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JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 113, B08315, doi:10.1029/2007JB005424, 2008

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Seismic structure of the crust and the upper mantle beneath the Himalayas: Evidence for eclogitization of lower crustal rocks in the Indian Plate G. Monsalve,1 A. Sheehan,1 C. Rowe,2 and S. Rajaure3 Received 5 October 2007; revised 14 March 2008; accepted 22 April 2008; published 22 August 2008.

[1] Variations in the seismic velocity structure of the Himalayan collision zone include

significant differences between its north and south portions, with transitions in physical properties across the Greater Himalaya. We combined P- and S-wave traveltimes from a temporary broadband seismic network in eastern Nepal and southern Tibet with arrival times at the permanent station network of the Department of Mines and Geology of Nepal to determine the seismic velocity structure across the Himalaya, using local earthquake tomography and traveltimes of regional earthquakes. The P-to-S velocity ratio (Vp/Vs) structure marks the difference between the Indian Plate and the overlying materials, with the Vp/Vs ratios being high for the former and low for the latter. We also found a significant increase in the uppermost mantle seismic velocities from south to north, reaching P-wave velocities (Vp) over 8.4 km/s north of the Greater Himalaya. These high Vp values do not seem to be the result of biases due to anisotropy in the upper mantle beneath the Greater and Tethyan Himalayas. Instead, we suggest that rocks in the lower crust of the underthrusting Indian Plate undergo metamorphism to eclogite as they plunge to greater depth beneath the mountain range, explaining the high seismic velocities. Citation: Monsalve, G., A. Sheehan, C. Rowe, and S. Rajaure (2008), Seismic structure of the crust and the upper mantle beneath the Himalayas: Evidence for eclogitization of lower crustal rocks in the Indian Plate, J. Geophys. Res., 113, B08315, doi:10.1029/2007JB005424.

1. Introduction [2] The Himalaya chain is one of the products of the continental plate collision between the Indian Plate and Eurasia, and is composed of detached slices of the crust of the Indian Plate [Molnar, 1984] that underthrusts the mountain chain and the southern Tibetan Plateau [Zhao et al., 1993]. As a result, the crust has been thickened across the collision zone, with differences in thickness greater than 30 km between southern Nepal and southernmost Tibet [Schulte-Pelkum et al., 2005]. The structure of the lower crust and the upper mantle is critical in understanding processes associated with mountain building. For instance, the pressure increase resulting from thickened continental crust can trigger metamorphic reactions at lower crustal levels; in particular, evidence for transformation to eclogite in the Indian Plate gives support to theories about mass transfer from the lithosphere to the underlying mantle. This phenomenon might also influence the topography of moun1

Department of Geological Sciences and Cooperative Institute for Research in Environmental Sciences, University of Colorado at Boulder, Boulder, Colorado, USA. 2 Los Alamos National Laboratory, Los Alamos, New Mexico, USA. 3 Department of Mines and Geology, National Seismological Centre, Kathmandu, Nepal. Copyright 2008 by the American Geophysical Union. 0148-0227/08/2007JB005424$09.00

tain ranges and plateaus [e.g., Henry et al., 1997; Sapin and Hirn, 1997; Bousquet et al., 1997; Boyd et al., 2004; Garzione et al., 2006; Kay and Kay, 1993; Beck and Zandt, 2002]. [3] In this study, in addition to producing a seismic image of the crust and upper mantle across the Himalaya, we emphasize the seismic velocities in the lowermost crust and uppermost mantle. We aim to determine the fundamental differences in the seismic structure beneath Nepal and the southern Tibetan Plateau. Using seismic data from local and regional earthquakes recorded at a temporary network and nearby permanent stations, we obtain a two-dimensional south-to-north model across the Himalayan Collision Zone, combining data from receivers on the plains of southern Nepal, the Himalayas and the Tibetan Plateau, south of the Indus-Tsangpo Suture Zone.

2. Tectonic Setting [4] The collision between the Indian Plate and Eurasia began between 45 and 55 Myr ago [Garzanti et al., 1996; Rowley, 1996; Zhu et al., 2005] and is ongoing, with a convergence rate of about 36 mm/yr [Chen et al., 2000; Holt et al., 2000; Wang et al., 2001]. The Indian Plate slides as a coherent unit beneath the Himalayan range [e.g., Molnar, 1984], which comprises pieces of Indian crust detached from the rest of the Indian lithosphere [e.g.,

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MONSALVE ET AL.: SEISMIC STRUCTURE OF THE HIMALAYA

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Figure 1. (left) Topographic map of part of Tibet and the Himalayan region of Central Asia. The rectangle marks our area of study. (right) Simplified geologic map of the study region, showing the main tectonic features and the principal geologic units (units taken from Hodges [2000]). Lyon-Caen and Molnar, 1983; Molnar, 1984]. The Moho dips a few degrees to the north beneath the Sub and Lesser Himalaya and steepens to about 15° beneath the greater Himalaya [Hauck et al., 1998; Schulte-Pelkum et al., 2005]. Convergence between the Indian Plate and southern Tibet has caused slip on the major thrust faults of the Himalaya with a southward progression of thrusting [Gansser, 1964; Alle´gre et al., 1984; DeCelles et al., 2001; Robinson et al., 2003]. These thrust faults have been interpreted as different splays from a common decollement [Schelling and Arita, 1991]. [5] A set of tectonostratigraphic units across the Himalayan collision zone, separated by major fault systems, has been proposed by Hodges [2000] and we use his nomenclature in this study (Figure 1). The Main Frontal Thrust in south Nepal represents the front of the Himalayan orogen and constitutes the northern limit of the sediments of the Indo-Gangetic plains (Figure 1). The sedimentary rocks of the Sub-Himalayan Zone are located between the Main Frontal Thrust and the Main Boundary Thrust fault systems (Figure 1). North of the Main Boundary Thrust is the Lesser Himalaya, composed of metasedimentary units limited to the north by the Main Central Thrust fault system, north of which is the Greater Himalayan Zone. The South Tibetan Fault Zone separates the metamorphic rocks of the Greater Himalaya from the sediments of the Tethyan Himalaya. The latter has its northern limit at the Indus-Tsanpo Suture Zone, which is considered to be the mark of the Indian PlateEurasia collision [Gansser, 1964].

