earthquake depth distributions in central asia, and their ...jacdev/pdf/sloan... · gji geodynamics...

Geophys. J. Int. (2011) 185, 1–29 doi: 10.1111/j.1365-246X.2010.04882.x

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Earthquake depth distributions in central Asia, and their relationswith lithosphere thickness, shortening and extension

R. A. Sloan, J. A. Jackson, D. McKenzie and K. PriestleyDepartment of Earth Sciences, University of Cambridge, Cambridge, UK. E-mail: [email protected]

Accepted 2010 November 2. Received 2010 October 27; in original form 2010 July 13

S U M M A R YThis paper examines the relationship between seismogenic thickness, lithosphere structure andrheology in central and northeastern Asia. We accurately determine earthquake depth distribu-tions which reveal important rheological variations in the lower crust. These variations exert afundamental control on the active tectonics and the morphological evolution of the continents.We consider 323 earthquakes across the Tibetan Plateau, the Tien Shan and their forelands aswell as the Baikal Rift, NE Siberia and the Laptev Sea and present the source parameters of94 of these here for the first time. These parameters have been determined through body waveinversion, the identification of depth phases or the modelling of regional waveforms. Lowercrustal earthquakes are found to be restricted to the forelands in areas undergoing shortening,and to locations where rifting coincides with abrupt changes in lithosphere thickness, suchas the NE Baikal Rift and W Laptev Sea. The lower crust in these areas is seismogenic attemperatures that may be as high as 600◦C, suggesting that it is anhydrous, and is likely tohave great long-term strength. Lower crustal earthquakes are therefore a useful proxy indi-cating strong lithosphere in places that are too small in areal extent for this to be confirmedindependently by estimating effective elastic thickness from gravity–topography relations.The variation in crustal rheology indicated by the distribution of lower-crustal earthquakeshas many implications ranging from the support of mountain belts and the formation of steepmountain fronts, to the localization and orientation of rifting. In combination, these processescan also be responsible for the separation of the front of the thin-skinned mountain belts fromtheir hinterlands when continents separate.

Key words: Earthquake source observations; Seismicity and tectonics; Continental neotec-tonics; Continental tectonics: compressional; Continental tectonics: extensional; Asia.

1 I N T RO D U C T I O N : R H E O L O G YA N D E A RT H Q UA K E D E P T H S

The last decade has seen a great deal of interest in the relationsbetween lithosphere rheology, structure and composition on thecontinents. Beginning with reassessments of effective elastic thick-ness (Te) estimates and earthquake depth distributions by McKenzie& Fairhead (1997) and Maggi et al. (2000a), it became clear thataccepted views which had served well and remained unchallengedfor 20 years, such as those of Brace & Kohlstedt (1980) and Chen& Molnar (1983), needed some modification. In particular, a singlegeneric view of the continental lithosphere could not account for thedifferences in earthquake depths, elastic thicknesses and geologicalhistories commonly observed between ancient shields (or cratons)and younger orogenic belts.

An important conclusion of Maggi et al. (2000a) was that thelong-term strength of the continental lithosphere resides only in itsupper part, which was contained wholly within the crust, and thatthere was little evidence for substantial long-term strength in thecontinental mantle, contrary to previously accepted views. The ef-

fort to understand the physical mechanisms responsible led quicklyto a series of developments in related subjects, including: (1) a re-examination of geotherms and thermal structure in both oceans andcontinents; (2) connections between seismic velocity and tempera-ture; (3) relations between mechanical properties and the metasta-bility of lower crustal rocks in mountain roots; (4) causes of thevariations between the ancient continental shields (or ‘cratons’) andyounger Phanerozoic orogenic belts and (5) the issue of how thecontinental cratons were created. The new views originating in thestudies of McKenzie & Fairhead (1997) and Maggi et al. (2000a)have, of course, been challenged; but because the subject has beenfast-moving, debate in publications has not been able to keep pacewith these new developments.

A recent review by Jackson et al. (2008) summarized the devel-opments above, in order to demonstrate that a coherent picture isemerging which reconciles observations from fields as diverse asseismology, gravity, heat flow, rock mechanics, metamorphic petrol-ogy and geochemistry. Furthermore, the insights, and agreement,that these widely differing disciplines offer on the same subject pro-duce an overall view that is more robust than any obtained from just

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2 R. A. Sloan et al.

one of those disciplines alone. A relatively simple view of the dis-tribution of lithosphere seismicity and its implications for rheologyis now emerging, which can be summarized as follows:

(1) Earthquakes in the mantle, in both oceans and continents,are confined to regions colder than about 600◦C. In most places onthe continents the uppermost mantle is expected to be at 600◦C ormore, and hence aseismic. However there are some circumstancesin which the continental mantle might be colder, and experiencesome earthquakes (e.g. Emmerson 2007; Priestley et al. 2008a),one example of which is in the Himalaya (below).

(2) With very few exceptions, earthquakes everywhere are con-fined to a single seismogenic layer, which in the oceans is limitedby the 600◦C isotherm, in young orogenic belts is typically lim-ited to the upper crust (≤ ∼350 ◦C), and in some regions, oftenin or adjacent to ancient shields, may include the whole crust. Animportant exception is in the Himalaya, where the seismogeniclower crust of the Indian shield underthrusts the seismogenic uppercrust of Tibet, giving an apparent bimodal depth distribution, butone that is not in steady-state and has no generic significance forcontinental rheology. In this particular example of the Himalaya,where the underthrusting Indian crust is unusually thin (∼35 km)for a Precambrian shield and the Moho temperature is correspond-ingly unusually cold (∼500 ◦C rather than ∼600 ◦C), earthquakesextend from the lower crust into the uppermost 10 km of the mantle(Monsalve et al. 2006; Priestley et al. 2008a).

(3) The great strength of some of the ancient shields is associatedwith lower crustal earthquakes and larger elastic thickness than inyounger continental lithosphere. In such shields two effects arelikely to be responsible: (a) the crust may be relatively cold becauseof a thick low-density mantle root and (b) the composition of thelower continental crust is probably dominated by a dry granulite-facies mineral assemblage.

(4) Lateral strength contrasts in the continents between ancientshields and young orogenic regions are important, and cannot berepresented by a laterally uniform continental rheology or com-position. In spite of the lower-than-average Moho temperatures inshields, lower crustal earthquakes within them occur in materiallikely to be at 500–600 ◦C, much hotter than the normal cut-offtemperature of ∼350 ◦C for earthquakes in younger regions; a situ-ation which probably requires the seismogenic lower crust to be dry.These lateral strength contrasts allow mountains to be supported bytheir adjacent forelands without requiring the mantle beneath theforelands to be strong.

(5) The stability and survival of the ancient shields and cratonsover Ga of geological time is related to both their strength andbuoyancy, neither of which can easily be changed.

During the period of these developments an important advancehas been the ability to map the lithosphere thickness on the conti-nents using surface wave tomography. This involves using funda-mental and higher-mode surface waves to spatially map shear wavevelocity as a function of depth Vs(z) (e.g. Debayle & Kennett 2000;Ritsema & van Heijst 2000; Priestley & Debayle 2003), conver-sion of Vs(z) to temperature as function of depth T(z), and then thefitting of geotherms to establish the depth of the rapid change intemperature gradient over the thickness of the thermal boundarylayer that corresponds to the base of the lithosphere (Priestley &McKenzie 2006). For reasons explained by Priestley & McKenzie(2006), Priestley et al. (2006) and McKenzie & Priestley (2008),technical limitations restrict such analyses to regions where thelithosphere is thicker than about 125 km, and impose horizontal andvertical resolutions of about 300 and 30 km, respectively. Nonethe-

Figure 1. (a) Lithospheric thickness (from Priestley et al. 2006) and regionalseismicity of eastern Asia. Values of lithospheric thickness are calculated ona 2◦ × 2◦ grid limiting the scale of reliably resolved structures. Pink areasare only constrained to have a lithospheric thickness of <125 km. For moredetails see Section 2.2. (b) Topography and regional seismicity of easternAsia. In both maps earthquakes with depths of less than 100 km from theEHB catalogue are shown as red circles.

less, these studies confirm that the continental lithosphere can reachthicknesses of 250 km or more, with dramatic variations in thick-ness within, and on the margins of, the deforming regions of centralAsia (Fig. 1). As Priestley & McKenzie (2006) emphasize, thereis no simple correlation between thick lithosphere and the ancientshields: not all shields have thick lithosphere, and not all regionswith thick lithosphere are associated with the Precambrian shields;nor can the age of the thick lithosphere be determined from seis-mology.

The purpose of this study is to examine the spatial relations be-tween earthquake depth distributions and lithosphere thickness inAsia. There are two principal reasons for doing this. One is to exam-ine the contrast in the behaviour between extending and shorteningregions and whether this differs in places where the lithosphere

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Earthquake depth distributions in central Asia 3

is thicker than normal. In principle, the relation between gravityand topography can be used to estimate effective elastic thick-ness and thereby identify regions that are unusually strong, suchas north India. However such analyses are not always possible orsensible: for example, when analysed in the space domain, flexuralsignals in strong regions have wavelengths of several hundred km(e.g. McKenzie & Fairhead 1997) and in intermontane basins evenas large as the western Tarim Basin (200–300 km across), it is notstraightforward to isolate the behaviour of one margin from another.Similarly, in the spectral domain, elastic thickness estimated fromadmittance or coherence studies of gravity and topography are nec-essarily averages over broad regions typically at least a thousandkilometres in dimension, which are likely to include substantialvariations in local properties and structure within them. Addition-ally, while gravity-topography analyses can be used to estimatethe long-term strength of a region, they are unable unambiguouslyto determine the depths at which this strength resides. HoweverJackson et al. (2008) point out that, where good estimates of bothelastic (Te) and seismogenic (Ts) thickness are available, Te ≤ Ts

and nowhere do the data require Te > Ts. The simplest explanationof this relationship is that long-term strength resides in the seismo-genic layer. In particular, Jackson et al. (2008) emphasize that lowercrustal earthquakes on the continents are probably associated withdry granulite terranes that are likely to be strong; in which case theiroccurrence may be a proxy for the likely anhydrous composition,and probable strength, at depth. The identification of lower crustalseismicity would then allow the examination of regions too smallfor gravity-based studies of Te to be appropriate and would providean important constraint on the depths at which long-term strengthresides.

A second aim of this paper is therefore to see whether lowercrustal earthquakes are associated with regions whose topographyor geological history suggest they are strong, even though directestimates of Te are unavailable or impossible, and thereby examinethe viability of this potential proxy.

We first examine the regions of shortening in the margins ofTibet, the Tien Shan and their forelands. We then investigate themargins of the Siberian shield, from the Baikal rift to the LaptevSea. The contrasts revealed by these regions give several insightsinto the geological development of continental rifts and mountainbelts.

2 M E T H O D S

2.1 Earthquake source parameters

This paper is concerned with accurate estimates of earthquakesource parameters, especially depth. The principal focus is on earth-quakes whose parameters are determined teleseismically, and whichare therefore larger than about Mw 5.2. For the better recorded eventswe use synthetic seismogram modelling to invert for all the sourceparameters of interest. For some smaller events, we fix all sourceparameters other than centroid depth and use synthetic seismogrammodelling to match the observed P, pP and sP depth-phase sep-aration. A number of other studies have used similar or identicalteleseismic methods to estimate source parameters (e.g. Huang &Chen 1986; Doser 1991a,b; Bayasgalan & Jackson 2002; Delouiset al. 2002; Brazier & Nyblade 2003; Bayasgalan et al. 2005a;Emmerson et al. 2006). These teleseismic methods are well es-tablished and only a brief description is included below. In someareas of interest to this paper other studies have used dense local

seismograph networks to constrain the depth distribution of smallerevents (e.g. Deverchere et al. 1991; Avetisov 1993; Roecker et al.1993; Kovachev et al. 1994; Schulte-Pelkum et al. 2005; Monsalveet al. 2006; Xu et al. 2006; Suvorov & Tubanov 2008) and these arediscussed in each region, as appropriate.

