Earth and Planetary Science Letters 366 (2013) 163–175

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Earth and Planetary Science Letters

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journal homepage: www.elsevier.com/locate/epsl

Dating the India–Eurasia collision through arc magmatic records

Pierre Bouilhol a,1, Oliver Jagoutz a,n, John M. Hanchar b, Francis O. Dudas a

a Department of Earth, Atmospheric, and Planetary Sciences, Massachusetts Institute of Technology, Cambridge, MA 02139, USAb Department of Earth Sciences, Memorial University of Newfoundland, St. John’s, Newfoundland, Canada A1B 3X

a r t i c l e i n f o

Article history:

Received 26 October 2012

Received in revised form

18 January 2013

Accepted 19 January 2013

Editor: T.M. Harrisonupper crust of the Kohistan Paleo-Island Arc that formed in the equatorial part of the Neo-Tethys Ocean.

Available online 29 March 2013

Keywords:

Himalaya

collision age

Kohistan–Ladakh Arc

zircon U–Pb/Hf isotopes

Nd/Sr isotopes

1X/$ - see front matter & 2013 Elsevier B.V.

x.doi.org/10.1016/j.epsl.2013.01.023

esponding author. Tel.: þ1 617 324 5514.

ail addresses: pierre.bouilhol@durham.ac.uk (

@mit.edu (O. Jagoutz), jhanchar@mun.ca (J.M

esent address: Department of Earth Sciences,

E, United Kingdom.

a b s t r a c t

The Himalayan orogeny, a result of the collision of India and Eurasia, provides direct evidence of strain

accommodation and large-scale rheological behavior of the continental lithosphere. Knowledge of the

timing of the India–Eurasia collision is essential to understand the physical processes involved in

collisional systems. Here we present a geochronological and multi-isotopic study on rocks from the

In situ U–Pb geochronology and Hf isotopes in zircon, and whole-rock Nd and Sr isotopic data of

plutonic rocks from the Kohistan-Ladakh Batholith, are used to construct a continuous record of the

isotopic evolution of the source region of these granitoids that are related to both the subduction of the

oceanic lithosphere and subsequent arc–continent collisions. We demonstrate that profound changes in

the source region of these rocks correspond to collisional events. Our dataset constrains that the

Kohistan–Ladakh Island Arc initially collided along the Indus suture zone with India at 50.271.5 Ma,

an age generally attributed to the final India–Eurasia collision for the entire Himalayan belt. In the

western Himalaya, the final collision between the assembled India/Arc and Eurasia however, occurred

�10 Ma later at 40.471.3 Ma along the so-called Shyok suture zone. We present evidence indicating

that a similar dual collision scenario can be extended to the east and conclude that a final India/Arc–

Eurasia collision at �40 Ma integrates crucial aspects of the magmatic, tectonic, and sedimentary

record of the whole Himalayan mountain belt.

& 2013 Elsevier B.V. All rights reserved.

1. Introduction

The continent–continent collision of India with Eurasia hassignificantly shaped the surface of the Earth by producing one ofthe largest known orogenic systems. To understand the impor-tance of the physical processes controlling crustal deformationand the influence of this mountain building event on the climaticsystem (Le Pichon et al., 1992; Molnar et al., 1993; Molnar andTapponnier, 1975; Royden et al., 2008; Tapponnier et al., 1982),an accurate knowledge of the plate convergence and the kine-matic history of the collision is essential. Key boundary para-meters include the timing of collision and plate reconstructions,which together constrain the absolute amount of plate conver-gence accommodated. Whereas modern paleomagnetic studiesagree on plate reconstruction within a few hundred kilometers(e.g., van Hinsbergen et al., 2012), the timing of the final India–Eurasia collision is debated. Estimates range from �70 to �25 Ma(e.g., van Hinsbergen et al., 2012; Yin and Harrison, 2000) withmost workers preferring a �52–50 Ma collision age (e.g., Najman

All rights reserved.

P. Bouilhol),

. Hanchar).

Durham University, Durham

et al., 2010). Recent paleomagnetic data suggest a complexcollision scenario that invoked multiple continental or arc frag-ments and increasingly casts doubt that the �50 Ma eventsconstrains the final India–Eurasia collision (Aitchison et al.,2007; Khan et al., 2009; van Hinsbergen et al., 2012). Thisuncertainty on the collision age translates to up to �3000 kmdifferences in estimates on the convergence accommodated in thesystem, which strongly hampers our understanding of the forma-tion of the India–Asia collisional system.

In this paper we present new results that constrain the timingof collision in the western Himalaya (Fig. 1). Our approach isderived from our understanding of currently active continent–island arc collision zones (see Section 3). In these systems, thesubduction of continental lithosphere below the arc can betracked by the differences in isotopic compositions betweenpre- and post-collisional arc magmas if significant isotopic differ-ences exist between the subducting continental lithosphere andthe obducting arc crust. Along the Himalayan belt, only theKohistan–Ladakh Paleo-Island Arc (KLA) that is exposed in thewestern Himalaya offers such an opportunity.

The main objective of this paper is to construct a continuousspatial and temporal record of the isotopic composition of the KLArelated rocks, spanning from the intra-oceanic subduction inthe early Cretaceous to the Miocene, thus covering all proposedcollision times. Kohistan–Ladakh Arc magmas that formed prior

Kargil

Besham

Shyok Suture

Indus Suture

Karakoram Plate

Gilgit

Drosh

Chilas Nidarophiolite

Karakoram Fault

Pangong Complex

80º00’78º00’76º00’74º00’

72º00’

34º00’

36º00’

36º00’

34º00’

100 Km

NangaParbat

Skardu

Islamabad

Indus SutureAccretionary Complex

Tsangpo Suture

Shyok SutureKarakoram Fault

LadakhKohistan

IndiaHimalayan Belt

Eurasia

ca. 500 km

N

Indus Suture

Northern / Central / Southern Batholith samples

/ /

Literature samples

εε

ε

ε ε

= Eclogitesε

Indian Plate

Lower Arc crust

Volcanic / sedimentary units

Arc mantle

Mid-upper crustal Batholith

N

Supra-subduction zoneremnants

Supra-subduction zoneremnants

Leh

Fig. 1. Sample location on a simplified geological map of the Kohistan–Ladakh Arc (KLA). White symbols indicate homogenous unradiogenic samples and yellow filled

symbols indicate isotopically evolved and/or heterogeneous granitoids. Inset shows the structural position of the KLA with respect to the different Himalayan suture zones.

(For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

P. Bouilhol et al. / Earth and Planetary Science Letters 366 (2013) 163–175164

to the 50–52 Ma collision of the arc with India have isotopiccharacteristics comparable to modern intra-oceanic arc volcanics(Figs. 2 and 3). Conversely, the post-collisional, late-Eocene/Oligocene granites are more variable in composition and trendto significantly more radiogenic isotopic compositions (Heubergeret al., 2007) indicating the involvement of old continental litho-sphere in their formation.

2. Geological setting

Three suture zones, the Indus, Shyok and Tsangpo, markthe Himalayan collision (Fig. 1). In the NW Himalaya, the�70,000 km2 KLA is wedged between India and Karakoram (theformer paleo-Eurasian continental margin). The Shyok sutureseparates the KLA in the north from the Karakoram, whereas inthe south the Indus suture divides the KLA from the Indiancontinent (Fig. 1). The Tsangpo (or Yarlung-Tsangpo) separatesIndia from Eurasia in the central and eastern part of the Himalaya(Burg and Chen, 1984; Gansser, 1980; Yin and Harrison, 2000).

2.1. The KLA and the Karakoram

The KLA (for a recent review see Burg, 2011) formed primarilyduring the late Mesozoic as an intra-oceanic arc, during thenorthward directed intra-oceanic subduction of the Neotethysocean basin (Bard, 1983; Coward et al., 1987; Tahirkheli, 1979).This interpretation is based on the absence of continental base-ment and continental derived sediments in the KLA (Bard, 1983;Jagoutz et al., 2009; Tahirkheli, 1979), and on paleomagnetic dataplacing the formation of the KLA near the equator within anoceanic basin (e.g., Khan et al., 2009) much further south than theEurasian margin during the Mesozoic (Schettino and Scotese,2005).