3. Experiment Description and Data [6] The 27-station HIMNT seismic network [de la Torre and Sheehan, 2005] in Eastern Nepal and Southern Tibet operated between October 2001 and March 2003 (Figure 2). This network was composed of three-component, broad-

band sensors, with sampling rates of 40 and 50 samples per second (sps). We measured first P and S arrival times for over 1600 local earthquakes, with uncertainties between 0.05 and 1 second for P arrivals and between 0.15 and 2 seconds for S arrivals. Additional arrival times of P and S waves to permanent stations in Nepal were obtained from a catalog of the Department of Mines and Geology (DMG) of Nepal [Pandey et al., 1999]. We used data from the eight easternmost stations of the Nepali network (Figure 2), which are short-period vertical-component, with a sampling rate of 50 sps.

4. Local Earthquake Tomography 4.1. Method [7] For the tomography study, we used a set of 542 local earthquakes within the study area with magnitudes between 1 and 5.5, which were previously located using a 1-D velocity model [Monsalve et al., 2006]. The selected events all have seven or more associated handpicked arrivals at stations within 200 km from the epicenter. We inverted 5767 P-wave arrivals and 4801 S-wave arrivals. Earthquake depths range from near surface to about 85 km below sea level (BSL). Figure 3 shows the epicentral location for these events. The main features of the seismicity are described by Monsalve et al. [2006]. [8] We used the traveltime inversion program SIMUL2000, based on an algorithm developed by Thurber [1983, 1993] and Eberhart-Phillips [1986] to invert P and S-P arrival times of hundreds of local earthquakes, solving for earthquake locations, P-wave velocity and P-to-S wave velocity ratio (Vp/Vs). The goal of the inversion is to minimize the arrival time residuals by finding models consistent with the data. The minimization is achieved through a least squares algorithm, iterating between the velocity parameters and the earthquake relocations, where we assign a damping (smoothing) value to the velocity model parameters

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Figure 2. Seismic stations in the study area. Inverted triangles indicate stations from the Himalaya Nepal-Tibet Seismic Experiment (HIMNT). Normally oriented triangles denote vertical component shortperiod stations that belong to the permanent seismic network of the Department of Mines and Geology of Nepal. [Thurber, 1983]. Arrival time weighting is defined taking into account the observation quality defined by the seismic analyst, the traveltime residual and the source-receiver distance. The traveltimes are determined using a ray-tracing method [Thurber, 1983] that searches for the approximate raypath with the minimum traveltime between source and receiver. [9 ] For the model parameterization we adopted an approach similar to that used by Husen et al. [2000], Eberhart-Phillips [1990], and Eberhart-Phillips and Michael [1998], starting with coarse grids of nodes and progressively making them finer, using the results of the previous coarser scale inversion as the initial velocity model for the next finer grid. Since we are mainly interested in the variations in the structure along a direction approximately normal to the Himalayan arc, we sought a 2-D south-tonorth velocity model. The starting velocity models (Tables 1 and 2) and hypocenter locations were taken from Monsalve et al. [2006]. Since the velocity structure obtained from the inversion is dependent on the input structure [Kissling et al., 1994], we ran the SIMUL2000 inversion routine using three starting velocity models: (1) separate models from Tables 1 and 2 with a transition at latitude 28°N, (2) an average of the two models for the whole region, and (3) a model with velocities as in Tables 1 and 2 but with the interfaces found by Schulte-Pelkum et al. [2005]. We also inverted the data

using starting velocity models resulting from random perturbations of those three cases (at least 12 for each case), allowing the changes to be up to 6% for Vp and 3% for Vp/Vs. The appropriate damping for each parameterization was found following Eberhart-Phillips [1986], seeking a value that allows a significant data variance reduction without compromising the smoothness of the solution. A total of four node configurations was tested (Figure 4). The depths at which we calculated velocity values ranged from 0 to 70 km BSL. Table 3 shows the results of the inversions with the smallest misfit for each type of grid, the average horizontal and vertical node spacings for each of the four geometries as well as the RMS of the time residuals after inversion. 4.2. Model Selection and Solution Quality [10] When performing seismic traveltime inversions, we seek a model that produces a small data misfit, makes physical sense, and matches the previously known properties of the model space. We also want a model parameterization that allows us to visualize the seismic velocity anomalies at an adequate scale, without excessive aliasing or oversampling. Information from the resolution matrix can help us choose the most appropriate parameterization. [11] The whole model resolution matrix was obtained after each inversion, and its analysis provides information