2.1.1 Long-period body wave inversion for all source parameters

For the larger earthquakes, we took broad-band seismograms fromthe Global Digital Seismograph Network (GDSN) and deconvolvedthem to change the response to that of a World-Wide StandardizedSeismograph Network (WWSSN) 15–100 s long-period instrument.The removal of the higher frequency parts of the spectrum reducessensitivity to complexities in local velocity structure and allowsevents of Mw 5.0–6.5 to be modelled as point (centroid) sources.We then use the MT5 version (Zwick et al. 1994) of the algorithm de-veloped by McCaffrey & Abers (1988) and McCaffrey et al. (1991)which inverts P and SH waveform data for source time-function,scalar moment, strike, dip, rake and centroid depth. The source isconstrained to be a pure double-couple. P, pP and sP phases aremodelled on vertical component seismograms with an epicentralrange spanning 30–90◦ and S and sS phases are modelled on trans-verse component seismograms in the range 30–80◦. Amplitudes arecorrected for the effects of geometrical spreading and anelastic at-tenuation using Futterman operators with a t* of 1.0 and 4.0 s forP and SH waves, respectively. Stations are weighted according toazimuthal density and the SH waveforms are weighted 0.5 relativeto P waveforms. When possible the waveforms are aligned usingtime picks from the broad-band records.

As the velocity structures of the areas studied are not well known,and probably show marked lateral variation, we have chosen to use asimple half-space velocity model with Vp 6.5 km s−1, Vs 3.7 km s−1

and density ρ = 2800 kg m−3 for all events. This approachallows a fairer comparison between events within this studyand those previously published (e.g. Bayasgalan et al. 2005a;Emmerson et al. 2006). Previous studies have used sensitivity anal-ysis to establish that the typical uncertainty in centroid depth is lessthan 4 km (e.g. Molnar & Lyon-Caen 1989; Emmerson et al. 2006),which is sufficient to conclusively separate upper crustal seismicityfrom lower crustal seismicity. Typical uncertainties in the other pa-rameters (e.g. strike, dip and rake of the focal planes) are discussedby other authors (Nabelek 1984; McCaffrey & Nabelek 1987;Molnar & Lyon-Caen 1989; Taymaz et al. 1990) and are not themain focus of this paper.

2.1.2 Modelling of depth phases

In cases where only a few stations were available, or where the long-period signal was not clearly distinguishable above backgroundnoise, we were unable to invert for all source parameters. For theseevents we fixed the strike, dip and rake to that of the best double-couple Global CMT solution and attempted to match synthetic seis-mograms to the early part of the vertical component broad-banddata. We used the program WKBJ3 (Chapman 1978; Chapmanet al. 1988), which traces rays through a spherical Earth by us-ing the WKBJ approximation for turning waves. Impulse responsesfor P, pP and sP phases are generated and convolved with each sta-tion’s broad-band response and an attenuation corresponding to t* =1.0 s. We use a version of the AK135 global velocity model(Kennett et al. 1995), modified to have a crustal thickness of 40km. The synthetic and broad-band data are aligned by picking the




maximum amplitude point in the first peak or trough (correspond-ing to the P phase). The depth is then varied until a satisfactoryvisual fit for the pP and sP is achieved. This method requires theconfident identification of these phases in order to give reliabledepths. This can be difficult when the centroid depth is shallowerthan about 10 km because the impulses from these phases over-lap, preventing definite identification. At depths of more than 10km there is a clear separation of phases, and the simple, impul-sive source time functions of the relatively small events mean thatthe fit between the synthetic and observed seismograms is signifi-cantly reduced by a 1–2 km change in depth. This means that themain source of uncertainty is the velocity structure and the error incentroid depth will be less than ±4 km. At depths shallower thanthis the main uncertainty is due to the possible misidentification ofdepth phases; however, the events can often be constrained to bewithin the uppermost crust, which is sufficient for the purposes ofthis paper. The results from this depth-phase modelling are listed inTable 2.

In the Tibetan Plateau and its forelands 123 events with reliablydetermined centroid depths are shown in Table 1. This includes 21events presented here for the first time, including 13 obtained us-ing a body wave inversion for all parameters. Table 2 shows 123earthquakes for which reliable centroid depths are known in theTien Shan. Of these, 51 were obtained through the inversion ofregional data including three from this study, and 69 are largerearthquakes studied using teleseismic data. This number includes50 presented in this study, 31 obtained from body wave inver-sion for all parameters and 19 through depth phase modelling. Inthe Lake Baikal Region (including the fold-and-thrust belt the tothe south) 49 events with reliably located depths are presented inTable 3. Fifteen of these are presented here for the first time in thisstudy including three for which a full body wave inversion has beenperformed. In northeastern Siberia 27 events with reliable centroiddepths are presented in Table 4. Four of these (including one dou-ble event) are presented here for the first time, all of which havebeen determined through a full body wave inversion. Some otherevents have been repeated using depth phase modelling. In totalthen we consider 323 earthquakes across central and eastern Asia,and present new results for 94 of these. 51 of these were obtainedthrough body wave inversion for all parameters and therefore alsoconstrain the mechanisms of these earthquakes to a greater degreethan previous studies, 40 were obtained through depth phase mod-elling and three through the inversion for all parameters of regionaldata.

2.1.3 Modelling of local and regional waveforms

In one area of particular interest in the Tien Shan foreland we usedregionally recorded broad-band waveforms from the GHENGIS net-work (∼150–400 km from the events studied) to model the depthsand focal mechanisms of events too small to analyse teleseismi-cally. We used the program FKCMTINV to perform this analysis.The method follows that of Kao et al. (1998a,b) and involves thecreation of a depth array (with 1 km spacing) to calculate the Greensfunctions in a 1-D model for each receiver. The algorithm used tocalculate the Greens functions is based on the technique of Yao &Harkrider (1983) which combines the reflectivity method (Kennett1980) and discrete wavenumber summation (Bouchon 1981). Ateach depth of the depth array a moment tensor inversion is per-formed to minimise the misfit between the observed and syntheticwaveforms. The optimum depth and moment tensor is then selected

by finding the global minimum over all the depths in the deptharray. Observed waveforms were bandpass filtered between 0.03and 0.08 Hz in order to minimize the effect of variation withinthe local velocity structure. It should be noted that nearby stationswithin the Tarim Basin (such as at Kashgar) had to be removedfrom the inversion due to high amplitude long-period noise. Thisresults in poor azimuthal coverage of the events studied, which islikely to increase uncertainties, particularly in mechanism. Misfitversus depth plots are shown in the Supporting Information al-lowing the range of minimum misfit to be assessed. There willalso be an error associated with the simple velocity model used(Vp 6.0 km s−1, Vs 3.45 km s−1 and density ρ = 2600 kg m−3 forthe top 40 km, Vp 6.96 km s−1, Vs 4.00 km s−1 and density ρ =2900 kg m−3 for the next 30 km, and Vp 8.35 km s−1, Vs 4.80 kms−1 and density ρ = 3640 kg m−3 below that). The misfit Vs depthplots have relatively broad minima (compared to the other methodsused) and so we estimate that in total the error is of the order of±5–6 km.

2.2 Lithosphere thickness maps

An important recent development is the availability of lithosphericthickness maps (e.g. Fig. 1) published by Priestley et al. (2006),produced by using the surface wave methods described in the In-troduction. The vertical resolution of ∼30 km is achieved throughthe use of the higher modes and the period range of 50–160 s usedin the analysis limits the horizontal resolution to 200–400 km. Themaps are created by estimating lithospheric thickness from Vs(z)at points on a 2◦ × 2◦ latitude-longitude grid and then contouringthe values. Uncertainties in crustal models, and the sensitivity oflong-period surface waves to the crustal velocity structure, restrictaccurate estimation of the lithosphere–asthenosphere transition toregions where it is deeper than ∼125 km, though areas where thetransition is shallower than this can be identified. These uncertain-ties and procedures limit the extent to which detail can be interpretedin these maps. The methods and procedures are discussed in moredetail in Priestley & McKenzie (2006), Priestley et al. (2006) andMcKenzie & Priestley (2008).

3 S H O RT E N I N G

3.1 The Tibetan Plateau and its forelands

3.1.1 India south of the Himalayan Front

Teleseismically determined source parameters of earthquakes inthe Himalayan foreland, and their depths compared to Moho depthsdetermined from receiver functions, have been reviewed by Maggiet al. (2000b), Jackson (2002), Mitra et al. (2005) and Priestleyet al. (2008a). The general pattern is of earthquakes throughoutthe crust of the Indian shield, with the lower crust being moreseismically active in the north than in the south of India (Figs 2 and3a). Notable features of this seismicity include the 1997 Jabalpurearthquake in central India with a centroid at 35 km (Battacharyaet al. 1997; Maggi et al. 2000b), and thrusts at depths of 35–50 km beneath the Ganges basin with normal faults at depths of∼20 km above them, consistent with bending of the shield south ofthe Himalaya (Jackson 2002).

That the Indian shield is seismogenic throughout its crust isconfirmed by local aftershock studies of the 2001 Bhuj (Gujarat)earthquake by Negishi et al. (2002), Bodin & Horton (2004),




Table 1. India, Tibet and its forelands. Earthquake source parameters from seismology. Epicentres and origin times between 1964-2007 are taken from theupdated catalogue of Engdahl et al. (1998). References from which source parameters have been taken are indicated by La (Langston 1976), Ch81 (Chen et al.1981), Zho (Zhou et al. 1983), MC (Molnar & Chen 1983), Ba (Baranowski et al. 1984), Jo (Jones et al. 1984), LF (Langston & Franco-Spera 1985), Ek(Ekstrom 1987), MLC (Molnar & Lyon-Caen 1989), FN (Fan & Ni 1989), CM90 (Chen & Molnar 1990 with depths recalculated using the velocity structureof Mitra et al. 2005), Chu (Chung 1993), CK (Chen & Kao 1996), Co (Cotton et al. 1996) Se (Seeber et al. 1996), ZH (Zhu & Helmberger 1996), Ja (Jackson2002), AD (Antolik & Dreger 2003), CY (Chen & Yang 2004), Mi (Mitra et al. 2005), Pr (Priestley et al. 2008a), El-a (Elliott et al. 2010a), El-b (Elliott et al.2010b) and TS (This Study). Earthquakes marked with an asterisk have had their depths determined by depth phases modeling (see Section 2.1.2), and theirmechanisms are taken from the Global CMT catalogue. Source parameters of the event marked Ek (Ekstrom & England 1989) were determined by a muchbroader-band CMT-type inversion. The source parameters of the remaining events were obtained through the inversion of teleseismic body waveforms. Focalmechanisms are plotted in Fig. 2b and the final column of this table with the compressional quadrants shaded grey (depth <20 km), red (depth = 20–50 km)or green (depth > 50 km).

Tibet and its Forelands: Waveform Modelling Results

Date Time Lat/◦ Long/◦ Depth/km Mw Focal Mechanism Reference

yyyy mm dd hh mm Strike/◦ dip/◦ Rake/◦ FPS

1962 05 21 12 02 37.13 95.73 11 6.6 285 39 74 MLC

1963 04 19 07 35 35.53 96.44 10 6.7 277 80 -10 MLC

1963 06 19 10 47 24.97 92.06 45 5.9 57 80 42 CM90*

1963 06 21 15 26 25.13 92.09 36 5.7 238 88 −70 CM90*

1963 06 26 14 09 36.38 76.61 89 6.0 133 38 −105 CY*

1964 03 16 01 05 36.99 95.62 10 5.3 70 77 50 MLC

1964 09 26 00 46 29.89 80.45 18 5.9 310 23 90 Ba

1964 10 21 23 09 28.04 93.76 15 6.0 265 3 90 Ba

1965 01 12 13 32 27.40 87.85 15 5.9 270 15 90 Ba

1965 06 22 05 49 36.20 77.57 91 4.9 164 64 −134 CY*

1966 03 06 02 15 31.41 80.47 8 6.3 0 45 −90 MC

1966 06 27 10 41 29.55 80.42 15 6.0 277 27 70 Ba

1966 08 15 02 15 28.67 78.84 20 5.6 132 31 −90 Mo

1966 10 14 01 04 36.51 87.51 8 5.4 25 66 −90 MC

1966 12 16 20 52 29.62 80.82 12 5.7 290 24 90 Ba

1967 02 20 15 18 33.55 75.32 10 5.5 341 55 105 Ba

1967 03 14 06 58 28.44 94.32 15 5.7 273 10 90 Ba

1967 08 30 04 22 31.62 100.23 8 6.1 245 45 −70 Zh

1967 08 30 11 08 31.52 100.24 10 5.3 233 50 −90 Zh

1967 12 10 22 51 17.39 73.77 5 6.3 16 67 −29 La

1967 12 12 06 18 17.22 73.78 4 5.7 100 40 −120 LF

1968 06 12 04 29 24.82 91.89 38 5.3 132 60 135 CM90*

1969 04 13 15 24 17.83 80.68 10 5.8 245 72 −2 Chu

1970 02 19 07 10 27.42 93.95 10 5.4 257 5 90 Mo

1970 02 24 02 07 30.64 103.08 7 5.6 276 47 134 MLC

1970 03 23 01 53 21.58 72.98 11 5.4 273 58 130 Chu

1971 02 02 07 59 23.72 91.64 38 5.4 119 36 90 CM90

1971 03 24 13 54 35.43 98.03 7 6.0 283 74 5 MLC

1971 04 03 04 49 32.16 95.01 9 5.8 260 79 −5 MLC

1971 05 22 20 03 32.39 92.12 8 5.6 58 90 3 MC

1971 05 03 00 33 30.77 84.34 8 5.5 190 58 −90 MC

1971 07 17 15 00 26.39 93.16 36 5.5 79 60 46 CM90

1972 07 22 16 41 31.43 91.49 8 5.9 212 65 −17 MC

1972 08 30 15 14 36.64 96.32 15 5.3 90 62 60 MLC

1972 08 30 18 47 36.57 96.31 19 5.3 91 58 38 MLC

1972 09 03 16 48 35.97 73.34 12 6.2 341 55 105 Ba

1973 02 06 10 37 31.39 100.47 10 5.9 305 87 −179 Zho




Table 1. (Continued.)




