Corresponding author: M. J. Streule, Department of Earth Sciences, University of Oxford, South Parks Road, Oxford OX1 3AN, UK. (email@example.com)
 New apatite and zircon fission track (FT) data from the summit slopes of Everest and along the Barun, Arun, Dudh Kosi, and Kangshung valleys that drain the Everest and Makalu massifs cover a vertical sample transect of almost 8000 m of the Eastern Nepal Greater Himalayan Sequence (GHS). Apatite FT ages range from 0.9 ± 0.3 Ma to 3.1 ± 0.3 Ma in the GHS with ages increasing systematically with elevation. Apatite FT ages in the Everest Series and summit Ordovician limestones are much older, up to 30.5 ± 5.1 Ma. Zircon FT ages from the GHS range from 3.8 ± 0.4 Ma to 16.3 ± 0.8 Ma. The brittle exhumation rates calculated from these data show the GHS was exhumed, since ∼9 Ma, at an average rate of 1.0 ± 0.2 mm/a. Pliocene exhumation rates are higher at 1.7 ± 0.3 mm/a. These values are not significantly different from the estimate of ductile exhumation rates of 1.8 mm/a recorded by metamorphic minerals undergoing decompression between 18.7 and 15.6 Ma but are well below the values (up to 10 mm/a) used by thermomechanical models for ductile channel flow in the GHS. If representative of the GHS these models will therefore require further ‘tuning’. Higher exhumation rates in the Pliocene have also been observed in other parts of the Himalaya and points to a regional cause, likely increased erosion due to the onset of late Pliocene–Pleistocene glaciation of the high Himalaya.
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 The presence of melt in the midcrust linked as a continuous layer to the GHS has been used in thermal-mechanical and numerical models to show that channelized extrusion is a viable process to accommodate ongoing crustal thickening providing it is linked to focused surface erosion at the Himalayan front [Beaumont et al., 2001, 2004, 2006; Jamieson et al., 2004, 2006]. Channel flow models predict exhumation rates of up to 10 mm/a and only a very small amount of this can be accommodated as surface uplift or net transfer of material to the upper crust; therefore erosion along the Himalaya would have to be intense enough for around 10 mm/a of material to also be removed. If these models are correct and there is continuity of the channel from surface to midcrust then the ductile process of flow at depth must be directly linked to exhumation in the brittle field. However, such models are based mostly on assumed, simplified or idealized physical starting parameters that can be adjusted until favorable results are produced.
 In order to fully evaluate models of channel flow and to understand the exhumation history in the Himalaya a comparison must be made between the rates of more recent brittle field exhumation to the intensity of erosion that facilitates the exhumation with the rates cited in the channel flow model. Are the extremely rapid rates of exhumation of the order 10 mm/a used in the channel flow model realistic? One of the best areas to study the GHS is the Everest and Makalu region of eastern Nepal and South Tibet due to the extreme relief and most northerly location of Barrovian facies GHS rocks exposed beneath the southern Tibet plateau. To examine the history of brittle field exhumation and assess the values used in the channel flow model we conducted a regional fission track study of GHS rocks in the Everest and Makalu massifs. The samples include north-south profiles along the Dudh Kosi–Everest-Kangshung valley in Nepal and South Tibet and the Arun valley leading up to Makalu in Nepal (Figure 1).
2. Geology of the Everest-Makalu Himalaya
 The peak of Mount Everest (Figure 3) straddles the upper units of the GHS with the summit rocks being unmetamorphosed Ordovician lime-mudstones that contain relics of crinoid stems and trilobite fragments [Searle et al., 2003, 2006]. The upper strand of the STD, the Qomolangma Detachment (QD) separates these rocks from the metamorphosed ‘Yellow Band’, a middle Cambrian calc-schist, in part mylonitised, that contains white mica laths and recrystallized calcite grains and no sedimentary structures, beneath. Below the Yellow Band the ‘Everest Series’ comprises black-colored schists occasionally containing garnet and staurolite metamorphosed at amphibolite (∼650°C; 6.2 kbar) [Jessup et al., 2008] to greenschist facies. Decompression to ∼3 kbar following prograde metamorphism is apparent from compositional zoning of garnets and staurolites with cordierite overgrowths. The Lhotse Detachment (LD) is a ductile shear zone that separates the Everest Series rocks above from sillimanite grade gneisses, migmatites and abundant leucogranites beneath [Searle, 1999a, 1999b, 2003; Searle et al., 2003, 2006; Jessup et al., 2008].
