.3, 6.0 wt.%) and MgO (1.4, 2.2 wt.%). Three groups of rocks can be distinguished based on REE chondrite normalized patterns (Fig. 4 a and b). Sample P8, has a flat HREE segment, a

Fig. 4. Trace elements composition of dated samples. (a) Chondrite normalized REE diagram. b) Primitive mantle normalized diagrams. Normalization values are from (Sun and McDonough, 1989). Grey field represents crustal metagabbros composition from Bouilhol et al. (2010).

negatively fractionated LREE segment and a small positive Eu anomaly. It differs from the other samples, which have a plagioclase-like REE pattern, with a pronounced Eu positive anomaly, a rather flat M-HREE segment and a slightly positively fractionated LREE segment. The fractionated HREE segment of G13 is consistent with the presence of garnet. KO8-18, from the Thoregah pluton is strikingly different, with a flat HREE segment and a M-LREE segment positively fractionated, reaching the highest concentration in LREE at about 50 ppm for La. Normalized to primitive mantle, KO8-18 is again clearly different from the other samples, with a small Ti negative anomaly, a clear Nb–Ta depletion and a positive Zr–Hf anomaly due to the presence of zircon (Fig. 4b). Additionally, this sample shows the highest enrichment in the most incompatible elements and has a slight Pb and Sr positive anomaly. The dykes show similar LILE enrichment, but clear HFSE differences. Sr, Pb and Ba show positive anomalies. Ti shows either a positive or a negative anomaly. Nb and Ta have different concentrations. Either Nb and Ta are depleted and show a chondritic to infrachodritic ratio (5.26 b Nb/Ta b 18.36; samples P63, D101, G12), or Nb is not depleted, and Nb–Ta ratios are closer to chondrite to suprachondritic (Nb/Ta = 13.52 and 17.36; G13 and P8 respectively). From these features, trace elements clearly reflect the

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cumulative origin of these rocks, P8 and KO8-18 being the plagioclasepoorer samples. Except for sample P63, which has the highest 87Sr/86Sr(i) (0.705026) for a low 143Nd/144Nd(i) (0.512784), the analysed tonalites show an increase in Ndi with increasing Sri (from 0.703723–0.512779 to 0.704381–0.512870, respectively; Table 2). G12 shows the lowest 87 Sr/86Sr(i) value (0.703722) whereas P8, D101 and G13 have similar 87 Sr/86Sr(i) (~0.7043). Sample P8 shows the highest 143Nd/144Nd(i), whereas G13 is intermediate (0.512842) and P63, G12, and D101 are similar (~0.512790). For all the samples, 208Pb/204Pb(i) and 206Pb/ 204 Pb(i) have similar values, with distinct ranges for P63 and D101 (208Pb/204Pb(i) ~ 38.83; 206Pb/204Pb(i) ~ 18.65), for P8 and G12 (208Pb/ 204 Pb(i) ~ 38.76; 206Pb/204Pb(i) ~ 18.50), and with the lowest values for G12 (208Pb/204Pb(i) = 38.45; 206Pb/204Pb(i) = 18.33). Samples P8 and D101 have the same 207Pb/204Pb values (15.63), P63 and G13 overlap at ~ 15.61, whereas G12 is possibly slightly lower (15.60). Due to very low concentrations in the elements of interest, the samples have not been leached prior to isotopic analysis. This implies that the isotopic ratios may have been affected by fluid percolation for the most mobile elements. Indeed, the Sr isotopic composition of the pyroxenites, which have the lowest Sr concentration, shows an inverse correlation with the Sr content. This simple relationship may witness metamorphic re-equilibration, so that 87Sr/86Sr above ≈ 0.7045 may not represent pristine composition. Nevertheless, there are no correlations between trace element concentration and Nd or Pb isotopic compositions, which indicate that our Nd and Pb isotopic dataset can be used to infer source composition. The isotopic compositions of the rocks are representative of a subduction zone environment with the contribution of a slab component (melt, fluid, or supercritical fluid) to the mantle wedge. More importantly, the measured range indicates that the source of the Complex implied different amounts of slab and mantle components. Compared to the two other Kohistan lower crustal sections, the Sapat Complex data are significantly more radiogenic in 143Nd/ 144 Nd(i) at more variable Pb isotopic compositions than the Chilas Complex (Bignold et al., 2006). They are mostly overlapping with the data points of the Jijal lower crust, but extend to higher 143Nd/ 144 Nd(i) and 206Pb/204Pb(i) ratios as well (Fig. 5, (Khan et al., 1997; Dhuime et al., 2007, 2009). Compared to modern island arc Basalts (e.g. Mariana and Aleutians), Kohistan lithologies generally show lower 143Nd/144Nd and seem to be derived from a mantle source that

