Geological Society, London, Special Publications Monsoon control over erosion patterns in the Western Himalaya: possible feed-back into the tectonic evolution Peter D. Clift, Liviu Giosan, Andrew Carter, Eduardo Garzanti, Valier Galy, Ali R. Tabrez, Malcolm Pringle, Ian H. Campbell, Christian France-Lanord, Jurek Blusztajn, Charlotte Allen, Anwar Alizai, Andreas Lückge, Mohammed Danish and M.M. Rabbani Geological Society, London, Special Publications 2010; v. 342; p. 185-218 doi:10.1144/SP342.12 Email alerting service click here to receive free email alerts when new articles cite this article Permission request click here to seek permission to re-use all or part of this article Subscribe click here to subscribe to Geological Society, London, Special Publications or the Lyell Collection Notes Downloaded by Massachusetts Institute of Technology on 3 September 2010 © 2010 Geological Society of London Monsoon control over erosion patterns in the Western Himalaya: possible feed-back into the tectonic evolution PETER D. CLIFT1*, LIVIU GIOSAN2, ANDREW CARTER3, EDUARDO GARZANTI4, VALIER GALY2, ALI R. TABREZ5, MALCOLM PRINGLE6, IAN H. CAMPBELL7, CHRISTIAN FRANCE-LANORD8, JUREK BLUSZTAJN2, CHARLOTTE ALLEN6, ANWAR ALIZAI1, ANDREAS LÜCKGE9, MOHAMMED DANISH5 & M.M. RABBANI5 1 School of Geosciences, University of Aberdeen, Aberdeen, AB24 3UE, UK 2 Woods Hole Oceanographic Institution, Woods Hole, MA 02543, USA 3 School of Earth Sciences, University and Birkbeck College London, Gower Street, London, WC1E 6BT, UK 4 Dipartimento Scienze Geologiche e Geotecnologie, Universita’ di Milano-Bicocca, Piazza della Scienza 4 – 20126 Milano, Italy 5 National Institute for Oceanography, Clifton, Karachi 75600, Pakistan 6 Department of Earth, Atmospheric and Planetary Sciences, Massachusetts Institute of Technology, Cambridge, Massachusetts, USA 7 Research School of Earth Sciences, The Australian National University, Canberra, A.C.T. 0200, Australia 8 CRPG-CNRS, BP 20, 15 rue Notre Dame des Pauvres, 54501 Vandoeuvre les Nancy, France 9 Bundesanstalt für Geowissenschaften und Rohstoffe (BGR), Stilleweg 2, D-30655 Hannover, Germany *Corresponding author (email: p.clift@abdn.ac.uk) Abstract: The Indus Delta is constructed of sediment eroded from the western Himalaya and since 20 ka has been subjected to strong variations in monsoon intensity. Provenance changes rapidly at 12–8 ka, although bulk and heavy mineral content remains relatively unchanged. Bulk sediment analyses shows more negative 1Nd and higher 87Sr/86Sr values, peaking around 8 –9 ka. Apatite fission track ages and biotite Ar–Ar ages show younger grains ages at 8– 9 ka compared to at the Last Glacial Maximum (LGM). At the same time d13C climbs from – 23 to – 20‰, suggestive of a shift from terrestrial to more marine organic carbon as Early Holocene sea level rose. U–Pb zircon ages suggest enhanced erosion of the Lesser Himalaya and a relative reduction in erosion from the Transhimalaya and Karakoram since the LGM. The shift in erosion to the south correlates with those regions now affected by the heaviest summer monsoon rains. The focused erosion along the southern edge of Tibet required by current tectonic models for the Greater Himalaya would be impossible to achieve without a strong summer monsoon. Our work supports the idea that although long-term monsoon strengthening is caused by uplift of the Tibetan Plateau, monsoon-driven erosion controls Himalayan tectonic evolution. Supplementary material: A table of the population breakdown for zircons in sands and the predicted Nd isotope composition of sediments based on the zircons compared to the measured whole rock value is available at http://www.geolsoc.org.uk/SUP18412 The relationships between climate, continental erosion and mountain building continue to be debated and are central to understanding how the solid planet and its atmosphere have interacted over long periods of geological time. In particular, the links between mountain building in Cenozoic Asia and the intensification of the monsoon are presently unclear, with the different processes feeding back on each other. Although climate modelling has shown that a wide, high Tibetan Plateau is important to From: Clift, P. D., Tada, R. & Zheng, H. (eds) Monsoon Evolution and Tectonics –Climate Linkage in Asia. Geological Society, London, Special Publications, 342, 185–218. DOI: 10.1144/SP342.12 0305-8719/10/$15.00 # The Geological Society of London 2010. 186 P. D. CLIFT ET AL. maintaining a strong monsoon circulation (An et al. 2001; Kitoh 2004) it is also clear that monsoon strength has varied greatly over millennial to orbital timescales, driven by solar factors either directly or via its influence on the intensity of northern hemispheric glaciation (Clemens et al. 1991; Wang et al. 2005). Feedbacks may work in both direction however. Recent geomorphological and thermochronometric work indicates that the location of active faulting in mountains and thus orogenic architecture is controlled by climate zonation (Hodges et al. 2004; Wobus et al. 2003), at least as much as by plate tectonic processes. Both wedge and channel-flow models proposed to explain the origin of the Greater Himalaya (Nelson et al. 1996; Robinson et al. 2003; Harris 2007) require focused erosion driven by monsoon rains to allow exhumation of deeply buried metamorphic rocks. If we are to understand what influence climate alone has on erosion (and thus on tectonics) then we need to isolate the climate signal from tectonic overprints. One way to do this is to study changes in erosion on timescales that are too short to allow tectonism to be an important influence. In this study we assess the response of the Indus river basin to the intensification of the South Asian monsoon since the Last Glacial Maximum (LGM), around 20 ka. Cave records in Arabia (Fleitmann et al. 2003), together with Indian lake sediments (Enzel et al. 1999) and marine cores (Staubwasser et al. 2002) now show that the summer monsoon strengthened as global climate warmed between the LGM (c. 20 ka) and the Early Holocene c. 8 ka. Earlier sedimentological work in the Bengal delta demonstrated that the intensification of the summer monsoon correlated with a great increase in the rate of sediment delivery to the delta during the Early Holocene, c. 8 ka (Goodbred & Kuehl 2000). This is suggestive of greatly enhanced erosion and sediment transport rates in the source regions, presumably driven by the influence of stronger summer monsoon rains, as the speed of response rules out a tectonic trigger. Indeed studies of landsliding in the western Himalaya show that large scale mass wasting correlates with periods of stronger monsoon (Bookhagen et al. 2005), consistent with the delta records. Recently we reported that the Indus delta also prograded seawards during the onset of the Holocene (Giosan et al. 2006), that is, at a time of rapid sea level rise. This observation requires a major increase in the flux of sediment to the delta to balance the increasing accommodation space caused by sea level rise. Crucially, the delta was prograding southwards during the 8–12 ka period when rates of eustatic sea level rise were greatest (Camoin et al. 2004). Most recently, Clift et al. (2008) used a combination of bulk sediment Nd isotopes, single grain U –Pb zircon ages and Ar –Ar muscovite mica ages to argue for a sharp increase in the relative erosional flux from the Lesser Himalaya during the Early Holocene. In this paper we test their hypothesis that intensification of the summer monsoon resulted in much heavier rain along the southern edge of the Greater Himalaya and especially over the Lesser Himalaya. We do this using a series of additional geochemical proxies to assess the source of sediment at any given time and the changing environmental conditions in the Indus basin, which can then be compared to established climate records. Sampling Samples were taken from four boreholes drilled in the delta (Fig. 1). Although total recovery was not high material was recovered from most parts of the drilled sections, thus allowing a relatively continuous erosion record to be reconstructed (Fig. 2). At Keti Bandar drilling penetrated the ravinement surface that forms the base of the modern Indus delta and recovered Pleistocene sand deposited during the LGM (Clift et al. 2008). At all other sites only the Holocene to Younger Dryas sections were recovered. Ages of deposition were calculated from accelerator mass spectrometer (AMS) 14C dating of organic materials made at the National Ocean Sciences Accelerator Mass Spectrometry facility (NOSAMS, Woods Hole Oceanographic Institution) (Clift et al. 2008). Radiocarbon ages from below the deltaic sediments were 28.7 and 38.9 ka, suggesting reworking and mixing of older sediment prior to transgression after 20 ka. The Keti Bandar section shows two coarsening-upward cycles, separated by a transgressive mud deposited after c. 8 ka (Fig. 2). The cores were sampled both for sands, which were mostly used for single grain thermochronological methods, and for clays that were used for organic carbon analysis. Analytical methods A number of different provenance methods were used in order to establish a matrix of constraints, since typically one or two methods were insufficient to define a sediment source area for any given sand sample. In this study, we analysed the sediments using classical sand petrography, Sr isotopes, apatite and zircon fission track, Ar –Ar single biotite grain dating and additional U –Pb zircon dating, beyond the samples already considered by Clift et al. (2008). In addition, we selected clay samples for organic carbon isotope geochemistry in order to constrain the general composition of the vegetation in the Indus basin. Here we describe MONSOON CONTROL OVER EROSION HK 187 KK K NP Tarbela Dam LA 35° N ZA Islamabad m Ch llu vi Ra ks Je GH ali Siw b a en STD LH MBT MCT Garhwal MFT Sutlej us Ind rt Sukkur ar De se Thatta Keti Bandar Th Karachi 30° N Gularchy Jati 25° N Arabian Sea 70°E 75°E Fig. 1. Shaded topographic map of the Indus drainage basin, showing the location of the boreholes in the delta, the major tributaries of the modern river and the principal sediment source terrains (HK, Hindu Kush; K, Kohistan; NP, Nanga Parbat; ZA, Zanskar; LA, Ladakh; KK, Karakoram; GH, Greater Himalaya; LH, Lesser Himalaya; MBT, Main Boundary Thrust; MFT, Main Frontal Thrust; MCT, Main Central Thrust; STD, South Tibet Detachment) [from Clift et al. (2008)]. White dots indicate boreholes. Black dots with white rim indicate place name mentioned in the text. Reproduced with the permission of the Geological Society of America. the methods employed to reconstruct the evolving character of the Indus delta. Mineralogy All nine samples analysed for petrography from the Keti Bandar and Thatta cores were very finegrained sands containing variable amounts of bioclasts. In each sample, 400 points were counted by the Gazzi-Dickinson method (Ingersoll et al. 1984). Data are presented in Tables 1 and 2. Traditional ternary parameters and plots (Dickinson 1985) were supplemented, specifically as lithic grains are concerned, by an extended spectrum of key indices. Metamorphic rock fragments were classified according to both composition and metamorphic rank, mainly inferred from degree of recrystallization of mica flakes (Garzanti & Vezzoli 2003). Very low to low-rank metamorphic lithics, for which protolith can still be inferred, were subdivided into metasedimentary (Lms) and metavolcanic (Lmv) categories. Medium to high-rank metamorphic lithics were subdivided instead into felsic (metapelite, metapsammite, metafelsite; Lmf) and mafic (metabasite; Lmb) categories. 