To the north of the arc system, at the northern most margin ofthe Neothethys, a second subduction system was active simulta-neously, resulting in the Jurassic to Tertiary Gangdese andKarakoram Batholiths (e.g., Debon et al., 1987; Gaetani, 1997;Heuberger et al., 2007; Searle et al., 2010). These Batholiths were

part of an active continental margin due to the northwardsubduction of Tethyan oceanic lithosphere beneath the southernEurasian margin (e.g., Le Fort, 1988; Rex et al., 1988; Yin andHarrison, 2000).

2.2. The Indus suture

In contrast to the Tsangpo suture, the Indus suture zonecontains eclogites and ultra-high pressure (UHP) assemblagesthat can be used to reliably constrain the timing of the formationof the Indus suture. The P-T-t path of the eclogites providesevidence for the beginning of the subduction of the Indian marginat �50 Ma (Kaneko et al., 2003; O’Brien et al., 2001; Parrish et al.,2006). In conjunction with the fact that the sedimentary recordplaces the transition from marine to non-marine sedimentationon the northern Indian Margin at �50 Ma (e.g., Beck et al., 1995;Garzanti et al., 1987; Green et al., 2008), the timing of the Indussuture, and the collision between India and the KLA, have beenaccepted to be �50 Ma.

2.3. The Shyok suture

The Shyok suture has a complex history and its formation ispoorly constrained. North verging folds and top-to-the-northdirected shear zones within the northern part of the KLA indicatethat the KLA was initially overthrusted on top of the Karakorammargin (Coward et al., 1986). During the early Miocene, parts ofthe Shyok suture were reactivated as a top-to-the-south directedthrust, overprinting the previous suture related structures(Brookfield and Reynolds, 1990; Coward et al., 1986).

Originally, the formation age of the Shyok suture was inter-preted to postdate the India–KLA collision (Achache et al., 1984;Andrewsspeed and Brookfield, 1982; Bard, 1983; Brookfield andReynolds, 1981). Subsequently, based on a proposed systematicrelationship between emplacement age of plutonic rocks fromthe northern KLA (determined by Rb–Sr whole rock and mineralisochrons) and the presence or absence of deformational fabric inthe rocks (i.e., younger ones being undeformed, older deformed),the Shyok suture has been inferred to have formed during the

0.710

87Sr/ 86Sr(i) Bulk rock composition

India/Arc - Karakoram

P. Bouilhol et al. / Earth and Planetary Science Letters 366 (2013) 163–175 165

Cretaceous (Petterson and Windley, 1985; Windley, 1988). How-ever, recent U–Pb zircon dates demonstrate that no systematicrelationship between intrusion ages and absence or presence ofdeformational fabric exists (Jagoutz et al., 2009), leaving the

20 40 60 80 100

15

30 50 70 90

India - Arc suturing50.2 ± 1.5 Ma

Inhomogeneous samplesSingle grains

LB09-11.5LB09-11.4

K08-13

εNd(i) Bulk rock composition110

10

5

0

-15

-10

-5

εHf(i) Zircon composition

InheritedPaleozoic

zircongrains

-10

-8

-6

-4

-2

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20εNd(i) Bulk rock composition

0

2

4

6

40 60 80 10030 50 70 90 110

15

10

5

0

-15

-10

-5

MesozoicKarakoram

Zirconcomposition

εHf(i) Zircon composition

Single grains

O1B-22O1A-07

Age Ma

InheritedMesozoic

zircongrains

India/Arc - Karakoramsuturing 40.4 ± 1.3 Ma

Inhomogeneous samples

Age Ma

LB11-O16LB11-O10

InheritedPrecambrian

grain

InheritedPaleozoic

grains

-6

-4

-2

0.704

0.706

0.708

20 30 40 60 70 80 90 100 110

0.704

0.706

0.708

0.710

0.712

0.714

Age Ma

50

suturing 40.4 ± 1.3 Ma

India - Arc suturing50.2 ± 1.5 Ma

Fig. 3. For the northern and southern sampling segment of the Kohistan–Ladakh

Batholith, U–Pb zircon crystallization ages of granitoids vs. their initial Sr isotopic

composition (87Sr/86Sr(i)). The green area represents the isotopic range know from

the KLA (Bignold et al., 2006; Bouilhol et al., 2011; Dhuime et al., 2009; Jagoutz

and Schmidt, 2012; Khan et al., 2004; Rolland et al., 2002b). The blue and orange

bars represent the age range at which the India–Arc and the India/Arc–Karakoram

collision occurred, respectively. (For interpretation of the references to color in

this figure legend, the reader is referred to the web version of this article.)

formation of the Shyok suture zone essentially unconstrained(Burg, 2011; Jagoutz et al., 2009). Similarly, paleomagneticdata indicate that the KLA was situated �3000 km south of theEurasian margin in the Tertiary, making a Cretaceous formationage of the Shyok suture highly unlikely (Ahmad et al., 2000; Khanet al., 2009).

3. Approach for dating the India–KLA–Eurasia collision events

The rationale of our method to constrain collision events of theKLA by investigating the isotopic composition of igneous rocksfollows from observations on modern continent–arc collision sys-tems and the documented impact of collision on the geochemistryof the arc-related magmas. The Australian–Sunda–Banda Arc colli-sion system, for example, (see detailed review by Harris, 2011)provides a well-documented modern analog for the India–KLAcollision. The Australian continent, drifting north–northeastward(NNE) at a rate of �70 mm/a (Nugroho et al., 2009), initially collided

Fig. 2. For the northern and southern sampling segment of the Kohistan–Ladakh

Batholith, U–Pb zircon crystallization ages of granitoids vs. (1) their initial zircon

epsilon Hafnium (eHf(i)) composition and (2) initial epsilon neodymium isotopic whole

rock composition (eNd(i)). For samples with homogeneous zircon population, the eHf(i)

of zircon grains is averaged. For inhomogeneous samples, the eHf(i) of each analyzed

zircon grain are shown and plotted at the crystallization age of the sample at the

exception of inherited zircon grains from the intra-oceanic history (50oage Mao110

and 6oeHf(i)o14; Table A.1) which are not shown. The green area represents the

isotopic range known from the KLA (Bignold et al., 2006; Bouilhol et al., 2011; Dhuime

et al., 2009; Heuberger et al., 2007; Jagoutz et al., 2006; Khan et al., 2004; Ravikant

et al., 2009; Rolland et al., 2002b; Schaltegger et al., 2002). The blue and orange bars

represent the age range at which the India–Arc and the India/Arc–Karakoram collision

occurred, respectively. (For interpretation of the references to color in this figure

legend, the reader is referred to the web version of this article.)

Table 1Bulk rock analyses of the southern, central, and northern section samples.

Sample Location (UTM) Rock type Age 2sig eHf(i) 2sig Sm 1sig Nd 1sig 147Sm/144Nd 2sig 143/144Nd 2sig 143/144Nd(i) eNd(i) Rb (ppm) Sr (ppm) 87/86Sr 2sig 87Rb/86Sre 87/86Sr(i)

Northing Easting Streckeisena (Ma)b (95%) (95%) (ppm) s.ec (ppm) s.ec s.ec s.ec ICPMS TIMS ICPMS TIMS s.e c

Southern section

PK06-11 3869906 509255 Granodiorite 102.1 2.1 12.1 0.3 2.60 7 16.82 4 0.0936 4 0.512825 8 0.512763 5.0 31.7 – 174 162 0.704506 7 0.526 0.703744

LB09-21.3A 3828865 594044 Tonalite 98.6 2.6 9.6 0.5 1.63 4 8.52 2 0.1159 5 0.512787 8 0.512712 3.9 30.2 – 528 509 0.704102 18 0.165 0.703871

KO8-3 3942112 362876 Tonalite 76.9 2.1 6.2 1.0 1.73 5 9.49 2 0.1102 5 0.512656 5 0.512601 1.2 21.1 – 303 292 0.704531 8 0.201 0.704311

LB09-22.7 3822477 636965 Monzogranite 72.0 1.5 9.1 1.1 4.16 12 17.80 5 0.1414 7 0.512740 7 0.512673 2.5 113 – 44.7 40.1 0.711811 13 7.297 0.704347

LB09-23.11A 3810450 659724 Monzogranite 62.7 1.4 10.0 1.1 2.97 8 14.06 4 0.1278 6 0.512757 8 0.512705 2.9 52.8 – 246 228 0.704743 8 0.622 0.704189

LB09-5.2B 3760463 758794 Monzogranite 61.8 1.1 8.5 0.7 2.91 7 17.30 4 0.1016 4 0.512718 8 0.512677 2.3 92.4 – 58.9 54.4 0.708706 9 4.541 0.704719