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Figure 3. Location of earthquake epicenters used in the traveltime inversion. Event locations were taken from Monsalve et al. [2006]. on trade-offs or linkages between any pair of model parameters [Soldati and Boschi, 2005]. We are not only interested in the magnitude of the diagonal element of the resolution matrix, but also in the position of the largest element of each row and the magnitude of the off-diagonal elements, as explained by Toomey and Foulger [1989]. Therefore we used the spread function, as defined by Menke [1989], in order to assess the quality of the solution. Table 3 shows the spread function divided by the number of parameters for each one of the tested models. The parameterization that gives the best combination of RMS time residual and spread function is the one defined by Model 2 (Table 3 and Figure 4b). Finer grids did not produce a decrease in misfit, and show some significant small-scale oscillation that it is likely an artifact arising from overparameterization. We therefore use the results obtained from Model 2 (Figure 4b) for further analysis. [12] We use the diagonal element of the resolution matrix and the spread function as indicators of how independently

a parameter is estimated and to identify places of potential smearing. Figure 5 shows contours of both diagonal resolution and spread function for Model 2. In order to use both indicators, we only show and interpret those results at grid points for which diagonal resolution is greater than 0.2 and spread function is less than 0.8. Since we are allowing grid points with low diagonal resolution to be part of our result model, we also examined the resolution row for each of the grid points that we include in our analysis to determine whether the diagonal element was the largest. If this was not the case, we looked for the grid point most strongly influencing the velocity at the node under evaluation. For the case of Model 2 in Table 3 and Figure 4b, the largest element of the row corresponding to grid points with diagonal resolution greater than 0.2 and spread function less than 0.8, was always the diagonal element. [13] Seismic velocity anomalies of different spatial scales appear consistently in the smallest misfit results. Some may be artifacts of the inversion. To check for their robustness,

Table 1. Velocity Model for the Nepal Region (South of 28°N) From Monsalve et al. [2006]a

Table 2. Velocity Model for South Tibet (North of 28°N) From Monsalve et al. [2006]a

Depth Range, km

P-wave Velocity, km/s

S-wave Velocity, km/s

Depth Range, km

P-wave Velocity, km/s

S-wave Velocity, km/s

Surface-3 3 – 23 23 – 55 >55

5.5 5.7 6.3 8.0

3.2 3.2 3.7 4.5

Surface-3 3 – 40 40 – 70 >70

5.8 5.8 6.9 8.6

3.3 3.5 4.0 4.9

a

a

Depths are relative to sea level.

Depths are relative to sea level.

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Figure 4. Different parameterizations of the two-dimensional model space that we used for the tomographic inversion. (a) Model 1. (b) Model 2. (c) Model 3. (d) Model 4. we reran the inversion routine, fixing seismic velocity at the nodes corresponding to such anomalies. Our procedure was to force the values at those grid points to be the same as those at the surroundings and Vp and/or Vp/Vs were not allowed to change at those nodes during the inversion; in other words, we forced the anomaly to disappear. We then evaluated how those fixed values affected our solution by looking at changes in seismic velocities at neighboring points and the misfit variations. For instance, in some cases, after running an inversion with fixed values at nodes where an anomaly was seen, the misfit did not vary significantly (the RMS was within 0.01 seconds from the original one) but the velocity anomaly appeared at neighboring points. In such cases, we concluded that the anomaly was robust and required by the data. [14] Additionally, to further estimate the reliability of the results and to better constrain the appropriate damping for the tomographic inversion, we used synthetic model reconstruction tests, using a methodology similar to that described by Laigle et al. [2000]. To test for specific Table 3. Different Node Spacings Used for Inversion for 2-D Velocity Structurea

Model

Horizontal Node Spacing, km

Number of Nodes in the Vertical Direction

RMS, s

Spread(R)/m

1 2 3 4

80 50 30 30

7 7 7 9

0.341 0.335 0.335 0.334

0.508 0.507 0.522 0.556

a Horizontal node spacing refers to separation between nodes at the center of the network. The number of parameters of the model is denoted by m.

Figure 5. Contours of diagonal resolution and spread function for the two-dimensional south-to-north tomographic model for P-wave velocities. (a) Diagonal resolution. (b) Spread function.

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anomalies in our results, we made synthetic arrival time data, using simple models where we place structures with similar geometry and velocity values to those of the tested anomaly, and inverted those synthetic data to evaluate the recovery of the synthetic structure. This was done for a range of damping values. In the next section, where specific features are discussed, we give some examples of such tests. 4.3. Tomography Results [15] Models of spatial distribution of Vp and Vp/Vs were obtained for the study region, at depths from zero to 70 km BSL, and we relocated the earthquakes using the new velocity structures. Since the number of S arrival times is similar to that of the P arrival times, our solutions for Vp and Vp/Vs have almost the same spatial coverage. Results from the inversion for a two-dimensional velocity structure are shown in Figure 6, where we plot values of Vp and Vp/Vs for grid points with an associated diagonal resolution greater than 0.2 and spread function less than 0.8. All the features present in Figure 6 appeared at every inversion we performed and passed the robustness tests. We account for only South-North variations in the velocity structure. [16] Since we assumed linear variations of seismic velocities (or seismic velocity ratios) from one grid point to another, abrupt changes due to interfaces in between nodes cannot be resolved with this technique. However, the Vp structure (Figure 6b) shows that upper mantle-like velocities (around 8 km/s) are reached at shallower depths beneath the Lesser Himalaya than beneath the Tethyan Himalaya, consistent with a crust-mantle boundary that is deeper beneath Southern Tibet than beneath the Lesser Himalaya, with a steepening underneath the Greater Himalaya. Previous work shows that the Moho beneath the Lesser Himalaya is between 40 and 50 km BSL, it steepens beneath the High Himalaya [Schulte-Pelkum et al., 2005; Lyon-Caen and Molnar, 1983] and flattens out again beneath the Tethyan Himalaya, reaching depths between 60 and 75 km BSL [Schulte-Pelkum et al., 2005; Zhao et al., 1993; Kind et al., 2002]. [17] A mid-crustal interface beneath the Lesser and the Greater Himalayas at around 20 km BSL dips gently to the north [Schulte-Pelkum et al., 2005], which may correlate with a seismic reflector found between 30 and 40 km depth beneath the Tethyan Himalaya [Zhao et al., 1993]. This discontinuity is thought to correspond to the top of the underthrusting Indian Plate. The Vp structure (Figure 6b) shows that velocities around 6.3 km/s are reached at about 20 km beneath the Nepal Himalayas and about 30 or 40 km beneath Tibet, consistent with a north-dipping mid-crustal interface. [18] The Vp/Vs structure (Figure 6c) shows a region of relatively low ratio north of latitude 27.5°N at depths from the surface to 40 km BSL. Values of Vp/Vs greater than 1.73 dominate almost everywhere else. Wadati diagrams of earthquakes shallower than 40 km BSL, and recorded at stations in the Greater and Tethyan Himalayas, also suggest low values of Vp/Vs, with a mean of 1.66 and a standard deviation of 0.06. This wedge of low Vp/Vs coincides with the thrust sheets overlying the Indian Plate (Figure 6c). The southern end of the low ratios is located in the region between the Lesser Himalaya and the High Himalaya, where a concentration of upper crustal earthquakes is seen