188 P. D. CLIFT ET AL. Fig. 2. Diagram shows the stratigraphy of the sediment cored at each of the drilling sites considered in this study, as shown in Figure 1. Each log shows the depth, proportion of the section recovered by drilling (black equates to recovery), the age control points constrained by 14C AMS dating of organic materials, as well as the location of samples selected from isotopic and thermochronological analysis. TD, total depth. Reproduced with the permission of the Geological Society of America. Heavy minerals were separated from the veryfine to fine sand fraction (63–250 mm), treated with acetic acid and sodium dithionite, by centrifuging in sodium metatungstate (density c. 2.9 g cm23). From each sample, 200– 250 transparent heavy minerals were counted in grain mounts by the ‘ribbon-counting’ or ‘Fleet’ methods (Mange & Maurer 1992). The abundance of total (transparent þ opaque þ turbid) and transparent heavy minerals in the sediment were expressed by the ‘Heavy Mineral Concentration’ (HMC) and ‘transparent Heavy Mineral Concentration’ (tHMC) indices (Garzanti & Andò 2007). Sr isotope analysis Samples from both the Keti Bandar and Gularchy boreholes were sampled for Sr analysis. Sr analysis was only performed on samples that had previously been analysed for Nd isotopes (Clift et al. 2008). Sr can be used for provenance work, but because it is mobilized during chemical weathering is less definitive than Nd isotopes. However, this characteristic can be used to effect because Sr variability in the absence of provenance changes has been interpreted to reflect changing intensities of chemical weathering (Palmer & Edmond 1991; Blum & Erel 1995). In particular, the Sr isotope composition of clay reflects the Sr characteristics of the groundwater in which they were formed, rather than the composition of the source rock (McBride 1994; Derry & France-Lanord 1996). Assuming that the clays have not been affected by diagenesis (unlikely in such young materials) the Sr isotopes can be used to track the evolving intensity of silicate weathering. However, in this setting the Nd isotope composition is known to change especially at 14 –8 ka and so changes in Sr at those times may be mostly provenance driven. Samples were accurately weighed into teflon screw-top beakers and dissolved using HF-HNO3HCl. Sr samples were separated in 2.5 N HCl Table 1. Bulk mineralogy data for selected sands from the Keti Bandar and Thatta boreholes derived from point counting of microcopic grain mounts Site n/a KB-5-2 TH-10-8 KB-20-1 KB-23-2 KB-26-2 KB-30-1 KB-34-4 KB-40-5 KB-41-2 Delta Keti Bandar Thatta Keti Bandar Keti Bandar Keti Bandar Keti Bandar Keti Bandar Keti Bandar Keti Bandar Depth (m) Age (ka) Mean grain size (micron) Q KF P 15 31 58 67 76 88 102 118 120 0 0.18 7.14 9.21 9.65 10.64 12.26 14.23 20.00 20.00 119 82 93 87 39 51 38 73 85 82 41 40 38 46 40 36 37 41 39 46 7 10 11 10 10 8 11 5 12 8 9 12 15 10 10 8 8 15 14 14 Lvf Lvm Lcc Lcd Lp Lch Lms Lmv Lmf Lmb Lu Mu Bi 0 0 1 1 0 0 2 1 1 1 0 0 0 0 0 0 0 0 1 1 8 8 6 11 6 7 7 6 0 0 3 3 3 3 1 2 2 2 4 4 3 1 1 1 2 0 1 5 1 0 1 0 0 0 0 0 0 1 0 0 7 3 4 4 4 2 2 5 4 6 1 1 4 1 1 1 1 4 2 2 3 5 5 5 5 2 3 4 5 5 2 0 1 0 0 0 0 0 1 0 0 0 0 0 0 0 0 0 0 0 1 4 2 2 4 7 6 4 4 2 2 11 8 5 15 25 19 8 9 7 Dense Total minerals 11 3 2 2 1 2 1 0 3 2 100 100 100 100 100 100 100 100 100 100 MONSOON CONTROL OVER EROSION Sample Abbreviations: Q, quartz; KF, K-felspar; P, plagioclase; Lvf, lithic felsic volcanic; Lvm, lithic mafic volcanic; Lcc, calcite carbonate lithic; Lcd, dolomite carbonate lithic; Lp, shale/silt; Lch, chert; Lmv, metavolcanic lithics; Lmf, felsic metamorphic lithics; Lmb, basic metamorphic lithics; Lu, ultramafic lithics; Mu, muscovites; Bi, biotites. 189 190 Table 2. Heavy mineral data for selected sands from the Keti Bandar and Thatta boreholes derived from point counting of microcopic grain mounts Sample Depth (m) Age (ka) % of heavy minerals in 63 – 250 mm fraction % of heavy minerals transparent Delta Keti Bandar Thatta Keti Bandar Keti Bandar Keti Bandar Keti Bandar Keti Bandar Keti Bandar Keti Bandar n/a KB-5-2 TH-10-8 KB-20-1 KB-23-2 KB-26-2 KB-30-1 KB-34-4 KB-40-5 KB-41-2 0 14 30 58 67 77 78 101 118 120 0 0.18 7.14 9.21 9.65 10.64 10.77 14.20 20.00 20.00 9.6 3.0 4.9 3.9 0.9 1.3 1.4 4.0 4.3 4.3 8.0 2.7 3.9 3.4 0.7 1.1 1.2 3.2 3.8 3.6 % % % Total zircon tourmaline rutile sphene transparent opaque turbid 83 91 79 87 83 88 84 81 88 84 10 1 6 3 2 1 4 3 5 3 8 9 15 10 15 11 12 16 8 13 100 100 100 100 100 100 100 100 100 100 0 0 0 1 0 0 0 0 0 2 2 3 2 3 2 2 2 3 2 0 1 1 0 1 2 0 1 0 0 0 bluegreen greenbrown/ green hornblende brown red hornblende hornblende hornblende 2 2 0 3 2 4 1 1 3 1 39 45 33 39 34 38 34 36 47 51 4 1 5 2 4 5 4 2 0 4 3 2 2 4 0 1 2 1 3 2 2 0 0 0 0 0 0 0 0 0 Site Sample tremolite actinolite green augite diopside/ hedembergite enstatite hypersthene epidote clinozoisite zoisite other epidotes chloritoid garnet staurolite kyanite sillimanite Total Delta Keti Bandar Thatta Keti Bandar Keti Bandar Keti Bandar Keti Bandar Keti Bandar Keti Bandar Keti Bandar n/a KB-5-2 TH-10-8 KB-20-1 KB-23-2 KB-26-2 KB-30-1 KB-34-4 KB-40-5 KB-41-2 0 0 0 0 2 2 1 0 0 0 2 3 8 6 13 7 8 5 3 4 0 0 0 0 0 0 0 2 0 0 2 4 7 5 4 4 8 4 0 1 0 0 0 0 0 0 1 0 0 0 1 3 2 2 1 1 1 1 0 1 21 28 22 26 19 25 28 25 32 24 1 1 1 1 1 0 0 0 0 0 3 0 0 1 0 2 0 1 1 1 0 0 0 1 0 0 0 1 1 0 0 0 0 2 1 0 0 0 0 0 12 2 13 4 5 5 3 11 4 4 1 1 0 0 1 1 0 1 0 0 1 1 1 3 3 1 0 0 0 1 1 0 1 1 2 1 0 2 0 0 100 100 100 100 100 100 100 100 100 100 P. D. CLIFT ET AL. Site MONSOON CONTROL OVER EROSION using Bio-Rad AG50W X8 200-400 mesh cation exchange resin using standard column methods. They were analysed on a by Finnigan ‘Neptune’ multi-collector inductively coupled plasma mass spectrometer (MC-ICP-MS) at Woods Hole Oceanographic Institution. Sample measurements were normalized to 86Sr/88Sr ¼ 0.1194 and referenced to a value of 0.710240 for NBS987 standard. Results are provided in Table 3. Organic carbon analysis Total organic carbon content (TOC) and stable isotopic composition of the bulk organic matter (d13C) were determined for 20 samples taken exclusively from the Keti Bandar borehole (Fig. 1). A detailed description of the analytical method can be found in Galy et al. (2007a). Indus delta sediments contain significant amounts of detrital carbonates including dolomite. TOC and isotopic measurements must therefore be performed on decarbonated sediments. Efficient dolomite dissolution was achieved through 1 hr leaching with 4 wt% HCl at 80 8C. After 50 8C oven drying, the organic carbon content and stable isotopic composition were determined by EA-IRMS (elemental analysis – isotope ratio mass spectrometry). organic carbon solubilization during acid treatment was taken into account following the approach described by Galy et al. (2007a). The overall 2s uncertainty associated with the TOC and d13C determination is respectively 0.02% and 0.25‰. Results of the carbon isotope analysis are provided in Table 4. Major and trace element analysis Major and trace element concentrations for the muds which were also measures for carbon isotopes were measured respectively by Inductively Coupled Plasma Atomic Emission Spectrophotometry (ICP-AES) and Inductively Coupled Plasma Mass Spectrometry (ICP-MS) at the Service d’Analyse des Roches et des Minéraux (CRPG-Nancy, France) on bulk sediment after lithium metaborate fusion. Results of the element analysis are provided in Table 4. Fission track analyses In this study we employed the fission-track method applied to apatite, which records cooling through c. 125 –60 8C over timescales of 1– 10 Ma and to zircon which records cooling through c. 200 8C (Green 1989). Fission-track analysis was performed at University College, London, UK using five samples, one of which, dating from the LGM, was analysed for both apatite and zircon. Polished grain mounts were etched with 5N HNO3 at 20 8C 191 for 20 seconds to reveal the spontaneous fission tracks. Subsequently, the uranium content of each crystal was determined by irradiation, which induced fission of 235U. The induced tracks were registered in external mica detectors. The samples for this study were irradiated in the thermal facility of the Hifar Reactor at Lucas Heights, Australia. The neutron flux was monitored by including Corning glass dosimeter CN-5, with a known uranium content of 11 ppm, at either end of the sample stack. After irradiation, sample and dosimeter mica detectors were etched in 48% HF at 20 8C for 45 minutes. Only crystals with sections parallel to the c-axis were counted, as these crystals have the lowest bulk etch rate. To avoid biasing results through preferred selection of apatite crystals, the samples were systematically scanned and each crystal encountered with the correct orientation was analysed, irrespective of track density. The results of the fission-track analysis are presented in Table 5. Ar –Ar mica dating Single crystal 40Ar/39Ar laser-fusion analyses were performed on biotite grains separated from two sands (KB-41-2, from the LGM deposits at Keti Bandar, and TH-4-6 from c. 6.4 ka sand at Thatta) at the Massachusetts Institute of Technology (MIT). Biotites from a modern sample of the Indus at Thatta were previously published by Clift et al. (2004). Prior to analysis, samples were irradiated in the C5 position of the McMaster University Nuclear Reactor, Canada, using 1 mm Cd shielding for four hours at a power level of 2 MW. After fusion with an Ar-ion laser, the released gases were purified for 10 minutes with two Al –Zr getters operated at 400 8C and room temperature, respectively, and then admitted to an MAP 215-50 mass spectrometer for Ar isotopic analysis using a Johnson MM-1 electronic multiplier operated at a gain of about 10 000. The conversion efficiency of 39K to 39Ar was monitored using sanidine from the Taylor Creek rhyolite (TCR-2a) assuming an age of 28.34 Ma (Renne et al. 1998), and is known to better than 0.3% (1s). Corrections for neutron-induced interferences, determined using Fe-doped kalsilite glass and optical CaF2, were 0.00039 for 40Ar/39ArK, 0.01243 for 38Ar/39ArK, 0.000672 for 39Ar/37ArCa, 0.000033 for 38Ar/ 37 ArCa and 0.00028 for 36Ar/37ArCa. Final data reduction was conducted with the program ArArCalc (Koppers 2002); results are shown in Table 6. U – Pb zircon dating U –Th –Pb isotopic compositions of zircon grains were analysed at the Australian National University, 192 Table 3. Sr isotopic data from Keti Bandar and Gularchy boreholes. Matching Nd isotope data and 14C ages are from Clift et al. (2008). See Figure 1 for locations Sample number Fine Sand Fine Sand Clay Clay Silt Clay Fine Sand Sand Sand Sand Location Depth (m) Keti Bandar Keti Bandar Keti Bandar Keti Bandar Keti Bandar Keti Bandar Keti Bandar Gularchy Gularchy Gularchy 14.0 55.0 61.0 92.0 97.0 108.0 117.0 11.3 26.2 47.5 14 C Age (ka) 0.21 8.72 9.42 12.20 12.91 13.73 28.70 2.65 6.42 10.94 Age 1s (ka) Nd143/Nd144 1Nd 0.00 0.15 0.20 0.32 0.32 0.37 0.00 0.08 0.08 0.18 0.511947 0.511858 0.511986 0.511919 0.512033 0.512000 0.512082 0.511911 0.511931 0.511962 213.5 215.2 212.7 214.0 211.8 212.4 210.8 214.2 213.8 213.2 87 Sr/86Sr 0.724512 0.727689 0.719290 0.719951 0.710671 0.714531 0.714764 0.721362 0.719811 0.725178 P. D. CLIFT ET AL. KB-5-2 KB-19-4 KB-21-1 KB-31-2 KB-34-4 KB-37-4 KB-40-5 GUL-ZP-5-1 GUL-ZP-10-2 GUL-ZP-17-1 Lithology Table 4. Carbon isotope and organic carbon data, together with associated major and trace element compositions from Holocene samples from Keti Bandar and Gularchy boreholes. See Figure 1 for locations Sample # Depth Age SiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O TiO2 P2O5 (%) (%) (%) (%) (%) (%) (%) (%) (%) (m) (a bp) (%) 7.5 20.0 26.0 26.5 37.0 44.0 55.0 59.0 67.0 70.5 92.0 96.0 100.0 100.5 107.0 110.0 112.0 117.0 118.0 119.0 112 299 1235 1314 4086 6336 8140 8560 9400 9768 12026 12446 12866 12918 13601 13916 14126 14651 14756 28700 54.23 55.93 51.48 51.16 54.02 50.67 53.44 65.28 54.32 63.43 54.09 51.49 50.27 50.67 52.36 49.58 51.51 50.89 48.04 50.07 13.47 13.42 15.04 15.22 13.69 14.84 13.67 10.94 13.72 11.04 13.60 13.78 14.13 14.49 14.04 14.84 14.18 14.58 14.86 15.23 5.46 5.35 6.42 6.51 5.64 6.38 5.62 4.00 5.67 4.79 5.66 6.11 6.04 6.21 5.88 6.42 6.11 6.35 6.68 6.63 0.09 0.09 0.10 0.10 0.09 0.10 0.10 0.07 0.10 0.08 0.09 0.09 0.09 0.09 0.09 0.09 0.09 0.10 0.10 0.10 2.94 2.87 3.32 3.39 3.04 3.42 3.00 2.22 3.04 2.35 3.03 3.20 3.46 3.38 3.48 3.45 3.56 3.61 4.36 3.49 8.48 7.27 7.28 7.17 7.92 7.71 8.14 6.72 7.26 6.81 8.38 8.76 9.39 8.28 8.30 8.64 8.09 7.91 7.67 7.69 1.74 2.14 1.83 1.88 1.83 1.73 1.79 1.87 2.35 1.94 1.82 1.90 1.91 1.87 2.02 1.81 1.96 1.98 1.88 1.74 2.54 2.54 2.94 2.98 2.53 2.88 2.64 2.09 2.49 2.08 2.62 2.68 2.71 2.83 2.76 2.89 2.82 2.89 2.95 2.97 0.66 0.69 0.74 0.75 0.81 0.72 0.71 0.60 0.71 0.63 0.68 0.69 0.70 0.71 0.70 0.71 0.70 0.70 0.72 0.72 0.14 0.16 0.16 0.16 0.19 0.15 0.15 0.14 0.15 0.16 0.15 0.16 0.16 0.15 0.16 0.15 0.16 0.16 0.14 0.15 Total (%) 10.05 9.39 10.38 10.68 10.11 11.04 10.29 6.56 10.19 7.08 10.29 11.12 11.67 11.34 10.68 11.70 10.43 11.75 12.80 11.41 99.79 99.86 99.67 99.98 99.86 99.64 99.55 100.48 100.00 100.39 100.39 99.99 100.53 100.00 100.46 100.28 99.61 100.92 100.20 100.21 As Ba Be Bi Cd Ce Co (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) 9.8 10.9 13.1 12.3 8.6 9.9 10.2 5.2 14.1 3.9 12.3 19.4 12.0 14.6 13.0 17.1 13.5 12.2 8.4 11.9 348.0 363.9 379.2 384.7 329.5 370.2 370.2 328.4 425.9 325.9 361.0 364.6 364.8 367.9 375.3 378.6 371.9 364.1 337.3 380.7 2.4 2.5 2.8 2.7 2.4 2.7 2.7 2.4 2.5 2.2 2.4 2.3 2.5 2.4 2.3 2.6 2.4 2.4 2.4 2.6 0.9 0.8 1.0 1.0 0.8 1.0 1.0 0.6 0.5 0.6 0.5 0.6 0.6 0.7 0.6 0.7 0.6 0.6 0.5 0.7 ,L.D. ,L.D. ,L.D. ,L.D. 0.287 ,L.D. ,L.D. ,L.D. ,L.D. 0.271 ,L.D. ,L.D. ,L.D. ,L.D. ,L.D. ,L.D. ,L.D. ,L.D. ,L.D. ,L.D. 60.8 71.7 68.0 65.8 92.7 64.5 72.3 66.2 66.2 75.0 62.5 62.1 62.6 62.2 61.3 66.2 61.4 61.6 62.7 65.3 14.8 13.9 16.2 16.7 14.8 16.9 15.5 10.9 15.5 10.5 15.2 15.5 16.2 16.1 15.6 17.8 17.1 17.6 16.9 17.2 MONSOON CONTROL OVER EROSION KB-3-3 KB-8-2 KB-10-2 KB-10-5 KB-13-2 KB-16-1 KB-19-3 KB-21-2 KB-23-4 KB-25-2 KB-31-1 KB-33-3 KB-34-2 KB-35-2 KB-37-4 KB-38-2 KB-38-3 KB-40-2 KB-40-3 KB-41-1 PF (%) (Continued ) 193 194 Table 4. Continued Cr Cs Cu Dy Er Eu Ga Gd Ge Hf Ho In La Lu Mo Nb Nd Ni Pb Pr Rb (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) KB-3-3 KB-8-2 KB-10-2 KB-10-5 KB-13-2 KB-16-1 KB-19-3 KB-21-2 KB-23-4 KB-25-2 KB-31-1 KB-33-3 KB-34-2 KB-35-2 KB-37-4 KB-38-2 KB-38-3 KB-40-2 KB-40-3 KB-41-1 100.9 92.2 107.6 107.9 112.9 110.4 105.2 83.1 97.8 84.2 99.0 104.4 105.2 108.7 108.3 114.7 111.6 113.0 121.8 114.7 8.2 7.8 10.1 10.0 8.2 10.1 9.3 6.2 6.9 5.9 8.3 8.4 7.8 8.9 7.8 9.1 8.2 8.7 8.2 10.7 25.6 24.4 31.4 33.5 24.8 34.0 27.8 13.8 27.1 15.5 26.4 28.5 31.0 32.3 28.7 37.4 30.5 33.6 52.6 33.6 4.0 4.5 4.3 4.1 5.9 4.1 4.8 4.3 4.4 4.7 4.0 4.0 4.0 4.0 4.0 4.1 4.1 4.0 4.2 4.1 2.2 2.5 2.4 2.3 3.2 2.2 2.6 2.4 2.4 2.5 2.1 2.2 2.2 2.2 2.2 2.2 2.2 2.2 2.3 2.3 1.1 1.2 1.1 1.1 1.4 1.1 1.2 1.1 1.1 1.2 1.0 1.1 1.1 1.1 1.1 1.1 1.1 1.1 1.1 1.1 17.5 17.5 20.3 20.0 18.7 20.2 19.3 14.4 18.5 14.4 17.8 18.4 18.7 19.1 17.5 19.5 18.6 19.3 20.0 20.4 4.4 5.0 4.7 4.6 6.5 4.5 5.2 4.8 4.9 5.3 4.5 4.5 4.4 4.4 4.5 4.7 4.6 4.5 4.6 4.6 1.6 1.6 1.7 1.7 1.7 1.7 1.7 1.5 1.6 1.6 1.6 1.6 1.6 1.5 1.6 1.6 1.6 1.6 1.7 1.7 3.6 5.1 4.0 3.8 7.7 3.5 4.7 6.1 4.5 6.8 4.1 4.1 3.9 3.8 4.0 3.6 4.2 4.0 3.6 3.5 0.8 0.9 0.8 0.8 1.1 0.8 0.9 