was already enriched compared to present-day DMM, prior to slab addition. A Kohistan Primitive Mantle source can be estimated considering the isotopic composition of the most primitive lavas found in Northern Kohistan (Chalt volcanics), the most primitive value found in the Jijal Mantle rocks and, most of all, considering that the Kohistan Arc has started its activity at 150 Ma on top of an Indian lithosphere the mantle of which was significantly more enriched than DMM, (e.g. Mahoney et al., 1998; Xu and Castillo, 2004). If we consider this Primitive Mantle source, Kohistan lithologies have a rather primitive overall composition consistent with an intraoceanic setting. 4.2. U/Pb ages and Hf isotope analyses on zircon 4.2.1. U/Pb on zircons U–Pb analyses of single zircon grains show overall general range of 206 Pb/238U dates spanning from 106 to 98 Ma, which is considered representative of the duration of the Sapat Complex formation (Table 3 and Fig. 6). Although most of the analyzed grains are concordant, some of them are slightly off the Concordia at higher 207 Pb/235U ratios. Since 207Pb/235U is extremely sensitive to common and blank Pb correction, and considering the very low radiogenic Pb– common Pb ratio (Pbrad/Pbcom) of the non-concordant grains, we consider the elevated 207Pb/235U ratios as due to inaccurate common Pb correction. Ages are discussed on the 206Pb/238U dates. 4.2.1.1. G13 (Fig. 6a). This sample shows the lowest levels of U and radiogenic Pb (Pbrad.) of all analyzed zircons and, in addition, contains a component of common Pb. Therefore, the analyzed grains show very low Pbrad/Pbcom (0.2–0.7) despite the fact that the amount of blank was low (0.3–0.5 pg). Nevertheless, with the exception of grain 1, all analyzed grains are analytically concordant. Analyses 1 to 6 show a large spread in their 206Pb/238U dates from 103.83 to 98.94 Ma. This trend is continuous, and is very likely biased by minute molecular interferences at masses 204, 205 and/or 207 at very low count rates (as a consequence of smallest radiogenic Pb quantities). These results may also indicate recrystallisation processes, a possibility that might be reflected in Th/U ratios decreasing with age. Moreover , G13 grains have the singularity to show oval-shaped pores that may also attest for re-crystallization (Fig. 3). We assume that sample G13 has a complex crystallization history that will be discussed further.

Fig. 5. εNd(i) vs 206Pb/204Pb (i) of the Kohistan lithologies. Tethyan MORB data are from Mahoney et al. (1998), Kohistan data from Dhuime et al. (2007, 2009) for the Jijal Complex lower crustal lithologies (southern lower crust), and from Khan et al. (1997) and Bignold et al. (2006) for volcanics and Chilas Complex. Mariana and Aleutians present-day island arc volcanics are taken from Georock (n = 389): Grey contoured scale approximates sample distribution. Indian Ocean Sediments from Ben Othman et al. (1989). See text for the approximation of the Kohistan Mantle source.