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(Figure 6c). We tested the ability of the tomographic inversion to image this kind of feature using a synthetic test (Figures 7a and 7b). The input model consists of a homogeneous space with Vp/Vs of 1.73 and a wedge with Vp/Vs values of 1.66 (Figure 7a). Inversion of the generated data indicates that the input structure is reasonably well recovered (Figure 7b). [19] There is one localized low in the Vp/Vs structure in the mid-crust beneath the Lesser and Sub-Himalayas, near latitude 27.2°N at a depth of 25 km BSL (Figure 6c), with a minimum value of around 1.65. A synthetic test for this anomaly, similar to the one described in the previous paragraph, shows a good recovery of the input structure (Figures 7c and 7d). It is striking that this anomaly in the underthrusting Indian crust is localized and isolated. This region of low Vp/Vs lies between the locus of earthquakes beneath the Sub-Himalaya in the area of the August 20, 1988, Magnitude 6.5 Udayapur earthquake [Pandey et al., 1999], and the region of high seismicity beneath the topographic front of the Himalaya. [20] The presence of upper mantle earthquakes beneath the Greater and Tethyan Himalayas [Monsalve et al., 2006] allowed us to infer the seismic velocities at subcrustal levels. The P-wave velocities we find at 70 km BSL beneath the Greater and Tethyan Himalayas are anomalously high, reaching values around 8.7 km/s. There is also a very pronounced vertical velocity gradient in this region between 55 and 70 km BSL, where velocities go from values between 6.5 and 7 km/s at a depth of 55 km to values greater than 8.5 km/s at 70 km. To make sure that this contrast is not an artifact of the model parameterization of the ray geometry, we ran several synthetic tests with a decreased contrast in the input structure; in no case was the contrast obtained after inversion greater than that of the input structure. This result suggests that this high vertical gradient is in fact required by the arrival time data. [21] High Vp in the upper mantle beneath the Himalayas and South Tibet has been also obtained by Hirn et al. [1984] and Hirn and Sapin [1984], who proposed Vp values from 8.5 to 8.7 km/s, Menke and Jacob [1976], with velocities around 8.5 km/s, and Beghoul et al. [1993], who found seismic velocities of around 8.4 km/s. Tomographic upper mantle velocity models by Reiter et al. [2005] give Vp values between 8.3 and 8.4 km/s beneath the Tethyan Himalaya in the HIMNT study area. Holt and Wallace [1990], McNamara et al. [1997], Liang et al. [2004], Liang and Song [2006], and Phillips et al. [2007] report Pn velocities between 8.2 and 8.3 km/s for the Tethyan Himalaya south of the Indus-Tsangpo Suture Zone in the area of the HIMNT experiment. Additional inversions are required in order to evaluate the robustness of the anomalously high seismic velocities we obtained, so we ran some inversions with fixed Vp values at those nodes at 70 km BSL where our model is best resolved. We tested values between 7.9 km/s and 9.0 km/s for two different model parameterizations (Models 1 and 2 from Table 3 and Figures 4a and 4b). Station-earthquake pairs used for these inversions are shown in Figure 8. In both cases, the minimum misfit is attained when the P-wave velocity is 8.7 km/s (Figures 9a and 9b, solid lines). [22] A map of earthquake-station pairs (Figure 8) suggests that most raypaths from earthquakes at depths greater

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Figure 6. Two-dimensional north-south structure of Vp and Vp/Vs. Earthquakes are projected onto an N-S cross-section. (a) North-south topographic cross-section at 86.5°E, red triangles denote projected station locations. (b) North-south Vp structure. (c) North-south Vp/Vs structure. than 70 km BSL have a predominant east-west component, so it is possible that those high values of Vp are the result of azimuthal anisotropy, with a nearly east-west fast direction, analogous to a case presented by Scherwath et al. [2002] for

the South Island of New Zealand, where they found variations from 7.7 to 8.6 km/s in Pn velocities at nearly perpendicular directions. To test for such a possibility in our study area, we eliminated the rays going from those deep

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Figure 7. Synthetic model reconstruction tests. (a) Synthetic structure for testing the wedge of low Vp/Vs beneath the Greater Himalaya and Southern Tibet. (b) Recovered Vp/Vs structure after inverting data generated using velocities in Figure 7a. (c) Synthetic structure for testing the localized low Vp/Vs beneath the Lesser Himalaya. (d) Recovered Vp/Vs structure after inverting data generated using velocities in Figure 7c. earthquakes to stations west of longitude 86°E, and performed the same robustness test described above. The resulting curves are flat (Figures 9a and 9b, dashed lines), and a large range of seismic velocities would fit the data almost equally well (8.2 to 9 km/s). [23] Given the flat curves displayed in Figure 9, it is hard to place narrow bounds for the P-wave velocity of the upper mantle beneath the High and the Tethyan Himalayas because the differences in misfit are not significant: If we compare the RMS of the time residuals when Vp is 8.1 km/s and when Vp reaches its minimum, the differences are always below 0.015 seconds. Therefore further tests are needed to more precisely constrain the magnitude of the upper mantle seismic velocities in this region, as well as to evaluate the possibility of seismic anisotropy in the uppermost mantle.