1973 02 07 16 06 31.45 100.29 10 5.9 210 60 −90 Zho

1973 07 14 13 39 35.19 86.60 7 5.7 37 68 −56 MC

1973 07 14 04 51 35.15 86.44 6 6.6 81 60 −35 MC

1973 08 01 14 05 29.57 89.14 81 4.9 220 60 −24 MC

1973 08 11 07 15 32.89 103.99 4 5.8 326 85 10 MLC

1973 08 16 08 02 33.21 86.81 8 5.4 160 55 −155 MC

1973 09 08 07 25 33.29 86.84 9 5.8 118 60 −161 MC

1974 03 24 14 16 27.65 86.01 16 5.4 275 2 90 Ba

1974 12 28 12 12 35.00 72.92 12 6.0 354 47 102 MC

1975 01 19 08 01 32.36 78.50 9 6.7 0 50 −90 MC

1975 04 28 11 06 35.77 79.90 7 5.9 169 62 −149 MC

1975 05 19 19 47 35.14 80.84 8 5.6 248 66 −50 MC

1975 05 05 05 18 33.11 92.84 7 5.9 250 78 −14 MC

1975 06 04 02 24 35.88 79.94 9 6.0 180 62 −121 MC

1975 07 19 06 10 31.91 78.64 6 5.2 180 50 −125 MC

1975 07 29 02 40 32.57 78.49 8 5.3 210 55 −90 MC

1976 08 16 14 06 32.75 104.08 12 6.1 165 63 40 Jo

1976 08 21 21 49 32.55 104.23 5 6.0 215 60 90 Jo

1976 08 23 03 30 32.45 104.17 8 6.2 165 65 40 Jo

1976 09 14 06 43 29.78 89.53 90 5.9 190 57 −90 Ch81

1976 09 22 20 07 40.02 106.32 8 5.5 230 75 −111 MLC

1976 10 01 11 27 35.98 77.40 86 5.1 336 41 −109 CY*

1977 01 01 21 39 38.18 90.99 8 5.8 288 36 82 MLC

1977 01 19 00 46 37.01 95.70 14 5.7 305 38 75 MLC

1977 11 18 05 20 32.64 88.40 11 6.3 236 68 −29 MLC

1978 04 04 00 40 32.99 82.31 11 5.8 327 78 −164 MLC

1978 07 31 11 55 35.42 82.11 6 5.5 236 77 −8 MLC

1979 03 29 07 07 32.49 97.17 12 5.8 270 84 −6 MLC

1979 05 20 22 59 29.92 80.27 16 5.5 251 16 53 MLC

1979 06 19 16 29 26.76 87.48 20 5.2 75 45 −83 MLC

1980 02 13 22 09 36.47 76.86 85 6.0 122 78 112 FN,Ch88

1980 02 22 03 02 30.50 88.61 6 6.2 188 48 −84 MLC

1980 06 01 06 18 38.97 95.60 12 5.4 128 53 48 MLC

1980 06 24 07 35 32.99 88.58 11 5.6 71 75 −15 MLC

1980 07 29 12 23 29.32 81.22 14 5.3 279 29 94 MLC

1980 07 29 14 58 29.64 81.05 18 6.4 288 25 86 MLC

1980 08 23 21 36 32.93 75.69 14 5.3 265 14 45 MLC

1980 08 23 21 50 32.92 75.71 13 5.3 320 5 90 MLC

1980 10 07 09 32 35.58 82.16 4 5.7 186 40 −77 MLC

1980 11 19 19 19 27.37 88.77 44 6.0 214 71 12 Ek

1981 01 23 21 13 30.97 101.17 7 6.5 322 80 5 MLC

1981 06 09 22 22 34.47 91.41 9 5.8 86 83 −6 MLC

1981 09 12 07 15 35.68 73.64 7 5.9 138 42 104 MLC

1982 01 23 17 37 31.67 82.27 9 6.3 210 68 −79 MLC








1982 06 15 23 24 31.83 99.87 9 5.7 55 70 −90 MLC

1985 05 20 15 11 35.54 87.22 8 5.7 234 77 3 MLC

1986 01 10 03 46 28.65 86.57 81 5.5 140 46 −163 Ek,Ch88

1986 04 26 07 35 32.17 76.42 13 5.3 254 16 22 MLC

1986 06 20 17 12 31.17 86.75 9 6.0 138 78 178 MLC

1986 07 16 22 03 31.08 78.06 13 5.1 305 8 90 MLC

1986 07 06 19 24 34.44 80.21 5 5.8 248 51 −27 MLC

1986 08 20 21 23 34.54 91.63 11 6.2 261 88 −11 MLC

1986 08 26 09 43 37.76 101.56 7 5.9 331 59 107 MLC

1988 02 06 14 50 24.68 91.51 31 5.9 225 76 5 CM90

1988 08 20 23 09 26.73 86.59 51 6.7 246 20 22 CK

1991 10 19 21 23 30.74 78.79 16 6.6 318 11 114 Co

1991 11 08 15 13 26.28 70.58 23 5.4 74 33 61 Ja

1991 12 21 19 52 27.88 87.98 66 4.7 112 82 −179 ZH

1992 01 24 05 04 35.51 74.55 71 − 80 90 0 Pr*

1992 03 07 00 00 24.99 89.37 76 4.3 350 68 −164 ZH

1992 04 04 17 43 28.15 87.96 76 4.9 46 66 −22 ZH

1992 07 30 08 24 29.57 90.16 10 5.9 175 52 −109 TS

1993 09 29 22 25 18.07 76.49 3 6.1 126 46 100 Se

1995 02 17 02 44 27.61 92.30 37 5.5 317 62 167 Mi

1996 04 01 08 08 31.50 73.44 38 5.5 98 33 101 Ma

1996 06 09 23 25 28.38 92.25 65 5.1 206 79 −79 CY*

1996 11 19 00 12 24.50 92.63 43 5.4 67 79 17 Mi

1996 11 19 10 44 35.34 78.12 10 6.8 95 87 −12 TS

1997 05 28 02 53 24.88 92.25 30 5.9 238 79 2 Mi

1997 05 21 22 51 23.10 80.12 35 5.7 65 64 75 Ba

1999 03 28 19 05 30.51 79.40 16 6.4 280 10 75 TS

2000 09 05 03 24 17.37 73.84 6 5.2 351 39 −94 Pr*

2001 01 26 03 16 23.42 70.23 20 7.5 281 42 107 Co

2001 07 16 16 07 33.01 73.12 47 5.1 273 61 109 Ja

2001 09 25 14 56 11.96 80.20 20 5.0 265 40 109 Pr*

2002 06 04 14 36 30.51 81.45 9 5.4 159 49 −130 TS*

2002 11 20 21 32 35.40 74.51 4 6.1 225 32 −87 TS

2003 03 25 18 51 27.31 89.41 37 5.4 37 87 −20 TS

2004 07 11 11 23 30.69 83.66 9 6.2 159 39 −90 TS

2004 07 28 22 22 30.63 83.61 3 5.1 9 47 −64 TS*

2004 10 26 02 11 30.98 81.07 8 5.6 112 73 −165 TS*

2005 04 07 20 04 30.48 83.66 5 6.2 169 40 −83 TS

2005 04 08 19 51 30.44 83.63 10 5.0 177 45 −84 TS*

2005 10 08 03 50 34.43 73.54 21 7.3 342 32 125 TS

2006 04 06 17 59 23.30 70.44 22 5.5 231 70 4 TS

2007 05 05 08 51 34.25 82.01 7 6.0 152 49 −125 TS

2008 01 09 08 26 32.34 85.30 6 6.4 208 47 −74 TS

2008 01 16 11 54 32.23 85.16 4 5.7 206 43 −89 TS








2008 01 17 07 09 32.42 85.26 8 4.9 7 31 −115 TS*

2008 03 20 22 33 35.54 81.37 7 7.1 206 45 −75 El-b

2008 05 12 20 08 31.37 104.08 17 5.5 199 38 109 TS

2008 08 25 13 22 30.89 83.51 8 6.6 25 46 −68 El-b

2008 08 29 09 43 30.58 83.52 5 4.9 171 39 −90 TS*

2008 10 06 08 30 29.80 90.34 7 6.2 185 51 −113 El-b

2008 10 06 12 10 29.56 90.53 9 5.2 173 43 −121 TS*

2008 10 08 14 07 29.76 90.57 7 5.5 69 62 −25 TS*

2008 11 10 01 22 37.51 95.75 18 6.2 248 35 60 El-a

2008 12 08 08 59 29.65 82.01 9 5.3 230 62 −17 TS*

2009 08 28 01 52 37.66 95.73 5 6.2 111 72 89 El-a

2009 08 31 10 15 37.59 95.86 4 5.7 291 34 89 El-a

Mandal & Pujol (2006) and Bhatt et al. (2009) and of the 1997Jabalpur earthquake by Acharyya et al. (1998). A geodetic analysisof the 1897 Assam earthquake, on the north side of the Shillongplateau, also required faulting to extend from the surface to the baseof the crust (Bilham & England 2001).

3.1.2 The High Himalaya

Teleseismic analysis shows that beneath the High Himalaya are anumber of gently north-dipping thrust-faulting earthquakes withcentroid depths in the range 10–20 km (Molnar & Lyon-Caen 1989,Figs 2 and 3c). These approximately coincide in depth and orien-tation with the Main Himalayan Thrust (MHT) (Mitra et al. 2005;Schulte-Pelkum et al. 2005; Nabelek et al. 2009), a ramp-and-flatstructure constituting the upper boundary of the Indian shield as it isthrust under Tibet. Beneath them, at depths of 35–45 km but abovethe Moho, are other large events, mostly with strike-slip mecha-nisms, that are clearly within the lower crust of the underthrustingIndian shield (Mitra et al. 2005, Fig. 3c). Deeper still are someearthquakes that are arguably just beneath the Moho of the Indianshield, such as the 1988 Nepal earthquake of Mw 6.7 at 51 km be-neath the Ganges basin (Chen & Kao 1996), the deepest earthquakein Fig. 3(a).