0.8 0.8 0.9 0.8 0.8 0.8 0.8 0.8 0.8 0.8 0.8 0.8 0.8 0.2 0.2 0.2 0.2 0.3 0.2 0.3 0.2 0.2 0.2 0.2 0.2 0.3 0.2 0.1 0.2 0.2 0.2 0.2 0.2 30.2 36.3 33.9 33.4 46.5 31.9 36.1 33.0 33.6 37.8 31.6 31.3 31.6 31.5 31.1 33.1 31.0 31.2 31.5 32.8 0.3 0.4 0.4 0.4 0.5 0.3 0.4 0.4 0.4 0.4 0.3 0.3 0.3 0.3 0.3 0.4 0.3 0.3 0.4 0.4 0.6 ,L.D. 0.7 0.7 0.5 0.8 0.6 ,L.D. 0.7 ,L.D. 0.6 1.3 0.7 0.9 0.6 0.7 0.6 0.6 0.8 0.7 12.2 13.0 14.2 14.4 15.1 13.4 13.8 11.3 13.9 11.6 12.7 12.9 13.1 12.9 12.5 13.4 12.8 13.0 13.2 13.8 26.2 30.9 28.6 28.5 40.1 27.5 31.0 28.6 28.6 32.3 26.5 26.8 26.7 26.8 26.5 28.3 26.4 26.9 27.3 27.7 53.5 48.1 60.7 59.2 53.3 61.0 54.4 36.9 53.9 34.4 54.7 55.6 59.4 62.3 56.6 64.9 61.4 65.8 71.8 62.9 17.9 16.7 22.1 20.6 19.1 20.3 18.4 19.9 11.1 17.6 18.1 15.9 14.7 16.6 14.4 15.1 15.8 14.9 14.0 20.5 6.9 8.3 7.7 7.6 10.6 7.3 8.3 7.7 7.7 8.7 7.2 7.2 7.2 7.1 7.1 7.6 7.1 7.1 7.2 7.5 117.7 117.0 141.6 136.1 119.7 133.7 127.8 97.8 115.7 95.0 119.8 121.1 121.2 125.8 120.1 131.7 123.2 125.8 124.2 143.0 (Continued ) P. D. CLIFT ET AL. Sample # Table 4. Continued Sb (ppm) Sm (ppm) Sn (ppm) Sr (ppm) Ta (ppm) Tb (ppm) Th (ppm) Tm (ppm) U (ppm) V (ppm) W (ppm) Y (ppm) Yb (ppm) Zn (ppm) Zr (ppm) TOC (%) d13C (‰) KB-3-3 KB-8-2 KB-10-2 KB-10-5 KB-13-2 KB-16-1 KB-19-3 KB-21-2 KB-23-4 KB-25-2 KB-31-1 KB-33-3 KB-34-2 KB-35-2 KB-37-4 KB-38-2 KB-38-3 KB-40-2 KB-40-3 KB-41-1 0.7 0.7 0.7 0.7 0.7 0.7 0.7 0.5 1.7 0.5 1.6 1.7 1.9 0.8 0.7 2.0 1.7 1.7 1.6 0.7 5.2 6.0 5.6 5.5 7.9 5.4 6.2 5.8 5.7 6.4 5.3 5.2 5.4 5.2 5.3 5.6 5.3 5.3 5.3 5.5 3.6 3.8 4.2 4.2 3.8 4.2 4.0 3.4 4.4 3.4 3.9 4.0 4.2 3.7 3.5 4.3 4.1 4.0 3.9 4.0 204.6 201.1 203.2 194.4 214.3 190.9 199.0 188.8 199.7 191.3 210.2 253.2 370.8 221.1 223.1 226.0 233.6 230.2 181.4 198.7 1.1 1.3 1.3 1.3 1.5 1.2 1.2 1.2 1.3 1.2 1.1 1.2 1.2 1.2 1.2 1.2 1.2 1.2 1.2 1.3 0.7 0.8 0.7 0.7 1.0 0.7 0.8 0.8 0.8 0.8 0.7 0.7 0.7 0.7 0.7 0.7 0.7 0.7 0.7 0.7 12.8 15.3 15.5 15.4 20.9 14.3 15.7 13.4 14.9 15.6 13.6 14.0 14.0 14.1 13.8 15.1 13.7 14.0 14.2 15.2 0.3 0.4 0.4 0.3 0.5 0.3 0.4 0.4 0.4 0.4 0.3 0.3 0.3 0.3 0.3 0.3 0.3 0.3 0.3 0.3 2.3 2.9 2.7 2.6 3.8 2.5 2.9 2.5 2.9 2.9 2.5 2.6 2.6 2.7 2.6 2.7 2.7 2.7 2.8 2.6 100.9 96.0 115.6 115.2 104.1 115.7 105.5 71.6 102.3 70.8 101.1 106.0 110.5 115.9 106.6 119.9 110.6 114.1 125.5 119.0 2.3 2.9 2.8 2.8 2.7 2.6 2.6 3.5 2.7 10.5 2.3 2.4 2.4 2.4 2.3 2.5 2.3 2.4 2.4 2.7 23.4 25.6 24.3 23.7 33.4 23.0 27.4 25.3 25.0 26.6 23.1 22.9 22.4 22.0 23.2 23.6 23.2 22.6 23.5 23.2 2.2 2.5 2.4 2.3 3.3 2.2 2.6 2.4 2.4 2.6 2.2 2.2 2.2 2.2 2.3 2.3 2.3 2.3 2.4 2.3 82.7 87.3 120.1 94.9 94.2 100.1 95.5 63.9 122.4 62.8 82.1 88.1 89.8 88.8 85.5 94.5 89.2 91.0 95.5 98.1 133.2 187.9 143.9 138.4 291.0 126.9 170.9 224.7 164.6 268.0 152.5 152.5 143.1 131.7 143.6 126.3 150.7 144.2 130.0 125.3 0.27 0.26 0.37 0.34 0.32 0.43 0.31 0.16 0.31 0.18 0.31 0.31 0.32 0.29 0.26 0.32 0.29 0.26 0.26 0.44 221.2 220.5 219.8 220.3 219.8 220.0 220.3 222.6 221.0 223.3 220.7 221.1 220.9 221.7 220.8 221.5 220.5 220.8 220.0 220.0 MONSOON CONTROL OVER EROSION Sample # 195 196 Table 5. Fission track analytical data for Holocene Indus delta sands Sample no./ Strat age/ No. of Field no. lithology crystals No apatite Apatite Apatite Apatite Apatite Zircon Apatite 44 15 16 70 28 61 rd Nd 1.317 1.317 1.220 1.317 0.493 1.220 3651 3651 6107 3651 3450 6107 Spontaneous rs Ns Induced ri Ni 0.101 135 5.538 6762 0.221 57 3.681 949 0.109 50 4.036 1857 0.162 661 4.122 16735 6.107 2404 5.414 2131 0.846 297 4.05 14219 Age dispersion Px2 RE% 0 3.0 5.3 0 0 0 Central age (Ma) + 1s 1st age comp. 2nd age comp. 3rd age comp. 4th age comp. 124.5 4.4 + 1.0 2.6 + 0.3 (40) 9 + 3 (1) 24 + 4 (3) 56.7 11.1 + 2.7 7.4 + 1.7(13) 23 + 4 (2) 40.4 5.5 + 1.0 1.8 + 0.7 (6) 7.7 + 1.2 (10) 45.5 8.9 + 0.7 3.5 + 0.4 (20) 8.7 + 0.5 (50) 64.7 34.9 + 4.4 17.1 + 0.9 (14) 42.3 + 4.2 (9) 96 + 11 (4) 79.8 4.4 + 0.6 2.2 + 0.3 (40) 9.6 + 0.9 (21) Note: (i) Track densities are (106 tr cm22) numbers of tracks counted (N) shown in brackets. Rd is the induced track density of the dosimeter. Nd is the number of tracks counted in the dosimeter/ Ns is the number of spontaneous tracks counted in the sample. Ni is the number of induced tracks counted in the sample. (ii) Analyses by external detector method using 0.5 for the 4p/2p geometry correction factor; (iii) Ages calculated using dosimeter glass CN-5; (apatite) zCN5 ¼338 + 4; CN-2 (zircon) zCN2 ¼127+4 calibrated by multiple analyses of IUGS apatite and zircon age standards (Hurford 1990); (iv) Px2 is the probability of obtaining x2 value for v degrees of freedom, where v ¼ no. crystals – 1. RE %, age dispersion (or the spread of the individual crystal data) is given by the % relative standard deviation of the central age; (v) Central age is a modal age, weighted for different precisions of individual crystals (Galbraith 1990). P. D. CLIFT ET AL. KB-34-4 KB-19-4 KB-23-3 KB-5-2 KB40-5 KB40-5 TH-10-8 Dosimeter MONSOON CONTROL OVER EROSION 197 Table 6. Ar – Ar analytical data for detrital biotite grains extracted from Holocene Indus delta sands KB-41-2 Age (Ma) 6.88 7.45 7.70 8.05 8.69 9.09 9.09 9.14 9.32 9.35 9.61 9.67 9.80 10.16 10.26 10.49 11.37 11.52 12.69 15.40 15.87 15.89 16.30 16.44 18.11 18.51 19.55 21.90 24.96 25.52 26.45 26.52 26.55 30.15 32.80 33.23 38.04 40.60 42.28 43.18 43.80 116.42 126.38 246.18 366.70 572.73 TH-4-6 Age (Ma) 0.22 0.41 0.48 0.82 0.97 1.29 1.36 1.39 1.71 +2s +1.64 +3.23 +0.44 +2.80 +0.66 +0.37 +0.21 +0.44 +0.18 +0.81 +3.81 +0.83 +0.39 +0.93 +3.69 +0.70 +0.65 +2.31 +1.71 +2.41 +0.79 +1.04 +0.87 +0.64 +0.51 +0.34 +0.74 +3.90 +1.98 +1.02 +1.48 +1.07 +0.54 +0.67 +0.55 +0.78 +1.14 +0.60 +1.07 +0.57 +0.39 +0.81 +0.93 +1.16 +1.99 +3.19 +2s +1.07 +0.91 +0.55 +0.67 +1.13 +0.37 +0.46 +0.64 +0.75 40 Ar(r) (%) 12.14 6.28 35.45 4.70 55.30 55.48 77.40 48.89 84.63 84.66 5.08 59.20 69.54 66.19 6.06 82.14 77.99 16.00 13.29 10.24 46.97 41.46 25.53 76.76 58.42 82.14 86.39 8.48 20.80 81.10 32.09 89.69 86.09 61.70 92.82 90.30 42.20 85.08 53.26 58.88 78.45 95.22 89.48 98.97 98.92 99.29 40 Ar(r) (%) 8.34 9.58 5.14 13.40 13.22 46.53 22.69 28.52 20.85 39 Ar(k) (%) TH-4-6 Age (Ma) 1.35 0.74 3.42 3.35 1.84 3.91 2.79 3.88 4.17 1.30 1.00 1.35 2.82 1.20 0.75 1.43 1.63 0.83 1.67 2.34 1.83 1.51 2.88 1.59 3.26 2.25 1.40 2.13 2.06 1.18 1.65 1.04 2.12 2.37 1.97 1.67 2.07 1.83 1.88 4.75 5.36 1.92 1.76 3.60 1.47 2.68 1.85 2.17 2.37 2.53 2.56 2.75 2.76 2.76 2.76 2.79 3.22 3.24 3.24 3.31 3.32 3.32 3.38 3.42 3.44 3.53 3.54 3.55 3.66 3.75 3.84 3.87 3.97 4.17 4.34 4.42 4.49 4.5 4.51 4.53 4.58 4.63 4.63 4.71 4.76 4.83 4.86 4.91 5.03 5.04 5.10 5.14 5.15 5.18 5.20 5.34 5.57 5.87 6.24 6.27 6.30 6.35 6.52 6.55 6.70 39 Ar(k) (%) 0.25 0.29 0.56 0.43 0.22 0.66 0.87 0.43 0.37 +2s +0.60 +0.94 +0.33 +0.54 +0.36 +0.56 +0.24 +0.63 +0.23 +0.75 +0.30 +0.79 +0.88 +0.64 +0.56 +0.52 +0.20 +0.84 +0.29 +0.69 +0.27 +0.59 +0.44 +0.56 +0.45 +0.37 +0.47 +0.45 +0.25 +0.67 +0.39 +0.50 +0.36 +0.49 +0.73 +0.53 +0.21 +0.38 +0.75 +0.55 +0.52 +0.39 +0.62 +0.16 +0.46 +0.63 +0.49 +0.46 +0.25 +0.16 +0.91 +0.71 +0.61 +0.36 +0.64 +0.51 +0.87 +0.38 +0.64 40 Ar(r) (%) 43.17 28.16 30.64 62.68 14.67 14.63 43.69 23.42 55.22 46.23 36.46 33.64 38.98 44.40 57.69 63.30 45.33 56.01 21.48 50.22 32.29 10.88 41.60 70.46 41.70 52.81 36.91 50.52 56.30 58.55 27.75 44.28 44.04 43.77 39.79 13.16 55.84 47.08 59.64 64.99 42.41 56.64 45.98 49.53 52.18 70.39 52.18 49.86 43.90 71.14 48.50 57.39 55.76 58.96 65.04 50.78 54.76 55.42 65.22 39 Ar(k) (%) 0.51 0.36 1.29 0.52 1.98 0.67 1.21 0.49 1.83 0.44 1.18 0.37 0.30 0.41 0.64 0.52 1.98 0.35 2.43 0.49 1.62 1.43 0.92 0.56 0.51 0.74 0.93 0.92 1.78 0.40 2.10 0.66 1.44 0.63 0.37 3.40 2.33 1.13 0.34 0.75 0.63 0.75 0.43 3.90 0.88 0.47 0.91 1.00 2.30 2.51 0.32 0.53 0.44 1.11 0.44 1.13 0.28 1.25 0.62 (Continued) 198 P. D. CLIFT ET AL. Table 6. Continued TH-4-6 Age (Ma) 7.47 7.82 8.21 8.27 8.37 8.39 8.62 8.87 9.42 11.29 11.61 12.73 14.18 16.09 16.63 17.25 17.86 18.13 18.67 19.12 19.86 20.08 21.91 25.48 27.75 29.01 31.37 41.00 53.50 92.92 111.32 +2s +0.69 +1.03 +0.59 +0.51 +0.69 +0.44 +0.47 +0.57 +0.17 +0.27 +0.30 +0.94 +0.21 +0.38 +0.34 +0.25 +0.80 +0.25 +0.54 +0.49 +0.57 +0.71 +0.32 +1.34 +0.25 +0.50 +0.45 +0.63 +0.71 +0.70 +0.62 40 Ar(r) (%) 41.98 61.43 40.77 24.57 60.82 75.78 52.41 61.82 62.58 56.83 65.97 80.57 68.42 85.05 79.19 81.98 91.50 75.91 80.59 85.30 71.92 82.82 79.92 75.56 87.03 75.20 76.57 76.73 92.32 95.18 97.04 39 Ar(k) (%) 0.62 0.25 0.62 2.36 0.61 0.58 0.92 0.42 2.73 2.71 0.97 0.27 3.69 1.36 1.26 1.60 0.31 2.37 0.58 0.84 0.49 0.41 1.36 0.23 1.94 1.16 1.12 0.82 0.64 0.57 1.26 Canberra, using Excimer Laser Ablation Inductively Coupled Plasma Mass Spectrometry (ELA-ICP-MS) employing a pulsed 193 mm ArF LambdaPhysik LPX 1201 UV Excimer laser and an Agilent 7500 quadrupole ICP-MS. The zircons were separated from the bulk sediment by conventional magnetic and heavy liquid separation techniques. In this study we analysed two samples (Jati-16-1 and KB-34-1), which were not considered by the earlier study of Clift et al. (2008). The extracted zircons were mounted in epoxy resin and polished. Dating by ELA-ICP-MS followed the procedure described in Campbell et al. (2005). Our method employs standard zircon TEMORA2 and NIST610 silicate glass (Pearce et al. 1997; Black et al. 2004) where the latter is used for concentration information and for U/Th determination. As we cannot measure common Pb (204Pb) directly because of systemic Hg, we use a 208Pb-based correction only when that correction makes the analysis more concordant than the uncorrected version. Once the data were compiled, an analysis was rejected for interpretation on the basis of the following: (a) the observed variance on 206Pb/238U or 207Pb/206Pb (depending if the grain is . or ,1200 Ma) is more than three times that calculated from counting statistics (this procedure omits grains that record mixed ages), or (b) the grain is deemed to be discordant. Analysis time drift corrections were applied to both analytical sessions. Results of the U – Pb dating are shown in the supplementary material. Overall uncertainty on an individual measurement is about 1–2%. Results Mineralogy Detrital modes in the Indus Holocene sands are typical of sediments derived from collision orogens (Garzanti & Andò 2007), with medium quartz content and equally abundant feldspars and lithic grains (quartz 49 + 4%, feldspar 26 + 3%, lithics 25 + 3%). Using a Dickinson ternary diagram (Fig. 3a) the sands mostly plot within the ‘recycled orogen’ field, with minor overlap in the ‘dissected arc’ range. In this respect they are similar to sandstones found in the Himalayan foreland basin, at least since the start of the Neogene (Najman & Garzanti 2000). Lithic grains comprise equally abundant sedimentary (limestone, dolostone, siltstone) and metamorphic types. However, volcanic, metavolcanic, and metabasite lithic grains are very minor, but decrease up-section in the lower part of the core (120 –60 m depth, 20 –9.5 ka; Fig. 4), with the greatest change between 12.2 and 10.6 ka. This trend suggests upward-decreasing relative contributions from Kohistan-like Transhimalayan sources and/or the West Pakistan ophiolites. The composition of the lithic grains defines the sands as typical of ‘suture belts’ in the ternary diagram (Fig. 3b). Conversely, carbonate grains, which are almost absent at the base of the Keti Bandar section (LGM) increase in the uppermost 90 m (since c. 12 ka), suggesting increasing contributions from sedimentary to metasedimentary strata and/or more arid climatic conditions that do not favour dissolution of these grains. The total heavy-mineral content is remarkably constant (4.2 + 0.7%) in the six very fine-grained sands analysed (KB-5-2, TH-10-8, KB-20-1, KB-34-4, KB-40-5, and KB-41-2), indicating that intrastratal solution is negligible throughout the cored sections. Heavy minerals are less abundant in the analysed 63 –250 mm fraction of the other three silt-sized samples (KB-23-2, KB-26-2, and KB-30-4), but this does not mean that bulk samples contain fewer heavy minerals, because denser detrital grains are markedly concentrated in the fine tail of the size distribution of each sample MONSOON CONTROL OVER EROSION (a) Tra ns. co nti ne nt Q Ba se me nt up lift Recycled Orogen Dissected arc Transitional arc Undissected arc F L (b) 199 (ZTR) 3 + 1% and all other heavy minerals 6 + 2%) (Garzanti et al. 2005). Slightly higher garnet content in modern bedload sand is ascribed to the selective entrainment of less dense grains and enrichment in denser grains in lag deposits on the channel bottom (Slingerland 1984). Most remarkable is the virtual lack of limestone grains and of pyroxene at the bottom of the Keti Bandar core (117–120 m depth; i.e. deposited at the LGM), which would indicate strong chemical weathering. This observation may be linked to a humid climate at the LGM, for which there is no evidence, or alternatively and more likely, due to prolonged exposure prior to the Holocene transgression. Lm Sr isotope evolution Suture belts Mixed arc and subduction complex Magmatic arc Lv Ls Fig. 3. (a) Triangular QFL plot (Q, quartz, L, lithics, F, feldspar) and (b) Lm (lithic metamorphic), Lv (lithic volcanic) Ls (lithic sedimentary) plots, with fields from Dickinson (1985) for the Indus Holocene sands. because of settling-equivalence effects (Garzanti et al. 2009). Amphiboles (mainly blue-green hornblende, comprising 52 + 5% of the total heavy mineral population) prevail over epidote (27 + 4%), subordinate garnet (6 + 4%) and clinopyroxene (6 + 3%), and minor tourmaline, titanite, hypersthene, kyanite, sillimanite, rutile, staurolite, chloritoid and zircon in order