 Leucogranites are present along the upper structural levels of the GHS along the entire range but are particularly thick in the Makalu region. The early garnet, two-mica tourmaline leucogranites along the Barun glacier and base of Everest can be traced west directly into the Nuptse and Everest leucogranites [Searle, 2007]. However, on Makalu these early granites have been crosscut and intruded by a later phase of cordierite-bearing leucogranite that makes up the upper pyramid of Makalu [Streule et al., 2010]. The lower GHS comprises a series of sillimanite grade gneisses with uncommon bands of kyanite grade gneisses down as far as the Main Central thrust zone of inverted metamorphism [Hubbard, 1989; Searle et al., 2003; Jessup et al., 2006].
 The timing of prograde metamorphism in the Everest region is well constrained by U-Th-Pb dating of monazites peaking as early as 38.9 ± 0.9 Ma (kyanite grade; 550–650°C and 0.8–1.0 GPa) and continuing to 28.0 ± 1.2 Ma (sillimanite grade; 650–750°C and 0.4–0.7 GPa) with ages becoming younger northward to ∼25.4–16.1 ± 0.1 Ma [Simpson et al., 2000; Jessup et al., 2008; Cottle et al., 2009a]. The youngest granulite facies metamorphism (∼750°C; 8–10 kbar;) associated with crustal melting occurred at 13.2 ± 1.4 Ma in the Ama Drime range, the northward prolongation of the GHS northeast of Everest [Cottle et al., 2009b]. At the highest temperatures crustal melting resulted in production of migmatites and leucogranite melts that crystallized sporadically from at least ∼21 Ma up to the youngest dykes dated at 11.6 ± 0.4 Ma [Cottle et al., 2009b]. Leucogranites above Everest base camp on the Nepal side have U-Pb monazite and xenotime ages of 21.3–20.5 Ma [Simpson et al., 2000] and from near base camp on the Tibetan side have U-Pb monazite age of 16.9 ± 0.2 Ma [Searle et al., 2003]. To the north of Mount Everest the ages of both metamorphism and leucogranite formation become progressively younger with high temperatures persisting until the final phases of leucogranite melt injection at 15.2 ± 0.2 Ma, 12.6 ± 0.2 Ma and 11.6 ± 0.4 Ma [Cottle et al., 2009a, 2009b].
 In the Makalu–Arun region the timing of regional metamorphism has also been well constrained by U-Pb dating of monazites growing during peak sillimanite grade metamorphism (∼710°C; 0.59 GPa) while the migmatitic Barun gneisses have been dated at 15.6 ± 0.2 Ma to 16.0 ± 0.6 Ma [Streule et al., 2010]. Muscovite dehydration melting resulted in numerous Grt-Tur-Ms-Bt leucogranites that crystallized mainly during the period 24–21 Ma [Schärer, 1984; Streule et al., 2010]. Decompression melting during the later stage resulted in cordierite-bearing leucogranites (comprising most of the upper structural levels of Makalu) formed simultaneously with melting in the source Barun gneisses at P-T conditions of ∼700°C and 4 kbar. All high-grade gneisses, migmatites and leucogranites along the GHS are bounded along the top by a thick ductile shear zone (top to north); part of the South Tibetan Detachment (STD) system of low-angle normal faults [Burchfiel et al., 1992; Carosi et al., 1998; Searle, 1999b, 2003; Searle et al., 2003, 2006; Cottle et al., 2007, 2011]. On Everest, the STD splays into two fault systems, a lower ductile shear zone, the Lhotse detachment and an upper brittle fault, the Qomolangma detachment [Searle, 1999b, 2003], which merge toward the north along the Rongbuk valley into one large-scale ductile shear zone without a significant brittle detachment [Cottle et al., 2007, 2011]. In summary, the structural, metamorphic, and geochronological data show that the GHS beneath the Everest and Makalu massifs maintained high temperatures during a protracted sillimanite grade event lasting at least from ∼39 Ma up to ∼15 Ma with final leucogranite dykes injected as young as 11.6 Ma.