LB08-12 3826627 603150 Tonalite 58.4 1.1 10.4 0.5 4.23 11 17.16 4 0.1491 6 0.512751 4 0.512694 2.6 42.7 42.7 391 374 0.704644 8 0.331 0.704382

LB08-7 3787912 723697 Monzogranite 58.1 1.8 8.6 1.1 1.82 5 13.66 3 0.0804 3 0.512721 5 0.512690 2.5 103 128.3 43.3 40.1 0.709976 11 9.252 0.704127

LB09-48 3768552 749877 Tonalite 53.4 1.9 8.9 0.8 4.46 12 20.31 5 0.1329 6 0.512727 5 0.512681 2.2 35.2 – 451 448 0.704421 13 0.226 0.704250

LB09-30 3808235 683825 Granodiorite 53.3 1.3 9.1 1.0 4.77 12 23.97 6 0.1202 5 0.512754 7 0.512712 2.8 118 – 408 385 0.704863 14 0.833 0.704233

LB09-24.5 3798089 695491 Granodiorite 51.5 1.6 9.6 0.7 2.86 8 16.59 4 0.1041 5 0.512678 10 0.512643 1.4 124 – 245 223 0.705598 10 1.469 0.704523

LB08-25 3774936 740850 Granodiorite 50.3 2.1 9.3 0.8 3.29 9 19.33 5 0.1028 4 0.512686 3 0.512652 1.5 110 125 268 249 0.705258 10 1.452 0.704221

KO8-13 3956536 379587 Granodiorite 50.4 1.6 Hetd – 4.34 60 27.2 7 0.097 4 0.512421 6 0.512389 �3.6 85.9 – 391 374 0.706317 10 0.636 0.705862

LB08-23 3824622 647205 Granodiorite 48.1 1.1 4.1 0.6 9.09 22 52.8 14 0.104 4 0.512281 4 0.512248 �6.4 124 157 1114 1306 0.710577 13 0.347 0.710340

LB09-11.4A 3714094 792703 Granodiorite 39.8 0.7 Het – 3.86 10 25.1 6 0.093 4 0.512301 7 0.512277 �6.0 253 – 473 453 0.709787 10 1.548 0.708912

LB09-11.5A 3716910 792261 Monzogranite 29.6 0.8 Het – 6.06 16 36.5 9 0.100 4 0.512124 7 0.512105 �9.7 306 – 263 236 0.714583 9 3.365 0.713170

Central Batholith

PGAK-26 4002072 386266 Granodiorite 72.1 1.4 9.2 0.9 4.15 10 18.01 5 0.1393 6 0.512701 8 0.512635 1.8 62.2 – 442 424 0.704924 8 0.407 0.704507

PGUP-15 3998723 347648 Monzogranite 58.8 1.4 9.4 1.1 2.76 7 13.30 3 0.1252 6 0.512789 8 0.512741 3.5 134 – 72.8 67.4 0.708570 9 5.304 0.704139

Pba-4 3988739 356574 Tonalite 61.3 1.6 10.7 0.7 1.47 4 6.82 2 0.1299 6 0.512787 8 0.512735 3.4 13.0 – 456 429 0.704276 10 0.082 0.704204

Pba-15 3998126 349040 Granodiorite 72.0 3.0 9.0 0.5 6.52 16 28.32 8 0.1392 6 0.512772 6 0.512706 3.1 76.8 – 356 330 0.704588 10 0.623 0.703950

PK05-20 3874950 538460 Kfs-Granitoid 42.1 0.9 4.0 0.6 – – – – – – – – – – – – – – – – – –

PK05-16 3882204 507148 Kfs-Granitoid 41.5 1.0 3.9 0.5 – – – – – – – – – – – – – – – – – –

PK06-36 3892534 556405 Monzogranite 46.3 1.3 6.8 0.7 4.74 12 24.07 7 0.1190 6 0.512674 8 0.512638 1.2 174 – 251 233 0.706374 8 2.010 0.705061

LB08-31a 3796128 739770 Granodiorite 59.5 1.9 9.1 0.5 2.59 7 13.66 3 0.1147 5 0.512734 4 0.512689 2.5 94.3 106 164 154 0.705452 10 1.977 0.703781

LB09-16.6A 3769896 769686 Granodiorite 60.9 1.6 8.5 0.6 3.60 9 14.08 4 0.1546 7 0.512753 5 0.512691 2.6 67.7 – 306 287 0.704715 8 0.641 0.704161

LB09-33 3819161 683181 Granodiorite 66.0 2.2 9.4 0.9 2.88 8 17.11 4 0.1019 4 0.512682 7 0.512638 1.7 110 – 207 189 0.705732 8 1.540 0.704288

KO8-4 3961913 354739 Granodiorite 31.7 0.8 Het – 2.98 80 21.18 5 0.0850 4 0.512487 8 0.512469 �2.5 139 – 797 759 0.706179 7 0.506 0.705952

PJA-4 3952766 466314 Granodiorite 30.1 1.3 Het – 2.87 7 20.10 5 0.0863 4 0.512670 6 0.512653 1.0 67.6 – 614 546 0.705095 10 0.318 0.704959

Northern section

PK05-39 3938738 528932 Tonalite 86.0 2.6 14.0 0.8 1.70 5 7.66 2 0.1340 6 0.512881 10 0.512806 5.4 39.0 – 514 467 0.704079 8 0.219 0.703811

PGUP-16 4006683 342237 Granodiorite 83.5 2.8 13.2 0.5 2.16 6 12.23 4 0.1068 5 0.512848 6 0.512790 5.1 53.0 – 485 452 0.704029 11 0.316 0.703654

PGLT-19 3989516 422611 Granodiorite 64.8 3.9 9.9 1.1 3.87 10 19.50 6 0.1200 6 0.512853 4 0.512802 4.8 81.8 – 443 405 0.704347 8 0.535 0.703855

PGUP-9 4003763 334840 Granodiorite 58.9 1.4 10.8 0.8 3.34 8 16.86 5 0.1197 5 0.512851 6 0.512805 4.7 80.0 – 202 189 0.705064 10 1.147 0.704105

PGAK-18.5 4008678 384781 Monzodiorite 51.5 2.7 11.8 0.8 3.39 8 18.51 5 0.1106 5 0.512790 10 0.512753 3.5 66.3 – 847 929 0.704510 10 0.226 0.704344

LB09-16.7 3774170 771835 Tonalite 51.5 2.2 9.0 0.9 3.45 9 17.03 4 0.1224 50 0.512724 6 0.512683 2.2 64.7 – 316 279 0.704696 8 0.592 0.704263

PGAK-13.1 4008937 384570 Tonalite 47.8 1.5 11.3 0.8 4.70 80 22.07 6 0.1289 5 0.512764 8 0.512724 2.9 49.0 – 496 457 0.704445 7 0.286 0.704251

01b-24 3922976 206395 Quartz-diorite 40.9 1.1 11.7 2.0 1.93 5 7.44 2 0.1566 7 0.512919 8 0.512877 5.7 – 5.3 393 396 0.704044 8 0.039 0.704022

01B-22 3931657 208900 Granodiorite 39.9 1.4 Het – 3.13 8 23.53 6 0.0805 3 0.512684 6 0.512663 1.5 65.8 – 396 381 0.705289 8 0.480 0.705017

O1A-7 3930468 209417 Granodiorite 37.2 1.1 Het – 3.87 10 23.57 6 0.0994 4 0.512702 10 0.512678 1.7 103 – 652 652 0.704767 8 0.457 0.704526

LB11-O10 3854378 665262 Tonalite 21.5 1.2 Het – 1.62 4 6.17 2 0.1583 7 0.512276 8 0.512254 �7.0 – 291 – 162.4 0.711286 7 5.192 0.709701

LB11-O16 3856699 672432 Granite 25.1 1.3 Het – 2.90 8 14.37 4 0.1218 5 0.512296 8 0.512276 �6.4 – 126 – 294.1 0.707863 8 1.238 0.707422

LB11-S36 3834088 719460 Granite 68.5 1.5 9.2 0.5 2.10 6 15.24 4 0.0832 4 0.512722 8 0.512685 2.6 – 67.7 – 17.42 0.714552 13 11.270 0.703582

a Streckeisen, A.L., 1974. Classification and Nomenclature of Plutonic Rocks. Recommendations of the IUGS Subcommission on the Systematics of Igneous Rocks. vol. 63, Geologische Rundschau, Internationale Zeitschrift fur Geologie,

Stuttgart, pp. 773–785.b Crystallization age calculated with Isoplot, see Fig. A1.c Error number corresponds to the last digits.d Het: Refers to samples that have a heterogenous Zircon population.e Values in Italic refer to Isotopic dilution TIMS (ID-TIMS) measurements.