of decreasing abundance (5 + 3%). Hypersthene increases slightly up-section, suggesting increasing contribution from the Kohistan Arc in the upper part of the core (depth ,60 m; since c. 9 ka), since this mineral is distinctive of erosion from arc units (Cerveny et al. 1989). Marked temporal changes in mineralogy are not apparent (Fig. 4), and sediment composition remains comparable to the modern Indus detrital modes (quartz 48 + 4%, feldspar 21 + 3%, lithics 32 + 5%) and heavy-mineral assemblages (amphibole 50 + 8%, epidote 25 + 7%, garnet 12 + 3%, pyroxene 4 + 3%, zircon þ tourmaline þ rutile Temporal evolution in Sr isotopes is shown together with changing Nd isotopes in Figure 5. What is apparent is that the two systems are closely correlated and that Sr increases rapidly from low values at the LGM and until c. 13 ka, after which there is a rapid rise to a high 87Sr/86Sr between 8 and 9 ka. Subsequently, 87Sr/86Sr values decrease slightly to the present day. The correlation suggests that Sr isotope composition is dominated by source and provenance rather than by chemical weathering intensity, although the total number of analyses is rather low. As with the Nd data there is no clear correlation between isotopic composition and grain size. Sediments of the same grain size show a variety of isotopic ratios (Fig. 5). The shift to higher isotope ratios is consistent with increased relative erosion from the radiogenic crust of the Lesser and Greater Himalaya, and away from the more primitive crust of Kohistan and the Transhimalaya (Trivedi et al. 1984; France-Lanord & Le Fort 1988; Scaillet et al. 1990; Ahmad et al. 2000). Because the change in isotope character occurs at a time of greatly increased sediment flux to the delta we rule out the possibility that the change in isotope character reflects decreased sediment flux from the arc sources lying north of the Himalayan ranges and in any case the increase in hypersthenes (but not metabasic rocks) up-section suggests more erosion from Kohistan, not less. However, the trend in the Nd isotope curve indicates that the influence of Kohistan on the total sediment composition is swamped by increased flux from Himalayan sources. Organic carbon Although the shift to higher 87Sr/86Sr values in the Early Holocene cannot be interpreted as indicating stronger chemical weathering under the influence of a stronger summer monsoon, organic carbon 200 P. D. CLIFT ET AL. Proportion of total mineralogy (%) (a) Age (ka) 0 10 20 30 40 Proportion of heavy mineral population (%) (b) 50 0 60 0 0 5 5 10 10 15 15 20 20 Quartz Feldspar Sedimentary lithic grains Metamorphic lithic grains Igneous lithic grains 10 20 30 40 50 Amphibole Group Pyroxene Group Epidote Group Zircon, tourmaline,Ti-oxides, titanite, apatite,monazite Chloritoid, staurolite, andalusite, kyanite, sillimanite Garnet Fig. 4. Plots showing the evolution in sand mineralogy at the Keti Bandar borehole since 20 ka. (a) Major mineral groups, and (b) only heavy minerals. Data are plotted from Tables 1 and 2. The plots confirm the overall lack of strong changes in mineralogy throughout the deglaciation process. analysis can be used to constrain environmental conditions. Total organic carbon (TOC) concentrations are low and range between 0.16% and 0.44% (Table 4), consistent with the generally arid conditions in the Indus basin. There is no obvious TOC variation with age of sedimentation and most of the samples remain in a narrow range around 0.3% (Fig. 6a). Bulk organic carbon d13C varies between – 23.3‰ and –19.7‰, but only two samples with very low TOC (0.16% and 0.18%) have d13C lower than –21.7‰ (Fig. 6b). These negative d13C values ( –22.6 and –23.3‰) are associated with the lowest TOC and correspond to sandy sediments, as indicated by low Al/Si ratios. In these two quartz-rich sands, rock-derived organic carbon is likely to be a major component of the total organic carbon content (Galy et al. 2008). Their stable isotopic composition is therefore not representative of modern organic carbon directly derived from the biosphere. Excluding these two samples, d13C shows a c. 2‰ variation around an average value of –20.6‰. In detail, MONSOON CONTROL OVER EROSION 201 Fig. 5. Diagram showing the variability in bulk sedimentary Nd and Sr isotope character since the Last Glacial Maximum. Sediments are from the Indus delta and the Indus Canyon. Letters indicate grain size of sediment: C, clay; SI, silt; FS, fine sand; S, sand. Black arrows indicating age control points and Nd data are from Clift et al. (2008). Sr data are from Table 3. d13C reaches a minimum after the LGM and likely until the Younger Dryas. d13C increases between 12 ka and 5 ka and finally decreases again up to present (Fig. 6b). Organic carbon (OC) can be derived from a mixture of different sources: terrestrial OC derived from vegetation, soils and autotrophic production in the river and marine organic carbon. Historically, the Indus River is characterized by high sediment concentration, which limits autotrophic productivity (Ittekkot & Arain 1986). Terrestrial inputs may therefore be derived from the vegetation present in the basin, either directly (plant debris) or indirectly (soil organic carbon). C3 plants have a considerable range in d13C. Arid ecosystems are enriched in 13C (as high as 222‰) (Farquhar et al. 1989), but closed canopy flora are depleted in 13C, with d13C values as low as 235‰ (van der Merwe & Medina 1989). However, the average C3 value is about – 26‰. In contrast, C4 plants have a much more 202 P. D. CLIFT ET AL. (a) (b) Total Organic Carbon (%) 0.15 0.25 0.35 –23 0 5 5 10 10 –22 –21 –20 Age (ka) 0 d13C (‰) 15 15 LGM LGM Fig. 6. Diagrams showing (a) the evolution in total organic carbon (TOC) and (b) carbon isotope character of bulk sediments from the Indus delta. See Table 4. TOC values are generally small and show little variation with time. Lowest TOC values correspond to negative d13C values associated with reworking of old organic matter. The 2‰ positive shift between 12 ka and 5 ka is significant and may be related to an increase of C4 plants input. See Table 4 for data. restricted d13C range, with an average d13C value around –13‰ (Deines 1980; Hattersley 1982; Collister et al. 1994). In the Arabian Sea, modern marine plankton has d13C values around –20‰ (Fontugne & Duplessy 1978) and its isotopic composition has likely remained fairly stable during the last 20 ka. The sedimentological study of this record clearly indicates a marine transgression during the Holocene. Therefore, the middle section of the record at Keti Bandar (15 –70 m) may have been influenced by marine organic carbon, whereas it might be expected to be negligible in the upper and lower sections. In these two terrestrial sections of the record, bulk organic carbon d13C indicates mixed C3/C4 vegetation in the Indus basin. The lowest part of the record (LGM to 12 ka) shows a slight decrease of the bulk organic carbon d13C from –20‰ to –21‰. This shift may be related to an increase of the C3 plants proportion in the basin during the time of deglaciation. The c. 2‰ positive shift between 12.9 ka and 4.1 ka probably reflects increasing contribution of marine OC, although it might also indicate an increase in the proportion of C4 plants in the basin. Conversely, the reversed trend from 4.1 ka to present likely reflects decreasing contribution of marine organic carbon, or an increase of the C3 plants proportion in the basin. Comparison with published organic carbon data from the Bengal Fan system indicates that the Indus is relatively low in TOC (Galy et al. 2007b). Maximum values range up to 0.44% compared to .1.1%. However, like the Bengal sediment our data show a rough first order correlation between Al/Si ratios and TOC (Fig. 7). This is consistent with a control of the organic carbon content by the sediment properties, specifically a preferential association of organic carbon with fine grain sediments enriched in phyllosilicates (clays and micas) (Galy et al. 2008). Figure 7 shows that the slope on the TOC v. Al/Si chart, which characterize the organic carbon loading, is lower in the Indus basin than in the Bengal Fan, reflecting its more arid environmental conditions. MONSOON CONTROL OVER EROSION although with only 19 grains this sample may not be representative (Clift et al. 2004). The Greater and Lesser Himalaya, together with the Karakoram have yielded abundant young AFT ages that would be consistent with a source in those ranges. The slightly older ages seen in the reworked Miocene foreland sedimentary rocks of the Siwaliks argues against them being important since 8.7 ka, although they may partly be responsible for the older age population seen at the LGM. Other possible sources for the older grains deposited at the LGM are the Transhimalaya or Kohistan. Further source characterization is possible using Ar –Ar and U – Pb methods. al Fan 0.35 Beng elt a 0.25 Ind us D TOC (%) 0.30 0.20 0.15 0.00 0.10 0.20 0.30 203 0.40 0.50 Al/Si Fig. 7. Diagram showing the relationships between mud Al/Si and TOC for the Indus Holocene. Trend shows much lower slope than that recognized for the Bengal Fan (Galy et al. 2007b), suggesting much lower organic productivity. Fission track analyses The results of the fission track analyses are shown graphically in Figure 8 in the form of radial plots that show the ages and uncertainties of single grain apatite and zircon grains (Galbraith 1990). Statistical analysis allows a central age to be assigned for each sample, with greater confidence for those samples with more abundant grains. In a complex system like the Indus 100 grains are needed for a robust result (Ruhl & Hodges 2005) and in several of these sands the numbers are so low that they are not useful. Three samples can be used to look at the general development in sediment source since the LGM. Sample KB-40-5 has a central age of 9.0 + 1 Ma, yet by 8.7 ka sample KB-19-4 shows a central age of only 4.4 + 1 Ma. A similar age is yielded by sample TH-10-8, deposited around 7 ka. The radial plots show that there is a minority population dating .10 Ma, but that in the younger sands in particular this is very minor. The change in apatite fission track (AFT) ages during the Holocene must reflect a change in provenance as the duration is not long enough for this to represent a change in source exhumation rates. Comparison of the age spectrum in the sediments with AFT ages from possible source terrains allows the changing erosion patterns to be constrained (Fig. 9). Probability density diagrams emphasize the young AFT ages of the younger sediment and show the ‘tail’ of grains older than 10 Ma seen in the LGM sediment but not since that time. Comparison with the modern Indus sediments shows a similar pattern to the recent sediments, Mica ages Ar –Ar cooling ages in biotite and muscovite micas document the age that these grains cooled below c. 280 8C and 350 8C respectively (Hodges 2003). As exhumation is diachronous across the Himalaya these ages can be used as powerful provenance tools in modern and ancient South Asian sediment (White et al. 2002; Clift et al. 2004). Figure 10 shows the range of biotite cooling ages for two core samples and one modern river sand sample. The age spectra for the modern and 6.4 ka sand differ in one key aspect from the LGM sand at Keti Bandar, in the abundance of grains ,10 Ma. All samples show minority populations with older, albeit Cenozoic cooling ages, At the LGM, the most common age lies around 9 Ma, compared to c. 4 Ma for the younger sediments. Comparison with the rather limited number of bedrock analyses suggests that the Karakoram makes a relatively good match as a possible source at the LGM, although there are no published data for the Transhimalaya or Lesser Himalaya. Cooling ages in the Greater Himalaya largely range 10–24 Ma and do not account for the up-surge of young ages in the Holocene delta, although there is a significant population of 10–14 Ma grains in the modern river that may be derived from this area. Nanga Parbat is a possible source of the 1– 7 Ma grains seen in both younger samples. Additional source constraints are possible using the muscovite Ar –Ar ages reported by Clift et al. (2008). This system has the advantage over biotite in being more widely measured in the potential source regions. Figure 11 shows that like the biotite data the 6.4 ka and modern sediments have several muscovite grains dating ,10 Ma, which the LGM sediment does not contain. A probability maximum around 18 Ma at 6.4 ka and at the LGM correlates with known sources in the Lesser and Greater Himalaya, although there is some overlap between these sources that makes their separation hard with this method. However, we note that 204 P. D. CLIFT ET AL. 40 KB-40-5 (Apatite), >20 ka, LGM Central Age: 9±1 Ma P(X2): 0.0% Relative Error: 46% Number of grains: 70 KB-40-5 (Zircon), >20 ka, LGM Central Age: 35±4 Ma P(X2): 0.0% Relative Error: 65% Number of grains: 28 30 20 +2 200 150 100 10 0 -2 5 50 +2 0 -2 % relative error 1 % relative error 70 0 28 7 0 10 20 30 40 50 10 60 10 12 20 30 Precision (1/sigma) Precision (1/sigma) KB-5-2 (Apatite), 210 a depositional age Central Age: 6±1 Ma P(X2): 5% Relative Error: 40% Number of grains: 16 20 16 12 +2 8 0 6 40 30 KB-19-4 (Apatite), 8.7 Ka depositional age Central Age: 4.4±1 Ma P(X2): 0.0% Relative Error: 124% Number of grains: 44 20 10 +2 5 0 -2 1 -2 1 % relative error 55 0 10 20 % relative error 89 11 0 30 10 20 30 8 40 50 Precision (1/sigma) Precision (1/sigma) KB-23-3 (Apatite), 9.3 ka depositional age Central Age: 11±3 Ma P(X2): 3% Relative Error: 57% Number of grains: 16 40 30 20 16 TH-10-8 (Apatite), 7.0 ka depositional age Central Age: 4.4±1 Ma P(X2): 0.0% Relative Error: 80% Number of grains: 61 12 8 20 +2 5 +2 0 10 -2 5 0 -2 1 % relative error 76 1 13 0 10 20 30 Precision (1/sigma) % relative error 63 6 0 10 20 30 40 50 60 Precision (1/sigma) Fig. 8. Radial plots (Galbraith 1990) showing the ages and uncertainties of single grain apatite and zircon grains within Holocene sands from the Indus delta. Locations of samples are shown in Figures 1 and 2. Greater Himalayan muscovite ages peaks around 20 Ma, whereas the limited data from the Lesser Himalaya peak around 16 –18 Ma. The greatest probability peak in the detrital grains is younger than 20 Ma, consistent with the Lesser Himalaya being the dominant source. Again the muscovite confirms that the Siwaliks are not dominant sediment sources because the cooling ages are generally MONSOON CONTROL OVER EROSION 205 Fig. 9. Probability density plots showing the range of apatite fission track central ages for the sediment samples analysed versus the ages found in a variety of possible source terrains. Siwalik data is from Van der Beek et al. (2006). Karakoram data is from Zeitler (1985), Poupeau et al. (1991) and Foster et al. (1994). Greater Himalayan data are from Kumar et al. (1995), Sorkhabi et al. (1996), Searle et al. (1999), Jain et al. (2000), Thiede et al. (2004) and Bojar et al. (2005). Lesser Himalayan data are from Vannay et al. (2004) and Thiede et al. (2004). Pakistan Himalayan data are from Zeitler (1985). Transhimalayan data are from Zeitler (1985), Zeilinger et al. (2001), Clift et al. (2002a) and Kirstein et al. (2006). Figure shows preference to younger grain ages with the onset of the Holocene, consistent with relatively more erosion from the Greater and Lesser Himalaya and less from the Siwaliks and Transhimalaya. See Table 5 for data. too young. As for the ,10 Ma muscovite grains bedrock data suggest either Nanga Parbat or the Lesser Himalaya as likely sources. A probability maximum at 15 –16 Ma in the modern river matches several known sources in the Lesser Himalaya but not Nanga Parbat. This population is less abundant in the 6.4 ka sample. We conclude that the mica dating argues for reduced erosion in the Karakoram and more erosion in the Lesser Himalaya or Nanga Parbat between the LGM and the Early Holocene. As these sources are both negative in 1Nd (Parrish & Hodges 1996; Whittington et al. 1999; Ahmad et al. 2000) increased relative flux in either could explain the observed bulk sediment Nd and Sr isotope evolution (Fig. 5). Zircon dating Zircon U –Pb dating has proven an effective provenance tool in South Asia because it preserves the original age of crystallization of the source rocks, which varies significantly across the Himalayas and into Tibet (DeCelles et al. 2000). In this study we augment the data presented by Clift et al. (2008) with two additional samples in order to define the Holocene provenance evolution better. Figure 12 shows the age spectra for the four core samples, plus modern river data compared with various source terrains. All the sediments show a large population with grains of ,150 Ma and a spread of other older grains. Some samples show particularly well developed groups. The LGM sands show many grains dated at 800–1100 Ma, reducing in number up-section. The modern river sand shows an especially large number of grains c. 1800 Ma compared to the older sediments. The new data are consistent with the older in suggesting significant erosion from the Karakoram, or the Transhimalaya, especially at the LGM. In addition, the zircon indicate increased erosion from the Lesser Himalaya going up section. Very few of the grains are young enough to match those measured from the Nanga Parbat gneiss. This resolves one of the ambiguities from the mica Ar –Ar data in separating the erosional flux from the Lesser Himalaya. Despite its dramatic exhumation history (Zeitler et al. 1993) it appears that Nanga Parbat is a modest contributor of sediment to the Indus. The provenance can be further quantifying by dividing up the detrital zircons into families. We choose 0–20 Ma grains to represent the flux from Nanga Parbat, 20 –55 Ma grains are rarely known outside the Karakoram Batholith. The range 55 –300 Ma is chosen as a suitable range for much of the activity in Kohistan and the Transhimalaya, 206 P. D. CLIFT ET AL. Greater Himalaya Nanga Parbat Karakoram 0 10 20 30 40 50 60 Modern Indus at Thatta N = 95 0 10 20 30 40 50 60 TH-4-6, Biotite, <6.4 ka N= 46 0 10 20 30 40 50 60 KB-41-2, Biotite, 25 ka N = 99 0 10 20 30 40 50 60 Age (Ma) Fig. 10. Probability density plots showing the range of Ar–Ar cooling ages in biotite grains from three sand samples from the Indus delta. Top section shows known range of possible source ages from the Greater Himalaya (Copeland et al. 1990; Searle et al. 1992; Metcalfe 1993; Inger 1998; Stüwe & Foster 2001; Godin et al. 2006; Wang et al. 2006), Nanga Parbat (Zeitler et al. 1989; Winslow et al. 1996; Treloar et al. 2000) and Karakoram (Searle et al. 1989; Brookfield & Reynolds 1990; Krol et al. 1996; Villa et al. 1996). See Table 6 for data. 300– 1400 Ma grains represent the Greater Himalaya and those older than 1400 Ma the Lesser Himalaya. Because of overlaps in age ranges such a budget is necessarily schematic but does show the general trends in provenance evolution because the different ranges have preferred ages that are typical if not unique to them. In Figure 13 we plot pie charts to show how the bulk composition MONSOON CONTROL OVER EROSION 207 Greater Himalaya Probability Nanga Parbat Siwaliks Lesser Himalaya Modern River at Thatta (48 grains) Probability 0-10 Ma (N. Parb.+L. Him.) = 15% 10-40 Ma (L.+Gr.. Him.) = 80% >40 Ma (Transhimalaya) = 4% TH-4-6, Muscovite, <6.4 ka (99 grains) Probability 0-10 Ma (N. Parb.+L. Him.) = 25% 10-40 Ma (L.+Gr. Him.) = 70% >40 Ma (Transhimalaya) = 5% KB-41-2, Muscovite, >20 ka (50 grains) Probability 0-10 Ma (N. Parb.+L. Him.) = 4% 10-40 Ma (L.+Gr. Him.) = 84% >40 Ma (Transhimalaya) = 12% 0 10 20 30 40 Age (Ma) Fig. 11. Probability density plots showing the range of Ar–Ar cooling ages in muscovite grains in the glacial sample, at ,6.4 ka and in the modern river (2004), compared to those in possible source regions. Top section shows the known range of possible source ages from the Greater Himalaya within the Indus basin (Searle et al. 1992; Metcalfe 1993; Inger 1998; Walker et al. 1999, Stephenson et al. 2001), Lesser Himalaya (Catlos et al. 2001; Bollinger et al. 2004; Vannay et al. 2004), Nanga Parbat (Smith et al. 1992; George et al. 1995; Treloar et al. 2000), and Siwaliks (White et al. 2002; Szulc et al. 2006). Reprinted with permission from the Geological Society of America. of the Indus sediments has changed from the LGM to the present day. Again the heavy influence of the Lesser Himalaya on the modern river is clear and contrasts even with the mid Holocene samples. Because of damming of the modern Indus, most notably at Tarbela (Fig. 1), some of the anomaly in the modern sample may be anthropogenic, although dams do exist on many of the Himalaya tributaries too (e.g. the Mangla Dam on the Jhelum). All samples show very little sediment from Nanga Parbat sources, but a consistent dominant flux from the Greater Himalaya. 208 P. D. CLIFT ET AL. Nanga Parbat Karakoram Lesser Himalaya Probability Greater Himalaya Siwaliks Probability Thatta, Modern Indus (130 grains) 2% Nanga Parbat 19% Karakoram+Transhim. 38% Greater Himalaya. 41% Lesser Himalaya Probability TH-10-1, Age = 7.0 ka (186 grains) 0% Nanga Parbat 28% Karakoram+Transhim. 38% Greater Himalaya. 34% Lesser Himalaya Jati-16-1, Age = 8.0 ka (74 grains) Probability 3% Nanga Parbat 44% Karakoram+Transhim. 31% Greater Himalaya. 22% Lesser Himalaya Probability KB-34-1, Age = 12.6 ka (117 grains) 2% Nanga Parbat 41% Karakoram+Transhim. 35% Greater Himalaya. 22% Lesser Himalaya Probability KB-40-1 and -41-2, Age > 20 ka (271 grains) 1% Nanga Parbat 40% Karakoram+Transhim. 42% Greater Himalaya. 16% Lesser Himalaya 0 1000 2000 3000 Age (Ma) Fig. 12. Probability density plots showing the range of U– Pb ages in detrital zircons compared with source terrain values. Top section shows known range of possible source ages from the Greater Himalaya (Gehrels et al. 2006), Karakoram (Le Fort et al. 1983; Parrish & Tirrul 1989; Schärer et al. 1990; Fraser et al. 2001; Heuberger et al. 2007), Lesser Himalaya (Parrish & Hodges 1996; DeCelles et al. 2000; Chambers et al. 2008), Nanga Parbat (Zeitler & Chamberlain 1991; Zeitler et al. 1993), and the Siwaliks (DeCelles et al. 2000; Bernet et al. 2006). See supplementary material for data. MONSOON CONTROL OVER EROSION 209 Fig. 13. Pie diagrams showing the changes in the relative proportions of different zircon U –Pb age populations in sands from the Indus delta. See Figures 1 and 2 for sample locations. Population 0 –20 Ma is a proxy for flux from Nanga Parbat, whereas the 20– 55 Ma are likely from the Karakoram Batholith. 55– 300 Ma grains are dominantly from the Transhimalaya (Ladakh and Kohistan batholith). Grains dated 300–1400 Ma are typical of sources in the Greater Himalaya, while older grains are likely derived (directly or indirectly) from the Lesser Himalaya. 210 P. D. CLIFT ET AL. Table 7. Predicted percentages of eroded material from a variety of western Himalayan source in five Indus River sands spanning the LGM to present. The mean 1Nd values for each of the sources is a modal number derived from published Nd isotope measurements from the bedrock Source 1Nd KB-40-1 KB-34-4 Jati-16 TH-10-8 TH-1 Depositional age (ka) Nanga Parbat % Karakoram and Transhimalaya % High Himalaya % Lesser Himalaya % 225 þ1 216 224 20 1 40 42 17 13 3 39 36 22 8 3 44 31 22 7 2 37 39 22 0 2 22 36 40 210.7 210.8 211.4 211.8 210.5 n/a 211.7 212.9 215.7 215.4 Predicted 1Nd Observed 1Nd Source: Nanga Parbat data is from Clift et al. (2002b) and Whittington et al. (1999), Greater Himalaya data is from Ahmad et al. (2000), Deniel et al. (1987), Stern et al. (1989), France-Lanord et al. (1993), Parrish & Hodges (1996), Searle et al. (1997), Harrison et al. (1999), Whittington et al. (1999). Lesser Himalaya data is from Ahmad et al. (2000) and Parrish & Hodges (1996). Transhimalayan data is from Khan et al. (1997), Clift et al. (2000). Karakoram data is from Clift et al. (2002b) and Schärer et al. (1990). Monsoon and erosion patterns The different provenance proxies can be combined to generate a ‘best-fit’ sediment budget for the Indus since the LGM. Because of the large number of grains and samples and the good degree of separation between sources we choose to base the budget on the U –Pb zircon grains (Fig. 13), but then cross-check this by calculating what the Nd isotope composition of such a sediment mixture would be given the known range of Nd isotope characteristics in the sources. These predicted 1Nd values can then be compared with the actual measured values of these sediments (Clift et al. 2008). Because the Nd analyses are bulk analyses they would be expected to yield good averages of the sediment flux. In practice there are some significant departures between predicted and observed 1Nd values. This may reflect use of an inappropriate 1Nd value for the sources, yet we consider this unlikely because the modal values often lie close to measured values from the major modern rivers, which themselves should sample and average wide areas of the possible sources (Clift et al. 2002b). Alternatively, we suggest that the zircons, which are interpreted simply in the pie diagrams of Figure 13 are not as accurate as might be hoped in characterizing the total mass flux because of age overlap between the populations and sources. There exists a further possibility that there is shortterm variability in the sediment provenance that results in measurable differences in zircon populations for sediments that were deposited close together in time. In Table 7 we show a proposed erosion budget for the Indus based on five sand samples. We use the zircon populations shown in Figure 13 as a starting model for estimating the flux from each of the major sediment sources, but we adjust the relative proportions from each source to provide a closer match with the measured 1Nd values. Because the flux from the Karakoram and the Transhimalaya are hard to resolve from one another we plot these together as a single source. In each case the percentage adjustment from the observed was not more than 3%, and usually +2% or less. Figure 14 shows how this synthesized mass flux varies with time. The evolutionary patterns defined by this synthesis budget reflect many of the trends seen in the single mineral plots. Greater Himalayan flux remains high, if slightly variable throughout the period, as might be expected. Erosion from Nanga Parbat and the Lesser Himalaya because stronger Nanga Parbat Karakoram and Transhimalaya Proportion of total sediment (%) Discussion Greater Himalaya Lesser Himalaya 50 40 30 20 10 0 0 5 10 15 20 Age (ka) Fig. 14. Plot showing the evolving flux in zircon populations during the Holocene. We highlight the fall in relative flux from the Karakoram and Transhimalaya, compared to a sharp rise in the Lesser Himalaya, especially since 8 ka. MONSOON CONTROL OVER EROSION after the LGM, with a further sharp rise between 7 ka and the present day. At the same time flux from the Karakoram and Transhimalaya suffered