5. Pn Studies [24] An alternative way to constrain the structure of the seismic velocity in the upper mantle is using arrival times

from regional earthquakes for which the first arrival is the Pn phase. This can help us estimate the P-wave velocities in the uppermost mantle and their directional variations. Pn arrivals and their corresponding uncertainties for regional earthquakes were visually determined. For spatially close earthquakes recorded at the same station, we used waveform similarity to refine the arrival time pick. This was done only for pairs of seismograms with average coherence (at frequencies below the Nyquist) greater than 0.75. We used two methods for the estimation of Pn velocities beneath the network: the first method uses differences in arrival times at station pairs, and the second consists of determining the slope of the distribution of Pn arrival time vs. distance from the source. To evaluate south-to-north changes in the Pn velocity structure, we divided the study region into two subregions, located north and south of latitude 28°N, which we will further refer to as South Tibet and Nepal respectively. The HIMNT seismic stations in Nepal operated for longer times than stations in South Tibet, so we have more data south of 28°N than north of it. Figure 10 shows the

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Figure 8. Event-station pairs for earthquakes deeper than 70 km BSL used for tomographic inversion. earthquakes used for both areas with different symbols for each of the methods. 5.1. The Two Station Method [25] We use the method of Beghoul and Barazangi [1989] (a simplified version of a method further developed by Phillips et al. [2005]) to estimate Pn velocities between pairs of stations that are on the same azimuth with a given seismic event. In order to estimate the Pn velocity, when a station pair and an earthquake epicenter are approximately colinear, the method uses the interstation distance (or the difference in epicentral distance from each of the stations to the earthquake) and divides it by the difference in arrival time at the two stations. Given the thick crust in our study

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area, we corrected the interstation distance for earth sphericity, so that the distance we use represents a better approximation of the ray trajectory along the Moho. Some of the advantages of this method are that it minimizes the effects of earthquake mislocation, and the obtained velocities are only representative of the upper mantle in the interstation path. [26] For the selection of the station pair-earthquake sets, we allowed events within 3 degrees off the azimuth of the station pair. Earthquakes with magnitudes greater than 4, at epicentral distances between 3.5 and 16 degrees from the center of the network were selected. Unfortunately, the event distribution with back-azimuth is not even, and large azimuthal gaps are present (Figure 10). [27] Small interstation distance leads to a large scatter in the results, so after an inspection of the variations of velocities and their uncertainties with interstation distance, we decided to keep the results for interstation distances greater than 90 km for the case of Nepal and 50 km for the case of South Tibet. Seismic velocities with uncertainties greater than 0.25 km/s were not included in our study. With Pn velocities calculated using this method, differences in crustal seismic velocity and thickness between two station locations can cause deviations from the real Pn velocity. For instance, for the case of stations located on equally fast portions of the crust, apparent velocities would be greater than the real ones for the case of rays going updip and vice versa. We applied a correction to prevent this effect, taking the crustal seismic velocities and thicknesses from this study and Schulte-Pelkum et al. [2005] respectively. [28] Figure 11a shows Pn velocities vs. back-azimuth for stations in Nepal, with their correspondent uncertainties resulting from propagating the standard deviations of the arrival time differences through the seismic velocity calculation. A total of 54 station pair-earthquake sets satisfied the selection criteria; however, events locate predominantly to the north-west and south-east of the station network, resulting in considerable azimuthal gaps. No dependence of Pn velocities with back-azimuth is suggested. The average of seismic velocities is 8.13 ± 0.17 km/s beneath Nepal.

Figure 9. RMS residual vs. P-wave velocity in the upper mantle obtained from tomographic inversion for different model parameterizations and raypaths. (a) Model 1, all available data (solid line) and after removing paths to the westernmost stations (dashed line). (b) Model 2, all available data (solid line) and after removing paths to the westernmost stations (dashed line). 9 of 16

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Figure 10. Epicenters of the events used for Pn velocity determination. Black box indicates the study area. Diamonds indicate events used for the two-station method only; squares indicate events used for determination of Pn velocities using slopes of time vs. distance distributions only; circles indicate events used for both methods. (a) Earthquakes used for stations in Nepal. (b) Earthquakes used for stations in South Tibet.

[29] We are especially interested in the results for South Tibet because the anomalously high P-wave velocities in the upper mantle are predominantly obtained in this region. A total of 13 station pair-earthquake sets satisfied our criteria. Figure 11b shows the obtained Pn velocities vs. backazimuth, showing no obvious dependence of one upon another. Velocities average 8.44 km/s with a standard deviation of 0.14 km/s. Even though these values are somewhat lower than the best fit velocities we get from the tomography, they still indicate the presence of materials with remarkably high P-wave velocities in the upper mantle, with no clear evidence of anisotropy. 5.2. Determining Pn Velocity From Time Versus Distance Distributions [30] In addition to the two-station method, we use arrival times of earthquakes recorded at five or more stations to

estimate the slope from arrival time vs. epicentral distance distributions. A weighted linear least squares method allows us to find the Pn velocity that best fits the collected arrival times and their uncertainties. Our goal is to estimate the Pn velocity for both Nepal and Tibet subregions. In this case, a correction for earth sphericity is applied when calculating differences of epicentral distance in between stations, so that the calculated velocity is representative of the uppermost mantle beneath the area of our interest. This method has the advantage that each velocity value we obtain is the result of a linear regression with several data points. Earthquakes at epicentral distances between 3.5 and 16 degrees from a location near the center of our network are used in order to apply this method. A total of 65 events recorded at stations in Nepal and 11 earthquakes at stations in South Tibet satisfies the criteria (Figure 10b). No correction for dipping Moho is applied in this case.