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P. Bouilhol et al. / Earth and Planetary Science Letters 366 (2013) 163–175 167

�2 m.yr. ago with the Sunda–Banda arc system, resulting in thesubduction of the Australian continental lithosphere below the arc.Though the details of that collision are complicated, one clearresult from the well-studied Australia–Sunda–Banda Arc system isessential to the interpretation of our India–KLA results; the arrival ofthe leading edge of the Australian plate in the subduction zone canbe traced by the isotopic composition of the upper plate magmas inspace and time (Elburg and Foden, 1998; Elburg et al., 2005, 2002;Herrington et al., 2011; van Bergen et al., 1993). Lavas that formedduring subduction of the oceanic lithosphere (e.g., during theintra-oceanic arc stage of the subduction zone), have isotopiccompositions reflecting the formation of the melts in a slabcontaminated mantle wedge similar to the current situation furtherwest in Java (Handley et al., 2007) and typical for intra-oceanic arcmagmas (e.g., 0.51278o143Nd/144Ndo0.51289). Where geophysi-cal data indicate that the tip of the continental lithosphere enteredthe subduction zone (e.g. collision occurred Fichtner et al., 2010),a pronounced change in the range of isotopic variations isobserved in magmas erupted on the overriding plate (e.g.,0.51193o143Nd/144Ndo0.51284) and is interpreted as the partici-pation of the Australian lithosphere in the formation of these melts(e.g., Elburg et al., 2005). In the back-arc side of the collision zone,the furthest from the trench, the isotopic composition of contem-poraneous magmas show little variation (143Nd/144Nd�0.5125(1)),because the Australian continent has not been subducted deepenough to influence the source region of these melts (Elburg et al.,2005). Accordingly, the subduction of the Australian lithosphere canbe closely tracked in space and time by the chemical composition ofthe related magmatic rocks. Based on these observations we haveinvestigated the possible geochemical effects of the India–KLA–Eurasia collision on the arc-related magmatic products.

Within the Himalayan system, only the KLA offers this oppor-tunity as it formed as a Cretaceous to Eocene intra-oceanic islandarc and accordingly lacks the continental basement of theKarakoram or Lhasa Terranes (e.g., Chung et al., 2009; Yin andHarrison, 2000 and references therein) allowing attribution ofisotopic changes to the deeper source region of the rocks studied,rather than, for example, changes in intracrustal assimilationprocesses.

4. Analytical methods

We used in situ zircon U–Pb geochronology and Hf isotopes(Table A.1; Fig. A.1), and whole-rock Nd and Sr tracer isotopes of41 granitoids samples from the KLA batholiths ranging in com-position from diorite to granite (Table 1). For each sample, U–Pbdates and Lu–Hf compositions were determined at the IncoInnovation Centre of the Memorial University of Newfoundland(St. John’s, Canada). U–Pb and Pb isotopic ratios were analyzedusing a Finnigan ELEMENT XR ICP-MS coupled to GeoLas 193 nmArF Excimer laser. A 10 mm laser beam was rastered over thezircons surface in a 40�40 mm grid. The in situ Lu–Hf isotopicanalyses were acquired using a Finnigan NEPTUNE multi-collectorICP-MS directly on top of the U–Pb raster using a 49 or 59 mmlaser spot using the same laser ablation system. Bulk-rockisotopic analyses were done at MIT using standard wet chemicalseparation techniques and the IsoProbe-T multicollector thermalionization mass spectrometer (TIMS). See Electronic supplementfor more detailed information about the analytical methods usedin this study.

To identify and constrain the possible end-members involvedin the formation of the granitoids, we modeled their isotopiccompositions (Table 2; Fig. 4). This was done in order to identifyvariations in the sources and the significance of any observedchanges. As shown below, simple mixing calculations of partial

melts (calculated as batch melts) derived from four differentsource rocks, successfully explain the observed variations(Table 2; Fig. 4): (1) Enriched–Depleted MORB Mantle (E–DMM);(2) Tethys-type MORB; (3) evolved continental crust; and (4) theKLA lower crust. This fourth potential source rock does notrepresent a different independent isotopic end-member compo-nent, as its composition lies on a mixing line between TethyanMORB and E–DMM. We chose E–DMM for the mantle componentbecause of the enriched isotopic character of the mantle in theIndian Ocean, which is reflected in the isotopic composition of thearc (Bouilhol et al., 2011; Jagoutz and Schmidt, 2012; Khan et al.,1998). On average, the granitoids composition are close to a meltcomposition and cumulates are rare (Jagoutz et al., 2011), thusallowing to compare directly their measured Sr and Nd concen-trations in our model. To model the Hf and Lu data we usedmineral/melt Kd data (1000 1C zircon Kd from Rubatto andHermann, 2007) to calculate the Lu and Hf concentrations of thecorresponding liquids in equilibrium with the analyzed zircongrains. With the exception of a depleted mantle component(E–DMM), all other source rocks have been constrained fromactual geochemical data (e.g., trace elements and radiogenicisotopes) from rocks exposed in the KLA and surrounding unitsor from relevant geographical areas, providing very specificconstraints for our model calculations. All parameters used forthe model calculations are provided in Table 2.

The spatial distribution of the samples investigated coversthe entire KLA batholith and define three sections (Fig. 1): (1)the southern section comprises the samples that are the closest tothe Indus suture; (2) the northern section consists of the samplesthe closest to the Shyok suture; and (3) the central section includeall other samples that are geographically in between the twopreviously defined sections. For each section, the sample’s iso-topic record (143Nd/144Nd, 87Sr/86Sr, and 176Hf/177Hf) are dis-cussed in relation to their U–Pb zircon crystallization age(Figs. 2, 3 and 5).

5. Results and discussion

In regard to their isotopic compositions, the KLA granitoidsdefine two broad groups: (1) an older depleted group that reflectsa juvenile character of the granitoid source; and (2) a youngermore isotopically evolved group indicating the contribution fromcontinental crust. As the transition between the different groupsoccurs at different times and at different places in the arc, wediscuss the results separately for the southern, northern, andcentral part of the Batholith.

5.1. The southern section

Along the Indus suture, the granitoids show contrasting iso-topic records that can be temporally divided into two suites ofsamples: samples that are older than �50 Ma, show an isotopicrecord that is typical of the intra-oceanic history, whereas thosethat are younger than �50 Ma present a wide range of composi-tion that involves an old continental component in their source.The results of these two suites of samples are thus shown anddiscussed separately.

5.1.1. Pre-50 Ma samples

From 102.172.1 to 50.372.1 Ma, each granitoid sample fromthe southern part has a homogeneous U–Pb zircon age populationthat constrains the crystallization age (Fig. 2). Some samplescontain rare inherited zircon grains ranging in age from 94 to62 Ma (Table A.1). The inter-grain zircon 176Hf/177Hf variations ina given sample, overlap within analytical uncertainty, allowing us

Table 2Parameters used for the calculation of the four melt end-members of the mixing model. See Appendix for details.