a major decline. These changes may in part be related to the damming of the trunk river at Tarbela, which would raise the relative flux from the eastern tributaries draining the Himalaya, although these too have been dammed. Nonetheless, this theory does not explain the shift to similar negative 1Nd values in the Early Holocene. In this case we infer a similar shift to greatly enhanced erosion of the Lesser Himalaya relative to the Karakoram, peaking around 9 ka. The Nd isotopes suggest a moderate fall in Lesser Himalayan erosion after that time and before the most recent increase in the past 250 years. The later change in zircon sources (after 7 ka) compared to the earlier changes in Nd isotopes (10–14 ka) may also reflect a real lag in the sediment transport process for the zircon crystals versus clay minerals. The large scale changes in provenance tracked by the Nd isotopes (Figs 5 and 15) reflect a shift from preferential erosion of terrains lying to the north, around the Indus Suture Zone at the LGM to more erosion of the frontal Lesser Himalayan ranges in the south by the early Holocene. We do not think that drainage reorganization is responsible for the changing sediment compositions, even though this process has been used to explain changes in Nd isotopes after around 5 Ma (Clift & Blusztajn 2005). However, in this case the shift to more negative 1Nd values between 12 and 8 ka would require gain of isotopically negative sources. The increase in sediment flux at that time rules out loss of sources with positive 1Nd values as an alternative explanation. There is no evidence that the Punjabi tributaries, the Ravi, Jellum, Chenab and Sutlej, were captured into the Indus as recently as this time. Indeed, studies of Holocene drainage on the eastern edge of the Indus catchment indicate loss of drainage from Himalayan sources (Ghose et al. 1979) since the LGM, which would drive the opposite provenance shift than that observed. We conclude that the changes are driven by changing rates of sediment supply, not the wholesale capture of the Punjabi tributaries. When we compare the observed change in erosional style with records of the SW monsoon then we see that intensification of the summer rains [as tracked by speleothems (Fleitmann et al. 2003; Sinha et al. 2005) and pollen assemblages (Herzschuh 2006)] correlates with the strong change in Sr and Nd between 12 and 9 ka. This is also the time of accelerated sediment flux to the delta. Because the sediment composition changes quickly and is quite different from the sand deposited at the LGM we can rule out the sedimentation pulse as being caused by enhanced transport of 211 older glacially eroded sediment under the influence of the strong Early Holocene monsoon. In any case the Lesser Himalaya were not glaciated during the LGM (Owen & Benn 2005) and so the deglaciation process should not have directly affected the erosion of these ranges. Instead the stronger monsoon appears to be generating new sediment by erosion under its precipitation maximum. Satellite data shows that precipitation maxima in the western Himalaya are focused over the topographic breaks in the Lesser and Greater Himalaya (Bookhagen & Burbank 2006) and that the change in provenance was probably caused by a strengthening of rain and erosion in those zones. Studies of landslides in the western Himalayan confirms that the Early Holocene was a period of significant mass wasting and thus sediment production within the Himalaya (Bookhagen et al. 2005). In contrast, erosion in the Karakoram appears to be largely glacially driven and would have been strong at the LGM, as well as today. Because the Karakoram lie in the rain shadow of the Himalaya and derive most of their water via the westerly jet (Karim & Veizer 2002) their erosion would not change significantly as the summer monsoon intensified. Tectonics and the monsoon The primary conclusion of this study is that patterns of erosion across the western Himalaya changed significantly since 20 ka, driven by changes in summer monsoon intensity. Controls on erosion are important when considering the tectonic evolution of the Greater Himalaya because focused erosion is considered to have been a key factor in allowing deep buried metamorphic rocks to be exposed at the surface. This is true whether channelflow (Beaumont et al. 2001; Hodges et al. 2001) or orogenic wedge models are employed (Hilley & Strecker 2004) to explain the origin of the Greater Himalaya. Although we recognize that exhumation of ultra-high pressures, such as the Tso Moriri eclogites along the Indus Suture Zone are not erosionally driven (de Sigoyer et al. 2004; Leech et al. 2005), a purely tectonic origin for the Greater Himalaya is not currently favoured. Thermochronological transects across the Himalayan front suggest that zones of heavier precipitation correlate with areas of faster exhumation (Thiede et al. 2004; Wobus et al. 2005), although some indicators have been used to argue that rock uplift rather than monsoon rains dominate as drivers of erosion (Burbank et al. 2003). Furthermore, climaticallyfocused erosion appears to guide the location of active faults along the Himalayan front (Wobus et al. 2003). Our study reinforces the role of the monsoon in controlling orogenic architecture in the Himalaya. Without a strong summer monsoon 212 P. D. CLIFT ET AL. Fig. 15. Diagram showing the variability in various sediment and environmental proxies since the Last Glacial Maximum. Sediments are from the Indus delta and the Indus Canyon. (b) Nd data are from Clift et al. (2008). (c) Sr data are from Table 3. The Nd and Sr record are compared with (a) the GISP2 ice core climate record (Stuiver & Grootes 2000), (d) the variations in organic carbon isotope composition and (e) the intensity of the SW monsoon traced by speleothem records from Qunf and Timta Caves (Fleitmann et al. 2003; Sinha et al. 2005) in Oman and by pollen (Herzschuh 2006) from across Asia (black line), and well as western Himalayan landslides (Bookhagen et al. 2005). Note rapid change from C3 to C4 flora in early Holocene. MONSOON CONTROL OVER EROSION erosion is preferentially located in the Karakoram, such as at the LGM. However, for the exhumation of the Greater Himalaya to occur in the ways recently proposed a summer monsoon is crucial because the focused erosion required by such models does not occur in its absence. Our work also has implications for erosion/ tectonic coupling on a smaller scale. We show that the Nanga Parbat metamorphic massif, located in the western Himalayan syntaxis is a much less impressive sediment producer than its eastern twin at Namche Barwe, Tibet. France-Lanord et al. (2006) and Stewart et al. (2008) used a combination of thermochronological and U – Pb zircon data to indicate that as much as 45% of the sediment in the Brahmaputra is derived from erosion of the Namche Barwe massif, representing only 2% of the drainage system. In contrast, Nanga Parbat does not seem to contribute more than c. 3% of the sediment reaching the Indus delta. Our data casts some doubt over the idea that erosional unroofing by the Indus is driving the exhumation of the deep buried rocks (Zeitler et al. 2001) and that tectonic exhumation processes might also be significant (Hubbard et al. 1995). Why the two syntaxes behave in such different ways is unclear, although it is noteworthy that the Indus basin is much drier and generally much less erosive than the eastern Himalaya. Even within that region the effect of monsoon rain strength is a primary control on erosion rates (Galy & France-Lanord 2001). Monsoon and the environment The organic carbon data now for the first time allow us to examine the changing environments in the Indus drainage basin. Other climate indicators, such as lake sediments (Enzel et al. 1999) from the edge of the Thar Desert show that after the Early Holocene maximum summer monsoon strength declined towards the present day. The d13C record from Keti Bandar shows significant correlation with the speleothem climate records (Fig. 15). There is a minimum in d13C following the Younger Dryas at c. 12 ka during the earliest Holocene, and then a rapid rise in the Early Holocene. The low d13C values seen at 8.6 and 9.8 ka are interpreted to indicate a dominant input from rock-derived organic carbon, but even excluding these points there is a rise d13C values during the Early Holocene. We interpret the change in d13C to reflect an increase in marine organic carbon flux into the sediment at Keti Bandar, as result of the marine transgression. Interestingly, d13C stays high well after the speleothem records start to decline but then decreases rapidly after c. 4 ka, when the sedimentation again becomes fluvial. This indicates rapid reduction in 213 marine organic material after 4 ka. The carbon isotope data suggest that the summer monsoons in western India and Pakistan may not perfectly track those affecting Arabia and recorded in the Oman speleothem records (Fleitmann et al. 2003). The change in carbon isotopes has occurred most dramatically in the last 4 ka, and does not decrease gradually from 8 ka, as seen in the speleothems. A more detailed study, especially one using biomarkers, would be required to determine in detail when floral changes occurred, yet it is noteworthy that lake records from the environs of the Thar Desert show the most intense drying after 4.2 ka and were preceded by a much wetter period prior to that (Enzel et al. 1999). Conclusions In this study we employed a series of petrographic, geochemical and isotopic methods to examine the effect on climate change since 20 ka on the nature of erosion and environmental conditions in the Indus river basin. Although the mineralogy of the sediments did not change much during this time there are coherent changes towards more radiogenic Sr and unradiogenic Nd isotopes that reflect increasing erosion of ancient crust between 12 and 8 ka. This represents the transition from Younger Dryas to Early Holocene and is known as a period when summer monsoon rains strengthened (Fleitmann et al. 2003). AFT data shows that since the LGM the Indus preferentially eroded sources that have younger ages compared to those at 20 ka (more rapidly eroded sources). Ar–Ar mica dates also show a shift to younger cooling ages at this time and together with U –Pb zircon dating demonstrates that the greatest change has been a relative decrease in erosion from north of the Indus Suture (i.e. from the Karakoram and Transhimalaya) and an increase in erosion from the Lesser Himalaya. As these ranges now lie in the zone of heaviest monsoon rains (Bookhagen & Burbank 2006) we infer that the change in provenance deposited at the delta is caused by erosion modulated by monsoon intensity. We can rule out remobilization of glacially eroded sediments. The close correlation with the climate history also indicates that sediment flux from source to delta was rapid and lower than the uncertainties in the 14C dating. Carbon isotope data also argue for a change in environmental conditions, with a sharp change to more positive marine organic carbon values during the start of the Holocene, as sea-level rose. Minimum 1Nd values are followed by a moderate increase after 9 ka, before a further decrease in the past 250 years. We interpret this to reflect damming of the main Indus at Tarbela blocking the flux of 214 P. D. CLIFT ET AL. sediment from the suture zone. A drop in d13C values starting around 4 ka indicates a shift towards more terrestrial sources of organic carbon in the Indus basin, as the coast prograded southwards. The sediment record in the delta shows that erosion of the western Himalaya is strongly regulated by monsoon climate. Because current tectonic models for the exhumation of high grade metamorphic rocks in the Greater Himalaya require focused erosion our study suggests that formation of that range is dependent not just on the presence of tectonically thickened crust under southern Tibet but also on the activity of a strong summer monsoon. The pre-Holocene record shows that without the summer rains erosion is focused within and north of the Indus Suture and that consequently Himalayan exhumation cannot occur, at least not in the ways presently favoured by structural geologists. We thank US National Science Foundation (Ocean Sciences) for support of this project. References Ahmad, T., Harris, N., Bickle, M., Chapman, H., Bunbury, J. & Prince, C. 2000. Isotopic constraints on the structural relationships between the Lesser Himalayan Series and the High Himalayan Crystalline Series, Garhwal Himalaya. Geological Society of America Bulletin, 112, 467– 477. An, Z., Kutzbach, J. E., Prell, W. L. & Porter, S. C. 2001. Evolution of Asian monsoons and phased uplift of the Himalaya-Tibetan Plateau since late Miocene times. Nature (London), 411, 62– 66. Beaumont, C., Jamieson, R. A., Nguyen, M. H. & Lee, B. 2001. Himalayan tectonics explained by extrusion of a low-viscosity crustal channel coupled to focused surface denudation. Nature (London), 414, 738– 742. Bernet, M., van der Beek, P., Pik, R., Huyghe, P., Mugnier, J.-L., Labrin, E. & Szulc, A. G. 2006. Miocene to Recent exhumation of the central Himalaya determined from combined detrital zircon fissiontrack and U/Pb analysis of Siwalik sediments, western Nepal. Basin Research, 18, 393–412, doi: 10.1111/ j.1365– 2117.2006.00303. Black, L. P., Kamo, S. L. et al. 2004. Improved Pb-206/ U-218 microprobe geochronology by the monitoring of a trace-element-related matrix effect; SHRIMP, ID-TIMS, ELA-ICP-MS and oxygen isotope documentation for a series of zircon standards. Chemical Geology, 205, 115– 140. Blum, J. D. & Erel, Y. 1995. A silicate weathering mechanism linking increase in marine 87Sr/86Sr with global glaciation. Nature, 373, 415 –418. Bojar, A.