This seismicity distribution is confirmed by data from local seis-mograph networks. Schulte-Pelkum et al. (2005) and Monsalveet al. (2006) show that beneath the seismicity associated with andabove the MHT, the lower crust of the Indian shield is seismogenic,and that microearthquakes extend about 20 km into the uppernmostmantle beneath the Moho. This is therefore a case where the seis-mogenic layer involves the whole crust and the uppermost uppermantle. However, as Priestley et al. (2008a) point out, this situationis consistent with the conclusion of McKenzie et al. (2005) that theupper mantle is seismogenic at temperatures of up to ∼600 ◦C. Innorth India, the shield has a crystalline crustal thickness (i.e. ignor-ing the sediments of the Ganges basin) of only about 35 km, andfor this reason the Moho temperature beneath the Himalaya mayonly be ∼500 ◦C (see McKenzie et al. 2005; Priestley et al. 2008a).In addition, Priestley et al. (2008a) pointed out that these mantleevents are seen at locations where the Indian Plate is flexing down

or flattening out beneath the load of the Himalaya in the ramp-and-flat structure observed by Schulte-Pelkum et al. (2005), and that inthese locations the higher strain rates may also cause the mantle tobe seismogenic.

3.1.3 The Tibetan Plateau

In the Tibetan Plateau itself, most of the earthquakes are normal orstrike-slip events with centroids shallower than 12 km (e.g. Molnar& Lyon-Caen 1989). On its margins, and at lower elevations, somethrust faulting earthquakes occur deeper, to about 20 km, as inQaidam in the NE Molnar & Lyon-Caen (1989) and Schezwan inthe E (Ji & Hayes 2008). All of these events are well within the uppercrust. Crustal thickness estimates from receiver function analysesvary from 70 to 90 km in the Tibetan Plateau (e.g. Wittlinger et al.2004; Mitra et al. 2005; Rai et al. 2006; Priestley et al. 2008,Nabelek et al. 2009) and are up to ∼55 km beneath the QaidamBasin (e.g. Herquel et al. 1995). No events with a magnitude greaterthan 5 (i.e. those large enough to be studied teleseismically) havebeen observed in the mid crust (12–70 km) in the central part of theplateau or below 20 km in the Qaidam Basin (Fig. 3b).

There are two areas beneath the Tibetan plateau where someearthquakes of magnitude 5–6 have occurred at depths of about70–90 km (green focal spheres in Fig. 2b, also Fig. 3b), and whichhave attracted much attention (Chen & Yang 2004; Mitra et al. 2005;Priestley et al. 2008). One area is in the southern part of the plateauand the other in the far NW beneath the Kun Lun mountains. Tele-seismic determination of these earthquake depths is by identificationof pP–P and sP–P times, and the crustal thickness in their vicinityis determined by receiver functions, where the crucial measurementis the delay between the direct P and the P-to-S conversion at theMoho. Converting these observations of time to depth, and doing sowith the same crustal velocity determined from joint surface waveand receiver-function studies, shows that these earthquakes occurvery close to the Moho, so that, allowing for errors in the estimatesof both centroid depths and crustal thickness, it is difficult to besure which side of the Moho they occur (Priestley et al. 2008). Thesignificance of these earthquakes is discussed at length by Jacksonet al. (2004) and Priestley et al. (2008). Key points are:




Table 2. The Tien Shan and its forelands. Earthquake source parameters from seismology. Epicentres and origin times between 1964-2007 are taken from theupdated catalogue of Engdahl et al. (1998). References from which source parameters have been taken are indicate by ChK (Chen & Kao 1996), N (Nelsonet al. 1987), Me (Mellors et al. 1997), Gh7 (Ghose et al. 1997), Gh (Ghose et al. 1998b), F (Fan et al. 1994), Ba (Bayasgalan et al. 2005b), De (Delouis et al.2002) and TS (this study).Earthquakes marked by a dagger have been modelled using regional data, all other earthquakes have been modelled using teleseismicdata. Earthquakes marked with an asterisk have had their depths determined by depth phase modeling (see Section 2.1.2), and their mechanisms are taken fromthe Global CMT catalogue. The source parameters of the remaining events were obtained through the inversion of teleseismic body waveforms. Earthquakesmarked by d were found to require a double event in order to satisfactorily fit the observed waveforms. Focal mechanisms are plotted in Fig. 4b and the finalcolumn of this table with the compressional quadrants shaded grey (depth < 15 km), green (depth = 15–30 km) or red (depth > 30 km).

Tien Shan and its Forelands: Waveform Modelling Results



1965 11 13 14 33 43.87 87.74 44 6.3 72 43 81 N

1969 02 11 22 08 41.42 79.24 10 6.1 65 41 83 N

1970 05 15 17 13 50.20 91.26 7 6.3 272 50 87 Ba

1970 06 05 04 53 42.48 78.71 17 6.3 69 41 68 N

1971 03 23 20 37 41.42 79.20 11 5.8 73 46 93 N

1971 05 10 14 51 42.85 71.29 15 5.4 37 48 57 N

1972 04 09 04 10 42.09 84.58 13 5.4 100 49 90 N

1973 06 02 23 57 44.14 83.60 26 5.5 148 28 133 N

1974 07 04 19 30 45.19 93.94 11 6.4 2 90 179 Ba

1977 01 31 14 26 40.11 70.86 12 5.8 81 40 100 N

1977 12 18 16 47 39.86 77.30 7 5.8 74 51 79 Fa

1979 03 29 02 21 41.95 83.38 13 5.5 71 53 86 N

1979 09 25 13 05 45.09 76.96 40 5.4 77 44 119 N

1982 05 06 15 42 40.15 71.54 20 5.5 70 36 95 N

1983 02 13 01 40 39.99 75.25 8 6.3 319 87 −178 Fa

1983 04 05 06 50 39.96 75.30 10 5.9 320 88 −176 Fa

1985 08 23 12 41 39.49 75.27 16 6.9 66 59 56 Fa

1987 09 18 21 58 47.23 89.53 14 5.2 147 83 173 Ba

1987 12 12 00 16 41.32 89.66 30 5.3 195 50 115 Ba

1988 01 06 15 31 39.63 75.53 37 5.3 92 63 102 Fa

1988 07 23 07 38 48.74 90.53 16 5.7 328 75 163 Ba

1990 06 14 12 47 47.87 85.08 34 6.4 116 89 −157 ChK

1990 08 03 09 15 47.96 84.96 33 5.9 119 36 178 ChK

1990 11 12 12 28 42.98 78.03 18 6.2 204 80 11 TS

1991 04 18 09 18 37.47 68.28 5 5.4 336 59 85 TS

1991 08 19 06 05 46.99 85.34 42 5.3 275 55 81 TS

1991 09 07 17 30 42.57 74.82 10 3.5 118 29 130 Gh†

1991 10 31 02 29 40.15 72.84 31 5.1 99 20 119 Gh†

1991 12 27 09 09 51.08 98.17 13 6.3 244 72 −15 Ba

1992 05 15 08 08 41.02 72.43 7 5.9 288 34 123 Gh†

1992 06 25 09 41 40.58 78.51 10 4.2 264 34 131 Gh†

1992 07 21 03 20 42.99 77.09 13 4.1 108 22 123 Gh†

1992 08 19 10 17 42.15 73.19 5 5.1 120 70 140 Gh7†

1992 08 19 13 44 42.17 73.36 16 4.2 0 35 −100 Gh7†

1992 08 19 14 17 42.17 73.43 15 4.7 145 70 170 Gh7†

1992 08 19 22 45 42.18 73.20 4 4.8 125 85 160 Gh7†

1992 08 19 02 04 42.10 73.30 14 7.0 221 46 42 Me

1992 08 19 08 51 42.14 73.18 17 4.1 120 70 −80 Gh7†

1992 08 20 12 21 42.13 73.33 7 4.3 100 45 130 Gh7†




Table 2. (Continued)