Bulk Kd Concentrations

MORB Sediments Mantle Arc-LC KK-C Tethyan

MORB

10% melt Indian

Sediments

30% melt E-DMM

Mantle

20% melt Arc-LC 12%

melt

Sr 0.046 0.0459 0.0047 0.2000 2.06 127 889 87.9 265 9.72 47.70 240 810

Rb 0.016 0.0003 0.0003 0.0153 0.180 6.69 58.4 157 524 0.108 0.539 0.5 3.7

Sm 0.775 0.245 0.016 0.227 0.134 2.52 3.16 6.89 14.6 0.273 1.28 1 3.11

Nd 0.313 0.109 0.010 0.079 0.097 11.4 29.9 35.6 94.6 0.703 3.38 4 20.9

Lu 28.6 3.43 0.046 4.30 1.87 0.350 0.014 0.13 0.048 0.058 0.245 0.3 0.077

Hf 0.321 0.970 0.022 0.240 3.60 3.00 7.72 6.60 6.74 0.190 0.855 0.5 1.51

Initial isotopic compositions 87Sr/86Sr 0.70405 0.80000 0.70323 0.70400144Nd/143Nd 0.51273 0.51174 0.51285 0.51280177Hf/176Hf 0.28293 0.28200 0.28302 0.28310

0.5116

0.5118

0.5120

0.5122

0.5124

0.5126

0.5128

0.5130

0 2 4 6 8 10 12 14

143Nd/144Nd (i)

1/Nd*100 87Sr/86Sr (i)

0.704 0.706 0.708 0.710 0.712 0.714

Calculated melts derived from northern Indian margin sediments

MORBMelt

E-DMMMelt +

Fractionation effect

10%

20%

30%

40%

10%

20%

30%

MORBrange

20%30%

E-DMMMelt

SedimentMelt+

North

South

Central

Juvenile Evolved SedimentMelt

Arc-LCMelt +

20%1%

2%MORBrange

E-DMM

Arc-LCrange

E-DMM

10%

16

calculated melts derived from Karakoram margin rocks

≥ 50.3 Ma ≤ 50.4 Ma

≥ 46.3 Ma ≤ 31.7 Ma

≤ 39.9 Ma≥ 40.9 Ma

PK06-36E-DMM

MeltSediment

Melt+

Lu/ Hf

Hf/ Hf (i)

0 0.002 0.004 0.006 0.008

Source inheritedzircon grains

South North

Density distribution of zircon Hf isotopic and Lu/Hf compositions

from the Karakoram

Indian derived zircon grains

5%10%20%

30%

10%

20%

30%

20%

10%

E-DMM

0.2820

0.2825

0.2830

0.7200

0.7400

0.7600

0.7800

0 5 10 15 20 25

Sr/ Sr (i)

1/Sr*1000

MORBMelt

E-DMMMelt +

5%

50%

30%

20%

10%

MORBrange

SedimentMelt

Arc-LCMelt +

30%

E-DMMMelt

SedimentMelt+

40%

70%

40%

10% E-DMM

PK06-36

Fig. 4. Whole rock (a,b,c) and zircon mineral (d) isotopic compositions from this study. Lines represent mixing lines between calculated modal batch melts derived from E-

DMM, Tethyian type MORB, and Arc Lower crust (Arc-LC) and Indian sediments. Additionally shown are melts calculated in equilibrium with different potential source rocks

from Northern Indian margin (grey Squares) and the Karakoram (blue squares). (a) Initial 144Nd/143Nd composition vs. 1/Ndn100. (b) Initial 87Sr/86Sr composition vs. 1/

Srn1000. (c) Initial 144Nd/143Nd composition vs. initial 87Sr/86Sr. (d) Initial 176Hf/177Hf vs. 176Lu/177Hf in zircon grains. Only the grains used to determine crystallization age and

those inherited from the source are plotted. Indian grains are from the proterozoic sediments of the Indian margin of the Sutlej valley (Richards et al., 2005), Karakoram zircon

grains from Ravikant et al. (2009) and Heuberger et al. (2007). See Table 2 and Appendix for details. (For interpretation of the references to color in this figure legend, the

reader is referred to the web version of this article.)

P. Bouilhol et al. / Earth and Planetary Science Letters 366 (2013) 163–175168

to calculate an average eHf(i) for each sample (Fig. 2; Table A.1).The eHf(i) variability between samples shows a restricted range of6.271 to 12.171.1 (Fig. 2) with an average of 9.470.7. Inheritedolder zircon grains exhibit a range in eHf(i) comparable to thevariability between samples and are attributed to reworkingand assimilation of older arc-crust. The whole rock 143Nd/144Ndcompositions of the same samples show an eNd(i) that ranges

from 1.2 to 5, with a weighted mean of 2.670.7 (Table 1;Figs. 2 and 3), and an initial 87Sr/86Sr(i) ranging from 0.703744to 0.704719. The eHf(i), eNd(i) and 87Sr/86Sr(i) observed during thistime interval is within the range of the isotopic compositionpreviously reported for the KLA rocks of this age.

The modeling results of the isotopic and trace element char-acteristics of these rocks indicate that the observed variations are

15

10

5

0

εHf(i)

20 40 60 80 10030 50 70 90 110

Age Ma

-2

0

2

4εNd(i)

Single grainsKO8-04Pja-04

Inhomogeneous samples

0.7040

0.7045

0.7050

0.7055

0.7060

20 30 40 50 60 70 80 90 100 110

87Sr/ 86Sr(i)

Age Ma

Bulk rock composition

Zircon composition

Fig. 5. For the Central Batholith samples, U–Pb zircon crystallization ages of

granitoids vs. (1) their initial epsilon Hafnium (eHf(i)) composition of zircon grains

and (2) initial epsilon neodymium isotopic whole rock composition (eNd(i)) and (3)

initial Sr isotopic composition. Legend as in Figs. 2 and 3.

P. Bouilhol et al. / Earth and Planetary Science Letters 366 (2013) 163–175 169

adequately explained by a mixture of components derived fromthe E–DMM and from subducted oceanic lithosphere (Fig. 4). Theinvolvement of these components in their source, typical forisland arcs, indicates their juvenile character and their formationduring the intra-oceanic history of the arc. They show largevariability in the parent and daughter trace elements (i.e., wholerock Sm, Nd, Sm/Nd, Rb, Sr, Rb/Sr, and zircon Lu/Hf ratios)at relatively constant depleted isotopic compositions (i.e., high144Nd/143Nd(i) and 176Hf/177Hf(i) and low 87Sr/86Sr(i)) that couldonly be partially explained by magmatic fractionation (Fig. 4), ifthey were derived from a similar parental melt composition(Fig. 4). Additionally, the major element chemistry of most ofthese granitoids is relatively similar, indicating only limitedvariations related to fractionation. Variable degrees of partialmelting of one source component of the granitoids cannot explainthe observed variations in parent and daughter element concen-trations. The scatter in the isotopic composition of the dataorthogonal to this mixing line (Fig. 4) could indicate a minimalcontribution (o5%) of an evolved continental crust component,but can also be adequately explained by the observed variationsin E–DMM and MORB end-members (Mahoney et al., 1998;Workman and Hart, 2005; Fig. 4). Therefore, the variationsobserved are best explained by variations in the basaltic parentalmagma composition, which reflects changes in the mantle sourceregion of the basalts due to mixing of different components.

5.1.2. Post-50 Ma samples

In contrast, samples from the southern part that are youngerthan 50.471.6 Ma have more evolved isotopic compositionscompared to the 4�50 Ma samples. Three of the four younger

samples have heterogeneous zircon populations dominated byinherited grains (Fig. 2). The youngest grains that were used tocalculate the crystallization age often show significant inter-grain176Hf/177Hf variations of up to 10 eHf(i) in a given sample (Fig. 2).These zircon populations trend to significantly lower, radiogeniceHf(i) compositions, with eHf(i) as low as �15. The eNd(i) foro50 Ma samples is negative and varies over 6 eNd(i) (�9.7oeNd(i)o�3.6), associated with very radiogenic Sr isotopiccomposition of 87/86Sr(i) ranging from 0.705862 to 0.713170(Figs. 2 and 3). The youngest heterogeneous sample (LB09-11.5A, 29.670.8 Ma), having the most isotopically evolved com-position, contains numerous inherited zircon grains thatyield generally discordant Paleozoic 206Pb/238U dates between�434 Ma and �384 Ma and 207Pb/206Pb dates between �940and �474 Ma (Table A.1). These older grains have maximumPaleozoic/Proterozoic eHf(i) values between �7 and �2, compar-able to that of the young zircon grains. The age and isotopiccomposition of the inherited zircon grains are similar to thosefound in the Indian margin (Veevers and Saeed, 2009), but areabsent from the 450 Ma KLA rocks. A significant shift of �10eHf(i) and �6 eNd(i) units between the more unradiogenic andthe isotopically evolved samples is observed at 50.271.5 Ma,together with a significant increase in Sr isotopic composition.This distinct change is coeval within error with the onset of theIndia–KLA collision (Garzanti et al., 1987; Kaneko et al., 2003).

In contrast to the 4�50 Ma samples, these granitoids display astrong variability in their isotopic composition for their limitedvariations in the parent and daughter elements concentrations(Fig. 4). Modeling of the trace element and isotopic compositionsof these isotopically evolved rocks does not suggest direct involve-ment of a depleted mantle component in their formation (Fig. 4), butrather require the involvement of an additional enriched component,such as old isotopically evolved crust. Additionally, the isotopicallyevolved composition necessitates the involvement of an arc compo-nent such as the lower KLA crust (Fig. 4). The isotopic compositionsof these o50 Ma rocks from the southern transect are best explainedas mixtures of melts of Indian-derived sediments or the Indian crustand the lower KLA crust (Fig. 4).