 Although the prograde metamorphic history and ages of magmatism in the Himalaya are reasonably well constrained, the later stages of exhumation and cooling are relatively unknown. Limited 40Ar/39Ar data from Everest suggest that a period of rapid cooling between 17 and 16 Ma immediately followed the main phase of leucogranite crystallization as the GHS midcrustal channel was exhumed toward the Earth's surface [Searle et al., 2003, 2006]. Additionally apatite and zircon fission track ages and 40Ar/39Ar data from one sample of the Yellow Band rocks, juxtaposed between the Lhotse and Qomolangma detachments near the summit of Everest was determined by Sakai et al.  and similarly showed increasingly rapid cooling of this unit from 24 to 14 Ma. What happened after this is not known and therefore to constrain the timing and rate of late Miocene and Pliocene exhumation in the Everest-Makalu area 22 samples were collected for fission track (FT) thermochronometry. The samples cover an elevation range from 8848 m to 886 m (Figure 3) and include eleven samples collected by Dana Coffield and analyzed by R. Donelick. These samples have the prefix DA-91 in Table 1, and the data have been previously reported in a conference abstract [Bergman et al., 1993]. The remaining samples (prefixes A and KG in Table 1) were collected and analyzed by the authors of this paper. The lowest-elevation sample (A99 at 886 m) is the furthest south sampled point (54 km). The majority of samples were collected either from Everest or nearby within 10–15 km.
Samples DA91 were analyzed by Donelick, all other samples counted by Streule and Carter. P indicates fission track densities and N is the numbers of tracks counted. Analyses were by the external detector method using 0.5 for the 4π/2π geometry correction factor. Ages calculated using the zeta calibration approached derived from multiple analyses of IUGS apatite and zircon age standards [Hurford, 1990] and corning dosimeter glasses. Zeta factors used were apatite 347 ± 14 (CN5, analyst Streule) and 339 ± 5 (CN5, analyst Carter), 118.9 ± 3.6 (CN1, analyst Donelick). For zircon: 127 ± 5 (CN2, analyst Carter), 135.0 ± 3.8 (CN1, analyst Donelick). Pχ2 is probability for obtaining χ2 value for n degrees of freedom, where n = no. crystals – 1.
Given as degrees, minutes, seconds.
 Samples were crushed and apatite and zircon were separated using standard heavy liquid and isodynamic techniques. Apatite and zircon separates were mounted, polished and etched following the standard procedures of the London Geochronology Center (apatite etch using 5M HNO3 at 20°C for 20 s and zircon etch using KOH-NaOH at 225°C with multiple Teflon mounts for variable durations up to 48 h). Mounts were irradiated in the FRM II thermal neutron facility at the University of Munich in Germany with muscovite external detectors and corning glass dosimeters. Fission track densities were measured using an optical microscope at 1250× magnification. Ages (±1σ) were calibrated by the zeta method [Hurford and Green, 1983; Hurford, 1990]. Only crystals with prismatic sections parallel to the crystallographic c axis were accepted for analysis.
 Results are shown in Table 1. Apatite FT data cover both the GHS and Everest Series rocks from above the STD. By contrast the zircon FT data are confined to the GHS. Low spontaneous track densities prevented measurement of confined track lengths in the younger apatite samples, however this was not problematic for the interpretation as these ages can only be explained by recent cooling. Compositional data were not incorporated into the final age calculation, as rapid cooling removes any resolvable compositional control on fission track annealing. FT ages are reported with 1σ error along with measures of grain age dispersion in Table 1.
 The youngest apatite FT ages occur within the GHS and range from 3.1 ± 0.3 Ma to 0.9 ± 0.3 Ma. The data show a linear relationship (R = 0.68) between age and sample elevation (higher-altitude samples are older than lower-altitude samples) up to 6000 m (Figure 4), consistent with the exhumation of a coherent rock unit. The five zircon FT ages range from 16.3 ± 0.8 Ma to 3.8 ± 0.4 Ma. Three of these samples cover a small elevation range from 2560 m and 3250 m and the other two samples come from similar elevations at 5663 m (A51) and 5620 m (DA91-78) but have contrasting zircon FT ages of 8.8 ± 0.4 and 16.3 ± 0.8 Ma respectively. Because there is a 2000 m data gap a linear correlation between age and elevation can be made to fit either of these two high-elevation zircon ages with the result that the exhumation rates will be different Therefore both options will be taken into account in the interpretation (Figure 4).