1992 08 20 16 30 42.10 73.67 16 4.3 95 65 110 Gh7†

1992 08 20 02 46 42.16 73.59 8 4.6 275 90 170 Gh7†

1992 08 21 04 14 42.17 73.56 4 4.7 95 45 130 Gh7†

1992 08 22 08 52 42.19 73.57 15 4.3 105 70 −160 Gh7†

1992 08 23 20 35 42.20 73.53 18 4.5 85 50 120 Gh7†

1992 08 23 09 04 42.17 73.57 10 4.9 285 80 140 Gh7†

1992 08 26 07 40 42.10 73.42 8 4.8 20 85 60 Gh7†

1992 11 06 07 21 41.05 72.51 10 4.9 320 35 145 Gh†

1992 12 24 05 09 42.22 72.23 8 5.1 104 39 77 Gh†

1993 02 02 16 05 42.20 86.14 32 5.0 90 35 90 TS*

1993 03 17 10 15 41.07 72.05 19 4.4 245 46 63 Gh†

1993 04 13 17 56 41.19 75.72 17 4.5 264 53 100 Gh†

1993 09 20 00 45 42.57 76.05 21 3.8 232 86 40 Gh†

1993 12 30 14 24 44.76 78.81 15 5.4 284 52 117 TS

1994 01 20 23 12 42.14 70.30 16 4.2 34 53 76 Gh†

1994 05 01 12 00 36.94 67.14 23 6.0 102 32 105 TS

1994 05 09 09 14 49.27 78.86 5 5.0 68 45 102 TS*

1994 08 23 14 18 40.21 78.77 5 5.0 85 35 134 TS*

1994 09 25 22 05 41.09 78.57 9 3.8 295 20 112 Gh†

1994 10 18 00 59 40.95 79.90 12 4.0 204 39 80 Gh†

1994 11 24 22 24 42.25 71.06 12 3.9 241 56 53 Gh†

1995 02 20 08 07 41.07 72.45 22 4.9 258 51 104 Gh†

1995 03 18 18 02 42.43 87.20 19 5.0 73 34 89 TS*

1995 03 29 15 52 41.83 79.45 8 4.0 268 40 65 Gh†

1995 05 02 11 48 43.80 84.69 21 5.4 99 72 171 TS

1995 06 22 01 01 50.35 89.96 10 5.4 89 31 82 Ba

1995 08 19 20 28 42.22 70.52 9 4.7 95 51 140 Gh†

1995 09 26 04 39 41.77 81.55 10 5.2 246 38 77 TS*

1995 10 08 08 55 41.05 72.15 18 5.6 237 40 59 Gh†

1995 11 02 23 16 42.16 75.08 16 4.0 116 33 140 Gh†

1995 11 03 11 09 40.02 73.70 20 4.4 71 22 146 Gh†

1995 11 04 18 35 42.22 75.05 15 4.0 68 45 75 Gh†

1996 01 14 18 16 42.69 74.88 5 4.2 102 23 106 Gh†

1996 01 18 09 33 41.81 77.50 16 4.8 262 36 101 Gh†

1996 03 19 15 00 40.04 76.62 34 6.0 234 16 87 TS

1996 03 20 00 14 40.09 76.77 6 4.5 268 20 76 Gh†

1996 03 22 08 26 40.00 76.75 6 5.2 260 18 78 Gh†

1996 04 02 02 28 41.23 77.65 16 4.1 242 59 128 Gh†

1996 05 14 12 45 41.54 75.03 10 4.4 222 48 69 Gh†

1996 06 01 17 08 43.10 74.94 13 3.6 70 50 70 Gh†

1996 06 01 18 03 41.44 76.02 22 4.2 234 23 69 Gh†

1996 06 14 22 45 42.48 72.93 4 4.9 348 61 106 Gh†

1996 12 28 07 40 42.95 78.08 12 4.4 204 57 49 Gh†

1997 01 21 01 48 39.54 76.88 12 5.4 317 85 177 TS








1997 01 09 13 43 41.03 74.28 15 5.7 284 28 107 Gh†

1997 03 01 06 04 39.45 76.82 14 5.6 180 80 −173 TS*

1997 04 11 05 34 39.54 76.89 20 6.0 226 42 −79 TS

1997 04 12 21 09 39.48 76.86 16 5.1 239 27 −74 TS

1997 04 15 18 19 39.58 76.93 18 5.7 177 64 −139 TS

1997 04 05 23 36 39.53 76.83 18 5.4 68 78 −21 TS

1997 04 06 12 59 39.52 76.90 13 5.1 210 38 −74 TS

1997 04 06 04 36 39.50 76.95 17 5.8 246 41 −51 TS

1998 03 19 13 51 40.05 76.68 15 5.4 243 5 79 TS*

1998 05 28 21 11 37.41 78.84 14 5.6 350 1 −32 TS

1998 05 29 22 49 41.17 75.59 18 5.3 238 42 62 TS*

1998 06 25 06 39 41.48 80.16 17 5.1 262 27 129 TS*

1998 07 28 04 51 41.77 81.51 10 5.5 243 40 79 TS

1998 08 02 04 40 39.55 77.00 15 5.5 173 40 −140 TS

1998 08 03 15 15 39.57 76.99 29 4.6 253 10 129 TS†

1998 08 27 09 03 39.57 77.30 15 6.3 57 80 1 TS

1998 09 03 06 43 39.59 77.22 10 4.8 179 59 178 TS†

1998 10 31 16 09 39.53 77.58 4.6 14 152 74 −164 TS†

1999 01 30 03 51 41.65 88.50 18 5.4 288 48 100 TS

1999 03 15 10 42 41.77 82.69 13 5.3 208 37 22 TS

2002 03 25 14 56 36.08 69.24 5 6.0 194 43 101 TS

2002 03 27 08 52 36.07 69.27 7 5.5 25 34 124 TS

2002 04 12 16 26 35.99 69.21 4 5.8 209 47 109 TS

2003 02 13 17 34 43.90 85.92 19 5.2 94 26 84 Ba

2003 02 24 02 03 39.52 77.20 5 6.2 280 17 115 TSd

2003 02 24 02 03 39.52 77.40 7 6.2 300 30 134 TSd

2003 03 12 04 47 39.38 77.41 7 5.0 245 33 73 TS*

2003 03 30 23 15 39.44 77.30 10 5.2 287 27 117 TS

2003 05 22 18 11 42.95 72.76 7 5.5 70 42 34 TS

2003 05 05 06 34 40.93 72.46 13 5.0 210 35 81 TS*

2003 06 04 16 28 39.35 77.56 10 5.2 90 36 87 TS*

2003 12 01 01 38 42.88 80.55 17 6.5 259 25 100 TS

2004 10 27 09 23 45.02 80.16 16 5.1 303 51 146 TS*

2005 02 14 23 38 41.68 79.36 25 6.0 279 13 109 TS

2005 04 06 08 44 41.35 78.70 20 5.1 288 41 107 TS*

2005 09 26 16 41 38.54 69.94 9 5.0 212 40 −103 TS*

2006 11 23 11 04 44.20 83.50 20 5.3 231 39 52 TS*

2007 07 20 10 06 42.91 82.35 14 5.3 316 37 149 TS

2007 12 26 04 45 40.34 73.06 16 5.1 273 43 72 TS*

2008 01 01 06 32 40.34 72.90 13 5.5 263 40 75 TS

2009 04 19 04 08 41.28 78.21 13 5.2 127 90 177 TS*




Table 3. The Baikal Rift Zone. Earthquake source parameters from seismology. Epicentres and origin times between 1964–2007 are taken from the updatedcatalogue of Engdahl et al. (1998). References from which source parameters have been taken are indicated by H (Huang & Chen 1986), E (Ekstrom & England1989), Da (Doser 1991a), DB (Doser 1991b), BJ (Bayasgalan & Jackson 2002), De (Delouis et al. 2002), BN (Brazier & Nyblade 2003), Ba (Bayasgalan et al.2005a), Em (Emmerson et al. 2006) and TS (this study). Earthquakes marked with an asterisk have had their depths determined by depth phase modeling (seeSection 2.1.2), and their mechanisms are taken from the Global CMT catalogue. Source parameters of the event marked E (Ekstrom & England 1989) weredetermined by a much broader-band CMT-type inversion. The source parameters of the remaining events were obtained through the inversion of teleseismicbody waveforms. Focal mechanisms are plotted in Fig. 6b and the final column of this table with the compressional quadrants shaded grey (depth < 20 km) orred (depth > 20 km).

The Baikal Rift Zone: Waveform Modelling Results



1917 04 29 11 55 56.17 114.62 16 6.6 340 70 −16 Db

1950 04 04 18 44 51.70 101.00 14 6.9 100 75 0 Da,De

1957 06 27 00 09 56.39 116.39 10 7.1 100 81 −25 Db

1958 01 05 11 30 56.51 121.11 8 5.8 257 50 −100 Db

1958 09 14 14 21 56.61 121.00 6 6.2 63 63 −75 Db

1959 08 29 17 03 52.64 106.90 14 6.2 248 53 −50 Da

1962 11 11 11 31 55.84 113 .22 5 5.7 215 58 −78 Db

1967 01 18 05 34 56.65 121.00 10 5.5 63 57 −117 Db

1967 01 20 01 57 48.23 103.14 8 6.4 319 42 102 H,BJ

1967 01 05 00 14 48.28 103.05 5 7.0 2 83 −179 H,BJ

1978 08 03 06 07 52.12 96.94 14 5.7 187 53 148 E

1989 04 20 22 59 57.17 122.10 29 6.2 114 71 45 Db,Em

1989 05 13 03 35 50.16 105.42 8 5.5 210 87 164 Ba

1989 05 17 05 04 57.08 122.13 28 5.8 13 62 −165 Db,Em

1989 10 25 20 29 57.53 118.88 28 5.3 68 40 −78 Em*

1991 09 12 00 33 54.90 111.14 22 5.0 235 25 −65 Em*

1991 12 27 09 09 51.07 98.17 13 6.4 244 72 −15 Ba

1992 02 14 08 18 53.95 108.91 15 5.3 249 33 −65 Em*

1994 04 26 18 59 56.74 117.97 14 5.4 81 24 −56 Em

1994 08 21 15 56 56.74 118.01 12 5.9 46 44 −99 Em

1995 06 29 23 02 51.91 103.19 16 5.7 73 41 −40 Ba,Em

1995 11 13 08 43 56.08 114.58 21 5.8 56 43 −59 Em*

1996 05 03 03 32 40.78 109.71 18 5.6 308 75 −12 TS*

1997 10 21 06 32 41.13 107.27 16 4.9 279 71 −21 TS*

1998 01 10 03 50 41.14 114.52 8 5.7 207 54 135 TS*

1998 09 24 18 53 46.26 106.34 27 5.5 97 76 9 Em

1999 03 21 16 16 55.93 110.31 3 5.7 267 22 −74 BN,Em

1999 03 21 16 17 55.99 110.29 3 5.8 222 20 −113 Em

1999 05 30 15 56 55.85 110.13 6 5.2 198 40 −108 Em

1999 09 08 02 38 57.49 120.25 6 5.1 266 46 −111 Em

1999 12 21 11 00 55.84 110.14 5 5.4 42 59 −97 Em

2003 09 16 11 24 56.06 111.37 15 5.5 38 40 −111 Em

2004 03 24 01 53 45.37 118.25 14 5.3 167 28 86 TS*

2005 04 27 07 36 51.25 98.25 11 5.3 342 79 −171 Em*

2005 07 20 18 06 43.04 109.28 8 5.0 13 41 125 TS*

2005 07 20 21 54 43.06 109.00 8 5.2 166 41 86 TS*

2005 11 10 19 29 57.44 120.50 7 5.8 96 52 77 Em

2005 12 11 15 54 57.46 120.76 7 5.6 265 45 −75 TS*

2006 04 30 00 43 44.52 102.38 6 5.7 353 46 141 TS*





The Baikal Rift Zone: Waveform Modelling Results



2006 06 15 23 04 45.34 97.41 10 5.0 318 36 113 TS*

2006 06 15 06 49 45.38 97.34 8 5.8 359 75 170 TS*

2006 12 04 09 13 55.78 110.09 10 5.2 210 35 −92 TS*

2007 07 04 01 23 55.46 110.36 5 5.3 198 26 −119 TS

2008 08 16 04 01 52.16 98.31 11 5.6 15 77 167 TS

2008 08 27 01 35 51.76 104.02 10 6.3 91 56 −47 TS

2009 01 26 21 30 57.48 120.67 8 5.4 261 39 −96 TS*

Table 4. NE Siberia. Earthquake source parameters from seismology. References from which source parameters have been taken are indicated by McM(McMullen 1985), Jem (Jemsek et al. 1986), RGL (Riegel 1994), Fr (Franke et al. 2000), Fuj (Fujita et al. 2009) and TS (This Study). Earthquakes marked withan asterisk have had their depths determined by depth phase modeling (see Section 2.1.2), and their mechanisms are taken from the Global CMT catalogue.Earthquakes marked by ‡ were not listed in the Global CMT Catalog and so the mechanism obtained by Franke et al. (2000) from first motions is used. Thesource parameters of the remaining events were obtained through the inversion of teleseismic body waveforms. Focal mechanisms are plotted in Fig. 7b andthe final column of this table with the compressional quadrants shaded grey (depth < 20 km) or red (depth > 20 km).

NE Siberia: Waveform Modelling Results



1962 04 19 23 16 69.76 138.88 12 6.0 120 40 100 RGL,Fuj

1964 07 21 09 56 72.06 130.16 11 5.4 338 50 -70 RGL,Fuj

1964 08 25 13 47 78.15 126.65 5 6.2 346 47 −89 Jem

1969 04 07 20 26 76.55 130.86 10 5.4 314 48 −106 Jem

1971 09 30 21 31 61.61 140.38 2 5.5 150 89 −1 RGL,Fuj

1980 02 01 17 30 73.06 122.61 25 5.4 315 55 −78 Fuj

1981 05 22 04 59 61.10 156.66 6 5.1 278 45 71 McM,Fuj

1981 11 08 21 56 61.82 153.71 14 5.6 127 36 20 Fr*

1983 06 10 02 13 75.50 122.69 22 5.6 144 74 −102 Jem

1984 11 22 13 52 68.49 140.87 12 5.4 341 45 158 Fr*

1988 03 21 23 26 77.59 125.49 25 6.0 178 34 −73 Fr*

1989 08 05 06 55 76.11 134.60 14 5.3 186 45 −90 Fr*

1989 10 03 23 09 80.62 121.71 12 5.2 134 56 −101 Fr*

1990 03 13 00 01 73.30 134.92 18 5.5 186 45 −90 Fr*

1990 11 02 21 05 64.95 146.70 24 4.6 90 66 −25 Fr*‡

1991 03 01 01 05 72.14 126.83 15 5.2 68 64 165 Fr*‡

1991 06 18 23 37 82.15 118.91 16 5.0 134 51 −125 Fr*

1992 02 15 04 06 75.93 125.10 15 5.0 128 61 −28 Fr*‡

1992 02 17 00 02 79.17 124.51 9 6.0 142 47 −99 TS

1992 06 08 09 30 81.25 121.08 9 5.1 1 51 −55 Fr*

1992 09 13 21 58 61.64 154.14 22 4.7 307 90 0 Fr*

1993 03 24 22 43 71.69 130.36 15 4.8 264 64 16 Fr*

1993 10 05 21 07 77.70 126.36 18 5.0 32 35 −65 Fr*

1996 06 22 16 47 75.79 134.68 14 5.7 357 58 −64 TS

1999 01 07 18 13 67.75 141.36 18 5.1 331 31 138 TS

2008 06 22a 23 56 67.71 141.43 18 5.8 344 60 131 TSd

2008 06 22b 23 56 67.71 141.43 17 5.9 336 52 121 TSd




Figure 2. Centroid depths, fault plane solutions, lithospheric thickness and topography of the Tibetan Plateau and its forelands. (a) Lithospheric thickness(from Priestley et al. 2006, see Fig. 1) and centroid depths. Earthquakes with centroid depths constrained by body wave modelling are shown as grey (0–20 km), red (20–50 km) or green (50–100 km) on both maps. See also depth distribution in Figs 3(a), (b) and (c). (b) Topography, centroid depths and faultplane solutions. Contours indicate the lithospheric thickness variation. SP and HK indicate the Shillong Plateau and Hindu Kush regions. QB and SB indicatethe positions of the Qaidam and Sichuan basins. Earthquakes marked A and B indicate normal faulting events which are discussed in the text. Earthquakesdeeper than 30 km in the Hindu Kush region, mainly south of 39◦N, have been excluded as they probably relate to subduction and are beyond the scope of thispaper.