-V., Fritz, H., Nicolescu, S., Bregar, M. & Gupta, R. P. 2005. Timing and mechanisms of Central Himalayan exhumation: discriminating between tectonic and erosion processes. Terra Nova, 17, 427–433, doi: 10.1111/j.1365-3121.2005.00629.x. Bollinger, L., Avouac, J. P., Beyssac, O., Catlos, E. J., Harrison, T. M., Grove, M., Goffe, B. & Sapkota, S. 2004. Thermal structure and exhumation history of the Lesser Himalaya in central Nepal. Tectonics, 23, 19, doi: 10.1029/2003TC001564. Bookhagen, B. & Burbank, D. W. 2006. Topography, relief, and TRMM-derived rainfall variations along the Himalaya. Geophysical Research Letters, 33, doi: 10.1029/2006GL026037. Bookhagen, B., Thiede, R. C. & Strecker, M. R. 2005. Late Quaternary intensified monsoon phases control landscape evolution in the northwest Himalaya. Geology, 33, 149 –152. Brookfield, M. E. & Reynolds, P. H. 1990. Miocene Ar-40/Ar-39 Ages from the Karakorum Batholith and Shyok Melange, Northern Pakistan, Indicate Late Tertiary Uplift and Southward Displacement. Tectonophysics, 172, 155– 167. Burbank, D. W., Blythe, A. E., Putkonen, J., Pratt-Sitaula, B., Gabet, E., Oskins, M., Barros, A. & Ojha, T. P. 2003. Decoupling of erosion and precipitation in the Himalayas. Nature, 426, 652–655. Camoin, G. F., Montaggioni, L. F. & Braithwaite, C. J. R. 2004. Late glacial to post glacial sea levels in the western Indian Ocean. Marine Geology, 206, 119–146, doi: 10.1016/j.margeo.2004.02.003. Campbell, I. H., Reiners, P. W., Allen, C. M., Nicolescu, S. & Upadhyay, R. 2005. He– Pb double dating of detrital zircons from the Ganges and Indus rivers; implication for quantifying sediment recycling and provenance studies. Earth and Planetary Science Letters, 237, 402– 432. Catlos, E. J., Harrison, T. M., Kohn, M. J., Grove, M., Ryerson, F. J., Manning, C. E. & Upreti, B. N. 2001. Geochronologic and thermobarometric constraints on the evolution of the Main Central Thrust, central Nepal Himalaya. Journal of Geophysical Research, 106, 16 177 –16 204. Cerveny, P. F., Johnson, N. M., Tahirkheli, R. A. K. & Bonis, N. R. 1989. Tectonic and geomorphic implications of Siwalik Group heavy minerals, Potwar Plateau, Pakistan. In: Malinconico, L. L. & Lillie, R. J. (eds) Tectonics of the Western Himalayas. Geological Society of America, Boulder, Colorado, Special Paper, 232, 129–136. Chambers, J. A., Argles, T., Horstwood, M. S. A., Harris, N. B. W., Parrish, R. & Ahmad, T. 2008. Tectonic implications of Palaeoproterozoic anatexis and Late Miocene metamorphism in the Lesser Himalayan Sequence, Sutlej Valley, NW India. Journal of the Geological Society, 165, 725– 737, doi: 10.1144/0016-76492007/090. Clemens, S., Prell, W., Murray, D., Shimmield, G. & Weedon, G. 1991. Forcing mechanisms of the indian-ocean monsoon. Nature, 353, 720– 725. Clift, P. D. & Blusztajn, J. S. 2005. Reorganization of the western Himalayan river system after five million years ago. Nature, 438, 1001– 1003. Clift, P. D., Degnan, P. J., Hannigan, R. & Blusztajn, J. 2000. Sedimentary and geochemical evolution of the Dras forearc basin, Indus suture, Ladakh Himalaya, MONSOON CONTROL OVER EROSION India. Geological Society of America Bulletin, 112, 450–466. Clift, P. D., Carter, A., Krol, M. & Kirby, E. 2002a. Constraints on India; Eurasia collision in the Arabian Sea region taken from the Indus Group, Ladakh Himalaya, India. In: Clift, P. D., Kroon, D., Gaedicke, C. & Craig, J. (eds) The Tectonic and Climatic Evolution of the Arabian Sea Region. Geological Society, London, Special Publications, 195, 97–116. Clift, P. D., Lee, J. I. et al. 2002b. Nd and Pb isotope variability in the Indus River system; implications for sediment provenance and crustal heterogeneity in the western Himalaya. Earth and Planetary Science Letters, 200, 91–106. Clift, P. D., Campbell, I. H. et al. 2004. Thermochronology of the modern Indus River bedload; new insight into the control on the marine stratigraphic record. Tectonics, 23, doi: 10.1029/2003TC001559. Clift, P., Giosan, L., Blusztajn, J. et al. 2008. Holocene erosion of the Lesser Himalaya triggered by intensified summer monsoon. Geology, 36, 79– 82, doi: 10.1130/G24315A.1. Collister, J. W., Lichtfouse, E., Hieshima, G. & Hayes, J. M. 1994. Partial resolution of n-alkanes in the saline portion of the Parachute Creek Member, Green River Formation (piceance Creek Basin, Colorado). Organic Geochemistry, 21, 645–659. Copeland, P., Harrison, M. & Le Fort, P. 1990. Age and cooling history of the Manaslu granite: implications for Himalayan tectonics. Journal of Volcanology and Geothermal Research, 44, 33– 50. De Sigoyer, J., Guillot, S. & Dick, P. 2004. Exhumation of the ultrahigh-pressure Tso Morari Unit in eastern Ladakh (NW Himalaya); a case study. Tectonics, 23, 18. DeCelles, P. G., Gehrels, G. E., Quade, J., LaReau, B. & Spurlin, M. 2000. Tectonic implications of U– Pb zircon ages of the Himalayan orogenic belt in Nepal. Science, 288, 497– 499. Deines, P. 1980. The isotopic composition of reduced organic carbon. In: Fritz, P. & Fontes, J. C. (eds) Handbook of Environmental Isotope Chemistry 1. Elsevier, Amsterdam, 329–406. Deniel, C., Vidal, P., Fernandez, A., Lefort, P. & Peucat, J. J. 1987. Isotopic Study of the Manaslu Granite (Himalaya, Nepal) – Inferences on the Age and Source of Himalayan Leukogranites. Contributions to Mineralogy and Petrology, 96, 78–92. Derry, L. & France-Lanord, C. 1996. Neogene Himalayan weathering history and river 87Sr/86Sr: impact on the marine Sr record. Earth and Planetary Science Letters, 142, 59–74. Dickinson, W. R. 1985. Interpreting provenance relations from detrital modes of sandstones. In: Zuffa, G. G. (ed.) Provenance of Arenites. Reidel, Dordrecht, NATO Advanced Studies Institute, 148, 333– 361. Enzel, Y., Ely, L. L. et al. 1999. High-resolution Holocene environmental changes in the Thar Desert, northwestern India. Science, 284, 125– 128. Farquhar, G. D., Ehleringer, J. R. & Hubrick, K. T. 1989. Carbon isotopic discrimination and photosynthesis. Annual Reviews of Plant Physiology and Molecular Biology, 40, 503– 537. 215 Fleitmann, D., Burns, S. J., Mudelsee, M., Neff, U., Kramers, J., Mangini, A. & Matter, A. 2003. Holocene forcing of the Indian monsoon recorded in a stalagmite from southern Oman. Science, 300, 1737– 1739. Fontugne, M. R. & Duplessy, J. C. 1978. Carbon isotope ratio of marine plankton related to surface water masses. Earth and Planetary Science Letters, 41, 365– 371. Foster, D. A., Gleadow, A. J. W. & Mortimer, G. 1994. Rapid Pliocene exhumation in the Karakoram (Pakistan), revealed by fission-track thermochronology of the K2 gneiss. Geology, 22, 19–22. France-Lanord, C. & Le Fort, P. 1988. Crustal melting and granite genesis during the Himalayan collision orogenesis. Transactions of the Royal Society of Edinburgh, Earth Sciences, 79, 183–195. France-Lanord, C., Derry, L. & Michard, A. 1993. Evolution of the Himalaya since Miocene time: isotopic and sedimentologic evidence from the Bengal Fan In: Treloar, P. J. & Searle, M. P. (eds) Himalayan Tectonics. Geological Society, London, Special Publications, 74, 603– 621. France-Lanord, C., Galy, V., Pik, R. & Singh, S. K. 2006. Extreme Erosion of the Eastern Himalayan Syntaxis Traced by Isotopic Compositions of River and Bengal Fan Sediments. Eos, Transactions, American Geophysical Union, 87, T32B-03. Fraser, J. E., Searle, M. P., Parrish, R. R. & Noble, S. R. 2001. Chronology of deformation, metamorphism, and magmatism in the southern Karakoram Mountains. Geological Society of America Bulletin, 113, 1443– 1455. Galbraith, R. F. 1990. The radial plot: graphical assessment of spread in ages Nuclear Tracks and Radiation Measurement, 17, 207 –214. Galy, A. & France-Lanord, C. 2001. Higher erosion rates in the Himalaya: geochemical constraints on riverine fluxes. Geology, 29, 23–26. Galy, V., Bouchez, J. & France-Lanord, C. 2007a. Determination of total organic carbon content and 13 C in carbonate-rich detrital sediments. Geostandards and Geoanalytical Research, 31, 199– 207, doi: 10.1111/j.1751-908X.2007.00864.x. Galy, V., France-Lanord, C., Beyssac, O., Faure, P., Kudrass, H.-R. & Palhol, F. 2007b. Efficient organic carbon burial in the Bengal fan sustained by the Himalayan erosional system. Nature, 450, 407– 411, doi: 10.1038/nature06273. Galy, V., France-Lanord, C. & Lartiges, B. 2008. Loading and fate of particulate organic carbon from the Himalaya to the Ganga–Brahmaputra delta. Geochimica et Cosmochimica Acta, 72, 1767–1787. Garzanti, E. & Andò, S. 2007. Heavy-mineral concentration in modern sands: implications for provenance interpretation. In: Mange, M. & Wright, D. (eds) Heavy Minerals in Use. Developments in Sedimentology Series, 58, Elsevier, Amsterdam, 517– 545. Garzanti, E. & Vezzoli, G. 2003. A classification of metamorphic grains in sands based on their composition and grade. Journal of Sedimentary Research, 73, 830 –837, doi: 10.1306/012203730830. 216 P. D. CLIFT ET AL. Garzanti, E., Vezzoli, G., Ando, S., Paparella, P. & Clift, P. D. 2005. Petrology of Indus River sands; a key to interpret erosion history of the western Himalayan syntaxis. Earth and Planetary Science Letters, 229, 287– 302. Garzanti, E., Andò, S. & Vezzoli, G. 2009. Grain-size dependence of sediment composition and environmental bias in provenance studies. Earth and Planetary Science Letters, 277, 422–432. Gehrels, G. E., DeCelles, P. G., Ojha, T. P. & Upreti, B. N. 2006. Geologic and U– Th–Pb geochronologic evidence for early Paleozoic tectonism in the Kathmandu thrust sheet, central Nepal Himalaya. Geological Society of America Bulletin, 118, 185– 198. George, M., Reddy, S. & Harris, N. 1995. Isotopic Constraints on the Cooling History of the Nanga Parbat-Haramosh Massif and Kohistan Arc, Western Himalaya. Tectonics, 14, 237– 252. Ghose, B., Kar, A. & Husain, Z. 1979. The lost courses of the Saraswati River in the Great Indian Desert; new evidence from Landsat imagery. Geographical Journal, 145, 446–451. Giosan, L., Clift, P. D., Blusztajn, J., Tabrez, A., Constantinescu, S. & Filip, F. 2006. On the control of climate- and human-modulated fluvial sediment delivery on river delta development: The Indus. Eos, Transactions, American Geophysical Union, 87, OS14A-04. Godin, L., Gleeson, T. P., Searle, M. P., Ullrich, T. D. & Parrish, R. R. 2006. Locking of southward extrusion in favour of rapid crustal-scale buckling of the Greater Himalayan Sequence, Nar Valley, central Nepal. In: Law, R. D., Searle, M. P. & Godin, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 269–292. Goodbred, S. L. Jr & Kuehl, S. A. 2000. Enormous Ganges-Brahmaputra sediment discharge during strengthened early Holocene monsoon. Geology (Boulder), 28, 1083–1086. Green, P. F. 1989. Thermal and tectonic history of the East Midlands shelf (onshore UK) and surrounding regions assessed by apatite fission track analysis. Journal of the Geological Society, 146, 755–773. Harris, N. 2007. Channel flow and the Himalayan-Tibetan orogen: a critical review. Journal of the Geological Society, 164, 511– 523, doi: 10.1144/001676492006-133. Harrison, T. M., Grove, M., McKeegan, K. D., Coath, C. D., Lovera, O. M. & Le Fort, P. 1999. Origin and episodic emplacement of the Manaslu intrusive complex, central Himalaya. Journal of Petrology, 40, 3– 19. Hattersley, P. W. 1982. d13C values of C4 types of grasses. Australian Journal of Plant Physiology, 9, 139– 154. Herzschuh, U. 2006. Palaeo-moisture evolution in monsoonal Central Asia during the last 50,000 years. Quaternary Science Reviews, 25, 163– 178. Heuberger, S., Schaltegger, U. et al. 2007. Age and isotopic constraints on magmatism along the Karakoram-Kohistan suture zone, NW Pakistan; evidence for subduction and continued convergence after India-Asia collision. Swiss Journal of Geosciences, 100, 85– 107. Hilley, G. E. & Strecker, M. R. 2004. Steady state erosion of critical Coulomb wedges with applications to Taiwan and the Himalaya. Journal of Geophysical Research, 109, doi: 10.1029/2002JB002284. Hodges, K. 2003. Geochronology and thermochronology in orogenic systems. In: Rudnick, R. (ed.) The Crust. Elsevier-Science, Amsterdam, 263– 292. Hodges, K., Hurtado, J. & Whipple, K. 2001. Southward extrusion of Tibetan crust and its effect on Himalayan tectonics. Tectonics, 20, 799– 809. Hodges, K. V., Wobus, C., Ruhl, K., Schildgen, T. & Whipple, K. 2004. Quaternary deformation, river steepening, and heavy precipitation at the front of the Higher Himalayan ranges. Earth and Planetary Science Letters, 220, 379– 389. Hubbard, M. S., Spencer, D. A. & West, D. P. 1995. Tectonic exhumation of the Nanga Parbat massif, northern Pakistan. Earth and Planetary Science Letters, 133, 213– 225. Hurford, A. 1990. Standardization of fission track dating calibration: recommendation by the Fission Track Working Group of the IUGS subcommission on geochronology. Chemical Geology, 80, 177– 178. Inger, S. 1998. Timing of an extensional detachment during convergent orogeny; new Rb-Sr geochronological data from the Zanskar shear zone, northwestern Himalaya. Geology, 26, 223–226. Ingersoll, R. V., Bullard, T. F., Ford, R. L., Grimm, J. P., Pickle, J. D. & Sares, S. W. 1984. The effect of grain size on detrital modes: a test of the GazziDickinson point-counting method. Journal of Sedimentary