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Figure 11. Pn velocity vs. back-azimuth for Nepal and South Tibet. Vertical bars represent uncertainties within one standard deviation. (a) Pn velocities beneath Nepal calculated using the two-station method. (b) Pn velocities beneath South Tibet calculated using the two-station method. (c) Pn velocities beneath Nepal calculated using slopes of time vs. distance plots. (d) Pn velocities beneath South Tibet calculated using slopes of time vs. distance plots.

[31] Figures 11c and 11d illustrate the variation of the estimations of Pn velocity with back-azimuth, for earthquakes recorded in Nepal and in South Tibet respectively. Figure 11c suggests a dependence of Pn apparent velocity with back-azimuth, where the highest Pn velocity values correspond to seismic rays coming from earthquakes located north of the study area (roughly, back-azimuths greater than 300° and less than 90°); Pn apparent velocities become lower as the back-azimuths get closer to south-to-north directions. These variations can be explained by the nearly north-dipping Moho in the Nepal portion of the network. Unfortunately, there are no seismic rays coming from the south, and the azimuthal gap is nearly 130°. We determined a theoretical curve showing the variations of Pn velocity with back-azimuth due to a dipping Moho (Figure 11c). To find the curve that best fits the data points, we performed a grid search over dip values (between 3° and 8° with increments of 0.2°) and Pn velocities (between 8 and 8.5 km/s with increments of 0.02 km/s). A combination of a dip of 4.2° and a Pn velocity of 8.24 km/s gives the minimum root mean square of the weighted residuals (RMS = 0.22 s). Additionally, if we only consider those back-azimuths that approximately correspond to raypaths along the Himalayan arc, so that ray trajectories in the uppermost mantle are nearly horizontal, we obtain an average Pn velocity of 8.13 km/s with a standard deviation of 0.14 km/s. [32] Variations in crustal thickness beneath the Tibetan part of our network (in the Tethyan Himalaya) are not as

large as they are beneath the Himalayas [Schulte-Pelkum et al., 2005]. In fact, Figure 11d does not suggest any dependence of Pn velocity with back-azimuth. An average of the results plotted in Figure 11d gives a value of 8.65 ± 0.24 km/s. Even though these calculations are not enough to determine the Pn velocities beneath the Tibetan part of our network, the results clearly suggest that there is an increase in the upper mantle P-wave velocities from the Himalayas of Nepal to the southern Tibetan Plateau.

6. Discussion [33] Despite the uneven spatial distribution of earthquakes, we are able to put some constraints in the structure of the crust and upper mantle beneath the Himalayan Collision zone. Figure 6b suggests that the Moho deepens from the Himalayas of Nepal to the southern Tibetan Plateau. At depths of 55 km BSL materials under Nepal show mantle-like Vp values (about 8 km/s) whereas in South Tibet we still find crustal velocities (about 6.5 km/s) at that depth. In the upper crust, we see a steeper vertical P-wave velocity gradient beneath Nepal than under South Tibet, which is consistent with the presence of Indian crust underthrusting the Himalayas and South Tibet and deepening from south-to-north. [34] A wedge of low Vp/Vs (with an average of around 1.65) marks the difference between the materials in the underthrusting Indian Plate and the thrust sheets above it,

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with such low values corresponding to the upper crust of the Greater Himalaya and the mid-upper crust of the Tethyan Himalaya, up to a depth of about 40 km BSL (Figure 6c). This is consistent with some results found in previous studies: Owens and Zandt [1997] report a decrease in the Poisson’s ratio from north to south in the crust of Tibet, reaching values of around 0.25 (approximately equivalent to a Vp/Vs of 1.73) in the southernmost Lhasa terrain. Kind et al. [2002] show variations of Vp/Vs across the Tibetan Plateau, with anomalously high values in the Lhasa terrane and an abrupt drop to values below 1.7 south of 29°N, coinciding with the location of our northernmost resolved region. These low values of Vp/Vs suggest that the crust beneath the Tethyan Himalaya is probably composed of highly felsic materials at least to depths of 40 km BSL, and that the presence of large amounts of partial melt in the midupper crust of the Tethyan Himalaya is unlikely. At deeper crustal levels, Vp/Vs is higher than in the upper crust (values greater than 1.76 at 55 km BSL). [35] The localized low Vp/Vs at a depth around 25 km and a latitude about 27.2°N (Figure 6c) could reflect a difference in composition with respect to the surrounding materials. Another possibility is that these observations of low Vp/Vs are due to the presence of aqueous pore fluids, as suggested from calculations by Nakajima et al. [2001] and Sato and Ito [2002]. This region of low Vp/Vs ratio roughly correlates with the location of the Main Himalayan Thrust, near where Pandey et al. [1995] hypothesize about the presence of a mid-crustal ramp; the associated seismicity clusters preferentially in its hanging wall [Monsalve et al., 2006]. The location of the low Vp/Vs values also correlates with a high conductivity zone found by Lemonnier et al. [1999], which is slightly to the south of the cluster of upper crustal earthquakes beneath the topographic front of the Himalaya and at depths between 20 and 30 km. The low Vp/Vs, the high conductivity and the earthquakes are consistent with the presence of fluids with high pore pressure in highly fractured materials. As Lemonnier et al. [1999] propose, the fluids can originate from dehydration reactions as the Indian crust underthrusts beneath the Himalayas. The presence of fluids may contribute to the seismic activity in the hanging wall by lowering the Coulomb effective stress and putting those materials at closer proximity to failure. [36] Perhaps the most interesting result of this study is the high seismic wave velocity that we obtained in the uppermost mantle for the northern part of our network. The local earthquake tomography gives P-wave velocity values of about 8.7 km/s at a depth of 70 km BSL (Figure 6b); however, small differences in misfit over a wide range of fixed Vp values at that depth (Figures 9a and 9b) make it difficult to support the robustness of such high values. The estimations of Pn velocities helped us to conclude that high Vp values in the uppermost mantle beneath the Tethyan Himalaya are plausible and that the increment in the Pn velocity magnitude from the Himalayas of Nepal to the southern Tibetan Plateau (Tethyan Himalaya) is greater than 0.2 km/s (Figure 11). The increase in seismic velocities of rocks due exclusively to depth differences of about 30 km is usually less than 0.1 km/s [Christensen and Mooney, 1995]; such differences in pressure therefore cannot account for the change in velocities that we see across the Himalayan arc. One possibility is that changes in the orientation of olivine