Figure 3. Histograms of well determined centroid depths for earthquakes across Central and NE Asia. Blue bars show the range of Moho depths calculated ineach area. References are discussed in the text. (a) Indian plate south of the Himalaya. The entire crust appears to be seismogenic. (b) Interior of the Tibetanplateau. Other than some deep events near the Moho (see also Fig. 2) the majority of the plateau has a seismogenic thickness of <12 km with some deeper(∼20 km) events in the Qaidam Basin. The midcrust is aseismic. (c) Thrust events, usually with low angle dips (see Fig. 2b), that follow the range front of theHimalaya. Most are associated with the interface between the Indian Plate and its overriding crust. (d) Tien Shan and its forelands. Comparison with Fig. 4shows that all lower crustal earthquakes are in the forelands. (e) Baikal Rift Zone. Comparison with Fig. 6 shows that all earthquakes deeper than 20 km (exceptfor one) are restricted to the NE Rift. (f) NE Siberia. Comparison with Fig. 7 shows that earthquakes deeper than 20 km in the Laptev Sea region are restrictedto the westernmost edge of seismicity.




(1) They occur in two specific locations, both of them up to about300 km north of the southern margin of the Tibetan Plateau. Theyare not general features of seismicity beneath the plateau. North ofthis 300 km limit, the interior of the plateau is aseismic beneathabout 12 km.

(2) The northern limit of the SE group corresponds with a changein the SKS anisotropy that is thought to indicate the northern limitof the underthrusting Indian shield (Chen & Ozalabey 1998; Huanget al. 2000).

(3) If they occur in Indian material thrust beneath the thick andhot crust of the Tibetan Plateau they do not represent any steady-state or generic behaviour of the continental lithosphere.

(4) Their very close proximity to the Moho may be related to aninfiltration of hydrous fluids and brittle fracture that accompaniesa transformation of metastable granulite to eclogite (see Austrheimet al. 1997; Jackson et al. 2004).

We examined teleseismic records from a number of previously un-modelled events within 400 km of the Himalayan range front andbetween these two clusters of known deep events; but all the eventswe studied had centroids shallower than 12 km. The geographicalrestriction of these unusually deep events could simply be an arte-fact of the relatively short period of time for which the instrumentalrecord is available or, alternatively, it may linked to the geometry ofthe Indian shield thrust beneath the Tibetan Plateau with the clus-ters of earthquakes occurring within projections or salients in theleading edge of the Indian Plate.

3.2 Tien Shan and its forelands

3.2.1 Within the Tien Shan

Teleseismic determinations of centroid depths for events in the TienShan region are listed in Table 1 and shown in Figs 3(d) and 4. Inthe interior of the range all the events studied have centroid depthsof less than 25 km, and receiver function studies have shown thatthe crustal thickness varies from 50 to 70 km (Vinnik et al. 2004;Oreshin et al. 2002), with locally thinner crust of 40–50 km beneaththe Naryn Basin. Thus all the teleseismically studied events studiedwithin the Tien Shan are restricted to the upper crust.

Local seismic network studies confirm the above conclusion, andadd much detail. Ghose et al. (1998b) used the regional KyrgyzstanBroadband Network (KNET) to relocate ∼440 events which oc-curred over 6 yr in the rectangle marked ‘Gh’ in Fig. 4, concludingthat that earthquakes in the foreland occurred at greater depths thanthose within the range (∼18 km compared to ∼15 km). The cross-sections in Ghose et al. (1998b) show that events generally extendto a maximum of ∼22 km in the forelands and ∼18 km beneaththe range, however a small number of isolated events (∼3) occur at∼25 km beneath the range. Those events are not discussed in theirtext, and no indication of the likely errors is given so it is difficultto assess their reliability.

A regional moment-tensor inversion analysis by Ghose et al.(1998a) using KNET data from 32 events with magnitudes 3.5–6.0across the western half of the Tien Shan found that all the eventswithin the range had centroid depths less than 25 km. The deepestevent had a centroid depth of 31 km, but was at the southern edgeof the Ferghana Valley. Roecker et al. (1993) used a regional arraywith 267 stations in the western and central Tien Shan and northernPamirs to locate 472 local events, mainly in the northern Tien Shan,finding that virtually all of them were shallower than 30 km. Cross-sections show one event at ∼33 km within the range and two events

at 30–35 km within the Ferghana Basin. The error in depth in thisstudy was estimated to be up to 10 km. Xu et al. (2006) usedKNET and GHENGIS, a temporary regional broad-band seismicnetwork, to accurately relocate 480 local events (magnitude up to∼3) over the period 1997–1998, with estimated depth errors of 3–4 km. In the interior of the range west of 75◦E they found all buttwo earthquakes were shallower than 25 km. These two events wereon either side of the Naryn Basin at 31–36 km. In the eastern-central Tien Shan seismicity extended a little deeper, to ∼35 km.Two especially deep events were located near 41.5◦N 78◦E (south ofLake Issyk Kul) at about 45 km depth, in an area where the crustalthickness is about 60 km (Vinnik et al. 2004). Finally, Mellors et al.(1997) relocated more than 1000 locally recorded aftershocks of theMw = 7.0 Suusamyr earthquake of 1992 (marked Su on Fig. 4),finding them all to be shallower than about 18 km.

3.2.2 The Tien Shan forelands

By contrast, events in the Tien Shan forelands (including the TarimBasin, Junggar Basin and the Kazakh Platform) have teleseismicallydetermined centroid depths as deep as 44 km (Figs 3d and 4). WhilstMoho depth is less well constrained in the forelands most receiver-function studies suggest a crustal thickness of 40–50 km ((Oreshinet al. 2002; Vinnik et al. 2004; Wittlinger et al. 2004). Some areasof the Tarim Basin immediately adjacent to the high topographyhave crustal thicknesses of up to 55 km (Vinnik et al. 2004). Theforelands of the Tien Shan are demonstrably capable of generatinglower crustal seismicity.

Again, local network data confirm this conclusion. The KNET-GHENGIS study by Xu et al. (2006) extended into the Tarim Basinand Kazakh Platform. They found 21 earthquakes in the Tarim basinat 30–46 km depth (Fig. 5) and one outlier of unknown quality at∼64 km. Xu et al. (2006) used probability density function analysisto show that the depth errors of three selected events, all at depthsgreater than 20 km, were less than ±4 km. However, these eventswere within the network whilst the events in the Tarim Basin aresituated just beyond its edge. This suggests that the entire crust islikely to be seismogenic in this area as receiver functions show thatMoho depth is ∼47 km to the ENE (station XIKR on Fig. 5 Vinniket al. 2004) and ∼50 to the SE (station YAR in Fig. 5 Wittlingeret al. 2004). In the Kazakh Platform to the north, relatively few(∼20) events were recorded, five of which were in the depth range30–35 km.

The earthquakes in the NW Tarim basin near Kashgar (Fig. 5),including the 1997 Jiashi sequence, are of special interest. Be-tween 1997 and 1998, 12 earthquakes of Mw 5.0–6.3 occurred, ofwhich 10 had depths that could be confirmed by teleseismic bodywave analysis. In addition, local earthquakes between 1997 and1998 were recorded by Xu et al. (2006) in this same area (alsoshown in Fig. 5), who determined hypocentres from arrival timesrecorded by the GHENGIS/KNET networks, with probable errorsof 4 km for the events deeper than 25 km (errors for shallowerevents are not discussed in the text). We used the local and region-ally recorded waveforms to estimate focal mechanisms for someof these earthquakes using the FKCMTINV program. The largestevents (Mw 5.7, 5.9, 6.0 and 6.3) involved normal faulting witha roughly NNE–SSW strike, or strike-slip events with a similarT-axis azimuth of WNW–ESE, all with upper crustal depths of 14–20 km. In addition, a number of mechanisms involving thrusting areidentifiable, five through teleseismic body wave analysis and onethrough local waveform modelling. Four of these thrusts are found




Figure 4. Centroid depths, fault plane solutions, lithospheric thickness and topography of the Tien Shan and its forelands. (a) Lithospheric thickness (fromPriestley et al. 2006, see Fig. 1) and centroid depths. Much of the area has thick lithosphere, but the details not believable with the current resolution (200–400 km) of the surface wave tomography. Earthquakes with centroid depths constrained by body wave modelling are shown as grey (0–15 km), green(15–30 km) or red (30–50 km) on both maps. (b) Topography, centroid depths and fault plane solutions. Contours indicate the lithospheric thickness variation.NB, IK and FV indicate the Naryn Basin, Lake Issyk Kul and the Fergana Valley respectively. Su indicates the area where Mellors et al. (1997) studied theaftershocks of the Suusamyr earthquake. The box marked Gh denotes the area covered by the local earthquake survey of Ghose et al. (1998b).

at relatively shallow depths (7–15 km) and are positioned to theeast of the normal faulting earthquakes. This area is directly to thesouth of the thin-skinned deformation of the Kepingtage fold-and-thrust belt (Allen et al. 1999) and these events probably representthe continuation of this deformation south into the Tarim Basin.

The obvious interpretation of the normal faulting earthquakesin Fig. 5(b) involves flexure of the Tarim basin crust as it is over-thrust by the adjacent mountains, by analogy with the situation inthe Himalayan foreland (Section 3.1.1); but the pattern here is notso simple. First, there is no clear separation of the mechanisms intoshallow normal faults and deeper thrusts (Fig. 5b, inset), and in factthe thrusts are generally shallower than the normal faults. Howeverthese shallow thrusts are spatially separated from the normal fault-ing events, lying further east, and may be related to the thin-skinneddeformation in the Kepingtage fold-and-thrust belt (described byAllen et al. 1999) propagating further into the Tarim Basin. Onedeeper thrust earthquake occurred on the 1993 March 19 (Table 2),

at 34 km depth, and is 50 km north of the normal faulting, beneaththe main range front (Fig. 5b and inset). (This event has unusu-ally complicated waveforms, and an alternative interpretation ofits waveforms as a multiple shallower event can not be completelyruled out.) A cross-section through the local seismicity (Fig. 5a, in-set) recorded by Xu et al. (2006) reveals the apparent concentrationof earthquakes into two crustal depth ranges of 0–20 and 30–45 km,as would be expected for extension and shortening either side of abending neutral fibre; but there are no focal mechanisms deeper than20 km to check this possibility. Second, the strikes of the normaland thrust faulting mechanisms are not parallel, as they are in theHimalayan foreland. However the geometry in this western cornerof the Tarim basin is much more 3-D than in the Himalayan forelandand, since the normal faulting is nearly equidistant from the TienShan to the north and the Pamir to the west, it is not clear in whichdirection the Tarim basin would be flexed by its surrounding loads atthis point (or whether there is a dominant direction at all). Bending




Figure 5. Local and Regional Seismicity in the Kashgar Area. (a) Locationand depth of local seismicity from Xu et al. (2006). Upper crustal earth-quakes are shown as grey circles and lower crustal earthquakes are shown inred. Inset shows a north-south cross-section (projected onto a line betweenpoints A and B) of events within the dashed box. There is a gap in seismic-ity between the upper and lower crustal events. Isoseismals from the 1902Kashgar earthquake are taken from Gu et al. (1989). (b) Enlarged versionof the dashed box in (a). Mechanisms in red are determined by teleseis-mic waveform modelling, mechanisms in grey are taken from the GlobalCMT catalogue and mechanisms in blue are determined using FKCMTINV,using data from GENGHIS, a regional network in the W Tien Shan. Insethistogram shows the teleseismically determined depth distribution of earth-quakes coloured by mechanism type. Thrusts are generally shallower thanthe normal and strike-slip faulting events. One thrust to the NW, identifiedon the map, is at ∼34 km and is discussed in the text (Section 4.2.2).

to the west, rather than to the north, would explain the orientation ofthe normal faults, and is consistent with the rapid westward deep-ening of Tertiary sedimentary isopaches in this region (Allen et al.1999), but does not explain the orientation of the adjacent thrustingevents (which may be unrelated to bending anyway). A final point toconsider is time-dependence: on the outer rises of oceanic trenches,large normal-faulting earthquakes, related to bending, typically oc-cur after major earthquakes on the subduction interface thrust, ina way expected by simple loading considerations (e.g. Christensen& Ruff 1988). It may be no accident that this 1997–1998 forelandsequence at Jiashi is directly in front of the epicentral region of theM 8.6 1902 Kashgar earthquake (Fig. 5a), usually assumed to be athrust at the mountain front (Molnar & Deng 1984). In spite of the

complicated geometrical configuration of the western Tarim basinand its surrounding loads, and of the earthquake mechanisms anddepths, the normal faulting and the apparent concentration of earth-quakes into two depth bands perhaps favours a flexural interpretationfor at least some of this foreland seismicity. Other interpretationsare, however, possible: for example, Allen et al. (1999) try to relatethe normal-faulting earthquakes to activity on postulated splays orextensions of other structures in the Tarim or the adjacent frontalfold-and-thrust belt of the Tien Shan.