 The oldest apatite FT ages come from samples above the STD with mean FT ages between 30.5 ± 5.1 Ma and 9.7 ± 2.1 Ma. These data also show a linear correlation (R = 0.92) with ages increasing with elevation. A fourth apatite sample, DA91-79 is included in this group. It was collected near the STD and should be from within the uppermost part of the GHS but the apatite FT age matches more closely the age-elevation trend of the samples from above the STD. The unmetamorphosed sample of Ordovician Limestone collected from the summit slopes of Everest had an apatite fission track age of 30.5 ± 5.1 Ma indicating the Tethyan sediments were exhumed in the Oligocene, during crustal thickening prior to the main phase of metamorphism and granite production. However, the metamorphosed samples DA91-80 and DA91-82, from the Everest Series have younger apatite ages of 9.7 ± 2.1 Ma and 21.9 ± 2.7 Ma respectively indicating more recent cooling. There appears to be no consistent variation due to sample locations around the massif, or any systematic shift in ages due to changing rock types; for example across the kyanite bearing pelite band (Figure 3b), or from leucogranites to country rocks.
 If the entire data set recorded the behavior of a single coherent unit the marked break in slope in apatite FT age with elevation, between 6500 m and 7000 m, would imply exhumation rates increased at about 5 Ma. However, the break in slope coincides with the Qomolangma Detachment, the main low-angle fault of the STD system and therefore the two age-elevation slopes for the apatite FT data in Figure 4 reflect the behavior of two different structural units. The slope of the apatite FT ages from the GHS indicates an exhumation rate of ca. 2.6 ± 0.2 mm/a (1.8 ± 0.2 Myr age difference over an elevation range of 4777 m). By contrast the samples above the STD indicate an exhumation rate of 0.1 ± 0.08 mm/a (23.4 ± 5 Myr age difference over an elevation range of 2348 m). Interpretation of the zircon FT data is less straightforward due to two samples at similar elevation: A51 is from 5663 m and has an age of 8.8 ± 0.4 Ma and DA91-78 from 5620 m has an age of 16.3 ± 0.8 Ma. Lack of data points over a 2 km elevation range between 3250 m and 5620 m means that it is not clear which age should be used to estimate exhumation rates. We therefore consider both cases.
 The age difference or lag between the apatite and zircon ages for the same elevation can be used to track exhumation over a longer timespan/crustal depth. The lowest-elevation sample with paired FT ages (DA91-72) records a 2.1 Myr difference between the apatite and zircon FT ages. The next sample with paired ages (DA91-74) is 690 m higher and records an age difference of 4.7 Myrs. The oldest zircon FT age (sample DA91-78) is from 5620 m and although there is no apatite age for this sample similar elevation apatite ages indicate a lag of ∼13 Myrs. By contrast the paired ages for sample A51, which is from a similar elevation to sample DA91-78 records a lag of 5.7 Myrs. Although different the two high-elevation zircon FT ages are consistent in that they indicate slower exhumation during the Miocene compared to the Pliocene. If exhumation did not change the lag between the zircon and apatite FT ages would remain the same.
 While fission track ages can be used as proxy for cooling and therefore rates of rock uplift and exhumation; this has been particularly well demonstrated for other locations of the GHS [Grujic et al., 2006; Blythe et al., 2007; Patel and Carter, 2009], the determination of changing exhumation rates (or not) through time is not always so straightforward. The combined affects of extreme topography, thermal advection, lateral advection and fluid flow mean that the closure isotherms will not be horizontal and rock uplift rate will not equal exhumation rate. Exhumation of the thrust bound units such as the GHS is further complicated by the lateral transport of the exhuming material [Blythe et al., 2007]. The implication is that the exhumation rate values based on age-elevation profiles are unlikely to be correct.