The temporal concurrence of the change in isotopic composi-tions in the granitoids from the southern KLA and the onset of theIndia–KLA collision, combined with the inherited zircon grainsand the isotopic record of the evolved rocks, identify the sub-ducting Indian margin as contributing to the source of theo�50 Ma rocks. Therefore, the shift in isotopic compositions at50.271.5 Ma reflects a profound change of the source regions ofthe KLA magmas associated with the India–KLA collision. Theexpected delay between the initial contact of India with the KLAand the formation of evolved granites is unresolvable within ouranalytical uncertainty and thus less than �1.5 Ma, correspondingto �150–200 km of subducted Indian lithosphere.

5.2. The northern section

In contrast to the southern part, samples from the northernsection maintain a more depleted isotopic signature until �40 Ma(Fig. 2): Samples that are 440 Ma show an isotopic record thatdefines their juvenile character, whereas those that are o40 Ma haveisotopic characteristics that requires the presence of an old con-tinental component in their source. Therefore, like for the southernsection, the granitoids samples from the northern section arepresented and discussed in two steps.

5.2.1. Pre-40 Ma samples

From 8672.6 Ma to 40.971.1 Ma, these samples have homo-geneous zircon U–Pb age populations resulting in tightly

P. Bouilhol et al. / Earth and Planetary Science Letters 366 (2013) 163–175170

constrained U–Pb zircon crystallization ages for each sample. Rareinherited zircon grains range in age between 108 and 47 Ma. Thezircon eHf(i) of the samples studied does not vary significantly andhas a weighted mean eHf(i) of 10.771.4 (Fig. 2), similar to therange observed in the inherited zircon grains. The bulk-rock eNd(i)

of these samples varies between 3.5 and 5.7, and show a87Sr/86Sr(i) between 0.703582 and 0.704344 (Figs. 2 and 3), similarto the pre-50 Ma samples of the southern section. Isotopicmodeling confirms juvenile depleted mantle and basaltic oceaniccrust components in the source region (Fig. 4), similarly to the450 Ma rocks of the southern section.

5.2.2. Post-40 Ma samples

Contrastingly to the 440Ma samples, zircon grains fromgranitoids that are 39.971.4 and 37.271.1 Ma show a widereHf(i) range, from -5.271.0 to 10.172.3, and a slightly moreevolved whole-rock eNd(i) (e.g., 1.5 and 1.7) and 87Sr/86Sr(i)

(0.705017 and 0.704526; Figs. 2 and 3). These granites containnumerous inherited zircon grains of Mesozoic age with generallyconcordant U–Pb dates ranging from �235 to �103 Ma, andeHf(i) ranging from �6.471.4 to 1.471.4. The ages and isotopiccomposition of the Mesozoic inherited zircon grains are unknownfor 440 Ma rocks from the KLA and from the northern margin ofIndia (Veevers and Saeed, 2009).

Although more evolved isotopically than the 440 Ma rocks,these samples are significantly less radiogenic than the o50 Marocks from the southern section (Fig. 2). It is noteworthy, thatunlike the 440 Ma rocks, the observed compositions do notindicate direct involvement of an E–DMM mantle component intheir formation (Fig. 4). Importantly, the inherited Mesozoiczircon grains are identical in their Lu/Hf and eHf(i) isotopiccompositions to zircon grains from Mesozoic plutonic rocks fromthe Karakoram (Figs. 2 and 4), suggesting that the Karakorammargin contributed to the source of these rocks. Indeed, theKarakoram block, of continental origin, has a multi-phase historyinvolving Andean-type and rift related magmatism (Debon et al.,1987; Rolland et al., 2000, 2002a). During these different geody-namic episodes, significant volumes of mantle-derived meltswere emplaced in the Karakoram block. Compared to the Gond-wana derived basement (e.g., Crawford and Searle, 1992; Yin andHarrison, 2000) into which they intruded, this magmatism has anunradiogenic composition, thereby changing the bulk isotopiccomposition of the Karakoram Terrane as a whole.

The several episodes of Andean-type magmatism and riftinghave generated rocks with a wide range of isotopic compositionswithin the Karakoram block. The presence of rocks with similarnon-radiogenic Nd and Sr isotopic compositions as the o40 Magranitoids within the Karakoram indicate that they may havecontributed to their source. In Fig. 4 we have presented the datathat exist for the Karakoram crust to demonstrate its compositeand complex nature. Due to the lack of data for what would havebeen the Karakoram margin prior to collision, and considering thewide range of compositions found in the Karakoram, it remainsdifficult to choose a unique end-member for the modeling of theo40 Ma granitoids from the North. As no whole rock geochemicaldata from the Karakoram crust exists in the literature from whichboth Lu and Hf concentrations and Hf isotopic data are presented,we are not able to reliably model the Lu–Hf system. However, weconsider the inherited zircon record of these granitoids, which areidentical in U–Pb age, Hf isotopic composition and Lu–Hf ratio tothose recorded in Karakoram granitoids, as the strongest argu-ment for the involvement of the Karakoram margin in the sourceof isotopically evolved rocks in the North (Fig. 4d).

Therefore, the isotopic shifts occurring at 40.471.3 Ma alongthe Shyok suture is best explained by the involvement of

Mesozoic rocks from the Karakoram and not the northern Indianmargin. This interpretation is further supported by regional studiesthat have documented that the Karakoram margin was initiallyunderthrusted below the KLA (e.g., Coward et al., 1986).

Along the Indus suture, and observed in modern collisionalsystems (Section 3), an abrupt change from an unradiogenic to amore evolved isotopic signal in magmatic rocks closely corre-sponds to the onset of collision and suturing in continent–arccollision systems. We conclude that the abrupt shift in isotopiccomposition of approximately 16 eHf(i) units and 4 eNd(i) unitsobserved in granitoids in the northern KLA corresponds withinerror to the formation age of the Shyok suture, and dates thefinal collision between the now assembled India/KLA with theKarakoram at 40.471.3 Ma. This final collision age implies that aZ1000 km wide oceanic basin previously situated between theKLA system and the Karakoram margin was subducted afterthe initial India-Arc assembly, which is in agreement with thepaleomagnetic data (Khan et al., 2009) and in accordance with themagmatic and metamorphic history of the Karakoram (Searleet al., 2010).

Finally, the youngest granitoids, dated at 25.171.3 and21.571.2 Ma show the most isotopically evolved compositionin the Northern section. They have a heterogeneous zirconpopulation, showing a wide range of eHf(i) between �5.2 and1.2 (Fig. 2), together with inherited grains that are Precambrianand Mesozoic in age resembling those found in the Karakoramblock. The bulk-rock eNd(i) data for these samples are clusteredaround �6, and they show an evolved 87Sr/86Sr(i) composition(0.707422 and 0.709701; Figs. 2 and 3). These two samples havesimilar isotopic compositions as the Miocene ages granites fromthe Karakoram (Maheo et al., 2009). This Miocene post-collisionalmagmatism is not only present in the Karakoram, but alsothroughout the Tibetan Plateau, and has been widely attributedto the reworking of the Eurasian crust by the influx of hot mantlematerial, possibly related to a loss of crustal rock at depth (e.g.,Gao et al., 2010).

5.3. The central section

Samples dated between 72.171.4 Ma and 58.871.4 Ma haveaverage eHf(i) of 9.270.6 (Fig. 5). They have a whole rockeNd(i) and 87Sr/86Sr(i) falling within the same isotopic range(Fig. 5) of the other transect for similar ages, therefore referr-ing to the intra-oceanic history of the arc (1.7oeNd(i)o3.5;0.703781o87Sr/86Sr(i)o0.704507). Three samples dated between46.371.3 and 41.571 Ma are intermediate in isotopic composi-tion, showing slightly lower eHf(i) (3.970.5oeHf(i)o6.870.7)and eNd(i) (1.11) and more evolved 87Sr/86Sr(i) (0.705061) thanthe 458 Ma samples. These samples do not show the pro-nounced isotopic shifts observed along the southern transect. Aclear appearance of an old crustal evolved component in thesource is only demonstrated by the two youngest granites, datedat 31.770.8 and 30.171.3 Ma showing a wide range of eHf(i)

from 14.171 to 1.8571.3, eNd(i) ranging from �2.50 to 1.05 and87Sr/86Sr(i) from 0.705952 to 0.704959, after the collision of theIndia–Arc assembly with Karakoram (Figs. 4 and 5).