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Fig. 6. U–Pb Concordia diagrams with age determinations with εHf of analysed zircon fractions and radiogenic–common lead ratios of zircon fractions. Details in Table 3. Ellipses denote uncertainty at the 2 sigma level of individual analyses; mean ages are at the 95% confidence level. The grey band reflects the uncertainty on the Concordia position (U decay uncertainty) a) Sample G13, b) Sample P63, c) Sample D101, d) Sample G12, e) Sample P8, f) Sample KO8-18.

4.2.1.2. P63 (Fig. 6b). Sample P63 zircons yielded a range of 206Pb/238U dates similar to those obtained for G13. However, in contrast to G13, P63 zircons are clean, homogeneous grains exempt of any re-crystallisation feature, indicating a pristine magmatic origin (Fig. 3). The analyzed grains have higher U (6–22 ppm) and lower Pbcom than G13, with higher Pbrad/Pbcom (0.6–1.7). Given the low amount of blank, no initial Pbcom was considered. The oldest concordant grain (#6) with a 206 Pb/238U date of 103.16 ± 0.41 Ma stands apart from the other

grains closer to 99–100 Ma. The youngest grains (#2 and #4) are overlapping at a 206Pb/238U mean date at 98.91 ± 0.29 Ma, which could be interpreted as the crystallization age of the sample. The slightly higher 206Pb/238U date of grains #3 and #1 may represent zircons crystallized early among the analysed grains, or earlier grains. 4.2.1.3. D101(Fig. 6c). The analysed zircons of D101 show similarities with P63 zircons. U and Pb contents are low, giving low Pbrad/Pbcom

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(0.3–1.8). The youngest grains #5 and #6 yield a mean 206Pb/238U date at 103.81 ± 0.58 Ma. Once again, this age could represent the crystallization of the sample. On the other hand, the concordant cluster of grains #1, 2 and 3, with higher Pbrad/Pbcom ratios, may indicate emplacement at 106.83 ± 0.24 Ma. In this case, the younger cluster would be difficult to interpret, since one does not expect Pb loss in so low-U zircons. 4.2.1.4. G12 (Fig. 6d). Even though the analyzed zircons have considerable concentrations of U (14–42 ppm), the high Pbcom concentrations rendered the Pbrad/Pbcom very low (0.1–0.4) and made an initial Pbcom correction necessary. The three analysed grains show overlapping errors, and yield a mean 206Pb/238U date of 104.38 ± 0.78 Ma. 4.2.1.5. P8 (Fig. 6e). Only two zircon grains have been analysed, the younger grain (#2) has the highest Pbrad/Pbcom (4.7) of all zircons analysed in this study. The two grains are concordant with 206Pb/238U ages of 105.15 ± 0.35 and 104.77 ± 0.13 Ma. 4.2.1.6. K08-18 (Fig. 6f). With high U contents (69–128 ppm) and despite high Pbcom (1.7–16.4 pg), the analysed zircon grains show much higher Pbrad/Pbcom (4.6–30) than the previous samples. Two groups of data can be distinguished. They are interpreted as representing protracted growth over 1.5 Ma or as mixing of two growth episodes: One for the oldest zircons (#1 and 6) with a mean 206 Pb/238U date of 95 ± 0.06 Ma and the last crystallization with the two youngest zircons at 93.58 ± 0.07 Ma. These data demonstrate that the Thoregah laccolith is younger than the Sapat Complex, in agreement with field relationships, and shows a complex history involving several magmatic episodes (with melt batches sampling previously crystallised antecrystic zircon). In summary, the oldest zircons (106.83 ± 0.24 Ma) are those found in segregation patches that crystallized in gabbro lenses within the mantle peridotites. The youngest grains found in these patches overlap with the ages of samples G12 and P8. The younger P63 grains, crystallized at ≈99 Ma, have similar ages as D101, G12 and P8. G13 with a more complicated age pattern is comprised within the previous range. We consequently postulate that the Sapat Complex formed between ~106 and 99 Ma. This is consistent with the upper age bound provided by the discordant, 93.5 Ma Thoregah laccolith (KO8-18). 4.2.2. Hf results Initial 176Hf/177Hf ratios of some of the dated zircons are given in Table 3. Owing to the typically very low 176Lu/177Hf values in zircon (Kinny and Maas, 2003), measured 176Hf/177Hf values are usually considered to represent initial Hf isotopic compositions at the time of zircon crystallization. Nonetheless, the measured 176Hf/177Hf has been corrected for radiogenic growth of 176Hf from 176Lu using the obtained zircon ages and a fixed 176Lu/177Hf value of 0.005 to constrain more tightly the zircon isotopic composition at the crystallization time (176Hf/177Hf(i)). We tried to analyze grains with the largest 206Pb/238U differences within each sample to capture the possible evolution of 176 Hf/177Hf(i), within the range of zircons in each sample. The 176Hf/177Hf(i) values of analyzed fractions show a restricted range (0.283110 b176Hf/177Hf(i) b 0.283162), which overall corresponds to a MORB-like mantle source (Salters and Hart, 1989; Salters and Hart, 1991; Salters, 1996; Nowell et al., 1998; Salters and White, 1998; Chauvel and Blichert-Toft, 2001). However, the dataset shows small scale variations that exceed analytical uncertainty, not only among different samples, but also among single fractions of the same sample. For example, two zircons in sample D101 yield 176Hf/177Hf(i), of 0.283119(±11) and 0.283150(±10). In sample P63, four zircons covering the entire age range show 176Hf/177Hf(i) values between 0.283146(±12) and 0.283111 (±4). The Thoregah sample KO8-18, shows a 176Hf/177Hf(i) value of 0.283094(±4) at the time of its crystallization, which is significantly lower than the values found for other samples. Together with trace