 Numerical modeling by Huntington et al.  found that in situations where exhumation rates are high, samples collected over a lateral transect in the direction of transport (i.e., north-south in the Himalaya) lead to an underestimation of the exhumation rate that is approximately balanced (within error of the data) by an overestimation in exhumation rate due to topographic (and subsurface isotherm) effects. As a result the complexities of studying exhumation in the GHS largely balance one another out and any minor imbalances are irresolvable in comparison to the errors associated with the individual fission track ages themselves [Huntington et al., 2007]. This conclusion is further supported by the 3-D thermal model of Whipp et al.  which found no observable difference in fission track ages in response to lateral transport on thrust faults with a dipping angle as low as 10°. In the area of Everest, the STD is well exposed both on the summit pyramid (Figure 3b) and in the Rongbuk valley to the north, where it has a minimum dip angle of 15° [Searle et al., 2003], close to the 10° minimum limits of Whipp et al. . Therefore the data determined here can be considered to represent a vertical transect of fission track ages where only the relative vertical thermal advection and thermal conduction need to be taken into account to determine exhumation history; a complex 3-D modeling approach is not required as the caveats provided by other models are satisfied.
 A 1-D (vertical) approach to thermal modeling was therefore used to derive average exhumation rates for each of the periods of time represented by the apatite and zircon fission track data set. The AGE2EDOT program is used to estimate the cooling age for a thermochronometer that was exhumed by steady erosion at a constant rate. The first step in the calculation requires an appropriate closure temperature defined using the CLOSURE program of Brandon et al. ; temperatures of 135°C was derived for apatite, and 240°C for zircon, which are appropriate for the rapid cooling rates that are observed in the Himalaya [Blythe et al., 2007]. These closure temperatures are then used in the AGE2EDOT program, as summarized by Ehlers et al. . Other input parameters include a basal heat flow of 550°C, a thermal diffusivity of 10−6 m2s−1, which is only slightly variable around this value by less than 0.5 × 10−6 for metamorphic and magmatic rocks at the temperatures concerned [Vosteen and Schellschmidt, 2003; Whittington et al., 2009]. For heat production the values reported by Whipp et al.  are used which range between 0.8 and 3.5 μWm−3. An initial geothermal gradient of 25°Ckm−1 was assumed, which corresponds to a relatively high value for the continental crust where heat flow is high [Chapman, 1986] and exhumation is dominated by erosional advection. Surface temperature was taken to be 0°C. Calculations were undertaken for maximum and minimum values of heat production as this is a relatively poorly constrained value, particularly so for the Himalaya, where radiogenic heat production has been invoked as one of the primary causes of anatexis in the GHS.
 Output from the AGE2EDOT program for the input parameters described above is shown graphically in Figure 5; the relationship between apatite and zircon fission track age to calculated exhumation rate is shown for the maximum and minimum values of heat production used. The ages of samples A51 and DA91-72 are also plotted, as these are the two samples that contain both apatite and zircon that are the furthest apart in FT age. A summary of the data and results of the thermal modeling is shown in Table 2. While the curves for maximum and minimum heat production (which therefore yield maximum and minimum exhumation rates from determined ages) differ, the errors on the determined FT ages are of the order of, or in the case of DA91-72 exceed, the differences between the two curves. Therefore at the level of precision afforded by fission track chronometry, any changes in heat production have relatively little influence on the final exhumation rate. Given that the values of heat production used are at relatively extreme ends of the scale, an average of the values can be taken to represent the exhumation rate.
Table 2. Results of AGE2EDOT Calculations for the Apatite and Zircon Fission Track Ages of Samples DA91-72 and A51a
Fission Track Ages (Ma)
Exhumation Rates From Apatite Age (mm/a)
Exhumation Rates From Zircon Age (mm/a)
If 0.8 μW m−3 Heat Production
If 3.5 μW m−3 Heat Production
If 0.8 μW m−3 Heat Production
If 3.5 μW m−3 Heat Production
Average Exhumation Rate (mm/a)
Input parameters are described in section 3. Errors for exhumation rate values are 2σ.