For the intermediate rock composition observed between �50and �40 Ma, we favor the explanation that the Indian lithospherewas initially steeply subducted into the mantle, and that the mantlewedge was still present below the central part of the arc system asshown by the isotopic composition of sample PK06-36 that involvesa mantle component (Fig. 4). This interpretation is supported by theP-T-t path recorded in the ultrahigh pressure metamorphism in thewestern Himalaya (Guillot et al., 2008), indicating that the Indiancontinent was initially subducted to 4150 km. These eclogitesstarted to be subducted at �50 Ma, reaching their metamorphic

P. Bouilhol et al. / Earth and Planetary Science Letters 366 (2013) 163–175 171

peak at 46 Ma, and were exhumed at 40 Ma (Kaneko et al., 2003;Parrish et al., 2006; Tonarini et al., 1993; Treloar et al., 2003). Theseobservations confirm that the subduction of buoyant crustal mate-rial, rather than the continent–continent collision, reduced theconvergence rate after 50 Ma (Capitanio et al., 2010).

5.3.1. Significance of source derived zircons

The first Indian derived zircon grains occur at �30 Ma in thesouthern section, contemporaneously with the crustal componentin the central section. As zircon crystals would most likely notsurvive in a crustal-derived melt that re-equilibrated with themantle wedge (Hanchar and Watson, 2003), the first appearanceof Indian derived zircon grains indicate the absence of a significantportion of the mantle wedge. This would be facilitated if the Tethyanoceanic slab was detached, allowing the Indian lithosphere to beunderthrusted. In this model, the appearance of Indian derivedzircon grains at �30 Ma constrains the timing of the underthrustingof the Indian Lithosphere underneath the KLA. This occurred�20 Ma later than the decrease in the convergence rate and thebeginning of the Indian Lithospheric subduction, and is in accor-dance with numerical models (van Hunen and Allen, 2011). Accord-ingly, for the Northern section, underthrusting, rather thansubduction, of the Karakoram margin would explain why Kara-koram derived zircon grains are immediately found at the time ofcollision, in contrast with the southern section.

5.4. Results in the framework of the Himalayan belt

Collectively, our data, are in good agreement with paleomagneticstudies (Khan et al., 2009), and document that the KLA formed as anintra-oceanic island arc south of the Eurasian margin until itsamalgamation to India at 50 Ma. This India–Arc collision has beenfollowed by the closure of the Shyok back-arc basin at 40 Ma,

~ 49-42 MaContinental slab

drives subduction

UHP peak: 46 Ma

50.2 ± 1.5 Ma:India - Arc suturing

Shyok1000-1200 km India/Arc Karakoram

> 50 Ma Intra-oceanic subduction

IndianLithosphere

Arc

oceanic lithosphere

KarakoramShyokTethys Ocean

Slabs remnants ?

~ 30 Ma underthrusting40.4 ± 1.3 Ma:India/Arc - Karakoram suturingFinal India-Eurasia collision

UHPexhumation

Fig. 6. (Left) Geodynamic sketch in a lithospheric cross section passing through the N

Map-view of the two possible large scale geodynamic scenarios discussed in the text.

corresponding to the final India/Arc–Karakoram collision in thenorthwestern Himalaya (Fig. 6). In the framework of the broaderevolution of the India–Eurasia collision, our results can be inter-preted in two main geodynamic scenarios (Fig. 6): (1) the KLA waspart of a lateral extensive chain of island arcs situated in theequatorial part of the Tethys and correlate to the supra-subduc-tion zone ophiolites associated with the Tsangpo suture zone(Aitchison et al., 2000; Hebert et al., 2012). In this scenario a secondsubduction zone was active to the north of the intra-oceanic system,along which the Shyok Sea was subducting underneath the Eurasianmargin producing the Karakoram and Gandgese continental arcs. Inthis model, the India–Arc collision occurred thorough the lengthof this chain at �50 Ma, and the final India/Arc–Eurasia collisionoccurred only at �40 Ma; alternatively, (2), the KLA and theGangdese formed along the same subduction zone (Yin andHarrison, 2000), similar to the modern Alaska–Aleutian subductionsystem. In this scenario the Shyok ‘‘Basin’’ formed as a marginal basinbehind the KLA, and terminated toward the east, but could haveextended further toward the west (Fig. 6). In this case, to explain theKarakoram continental magmatism, a second subduction zone northof the KLA–Gangdese system must have existed in the west, wherethe Shyok Basin was subducted underneath the Karakoram marginthat has no continuation toward the east. In this model the India–Arccollision recorded in NW Himalaya at 50 Ma would correspond tothe main collision to the East, and the 40 Ma reflects the closure ofthe marginal Shyok Basin. The subsequent discussion focuses onthree main observations that have been used to constrain collisionand might help to decipher between the two different scenarios.

5.4.1. Origin of the Tsangpo suture ophiolites

A fundamental differences between these two scenarios concernsthe origin of the frequent ophiolites along the Tsangpo suture zonethat formed in a supra-subduction zone setting (e.g., Dai et al., 2011;

Eurasia

GangdeseKarakoram

Shyok Sea

KLA

India

Eurasia

GangdeseKLA

Shyok Basin

Karakoram

India

50 Ma

Eurasia

GangdeseKarakoram

Shyok Sea

KLA

Tethys

Scenario 1Eurasia

GangdeseKLA

Shyok Basin

Karakoram

Tethys

Scenario 2N> 50 Ma

Eurasia

GangdeseKarakoram KKF(?)

KLA

India

KLA

Eurasia

GangdeseKarakoram KKF(?)

KLA

India

40 Ma

W Himalaya representing the two collisional events recorded in our data. (Right)

KKF¼Future (?) location of the Karakoram Fault.

Post-Ypresian

1

2

3

4

5

6

5

6

7

Relative probability %

Ophiolite (Arc remnants)n=38

KLAn=1251

Lhassan=1893

1

2

3

4

7

5

10

P. Bouilhol et al. / Earth and Planetary Science Letters 366 (2013) 163–175172

Hebert et al., 2012) (Fig. 6). The few reliable zircon U–Pb agesranging from upper Jurassic to upper Cretaceous and Nd–Smisochrons indicate that the main crustal forming processes for theseophiolites was �120–125 Ma (review in Hebert et al., 2012),contemporaneous with the Gangdese arc magmatism. In the classi-cal model (scenario 2) these ophiolites are considered to representthe fore-arc of the Gangdese arc (e.g., Xia et al., 2003). Thisinterpretation is however at odds with several geological observa-tions: (1) many of these ophiolites are associated with deep seashales, limestones, and radiolarians indicating that at least some ofthese rocks formed far from any continental source (e.g. Ziabrevet al., 2003). (2) Different ophiolitic bodies are considered torepresent fore-arc (e.g., Luobuza), intra-oceanic arc (e.g., Jinlu) andback-arc (e.g., Yungbwa), constituting a complete supra-subductionzone system (Dubois-Cote et al., 2005; Hebert et al., 2012). If theserocks are indeed associated with the same subduction system as theGangdese arc, that would indicate that there would have been asecond volcanic chain in the fore-arc, separated by a ‘‘back arc’’ basinfrom the main continental arc. No currently active continentalmargin subduction zone system shows such a double volcanicchain; (3) finally, paleomagnetic data suggest that these ophioliteswere in an equatorial position in the mid Cretaceous (Abrajevitchet al., 2005), approximately at the same latitude as the KLA (Khanet al., 2009), but �2500 km south of the Gangdese arc that formedat �25 N (Dupont-Nivet et al., 2010; Tan et al., 2010; vanHinsbergen et al., 2012).

40 60 80 100 120 140 160 180 200

Zircon 238U/206Pb date Ma

Sedimentsn=1354

1

2

3

4

Pre-Ypresian Sedimentsn=1091

Fig. 7. Comparison of U–Pb Zircon age histograms from different terranes:

(a) zircons from the KLA (This study and Bosch et al., 2011; Bouilhol et al.,

2011; Heuberger et al., 2007; Jagoutz et al., 2009; Ravikant et al., 2009;

Schaltegger et al., 2002) and from the plutonic rocks of ophiolitic arc remnants

(Malpas et al., 2003; McDermid et al., 2002; Pedersen et al., 2001; Wei et al.,

2007); (b) zircon grains from Lhasa (Chu et al., 2006; Ji et al., 2009; Zhu et al.,

2009a, 2009b, 2011); and (c) zircon record from pre- and post-Ypresian sediments

in the Tethyan Himalaya (Aitchison et al., 2011; Henderson et al., 2010, 2011; Hu

et al., 2010, 2012; Wang et al., 2011; Wu et al., 2010). Only zircon grains Z40 and

o200 Ma are shown.