3.3 Patterns: seismicity and lithosphere thicknessin shortening regions

Figs 2(a) and 4(a) show how the earthquake depths beneath theTibetan Plateau and the Tien Shan are related to the lithospherethickness variations determined by Priestley et al. (2006). The wholeregion is characterized by lithosphere that is thicker than in mostcontinental regions (∼125 km or less), and extreme thicknesses areobserved beneath the Tibetan Plateau (>250 km) and in the area ofthe Tien Shan (>200 km). The horizontal resolution of the methodsused by Priestley et al. (2006) is however ∼200–400 km, greaterthan the width of the Tien Shan, and so there may be unresolvedstructural features in the mantle lithosphere that correlate with therange. In particular a number of regional studies have used S-to-Preceiver function analysis (e.g. Oreshin et al. 2002; Vinnik et al.2004; Kumar et al. 2005) or traveltime tomography (e.g. Roeckeret al. 1993) to suggest that there is an area of thin, ∼100 km, litho-sphere beneath the central Tien Shan. These regional methods havea higher horizontal resolution than those of Priestley et al. (2006)and so their results are not necessarily contradictory, however cau-tion is required. It is controversial whether receiver function studiescan image the lithosphere–asthenosphere transition (e.g. Rychertet al. 2005; McKenzie & Priestley 2008) rather than boundarieswithin the lithosphere (Yuan & Romanowicz 2010) and body-wavetomography has limited vertical resolution (Priestley & McKenzie2002).

As McKenzie & Priestley (2008) point out, the very large thick-nesses are likely to be caused by shortening of mantle lithospherethat was previously depleted in garnet and iron through melting andsufficiently buoyant to resist instability and delamination; thoughthe manner in which the thickening occurred (whether by wholesaleunderplating of one region by another, or by uniform shortening)is unclear. In Tibet, the thickest lithosphere is under the highesttopography, where the crust is also thickest, at more than 60 km.However, the deepest earthquakes in these regions are not, on thewhole, under the highest topography, but close to the edges of theTibetan plateau and the Tien Shan, either in the adjacent flat fore-lands (red in Figs 2 and 4; green in Fig. 5a) or close to the edge ofthe Tibetan plateau (green in Fig. 2), where they can reasonably beassociated with the underthrust Indian shield. In the interior of thehigh mountains or plateau, earthquakes are generally restricted tothe upper half of the crust (<30 km in the Tien Shan and <20 kmin Tibet).

This pattern is an important advance on the earlier generaliza-tions made by Maggi et al. (2000a), Jackson (2002) and Jacksonet al. (2008). Those papers recognized that earthquakes occurredthroughout the crust in the Tien Shan region, but did not distinguishthe interior of the range from its forelands, as is done in Fig. 3(d).The distinction is significant. The foreland of the Himalaya and thenorthern foreland of the Tien Shan are known to be stronger thanmost continental areas, from their unusually large elastic thickness




estimates (McKenzie & Fairhead 1997; Maggi et al. 2000a; Jackson2002). Both forelands are associated with lower crustal earthquakes.The obvious inference is that earthquakes in the lower crust of theforelands are proxies for strength, and that the other forelands ofhigh mountains that have such earthquakes, such as the Junggar,Fergana and Tarim basins, are also strong, though they are too smallfor this to be confirmed independently by gravity analysis. The im-plication, in turn, is that the steep mountain fronts are maintained byviscous gravity currents as the hot weak middle crust of the moun-tains flows over the rigid lower crust of the forelands, as suggestedfor the Himalaya by Copley & McKenzie (2007). The importantconclusion is that the high mountains are supported, not by thestrength of the adjacent mantle lithosphere, as would be necessaryin a laterally homogeneous rheology, but by the rigid lower crustof the foreland that is thrust beneath them. From the distribution oflower crustal earthquakes in Figs 2, 4 and 5, this conclusion appliesto many of the range fronts in central Asia.

This behaviour is not difficult to explain. The forelands of Tibetand the Tien Shan are associated with lithosphere that is thicker thannormal, but not as thick as it is beneath the Tibetan Plateau, whereit has been thickened further, presumably by tectonic shortening.For reasons summarized earlier, and discussed by Jackson et al.(2008), the lower crustal earthquakes in those forelands are likelyto occur in dry material with a granulite-facies mineral assemblage,which will be strong and seismogenic at temperatures up to about600 ◦C. Within the mountains, tectonic shortening and thickening ofthe crust, combined with internal heating, will eventually producemuch higher temperatures in the middle and lower crust, whateverits composition (McKenzie & Priestley 2008), and it will becomeaseismic. In this way the underthrusting material of the forelandwill lose its strength and its seismicity within the range, though thismay take tens of millions of years to happen (McKenzie & Priestley2008).

4 E X T E N S I O N

4.1 The Baikal rift zone

The Baikal rift zone is an important region of active extensionstretching from northern Mongolia NE through Lake Baikal to theMuyu region (Fig. 6). It lies between the Amurian Palaeozoic foldand thrust belt to the south and the Siberian Shield to the north.

The 4 mm yr−1 of extension across the rift observed throughGPS (Calais et al. 2003) may result from the far-field effects ofthe Indian Eurasian collision (Molnar & Tapponnier 1975), the up-welling of a mantle plume beneath the rift (e.g. Gao et al. 1994a,b)or a combination of the two (Petit et al. 1998). The rift borders theSiberian Shield, and the underlying lithospheric structure of thisarea, and its relation with rifting has been debated in recent pa-pers (Emmerson et al. 2006; Petit & Deverchere 2006; Petit et al.2008). Seismicity can provide important insights to these questionsand there are numerous studies of earthquakes observed both lo-cally (Vertlib 1981; Deverchere et al. 1991; Deverchere et al. 1993;Gileva et al. 2000; Deverchere et al. 2001; Melnikova et al. 2007;Suvorov & Tubanov 2008) and teleseismically (Huang & Chen1986; Bayasgalan & Jackson 2002; Bayasgalan et al. 2005a; Doser1991a,b; Delouis et al. 2002; Brazier & Nyblade 2003; Emmersonet al. 2006). In particular, the reported presence of lower crustalseismicity in many of these studies may provide insights into thevariation in rheology along the rift zone. Gao et al. (2004) andEmmerson et al. (2006) have presented receiver functions demon-strating that crustal thickness in the area varies from ∼35 km be-

neath the Lake Baikal Rift to ∼45 km to the south beneath theAmurian fold and thrust belt. Ten Brink & Taylor (2001) analysedseismic refraction profiles within Lake Baikal and concluded thatthe crustal thickness there was ∼40 km.

Numerous studies have examined the geomorphological expres-sion of faulting in the Baikal rift zone, from Florensov (1969) andSherman (1978), who described the overall patterns observed in therift system, to more recent studies such as Sankov et al. (2000) inthe NE rift and Arjannikova et al. (2004), who concentrated on thefaulting to the west of Lake Baikal.

Of particular interest in this paper, is that the zone is close to theedge of the surface outcrop of the Siberian shield, where it abutsthe Sayan-Baikal fold belt, and also close to the edge of a dramaticgradient in lithosphere thickness, changing from ∼125 km or lessin the SE to more than 200 km beneath the Siberian shield (Fig. 6).We refer to the area E of Lake Baikal as the NE Rift. The zonewhere the extensional faulting and seismicity are concentrated in aNNE–SSW band ∼300 km long that is roughly coincident with thelake, we refer to as the Lake Baikal Rift.

4.1.1 Seismicity

Teleseismic earthquake focal mechanisms and centroid depths, fromHuang & Chen (1986), Bayasgalan & Jackson (2002), Bayasgalanet al. (2005a), Doser (1991a), Doser (1991b), Delouis et al. (2002),Brazier & Nyblade (2003) and Emmerson et al. (2006) as well asthis study, are shown in Figs 3(e) and 6(b).

East of 105◦E, earthquakes with mb larger than 5.0 (shown aswhite dots on Figs 6a and b) are concentrated in a relatively narrowband 100–200 km wide through Lake Baikal. At the eastern endof the lake, the strike of this band of seismicity abruptly changesdirection where the active rift meets a rapid change in lithosphericthickness. The rifting does not continue NNE into the thick litho-sphere, but instead continues to the ENE following the edge of thesteep gradient in lithospheric thickness. Where the strike of the riftchanges, the width of the zone of moderate-to-high magnitude seis-micity increases to ∼300 km. This zone narrows again further eastand continues until 122◦E, beyond which there is no longer a clearnarrow rift zone defined by seismicity. Throughout this region thereis a sharp cut-off in seismicity to the north, with no moderate-to-large events occurring above the thick lithosphere, suggesting thatrifting is limited to the thinner material in the south.

Centroid depths also change where the rift meets and the followsthe edge of the thick lithosphere. In the Lake Baikal region to thewest, all accurately determined centroid depths for moderate-to-large magnitude events are less than 16 km, whereas in the NE Riftthere are five events with a centroid depth greater than 20 km, to amaximum of 29 km (Emmerson et al. 2006). This is significantlygreater than the more common depth range of 15–20 km found inmost continental rift systems (Chen & Molnar 1983; Jackson &White 1989; Maggi et al. 2000b). The three earthquakes with cen-troid depths of 28 km or more are all at the northern edge of theseismic zone, adjacent to the thick lithosphere. Another key obser-vation, also noted by Emmerson et al. (2006), is that the seismicityin the NE Rift does not follow the edge of the high topography,but instead cuts across the Vitim Embayment 300–400 km south ofthe range front. Priestley et al. (2006) and Emmerson et al. (2006)connected this line with a rapid lateral gradient in Vs, and Fig. 6shows we can now confirm that the high topography extends overthe area of thick lithosphere to the north, whereas the seismicitystops at its edge.




Figure 6. Centroid depths, fault plane solutions, lithospheric thickness and topography of the Baikal Rift Zone. (a) Lithospheric thickness (from Priestleyet al. 2006, see Fig. 1) and centroid depths. Earthquakes with centroid depths constrained by body wave modelling are shown as grey (0–20 km) or red (20–30 km) circles. The thick dashed line indicates the approximate geologically mapped boundary of the Siberian craton after Goodwin (1991). The edge of thethick lithosphere differs significantly from the geologically mapped boundary of the Siberian Shield. (b) Topography, centroid depths and fault plane solutions.Contours indicate the lithospheric thickness variation. White circles show all seismicity listed in the EHB catalogue. White boxes display the centroid depthsof earthquakes which are deeper than 20 km. Earthquakes in the lower crust are limited to the NE Rift. VE marks the Vitim Embayment. V and T indicate theVitim and Tariat and u and o on the inset location map indicate Udachnaya and Obnazhennaya. M is the Muya region.

A number of groups have studied local and regional seismicityin the Baikal rift zone and, whilst some features are still debated,a similar pattern to that observed teleseismically is now emerging.This evidence is, however, more difficult to evaluate because ofvariations in station density and limited discussion of errors.

The distribution of local events routinely located by Russianpermanent stations (shown, for example, in Petit et al. 1996) broadlymatches the observations above. A narrow band of concentrated

seismicity in Lake Baikal abruptly changes strike at the NE end ofthe lake, widening here as earthquakes follow a number of smallerrifts cutting across this corner, before narrowing again by 115◦E.There is a sharp cut-off of seismicity to the north at the edge ofthe thick lithosphere, but scattered background seismicity in thefold-and-thrust belt to the south.

Other groups have located events in limited areas, allowing usto compare the values of seismogenic thickness they record with




the teleseismic estimates, and these studies have been recently re-viewed by Radziminovich (2010). Reported values of seismogenicthickness vary significantly between 30 and 60 km in the NE Riftand 25–37 km in Lake Baikal, and in some cases (e.g. Deverchereet al. 1991) earthquakes have been reported in the upper mantle. AsEmmerson et al. (2006) and Radziminovich (2010) point out, stud-ies locating earthquakes deeper than 30 km (the maximum depthof teleseismically confirmed earthquakes) all use sparse regionalnetworks (typically >100 km station spacing), which limits theaccuracy of focal depth determination (Vertlib 1981; Deverchereet al. 1991; Gileva et al. 2000; Deverchere et al. 2001; Melnikovaet al. 2007). Deverchere et al. (1993), using a denser local network(minimum station spacing ∼20 km) in the north Muya region (Min Fig. 6), found that the cut-off for seismicity was at 30 km, inagreement with teleseismic observations elsewhere in the NE Rift.