3.1 ± 0.3
8.8 ± 0.4
1.0 ± 0.2
1.7 ± 0.7
3.8 ± 0.4
1.7 ± 0.3
 Using the constraints from the thermal modeling, an average exhumation rate during the period between the recorded zircon and apatite and ages (8.8 Ma to 3.1 Ma) of 1.0 ± 0.2 mm/a is determined for the highest-altitude sample, A51. If sample DA91-78 were used (ZFT age 16.3 Ma) instead of sample A51 this value would reduce further to ca. 0.6 mm/a. An average exhumation rate of 1.7 ± 0.3 mm/a is determined for the lowest-altitude sample, DA91-72 that records the most recent exhumation. These results show the exhumation rate has increased in this area of the Himalaya. Although a resolvable difference in exhumation rate is observed in the Pliocene the narrowing lag time between the AFT and ZFT data seen in Figure 4 suggest that this increase may have begun earlier. Both exhumation rates are lower than the 2.6 mm/a calculated from the age-elevation data but not to the extent that renders age-elevation derived exhumation rate as meaningless. The 1-D modeling does not alter the timing of any changes of rates of exhumation and on the basis of this exercise appears to only fine tune the exhumation rate data.
 The apatite and zircon fission track ages reported in this paper show a similar distribution to those recently documented from structurally equivalent units along strike from the whole of the Indian and Nepalese Himalaya [Schlup et al., 2003; Blythe et al., 2007; Thiede et al., 2009]. Following on from the main phase of melting and metamorphism in the GHS during the Miocene from ca. 21 to 16 Ma [Hodges, 2000; Simpson et al., 2000; Searle et al., 2003; Cottle et al., 2007, 2009a; Streule et al., 2010], exhumation in Eastern Nepal in the brittle field is determined here to have occurred at 1.0 mm/a from the late Miocene; this exhumation rate is thought to represent predominantly erosion driven exhumation after cessation of movement on the STDS in the Everest area at 11–13 Ma [Cottle et al., 2007, 2011; Jessup et al., 2008; Leloup et al., 2010]. However, given that the GHS is linked as a coherent unit from the surface to the midcrust the brittle exhumation rates can be related to the exhumation rates deduced from ductile processes occurring in the mid crust. An exhumation rate of 1.0 ± 0.2 mm/a is comparable (within error) to the 1.8 ± 0.5 mm/a exhumation rate derived from metamorphic pseudosection modeling for the time interval between 18.7 and 15.6 Ma [Streule et al., 2010] while the MCT and STD were still active and accommodating arc-perpendicular extrusion of the GHS footwall.
 Models of channel flow for the Himalaya require a focused surface denudation front to facilitate exhumation at calculated values of up to 10 mm/a [Beaumont et al., 2001, 2004], almost ten times the rates determined here, although Beaumont et al.  acknowledge that this is probably an overestimation by the models. Our results show that since ∼9 Ma exhumation rates were between 1 and 2 mm/a, well below that required by the models to accommodate pressure-driven, ductile channel flow in the GHS. The modest rise in the rate of exhumation since or before the Pliocene is also still not sufficient to satisfy the simulations of the channel flow models. Does this mean that channel flow model values for erosion are unrealistic and therefore call into question some of the model parameters? As yet there is no evidence for orogen-wide rates of erosion ca. 10 mm/a. Only the two syntaxial regions; Nanga Parbat in the western Himalaya [Crowley et al., 2009], and the Namche Barwa massif in the east, exhibit erosion rates approaching ca. 10 mm/a. Recent work on Namche Barwa has produced long-term exhumation rates of up to 3 mm/yr based on petrological data of midcrustal rocks and decompressional melts [Booth et al., 2009]. For the same region zircon FT, Ar/Ar biotite, and zircon (U-Th)/He data suggest faster modern denudation rates of 5–17 mm/yr and longer-term rates of 7–9 mm/yr for the past 2 Myr [Stewart et al., 2008; Enkelmann et al., 2011]. These high rates exist in part because of the extreme relief and precipitation in these areas and the local nature of the drainage system. Local climate and orographic forcing is clearly important. Is it possible the past exhumation rates could have been higher due to a wetter climate?