element concentrations, these data support field observation that this laccolith has a different origin than the Sapat Complex. 5. Discussion 5.1. Building of the Sapat juvenile crust Because primitive arc rocks do not reach zircon saturation during the beginning of their fractionation process, dating juvenile arc crust is challenging. In the case of Sapat, we found zircons in specific petrostructural positions, the plagioclase-rich segregates, and not necessarily in the most differentiated rocks i.e. the tonalites. This indicates that the segregates are somehow the most evolved magmatic products where incompatible elements concentrate, allowing zircon saturation. The presence of young populations associated with older grains may indicate that the segregates are reworked over time by multiple melt injection. Our high precision U/Pb ages reveal the quasisystematic presence of xenocrysts or antecrysts within the samples calling for a complex magmatic process that seems to unfold a long lasting magmatic history. In the same way as U/Pb ages from single pluton reveal different stages of growth (e.g. Miller et al., 2007), here examplified by the zircons population of the Thoregah laccolith, the plagioclase segregates individually contain different populations. It is unclear whether these zircons represent assimilated grains or reflect a protracted event of crystallization and segregation. Even if the high precision dating renders difficult to attribute an age of crystallization for the samples, and since the dated samples were collected from the mantle, lower crustal gabbros and magmatic conduits of the Sapat Complex the age difference between and within individual samples allows further insight into the crust formation. Sample D101, contains the oldest zircons found in Sapat, clustering at 106.83 ± 0.24 Ma and has a younger cluster at 103.81 ± 0.58 Ma. Sample G12, also segregated from its host gabbro, has an age overlapping with that of D101 and sample P8. Sample P63 dated at ca. 99 Ma reflects the latest stage of magma percolation and magmatic activity in the Sapat Complex, and has zircons with similar ages as samples G12, P8 and D101 (≈104 Ma). Sample G13, has a more complex history, with concordant zircon ages ranging from 104 to 98 Ma. Considering that those zircons show signs of possible recrystallization (Fig. 3), the youngest dated grains likely represent rejuvenated grains. One may argue that recrystallization may occur during a lower temperature, fluid-assisted diffusion process, but the concomitant decrease of Th/U ratio with age and the lack of profound disrupted magmatic zoning pattern do not support it. On the other hand, zircons show decreasing concordant ages coupled with a Th/U decrease indicating a dissolution–reprecipitation process (Geisler et al., 2007). Since the youngest and lowest Th/U grain of G13 has the same age as the crystallization age of P63, we propose that, recrystallization occurred while the magmatic system of the Sapat Complex was still active. The G13 leucocratic dyke was feeding a shear zone favouring long lasting emplacement of multiple batches forming the dyke. Dissolution–reprecipitation of G13 zircons in the presence of residual melt cannot be excluded. Therefore, the age of G13 records continuous magmatic activity between 103 and 100 Ma. These interpretations, considering the zircon population in its whole, have several implications on the magmatic history that created the new, juvenile crust: The lower-crustal metagabbros first emplaced between 107 and 104 Ma and were fed through dunitic mantle channels found in the Sapat mantle rocks (Bouilhol et al. 2009). The piling of gabbroic protolith sills was followed by pipe intrusion, more or less cogenetic with the pegmatitic metagabbros (Bouilhol et al., 2010). Since pipe1 and pipe 2 are cogenetic and structurally linked by pegmatitic meta-gabbros, we consider that the zircons from sample P8, intruding Pipe 1, reflect only a small population of the zircons, and that this sample may likely contain younger zircons as for P63. Accordingly, melt transfer within the pipes lasted until 99 Ma, when