Petrology, 54, 103–116. Ittekkot, V. & Arain, R. 1986. Nature of particulate organic matter in the River Indus, Pakistan. Geochimica et Cosmochimica Acta, 50, 1643– 1653. Jain, A. K., Kumar, D., Singh, S., Kumar, A. & Lal, N. 2000. Timing, quantification and tectonic modelling of Pliocene– Quaternary movements in the NW Himalaya; evidence from fission track dating. Earth and Planetary Science Letters, 179, 437–451. Karim, A. & Veizer, J. 2002. Water balance of the Indus river basin and moisture source in the Karakoram and western Himalayas: implications from hydrogen and oxygen isotopes river water. Journal of Geophysical Research, 107, doi: 10.1029/2000JD000253. Khan, M. A., Stern, R. J., Gribble, R. F. & Windley, B. F. 1997. Geochemical and isotopic constraints on subduction polarity, magma sources, and palaeogeography of the Kohistan intra-oceanic arc, northern Pakistan Himalaya. Journal of the Geological Society, 154, 935–946. Kirstein, L. A., Sinclair, H. D., Stuart, F. M. & Dobson, K. 2006. Rapid early Miocene exhumation of the Ladakh batholith, western Himalaya. Geology, 34, 1049–1052, doi: 10.1130/G22857A. Kitoh, A. 2004. Effects of mountain uplift on East Asian summer climate investigated by a coupled atmosphere-ocean GCM. Journal of Climatology, 17, 783– 802. Koppers, A. A. P. 2002. ArArCALC - software for Ar-40/ Ar-39 age calculations. Computers & Geosciences, 28, 605–619. Krol, M. A., Zeitler, P. K., Poupeau, G. & Pecher, A. 1996. Temporal variations in the cooling and denudation history of the Hunza plutonic complex, Karakoram MONSOON CONTROL OVER EROSION Batholith, revealed by Ar-40/Ar-39 thermochronology. Tectonics, 15, 403–415. Kumar, A., Lal, N., Jain, A. K. & Sorkhabi, R. B. 1995. Late Cenozoic-Quaternary thermo-tectonic history of Higher Himalayan Crystalline (HHC) in KishtwarPadar-Zanskar region, NW Himalaya; evidence from fission track ages. Journal of the Geological Society of India, 45, 375 –391. Le Fort, P., Debon, F. & Sonet, J. 1983. Petrography, geochemistry and geochronology of some samples from the Karakoram Batholith (N. Pakistan). In: Shams, F. A. (eds) Granites of the Himalayas, Karakoram and Hindu Kush. Punjab University, Lahore, 377– 387. Leech, M. L., Singh, S., Jain, A. K., Klemperer, S. L. & Manickavasagam, R. M. 2005. The onset of India-Asia continental collision: early, steep subduction required by the timing of UHP metamorphism in the western Himalaya. Earth and Planetary Science Letters, 234, 83–97. Mange, M. A. & Maurer, H. F. W. 1992. Heavy Minerals in Colour. Chapman and Hall, London. McBride, M. B. 1994. Environmental Chemistry of Soils. Oxford University Press, New York. Metcalfe, R. P. 1993. Pressure, temperature and time constraints on metamorphism across the Main Central Thrust zone and high Himalayan slab in the Garhwal Himalaya. In: Treloar, P. J. & Searle, M. P. (eds) Himalayan Tectonics. Geological Society, London, Special Publications, 74, 485–509. Najman, Y. & Garzanti, E. 2000. Reconstructing early Himalayan tectonic evolution and paleogeography from Tertiary foreland basin sedimentary rocks, northern India. Geological Society of America Bulletin, 112, 435– 449. Nelson, K. D., Zhao, W., Brown, L. D. et al. 1996. Partially molten middle crust beneath southern Tibet; synthesis of Project INDEPTH results. Science, 274, 1684–1688. Owen, L. A. & Benn, D. I. 2005. Equilibrium-line altitudes of the last glacial maximum for the Himalaya and Tibet; an assessment and evaluation of results. Quaternary International, 138– 139, 55–78. Palmer, M. R. & Edmond, J. M. 1991. Controls over the strontium isotope composition of river water. Geochimica et Cosmochimica Acta, 56, 2099–2111. Parrish, R. R. & Hodges, K. V. 1996. Isotopic constraints on the age and provenance of the Lesser and Greater Himalayan sequences, Nepalese Himalaya. Geological Society of America Bulletin, 108, 904– 911. Parrish, R. R. & Tirrul, R. 1989. U–Pb age of the Baltoro Granite, Northwest Himalaya, and implications for monazite U –Pb systematics. Geology, 17, 1076–1079. Pearce, N. J. G., Perkins, W. T., Westgate, J. A., Gorton, M. P., Jackson, S. E., Neal, C. R. & Chenery, S. P. 1997. A compilation of new and published major and trace element data for NIST SRM 610 and NIST SRM 612 glass reference materials. Geostandards Newsletter – the Journal of Geostandards and Geoanalysis, 21, 115–144. Poupeau, G., Pecher, A., Benharbit, M. & Noyan, O. F. 1991. Ages traces de fission sur apatites et taux de denudation plio-quaternaires au Karakorum 217 central. Comptes Rendus de l’Academie des Sciences, Serie II Sciences de la Terre et des Planètes, 313, 917– 922. Renne, P. R., Swisher, C. C., Deino, A. L., Karner, D. B., Owens, T. L. & DePaolo, D. J. 1998. Intercalibration of standards, absolute ages and uncertainties in Ar-40/Ar-39 dating. Chemical Geology, 145, 117–152. Robinson, D. M., DeCelles, P. G., Garzione, C. N., Pearson, O. N., Harrison, T. M. & Catlos, E. J. 2003. Kinematic model for the Main Central Thrust in Nepal. Geology, 31, 359–362. Ruhl, K. W. & Hodges, K. V. 2005. The use of detrital mineral cooling ages to evaluate steady state assumptions in active orogens: an example from the central Nepalese Himalaya. Tectonics, 24, doi: 10.1029/ 2004TC001712. Scaillet, B., France-Lanord, C. & Le Fort, P. 1990. Badrinath-Gangotri plutons (Garhwal, India): petrological and geochemical evidence for fractionation processes in a High Himalayan leucogranite. Journal of Volcanology and Geothermal Research, 44, 163– 188. Schärer, U., Copeland, P., Harrison, T. M. & Searle, M. P. 1990. Age, cooling history, and origin of postcollisional leucogranites in the Karakoram Batholith; a multi-system isotope study. Journal of Geology, 98, 233 –251. Searle, M. P., Rex, A. J., Tirrul, R., Rex, D. C., Barnicoat, A. & Windley, B. F. 1989. Metamorphic, magmatic and tectonic evolution of the Central Karakoram in the Biafo-Baltoro-Hushe regions of north Pakistan. Geological Society of America Special Paper, 232, 47–73. Searle, M. P., Waters, D. J., Rex, D. C. & Wilson, R. N. 1992. Pressure, temperature, and time constraints on Himalayan metamorphism from eastern Kashmir and western Zanskar. Journal of the Geological Society, London, 149, 753–773. Searle, M. P., Parrish, R. R., Hodges, K. V., Hurford, A., Ayres, M. W. & Whitehouse, M. J. 1997. Shisha Pangma leucogranite, South Tibetan Himalaya; field relations, geochemistry, age, origin, and emplacement. Journal of Geology, 105, 295– 317. Searle, M. P., Noble, S. R., Hurford, A. J. & Rex, D. C. 1999. Age of crustal melting, emplacement and exhumation history of the Shivling leucogranite, Garhwal Himalaya. Geological Magazine, 136, 513–525. Sinha, A., Cannariato, K. G. et al. 2005. Variability of southwest Indian summer monsoon precipitation during the Bolling-Allerod. Geology, 33, 813– 816. Slingerland, R. 1984. Role of hydraulic sorting in the origin of fluvial placers. Journal of Sedimentary Research, 54, 137–150. Smith, H. A., Chamberlain, C. P. & Zeitler, P. K. 1992. Documentation of Neogene regional metamorphism in the Himalayas of Pakistan using U–Pb in monazite. Earth and Planetary Science Letters, 113, 93– 105. Sorkhabi, R. B., Stump, E., Foland, K. & Jain, A. K. 1996. Fission-track and 40Ar/39Ar evidence for episodic denudation of the Gangotri granites in the Garhwal Higher Himalaya, India. Tectonophysics, 260, 187–199. Staubwasser, M., Sirocko, F., Grootes, P. M. & Erlenkeuser, H. 2002. South Asian monsoon climate change and radiocarbon in the Arabian Sea 218 P. D. CLIFT ET AL. during early and middle Holocene. Paleoceanography, 17, doi: 10.1029/2000PA000608. Stephenson, B. J., Searle, M. P., Waters, D. J. & Rex, D. C. 2001. Structure of the Main Central Thrust zone and extrusion of the High Himalayan deep crustal wedge, Kishtwar-Zanskar Himalaya. Journal of the Geological Society of London, 158, 637– 652. Stern, C. R., Kligfield, R., Schelling, D., Virdi, N. S., Futa, K. & Peterman, Z. E. 1989. The Bhagirathi leucogranite of High Himalaya (Garhwal, India): age, petrogenesis, and tectonic implications. In: Malinconico, L. L. (ed.) Tectonics of the Western Himalayas. Geological Society of America, Boulder, Special paper, 232, 33–45. Stewart, R. J., Hallet, B., Zeitler, P. K., Malloy, M. A., Allen, C. M. & Trippett, D. 2008. Brahmaputra sediment flux dominated by highly localized rapid erosion from the easternmost Himalaya. Geology, 36, 711– 714, doi: 10.1130/G24890A.1. Stuiver, M. & Grootes, P. M. 2000. GISP2 oxygen isotope ratios. Quaternary Research, 53, 277–284. Stüwe, K. & Foster, D. 2001. 40Ar/39Ar, pressure, temperature, and fission track constraints on the age and nature of metamorphism around the Main Central thrust in eastern Bhutan Himalaya. Journal of Asian Earth Sciences, 19, 85–95. Szulc, A. G., Najman, Y. et al. 2006. Tectonic evolution of the Himalaya constrained by detrital 40Ar/39Ar, Sm/Nd and petrographic data from the Siwalik foreland basin succession, SW Nepal. Basin Research, 18, 375–391. Thiede, R. C., Bookhagen, B., Arrowsmith, J. R., Sobel, E. R. & Strecker, M. R. 2004. Climatic control on rapid exhumation along the Southern Himalayan Front. Earth and Planetary Science Letters, 222, 791–806. Treloar, P. J., Rex, D. C., Guise, P. G., Wheeler, J., Hurford, A. J. & Carter, A. 2000. Geochronological constraints on the evolution of the Nanga Parbat syntaxis, Pakistan Himalaya. In: Khan, M. A., Treloar, P. J., Searle, M. P. & Jan, M. Q. (eds) Tectonics of the Nanga Parbat Syntaxis and the Western Himalaya. Geological Society, London, Special Publications, 170, 137–162. Trivedi, J. R., Gopalan, K. & Valdiya, K. S. 1984. Rb-Sr ages of granitic rocks within the Lesser Himalayan nappes, Kumaun, India. Geological Society of India Journal, 25, 641– 654. van der Beek, P., Robert, X., Mugnier, J.-L., Bernet, M., Huyghe, P. & Labrin, E. 2006. Late MioceneRecent exhumation of the central Himalaya and recycling in the foreland basin assessed by apatite fission-track thermochronology of Siwalik sediments, Nepal. Basin Research, 18, 413– 434. van der Merwe, N. J. & Medina, E. 1989. Photosynthesis and 13C/12C ratios in Amazon rain forests. Geochimica et Cosmochimica Acta, 53, 1091– 1094. Vannay, J.-C., Grasemann, B., Rahn, M., Frank, W., Carter, A., Baudraz, V. & Cosca, M. 2004. Miocene to Holocene exhumation of metamorphic crustal wedges in the NW Himalaya; evidence for tectonic extrusion coupled to fluvial erosion. Tectonics, 23, 24. Villa, I. M., Lemennicier, Y. & LeFort, P. 1996. Late Miocene to Early Pliocene tectonometamorphism and cooling in south-central Karakorum and IndusTsangpo suture, Chogo Lungma area (NE Pakistan). Tectonophysics, 260, 201– 214. Walker, J., Martin, M. W., Bowring, S. A., Searle, M., Waters, D. J. & Hodges, K. 1999. Metamorphism, melting, and extension: age constraints from the High Himalayan Slab of southeast Zanskar and northwest Lahaul. Journal of Geology, 107, 473–495. Wang, Y., Cheng, H., Edwards, R. L. et al. 2005. The Holocene Asian monsoon; links to solar changes and North Atlantic climate. Science, 308, 854–857. Wang, Y., Li, Q. & Qu, G. 2006. 40Ar/39Ar thermochronological constraints on the cooling and exhumation history of the South Tibetan Detachment System, Nyalam area, southern Tibet. In: Law, R. D., Searle, M. P. & Godin, L. (eds) Channel Flow, Ductile Extrusion, and Exhumation of Lower-Middle Crust in Continental Collision Zones. Geological Society, London, Special Publications, 268, 327– 354. White, N. M., Pringle, M., Garzanti, E., Bickle, M., Najman, Y., Chapman, H. & Friend, P. 2002. Constraints on the exhumation and erosion of the High Himalayan Slab, NW India, from foreland basin deposits. Earth and Planetary Science Letters, 195, 29–44. Whittington, A., Foster, G., Harris, N., Vance, D. & Ayres, M. 1999. Lithostratigraphic correlations in the western Himalaya – An isotopic approach. Geology, 27, 585 –588. Winslow, D. M., Zeitler, P. K., Chamberlain, C. P. & Williams, I. S. 1996. Geochronologic constraints on syntaxial development in the Nanga Parbat region, Pakistan. Tectonics, 15, 1292– 1308. Wobus, C. W., Hodges, K. V. & Whipple, K. X. 2003. Has focused denudation sustained active thrusting at the Himalayan topographic front? Geology, 31, 861–864. Wobus, C., Heimsath, A., Whipple, K. & Hodges, K. 2005. Active out-of-sequence thrust faulting in the central Nepalese Himalaya. Nature, 434, 1008–1011. Zeilinger, G., Burg, J. P., Schaltegger, U. & Seward, D. 2001. New U/Pb and fission track ages and their implication for the tectonic history of the lower Kohistan Arc Complex, northern Pakistan. Journal of Asian Earth Sciences, 19, 79–81. Zeitler, P. K. 1985. Cooling History of the NW Himalaya, Pakistan. Tectonics, 4, 127– 151. Zeitler, P. K. & Chamberlain, C. P. 1991. Petrogenetic and Tectonic Significance of Young Leukogranites from the Northwestern Himalaya, Pakistan. Tectonics, 10, 729– 741. Zeitler, P. K., Sutter, J. F., Williams, I. S., Zartman, R. E. & Tahirkheli, R. A. K. 1989. Geochronology and temperature history of the Nanga ParbatHaramosh Massif, Pakistan. In: Malinconico, L. L. & Lillie, R. J. (eds) Tectonics of the Western Himalayas. Geological Society of America, Boulder, 232, 1– 22. Zeitler, P. K., Chamberlain, C. P. & Smith, H. A. 1993. Synchronous Anatexis, Metamorphism, and Rapid Denudation at Nanga-Parbat (Pakistan Himalaya). Geology, 21, 347 –350. Zeitler, P. K., Koons, P. O. et al. 2001. Crustal reworking at Nanga Parbat, Pakistan: metamorphic consequences of thermal-mechanical coupling facilitated by erosion. Tectonics, 20, 712– 728.