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cause anisotropy beneath South Tibet, and therefore, account for these variations; changes in upper mantle composition from the Himalayas of Nepal to the southern Tibetan Plateau could also produce this difference in seismic velocity. [37] Hirn et al. [1995] proposed the presence of anisotropy in the mantle beneath the Tethyan Himalaya using SKS phases and suggested a NNW fast polarization direction. Refraction and wide angle reflection data [Sapin and Hirn, 1997] indicated an upper mantle P-wave velocity of 8.7 km/s beneath the Tethyan Himalaya in the region of our experiment; the data they used were taken along an approximately east-west oriented line. Sapin and Hirn [1997] concluded that the unusually high upper mantle velocities are not due to azimuthal anisotropy because the orientation of the fast polarization direction does not coincide with the nearly east-west direction expected from the high velocities suggested by the explosion soundings. The delay times in this region reported by Hirn et al. [1995] were of the order of 0.25 seconds and obtained from a single earthquake; as stated by Sandvol et al. [1997], this kind of delay can also be explained by anisotropy in the crust, so that it does not prove that the upper mantle beneath this region is anisotropic. Sandvol et al. [1997] used ScS, SKS and SKKS phases for earthquakes recorded at stations in the Tethyan Himalayas and found no evidence of shear wave splitting. A recent study by Pei et al. [2007] suggests very low Pn anisotropy in the HIMNT study area, with velocity differences between the fast and slow axes of less than 0.1 km/s. [38] Despite the presence of azimuthal gaps greater than 100°, the distribution of Pn velocities measured by two methods beneath stations in South Tibet (Figures 11b and 11d) does not indicate any dependence of the velocities with back-azimuth. Seismic rays coming to South Tibet from the north and south-east of the network give high Pn velocities, with values of the order of 8.4 to 8.7 km/s, which weakens a possible argument in favor of anisotropy with a nearly eastwest fast direction. Thus we claim that azimuthal anisotropy in the upper mantle beneath the Tethyan Himalaya in the region of our network is unlikely and cannot be used to explain the high seismic velocities. [39] Another possibility is radial anisotropy as the cause of the high seismic velocities in the uppermost mantle beneath the Tethyan Himalaya, with elongated olivine crystals oriented in random directions in a nearly horizontal plane (associated with a fast direction) in the uppermost mantle. After inversion of surface wave dispersion data, Shapiro and Ritzwoller [2002] constructed a model of strength of the radial anisotropy at the top of the mantle, where they report low values in the Tethyan Himalaya for our study area (values between 0 and 3%). According to Ribe [1992], when the lattice preferred orientation induced by deformation causes anisotropy, the direction of the maximum compressional wave velocity coincides with the orientation of the maximum extension, and the minimum P-wave velocity aligns with the direction of maximum compression. In order to have radial anisotropy with a vertical slow axis and high velocities on a horizontal plane with no azimuthal anisotropy, the two principal axes of the finite strain ellipsoid in a horizontal plane should have a length greater than one, indicating extension in the two perpendicular directions. This situation is inconsistent with the tectonic setting of the study area; focal mechanisms

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determined by de la Torre et al. [2007] indicate that the P-axis is approximately horizontal with a north-south direction, and the T-axis has a nearly east-west trend, in agreement with the occurrence of maximum compression along an orientation perpendicular to the Himalayan arc and maximum extension parallel to it. Therefore we conclude that radial anisotropy is not the cause of the high Pn seismic velocity in the uppermost mantle. [40] An alternative explanation for these high velocities in the upper mantle is the presence of eclogite under the Moho beneath the Tethyan Himalaya, with lower crustal rocks transforming into higher density metamorphic rocks so that their new seismic velocities are upper mantle like, as proposed by Sapin and Hirn [1997]. Modeling by Henry et al. [1997] predicts the presence of an eclogitic body in the lowermost crust beneath the transition between the High Himalaya and the southern Tibetan Plateau (the Tethyan Himalaya). Anisotropy in eclogites is usually less than 2% [Christensen and Mooney, 1995]. An increase in pressure can drive lower crustal materials into the eclogite facies field, which is usually accompanied by the formation of clinopyroxene (omphacite) and the disappearance of plagioclase [Miyashiro, 1994], with an associated increment in density and seismic velocity. Christensen and Mooney [1995] report compressional wave velocities between 7.9 and 8.5 km/s for rocks in the eclogite facies at 20 km depth. Estimates of eclogite seismic velocities by Gubbins et al. [1994] give Vp values between 8.46 and 8.65 km/s at pressures between 0 and 3 GPa and temperatures between 0 and 750 °C. Calculations by Hacker et al. [2003] predict P-wave velocities of 8.4 and 8.5 km/s for eclogites at pressures between 2 and 3 GPa (roughly what we would expect in the upper mantle beneath South Tibet), and temperatures greater than 450 °C. Recently, Zhang and Green [2006] reported seismic velocities for eclogites between 8.25 and 8.75 km/s, with Vp anisotropies of less than 2%. These values of Vp are plausible according to our tomography results and Pn velocity estimations. The obtained Vp/Vs values at 70 km BSL from the tomographic inversion suggest Poisson’s ratios of around 0.25, consistent with unaltered eclogites [Gao et al., 2001]. For peridotites, higher Poisson’s ratios would be expected at upper mantle depths, typically between 0.28 and 0.29 [Wang et al., 2006]. In the case of the Himalayan collision, the eclogite could be the result of mineralogical transformations in the lower crust, after changes in depth as large as 30 km (roughly from 40 to 70 km) beneath the region of continental collision turned lower crustal rocks into upper mantle materials. Volume budgets estimated by LePichon et al. [1992] suggest that there is a mass loss in the Indian crust when it underthrusts the Himalayas and southern Tibet. They attribute this mass deficit to transferring of materials from the lower crust to the mantle due to eclogitization. [41] We see a high vertical gradient in the Vp distribution between 55 and 70 km BSL beneath South Tibet, where velocity magnitudes change by more than 2 km/s (Figure 6b). This gradient suggests that the Moho is somewhere in between. Our resolution is too poor to constrain the velocities in such a thin depth range in the lowermost crust of the Tethyan Himalaya; however, receiver functions beneath the Tethyan Himalaya, obtained using data from this experiment [Schulte-Pelkum et al., 2005], indicate a velocity