P. Bouilhol et al. / Chemical Geology 280 (2010) 243–256

the youngest Zircons in P63 witnessed the latest/coolest melt that went through it. The entire age dataset shows that magmatism took place at least between 107 and 99 Ma. Taking into account all dated grains, the data display a continuous history, showing that the Sapat Complex formed through a protracted, continuous period of magma emplacement, which lasted approximately 7 Ma. 5.2. Source evolution The bulk compositions of the different rock units of the Sapat Complex suggested that the lithological units shared a common parental melt (Bouilhol et al. 2009, 2010). However, the Nd isotopic data are clustered, whereas the Pb isotopic composition show a wide compositional range, suggesting that the melting source was heterogeneous and or evolved during the formation of the Complex. Our zircon populations cover 7 Ma of magmatic activity, and even if hafnium isotopic concentrations show a restricted range with typical MORB-like values, the high level of precision of ages and 176Hf/177Hf determinations shows resolvable temporal variations of 176Hf/177Hf(i) in the crystallizing zircons (Fig. 7). As discussed in the previous section, we rely on the fact that none of the age variations is due to lead loss. For sample G13, we consider that its population recorded a magmatic evolution from 103 to 100 Ma, and the Hf composition of two grains within this sample concerns grains of pristine origin, which do not show very low Th/U. We attribute the observed trend to a sharp 176Hf/177Hf(i) increase between 105.24 and 103.63 Ma (D101#4 to G12#2), and a linear 176 Hf/177Hf(i) decrease with decreasing zircon age, from P8#1 (176Hf/ 177 Hf(i) = 0.283154 at 105.15 ± 0.35 Ma) to P63#4 (176Hf/177Hf(i) = 0.283110 at 98.91 ± 0.36 Ma). Such a trend indicates that during the onset of melting at ~106 Ma, the source region consisted of components with small but distinct differences in initial Hf isotope composition. It also shows that these heterogeneities vanished within 1 Ma, with the Hf isotopic system showing a more depleted signal. Then, within 4 Ma, the source region evolved toward enriched values and the younger zircon shows 176Hf/177Hf (i) values identical to those found in the Sarangar gabbro (part of the Kohistan lower crust in the Jijal-Patan Complex, Burg et al., 2005) dated at 98.9 ± 0.9 Ma (Schaltegger et al., 2002). To summarize, initial Hf isotopic compositions of dated single zircon grains reveal small-scale variations within the melting source region. The zircon Hf isotope data indicate a small-scale heterogeneous source region at the onset of melting. This source evolved