 Reconstructions of the Asian monsoonal system show a more intense monsoon existed in the middle Miocene and became drier in the late Miocene before returning to wetter conditions in the Pliocene [Clift et al., 2008]. It is not clear if the middle Miocene monsoon was more intense than the present but detrital FT studies of the Siwalik sedimentary rocks in Nepal that record exhumation of the Himalayan source rocks extending to the study area, give average exhumation rates of ∼1.4–1.8 mm/a since ∼15 Ma [van der Beek et al., 2006; Bernet et al., 2006]. Similar studies on slightly older rocks from two ODP sites in the distal Bengal Fan suggest even lower exhumation rates [Corrigan and Crowley, 1990]. As a consequence the ca. 10 mm/a rates suggested by the channel flow models of Beaumont et al. [2001, 2004] are unrealistic.
 The higher rate of exhumation seen in the Pliocene (1.7 ± 0.3 mm/a) is comparable to Pliocene rates determined in other parts of the Himalaya; rates of 1.1 to 1.4 mm/a were determined in the Sutlej valley for the period 1–5 Ma [Huntington et al., 2006; Thiede et al., 2004, 2009] which also correlates to rates of 1.5 mm/a determined in the Annapurna area [Blythe et al., 2007] in western Nepal. These observed changes in exhumation assumes no systematic change in groundwater fluid flow occurred at this time; such a definite and systematic change in this aspect across the entire Himalaya seems unlikely. Pliocene increases in exhumation, erosion and sedimentation have been attributed to the onset of Northern Hemisphere glaciation [Peizhen et al., 2001], which occurred at 3.2 to 2.4 Ma [Raymo, 1994]. Given that even small areas of ice cover can produce a lot of sediment as well as highest erosion rates [Koppes and Montgomery, 2009] the Pliocene rise in erosion rates due to the expansion and shrinking of ice seems to be the important influence on exhumation. Local tectonics might be invoked as a cause but current evidence points to this only occurring on small local scales, for example as seen in the Garhwal Himalaya [Patel and Carter, 2009] and not on the scale of a whole orogen, as would be the result of a glaciation related cause.
 Although the prograde metamorphic history and ages of magmatism in the Himalaya are reasonably well constrained, the later stages of exhumation and cooling were relatively unknown. Limited 40Ar/39Ar data from Everest suggested that a period of rapid cooling between 17 and 16 Ma immediately followed the main phase of leucogranite crystallization as the GHS midcrustal channel was exhumed toward the Earth's surface [Searle et al., 2003, 2006]. Additionally apatite and zircon fission track ages and 40Ar/39Ar data from one sample of the Yellow Band rocks, juxtaposed between the Lhotse and Qomolangma detachments near the summit of Everest was determined by Sakai et al.  and similarly showed increasingly rapid cooling of this unit from 24 to 14 Ma. Our results show that since ∼9 Ma brittle field exhumation rates for the GHS were between 1 and 2 mm/a.
 When the results from this area are put into context of the tectonics of the area, the kinematics of the region can be summarized as follows: during the early to middle Miocene exhumation was driven by high temperature ductile processes in the crust extruding the GHS as a single coherent unit via coeval movement on the MCT and STD. Exhumation was driven by tectonic processes, and accelerated due to Miocene melting in the mid crust during syncompressional arc-perpendicular extrusion of the GHS. Upon cessation of movement on the STD at 11–13 Ma in this area, syncompressional extension was transferred to an arc-parallel direction and normal faulting in the Ama Drime area became active. As a result a progression of tectonic effects of collision is demonstrated starting with the GHS and transferring to Ama Drime where ‘core complex’ type deformation is occurring and uplift, exhumation and metamorphism is all younger. In the Everest-Makalu region this progression was coincident with a reduction in exhumation rate and the shift from India-Asia convergence being accommodated by exhumation and crustal thickening concentrated within the GHS to the convergence being accommodated by wider-scale thickening and erosion.
 This work was carried out using NERC PhD grant NER/S/A/2005/13352 to M.J.S. and NERC grant NER/K/S/2000/00951 to M.P.S. We thank Tashi Sherpa for assistance in the field and Shiva Dhakal for logistical help in Nepal. Ray Donelick is thanked for permission to use his data. M.P.S. and M.J.S. are grateful for the RGS Peter Fleming award which provided financial assistance for fieldwork in the Barun Valley. We thank M. Jolivet and an anonymous reviewer whose input lead to a greatly improved version of this manuscript.