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contrast that they interpreted to be the Moho, but such a contrast is weaker than it is at the Moho conversion beneath Nepal. The high gradient and the weak velocity contrast suggest the presence of a high velocity layer in the lower crust of the Tethyan Himalaya (with a thickness of 15 km or less). Partial transformation to eclogite might be occurring in that layer and the seismically defined Moho can be now at the location of what was previously lower Indian crust. [42] The Himalayas are not in isostatic equilibrium [Molnar, 1988; Cattin et al., 2001]. Lyon-Caen and Molnar [1983] show that gravity anomalies require the Himalayan topography to be supported by the Indian Plate and that there is a need for a bending moment to be applied by an additional force at the northern end of the plate. The existence of lower crust-upper mantle materials in the eclogite facies in the Indian underthrusting plate could provide such an additional force, so that the Indian Plate flexes upward beneath the Lesser Himalaya and the IndoGangetic plains, and flexes downward under the Greater Himalaya [Molnar, 1988]. This force would be applied as a distributed load beneath the High and the Tethyan Himalaya. Cattin et al. [2001] fit the gravity data with a density contrast in the lower crust beneath South Tibet due to transformation to eclogite. It is possible that the density contrast in the lower crust is not as large as they suggest, since eclogitization in the lower crust is only partial, as we propose, and completely eclogitized rocks are located below the seismically defined Moho at a depth of about 70 km BSL [Schulte-Pelkum et al., 2005]. A topographic low is predicted for areas where transition into higher density rocks in the eclogite facies has occurred at the base of a thickened crust [Richardson and England, 1979]. Henry et al. [1997] interpreted the presence of a topographic depression of about 1000 m in the Tethyan Himalaya as evidence for the existence of an eclogitic body in the lower crust.

7. Conclusions [43] Figure 12 presents a summary of the 2-D south-tonorth seismic velocity model across the Himalayan continental collision. It reveals further evidence for the deepening of the Moho discontinuity and the south-to-north thickening of the crust. A wedge of low values of Vp/Vs marks the thrust sheet overlying the Indian Plate in the Greater and Tethyan Himalayas, corresponding with high S-wave velocity. This situation is inconsistent with the presence of large amounts of partial melt at mid-upper levels of the crust south of the Indus Tsangpo Suture Zone, and suggests that felsic rocks are predominant at such levels. Low values of Vp/Vs in the vicinity of the Main Himalayan Thrust may indicate the presence of fluids originated by metamorphic reactions due to Indian Plate underthrusting. [44] We found evidence of an increase in the uppermost mantle Vp of more than 0.2 km/s from the Himalayas of Nepal to South Tibet. Upper mantle P-wave velocities beneath the Tethyan Himalaya are about 8.4 km/s or greater, and they do not appear to be the result of anisotropy. Such south to north increase in upper mantle seismic velocities and the suggestion of a thin high velocity layer in the lowermost crust of the Tethyan Himalaya favors the hypothesis of metamorphic transformations to eclogite of rocks in the lower crust of the Indian Plate as they

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Figure 12. Summary of 2-D velocity structure interpretation. Discontinuities are taken from SchultePelkum et al. [2005], and the thicknesses represent their depth uncertainties. MHT: Main Himalayan Thrust. underthrust the High and the Tethyan Himalaya, so that materials in the uppermost mantle beneath South Tibet can be derived from rocks previously in the lower crust of the Indian Plate. [45] Depending on the density of the resulting high grade metamorphic rock, the presence of an eclogitic body in the mantle beneath the southern Tibetan Plateau can possibly generate a gravitational instability, favoring models of removal (or delamination) of the lithosphere, where the detached materials include lower crustal rocks, as proposed by Kay and Kay [1993]. However, even though we see an increase in seismic velocities that may correspond to eclogite in the lowermost crust and uppermost mantle, it is possible that the resulting eclogite densities are insufficient to produce gravitational instabilities at the location of our seismic experiment. Such instabilities may have been generated further north, with or without eclogitization, and have facilitated the convective removal of part of the lithosphere [e.g., England and Houseman, 1989; Molnar et al., 1993]. [46] Acknowledgments. We would like to thank the Department of Mines and Geology of Nepal for their help and assistance during the experiment execution and for generously sharing their seismic data with us. The authors also thank Peter Molnar and Craig Jones for the constant criticism and suggestions. This work was supported by grants from the National Science Foundation and the Department of Energy.

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