253

within 1 Ma toward more depleted values before showing a continuous trend towards more enriched values with time. We may expect this evolution to be reflected in the bulk Nd and Pb isotopic composition, but the difficulty to unravel the age of crystallization due to a long-lasting melt transfer renders the bulk-rock composition evolution difficult to apprehend. To understand the time variation of 176Hf/177Hf (i) in terms of geodynamics, we take into consideration the depleted character of the Sapat rocks in terms of major and trace element, the small variability of the depleted Nd isotopic composition compared to Pb, and, most of all, the fact that the Sapat parental melts were derived from a refractory source formed during the melting of a depleted mantle region in front of the Kohistan arc (Bouilhol et al., 2009, 2010). All of these constraints concur to a trenchward migration of the isotherms, tapping into depleted mantle that favored higher εHf values at the beginning of melting. This interpretation is consistent with values obtained in modern subduction systems such as the Mariana where high εHf values are linked to subduction initiation (Pearce et al., 1999); indeed, one may consider that subduction initiation involves the same mantle processes as forearc magmatism i.e. melting of a previously depleted mantle region through flux melting and production of highly depleted primitive arc-melt. The short term increase of 176 Hf/177Hf(i) at the beginning of Sapat evolution would reflect a steepening of the slab dip and a subsequent trenchward migration of the melting source. A depleted forearc region mantle would satisfy the primitive Hf and Nd isotopic compositions together with a depleted character of the melts. The wide range of Pb isotopic compositions would also fit this interpretation, where the most mobile, slab-derived fluid component is preferentially extracted in the forearc region. In this hypothesis, the evolution of εHf from +16 to +14 would correspond to either a greater slab contribution, where melting affects the tip of the mantle wedge that has been flushed trough time by different slab component, or a smaller contribution of the depleted mantle. In the first case, melt transfer occurs through a system of channels that form in a depleted mantle with high εHf to confine lower εHf melts with a slab component becoming more visible with time. In the second case, the amount of mantle melting may diminish with time, lowering the mantle contribution, thus mirroring lower εHf values. Because the Sapat Complex displays isotopic similarities with the Jijal Complex (Fig. 5), yet with higher Nd isotopic compositions (i.e. more depleted source) for some samples, our interpretation is in line with a model in which Sapat formed during a hot mantle influx towards the trench, dragging a “Jijal like component” into a more depleted region of the front of the Kohistan Arc in Cretaceous times. 5.3. Constraints on the Kohistan arc evolution

Fig. 7. Isotopic age evolution in Sapat. Initial Hf single zircon grains plotted against age of each grain. Ages and single grain zircons from the Jijal lower crust are from Schaltegger et al. (2002). All errors are 2σ, and shown when bigger than plot symbol. Numbers next to each point refer to analyzed grains numbers.

To place the formation of the Sapat Complex in the frame of the Kohistan Arc evolution, and in order to compare datasets of equal precision and accuracy, we now discuss U–Pb ages obtained by IDTIMS techniques and Hf isotope analyses from dated zircon grains only. Such data cover most of the geological units of the Kohistan Arc (Schaltegger et al., 2002; Schaltegger et al., 2003; Heuberger et al., 2007). We use this dataset to track the evolution of the arc source with time (Fig. 8). Within the “Kohistan Batholith”, i.e. the mostly plutonic mid-level of the arc, the 155 Ma Matum-Das tonalite is so-far the oldest rock (Schaltegger et al., 2003). It also has the highest εHf value of + 21. In the same way that very high εHf is found nowadays at the Southeast Indian Ridge (Hanan et al., 2004), the εHf value of Matum-Das may represent melts extracted from a primitive mantle unaffected by a subduction component and characterized by a mixed composition inherited from the rifting of Gondwana during opening of the Tethys Ocean. Therefore, the Matum-Das tonalite may represent the initiation of the Kohistan Arc. No available data allows tracing εHf values between this +21 at 155 Ma and the εHf ≈ 14 at

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Fig. 8. εHf versus age for dated zircons from different Kohistan lithologies. Symbols type represent lithologies, symbol color literature source: Yellow from Schaltegger et al. (2002); red from Schaltegger et al. (2003); blue from Heuberger et al. (2007); green = this study. Also, unscaled plate model for the proposed interpretation of εHf signals in terms of geodynamic evolution.