5.4.2. Detrital zircon record

The appearance of Eurasian derived detritus (e.g., detrital zircongrains) within the Tethyan Himalayan series is used to constrain thetiming of the India–Eurasia collision. In this approach, the miner-alogy of the sediments is used to constrain the source regions, andrecent studies have focused on a pronounced shift in U–Pb agespectra of the detrital zircons at the Ypresian. Whereas pre-Ypresiansamples show a rather simple age spectra dominated by a �120 Maage peak, post-Ypresian rocks show an age spectra characterized bya peak at �55–60 Ma and a broad double peak at �90–100 Ma(Fig. 7). The source of these zircons grains is actively debated:(1) one model considers these zircon crystals to be derived from theintra-oceanic arc system (e.g., Aitchison et al., 2011), in accordancewith our proposed scenario 1 above; whereas, (2) others considerthese zircons to be derived from Eurasia (e.g., Najman et al., 2010the scenario 2; Wang et al., 2011). The previous studies are mainlybased on apparent matches between age peaks observed in theGangdese arc, the ophiolitic remnants, and in the detrital zircon agespectra. Because the dataset for the arc was very limited prior to ourstudy, and that the extent and origin of the sedimentary units is stilldisputed (Aitchison et al., 2011; Clift et al., 2002; Henderson et al.,2011; Wu et al., 2010), we re-evaluate the two possible scenarios bycomparing in Fig. 7, the detrital zircon record from pre- and post-Ypresian Tethyan sediments with the igneous zircon record of theEurasian Margin, and our new data from the KLA and other arcfragments found within and south of the Tsangpo suture zone.

Probably the most obvious conclusion from Fig. 7 is that neitherthe data set from the Eurasian margin, nor the intra-oceanic arcsystem, fully matches the detrital zircon age record. The 55–60 Mapeak in the post-Ypresian samples is present in both the oceanic arcand continental margin dataset, thus inconclusive. A lack of zirconsdates �75 Ma in the continental margin dataset could mirror thefew zircon crystals of that age in the post-Ypresian samples. Thepeaks at �90–100 Ma in the detrital zircon are very well repre-sented in the intra-oceanic dataset but not recorded in the con-tinental margin.

The correlation of the detrital U–Pb zircon record to igneousactivity recorded in the source region is complicated. For example,

the preservation likelihood of less mature oceanic arc systems is lowand it is possible that part of the source region of the zircon grainswere subducted together with the leading edge of India duringthe final India–Eurasia collision (Boutelier and Chemenda, 2011;Boutelier et al., 2003; Burg, 2006). Additionally, it is in generaldifficult to compare igneous U–Pb zircon dates to detrital U–Pbzircon dates because it is unknown if age peaks are truly related tolarge volumes of temporally constrained igneous pulses. Finally, agepeaks could be influenced by the analytical method used which relyon significantly different number of analyses, possibly resulting in astrong bias in ‘‘age peaks’’ toward analyses done by LA-ICPMS.Altogether, this indicates that an argumentation based on thecomparison of age peaks between the Eurasian margin, the arc,and the detrital zircon alone, does not allow deciphering the twomain geodynamic scenarios.

5.4.3. Change from localized calk-alkaline magmatism to

widespread alkaline magmatism

The continental margin calk-alkaline magmatism recorded in theLhasa terrane, that formed the Gangdese batholith, as well asthe Linzizong volcanics, ended at �40 Ma (e.g., Wen et al., 2008).

P. Bouilhol et al. / Earth and Planetary Science Letters 366 (2013) 163–175 173

Subsequently, the magmatic activity switched to widespread alka-line dominated intra-plate magmatism distributed throughout theTibetan plateau (Ding et al., 2007). This change in the location andchemical characteristics of magmatism at �40 Ma in the TibetanPlateau is readily explained by the geodynamic scenario 1 above.In this model the calk-alkaline magmatism is related to oceanicsubduction, whereas widespread alkaline magmatism is post-collisional in accordance with the frequent interpretation of alkalinemagmatism worldwide. However, in the case of tectonic scenario2 above, the switch in magmatism at �40 Ma can only be poorlyexplained. For example, an ad hoc slab break-off model, where theoceanic slab detached from the subducting Indian continentallithosphere right after the onset of collision, has been proposed toexplain the presence of the �50–40 Ma mantle-derived melts withsubduction characteristics (e.g.,Wen et al., 2008). Yet, numericalmodel results indicate that slab break-off is unlikely to occursimultaneously with collision (van Hunen and Allen, 2011) andthe magmatic record where slab tears and break-offs currentlyoccurs, e.g., in the Mediterranean region, is generally highly diverseand not ‘‘simple’’ calk-alkaline melts.

Based on these observations and the fact that both the Gangdeseand Karakoram batholith record subduction related magmatic activ-ity until �40 Ma (Searle et al., 2010; Wen et al., 2008), we favor ageodynamic scenario involving two subduction zone systems (sce-nario 1; Fig. 6) as the one most compatible with our currentunderstanding of the broader geology (e.g. Aitchison and Davis,2004). We speculate that the KLA and the opholitic fragments ofthe Tsangpo were part of the same intraoceanic arc chain present inthe Neo-Tethys south of the Eurasian margin. The multiple lowvelocity zones observed in the upper mantle and interpreted as slabremnants could reflect these two subduction systems (Hafkenscheidet al., 2006; Replumaz et al., 2010; Van der Voo et al., 1999). In thismodel the extension toward the west of what we call the Shyok Sea,between the two subduction systems, could be the Cretaceous–Miocene Pishin sedimentary belt (Khan et al., 2012) that have beenobserved behind the ophiolites in the western extension of the KLA(Kakar et al., 2012).

Our scenario implies that two suture zones may be present in thecentral and eastern Himalaya that have not been fully identified sofar. This could stem from the fact that both suture zones are nowrepresented as the apparent single Tsangpo suture, that often showduplications (e.g. Gansser, 1980), and are not yet resolved. Alter-natively, an unidentified major suture zone exits north or south ofthe Tsangpo: either within the Lhasa Terrane, but there are noevidence for a major fault zone and/or tertiary oceanic sediments;or, included in the poorly understood Tibetan Himalaya sequencethat is composed of highly variable sedimentary units of differentorigin (e.g. Burg and Chen, 1984).

6. Conclusions

We have presented results used to reconstruct the temporal andspatial isotopic record of the Kohistan–Ladakh Paleo-Island Arcgranitoids, and showed that India collided first with the Arc at50.271.5 Ma and that the final India/Arc–Eurasia collision occurredat 40.471.3 Ma in northwestern Himalaya. The geological record ofthe whole belt is in accordance with a scenario of a final India–Eurasia collision at �40 Ma. Considering a post 40 Ma Indian platevelocity of 50 mm per annum (Molnar and Stock, 2009), the totallength of the Indian lithosphere accommodated in the orogen sincethe final collision is �2000 km. This estimate is �1000–1500 kmless than the convergence that would have had to be accommodatedif collision occurred at 50 Ma (e.g., �3000–3500 km). A total post-collision India–Eurasia convergence of �2000 km resolves a long-lasting debate that pertains to the importance of different

deformation processes that accommodated convergence (e.g.,Molnar and Tapponnier, 1975; Royden et al., 2008; Tapponnieret al., 1982). Estimates on the shortening recorded in the Himalayaand Eurasia (DeCelles et al., 2002; van Hinsbergen et al., 2011) arebroadly consistent with the reduced amount of convergence due toour 40 Ma collision scenario.

Acknowledgments

This research was supported by the National Science FoundationEAR 0910644 and the Swiss National Fund PBEZP2-122848. We aregrateful to J.P. Burg and P. van der Beek for sharing samples. Z. Zhaoand P. Kapp provided useful reviews and we thank T.M. Harrison forhis editorial handling of the manuscript.

Appendix A. Supplementary materials

Supplementary data associated with this article can be found inthe online version at http://dx.doi.org/10.1016/j.epsl.2013.01.023.

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