110 Ma (Heuberger et al., 2007). We contend that the εHf values dropped to εHf ≈ 14 with stabilization of the subduction regime prior to 110 Ma. From 110 Ma to 80 Ma, the Kohistan baseline εHf composition is defined at a near-constant εHf of 14 (Schaltegger et al., 2002; Schaltegger et al., 2003). With respect to this baseline, two events can be pointed out:

migration of the mantle and the formation of the Sapat Complex in a frontal position may represent slab retreat, which later favored and set-off arc rifting (Burg et al., 2006; Jagoutz et al., 2007). Sample KO818, which has a εHf = 13.2 at 93 Ma, would fall on the εHf trend toward the Chilas composition, linking in time the two events. 6. Conclusions

(1) Formation of the Sapat Complex at 105 Ma records an increase in εHf values toward a more depleted mantle source, with a return to the Kohistan baseline within 7 Ma. This increase is attributed to the migration of the melting region towards the trench. (2) Intrusion of the Chilas Complex at 85 Ma, which triggered a decrease in εHf values. The lower Hf isotopic composition can either represent an increase in the amount of slab component and/or a change in the mantle source. The Chilas complex is related to the splitting of the arc (Burg et al., 2006; Jagoutz et al., 2007). For this reason, the lower εHf values have been tentatively attributed to a Karakoram mantle component dragged into the Chilas source region (Schaltegger et al., 2002). However, recent data (Heuberger et al., 2007) show that a series of intrusions that formed at ≈105 Ma within the backarc region between the Kohistan Arc and the Karakoram continental margin has εHf ≈ 10. This may indicate that the Chilas source originated from a mantle influx from the rear-arc. After collision with the Indian continent at ≈50 Ma (Patriat and Achache, 1984; Gaetani and Garzanti, 1991; Clift et al., 2002; Kaneko et al., 2003) εHf values point to a less depleted melt source (+11; +9). Inherited zircons within the samples yield εHf values down to −6, indicating that melts have sampled and assimilated different types of continental crust (Heuberger et al., 2007). In summary, by tracking precisely the εHf evolution of the Kohistan Arc, we gain important insights into the geodynamic evolution of the Kohistan subduction system between 110 Ma and 80 Ma. We recognize a trenchward melt source migration for the formation of the Sapat Complex at 104 Ma, and splitting of the arc with the formation of the Chilas Complex at 85 Ma. The trenchward

The Sapat Complex documents the formation of a juvenile lower crust in the Kohistan paleo-arc. The combination of U–Pb dating and Hf isotope composition of zircons in plagioclase segregates and tonalites found at different structural positions within the Complex allowed unfolding the formation of the lower crust entity and tracking the coeval evolution of the mantle, between 105 and 99 Ma. This lower arc crust formed through protracted intrusion of sills that have been immediately intruded by magmatic pipes, suggesting that the pipe fed some of the sills and formed their own future host. Few zircons were found in of the Sapat juvenile crust, but they cover the range of the magmatic activity. This makes difficult the determination of a final crystallization age and the definition of a reliable evolution of bulk-rock isotopic compositions. On the other hand, our high precision U–Pb and Hf measurements of several zircon generations in single samples offer precise insight into the temporal evolution of the melt sources. Our data further constrain the long-term evolution of the Kohistan Arc during its Cretaceous intraoceanic history, when its mantle source region had an Hf isotopic composition baseline of εHf≈ 14. Deviations from this baseline are interpreted in terms of geodynamic events. In particular, the Sapat primitive parental melt formed in the frontal part of the arc at the favor of a trenchward migration of the mantle wedge, bringing hot mantle towards the wedge tip. Acknowledgements ETH-grant number 0-20220-04 supported this work. The Swiss National Fond is acknowledged for continued support of the University of Geneva radiogenic isotope laboratory. We sincerely thank Ulmer P., Zeilinder L. and Reusser E. for analytical support at

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