Emmerson et al. (2006) and Radziminovich (2010) concludethat, whilst focal depths as great as 46 km (Melnikova et al. 2007),60 km (Vertlib 1981) and 63 km (Deverchere et al. 1991) have beenreported in the NE Rift, the uncertainties involved mean that it isimpossible to be confident that locally observed seismicity extendsinto the upper mantle. All studies (both local and teleseismic) findearthquakes down to depths of at least ∼30 km in the NE Rift whichborders the thick lithosphere. This indicates that at least the lowercrust in this area is seismogenic, and that earthquakes occur to agreater depth than the 15–20 km range, seen in most continentalrifts (Chen & Molnar 1983; Jackson & White 1989; Maggi et al.2000b).

To the west, in the Lake Baikal Rift, Radziminovitch et al. (2005)and Suvorov & Tubanov (2008) found that the maximum hypocen-tral depths were at 25 km in south and central Lake Baikal. Suvorov& Tubanov (2008) also noted that only very rare unrepresentativeevents exceeded 22 km in depth. These values are shallower thanthat found by all local studies in the NE Rift, though still somewhatdeeper than the earthquakes in Lake Baikal studied teleseismically(16 km). By contrast, Gileva et al. (2000) and Deverchere et al.(2001) studied events across the entire rift system and concludedthat generally the seismogenic thickness was 35–40 km and was asmuch as 37 km in central Lake Baikal. Those studies, however, useda sparse regional network, and both Suvorov & Tubanov (2008) andRadziminovich (2010) suggest that the simple one layer velocitymodel assumed by Gileva et al. (2000) and Deverchere et al. (2001)in Lake Baikal leads to significantly greater depths than a morerealistic layered velocity model constrained by deep seismic sound-ing (Suvorov et al. 2002), which was used by Suvorov & Tubanov(2008).

In general then, whilst there is not absolute consensus, theselocal studies define a seismogenic thickness slightly larger than thatfound in teleseismic studies and, as in the teleseismic studies, thereare some indications that the seismogenic thickness is greater in theNE Rift (∼30–40 km) than in Central Lake Baikal (∼25 km).

4.1.2 Relations with structure

Emmerson et al. (2006) found that, in the Baikal rift zone as a whole,the teleseismically confirmed lower crustal earthquakes (shown inred on Fig. 6) are restricted to the NE Rift. The additional earth-quakes presented in this study support this conclusion, and Fig. 6(a)shows that the NE Rift closely borders the edge of the thick litho-sphere (∼200 km) which is characteristic of the bulk of the SiberianShield. In contrast the Lake Baikal Rift appears to cut through mate-rial underlain by thin (<125 km) lithosphere some distance from this

edge. Furthermore, the seismicity in the NE Rift does not follow thegeologically mapped boundary of the Siberian shield at the surface(shown as a dashed line in Fig. 6), but instead follows the steep gra-dient in lithospheric thickness ∼300 km to the south. In this area thegeologically mapped edge of the exposed shield follows the steeptopographic front of the Vitim Embayment. However if this frontis typical of the thin-skinned deformation that is often observed inthrust belts, it is unlikely that the extent to which the shield underliesthe Vitim Embayment can be determined through surface mapping.Emmerson et al. (2006) also suggest that the range-parallel nega-tive gravity anomaly found in the Siberian Shield to the north ofthe Vitim Embayment, together with the arcuate post-thrust faultswithin the Vitim Embayment, support this thin-skinned hypothesis.

The relationships between lithosphere structure and deformationin this area have also been explored by Petit et al. (2008). Theirinterpretation differs from ours because they delineate a differentboundary for the edge of the thick lithosphere. They use the modelof Moho depth and lithosphere–asthenosphere transition depth ob-tained by Petit & Deverchere (2006) from forward and inverse mod-els of Bouguer gravity anomalies. According to their analysis, theedge of the thick lithosphere lies more than 100 km south of centralLake Baikal, but more than 200 km to the north of the NE Rift,forming a sinusoidal edge similar in shape to the outcrop boundaryof shield rocks at the surface. They therefore conclude that in theLake Baikal region the rift cuts through thick lithosphere, whereasin the NE Rift the active rifting moves into thinner lithosphere.

There is, however, a fundamental problem with obtainingtwo subsurface boundaries (both Moho depth and lithosphere–asthenosphere transition depth) from gravity data, as the solutionis non-unique. Petit & Deverchere (2006) attempt to circumventthis by assuming the wavelength of signals from near-surface,crustal and lithospheric boundaries will increase as the depth ofthe interfaces increase. However, whilst a deep feature must havea long-wavelength signal when observed at the surface, a shallowlong-wavelength feature will also create a long-wavelength signal.Indeed it is notable that the variation in lithosphere thickness ob-tained from this method correlates closely with long-wavelengthtopographic features (such as the Vitim Embayment).

We regard the lithosphere thickness gradient obtained from thesurface wave tomography to be more robust than that obtained fromgravity. The association of the lower crustal earthquakes (red inFig. 6) with the colder material of the thick lithosphere is then obvi-ous. There is then no particular association between rift width andlithosphere thickness. The width of rift zone is relatively restrictedin both Lake Baikal and the eastern part of the NE Rift, and morediffuse in the central region where the strike of the zone changes.

4.2 Laptev Sea region

4.2.1 Geological setting

The Laptev Sea region (Fig. 7) represents a rare example of anocean rift propagating into a continental shelf, and has thereforebeen the focus of considerable interest. The rift in question is theultraslow-spreading Gakkel Ridge (Fig. 8), spreading at a rate only∼6 mm yr−1 off the Laptev Shelf (DeMets et al. 1994). As one ofthe slowest-spreading ocean rifts on the planet, its morphology hasbeen the focus of a number of recent studies (e.g. Cochran et al.2003; Dick et al. 2003). On the continental shelf, seismic reflectionand refraction lines (Kogan 1974; Vinogradov 1984; Ivanova et al.1990; Alekseev et al. 1992; Drachev et al. 1998, 2003) and gravity




Figure 7. Centroid depths, fault plane solutions, lithospheric thickness and topography of north eastern Siberia. (a) Lithospheric thickness (from Priestleyet al. 2006, see Fig. 1) and centroid depths. Earthquakes with centroid depths constrained by body wave modelling are shown as grey (0–20 km) or red (20–30 km) circles. The thick dashed line indicates the approximate geologically mapped boundary of the Siberian craton after Goodwin (1991). (b) Topography,centroid depths and fault plane solutions. White circles show all seismicity listed in the EHB catalogue. White boxes display the centroid depths of earthquakeswhich are deeper than 20 km. Light blue triangles mark the wander path of the pole of rotation between the North American and Eurasian plates (Gaina et al.2002, see text). Blue boxes give the age in Ma of the indicated pole. The red triangle marks the present-day pole of rotation between the Eurasian and NorthAmerican plates determined by GPS observations (Kogan & Steblov 2008). Dashed yellow lines labelled ULG indicate the approximate position of the Ust’Lena graben (after Drachev et al. 1998). LD indicates the location of the Lena river delta, ORB marks the Omoloi river basin and the Moma rift is labelledMR. The New Siberian Islands are indicated by NSI.

anomalies (Gaponenko et al. 1968; Laxon & McAdoo 1994, 1998)have been used to map the rifting as it splits into a series of sediment-filled graben. This information has also been supplemented by heat-flow studies (Drachev et al. 2003).

The continuation of the rift on-shore and the nature of the bound-ary between the North American and Eurasian plates have alsobeen explored in the literature over the last 40 yr. Grachev (1973)proposed that rifting continued to the south in the form of themapped Cenozoic Moma Rift. However, Cook et al. (1986) ob-served that, although seismicity between the two plates formed aband through the Chersky Mountains (including the Moma Rift),earthquake focal mechanisms south of 70◦N change to thrust andstrike-slip. They concluded that rifting had ceased in the MomaRift, and that the boundary was now in compression, probablydue to the northward migration of the pole of rotation betweenthe two plates. This hypothesis was supported by Gaina et al.(2002), who tracked the migration of the pole over time usinggravity data and magnetic anomalies in the North Atlantic andEurasian Basin (shown as blue triangles in Fig. 7b), and this agreeswith modern geodetic constraints on the pole which place it im-mediately to the south of the Lena River Delta, marked by thered triangle in Fig. 7(b) (Prawirodirdjo & Bock 2004; Kogan &Steblov 2008).

Mackey et al. (1998), following the method of Ruff et al. (1994),used regionally recorded Pg and Pn traveltimes to estimate crustalthickness at a number of seismic stations in the area. They foundcrustal thicknesses of ∼40 km in the Siberian Shield, 32–47 km inthe Chersky mountains, as low as ∼20 km near the Lena River delta.Within the Laptev Sea Avetisov & Guseva (1991) used convertedphases at closely spaced receivers to estimate a crustal thicknessof 29–31 km, agreeing with the estimate of 29–30 km reported byKogan (1974) from a refraction survey.

4.2.2 Seismicity

Fault-plane solutions and centroid depths from Jemsek et al. (1986);Franke et al. (2000); Fujita et al. (2009) and this study are shown inFig. 7. Lower crustal earthquakes, with depths greater than 20 km,are coloured red. North of 70◦N, seismicity forms a narrow bandon the ultraslow spreading Gakkel Ridge (Fig. 8), in contrast tothe more diffuse pattern on the continental shelf of the Laptev Sea.In this area deformation is extensional, with predominantly normaland some strike-slip focal mechanisms.

The westernmost limit of seismicity in the Laptev Sea closelyfollows the edge of the thick lithosphere, however, earthquakesoccur as far east as the New Siberian Islands. Lower crustal events




Figure 8. Bathymetry, lithospheric thickness and seismicity of the Eurasian Basin and its margins. (a) Lithospheric thickness (from Priestley et al. 2006, seeFig. 1) and seismicity. Earthquakes on both maps are shown as yellow circles and are taken from the updated EHB catalogue. (b)Bathymetry and seismicity.Contours show lithospheric thickness variation. Seismicity is concentrated in a narrow band along the Gakkel Ridge (the spreading centre of the EurasianBasin) until it reaches the continental shelf (Laptev Sea) where it spreads out. The boundary between thick (>200 km) and thin (<100 km) lithosphere closelyfollows the edge of the Eurasian continental shelf. If the Eurasian Basin were to be closed about the Gakkel Ridge, the Lomonosov ridge would approximatelymatch the shape of the Eurasian continental margin.

(down to 25 km) are only found in the western part of the area(Figs 3f and 7), adjacent to the thick lithosphere. These observationsare reminiscent of the NE Rift in the Baikal rift zone.

The remote nature of this region means that only limited studies oflocal seismicity have been undertaken. Two studies (Avetisov 1993;Kovachev et al. 1994) report the results of local-network campaignsnear the Lena River delta, both finding possible lower crustal andmantle earthquakes. In the land-based 1985–88 study of Avetisov(1993), most hypocentral depths were found to be shallower than 28km, although events were recorded with depths of up to 42 km, andone was reported at 55 km. The reliability of this deepest location isnot discussed, so it is difficult to assess. Kovachev et al. (1994), usinga network of 10 ocean-bottom seismometers (OBS) in the LaptevSea over one month in 1989, found most events were at depths ofless than 20 km. Two events, however, were found at 29 and 40 kmdepths immediately south of the OBS deployment. In addition, fourother events were reported at depths between 60 and 85 km, butas these were all more than 150 km from the nearest OBS, theirdepths are not likely to be reliable and Kovachev et al. (1994) notedthat they were highly dependent on the velocity structure used. Asshown in Fig. 7(a) the edge of the thick lithosphere lies close tothe Lena River delta and so lower-crustal earthquakes in this areawould match the teleseismically observed pattern. The complexcrustal velocity structure in the delta and the limited extent of thelocal networks mean that the existence of upper-mantle seismicityremains unproven.

4.2.3 Relations with structure

The patterns of seismicity described above are similar to thoseobserved in the Baikal rift zone in Section 4.1. Extensional lower-crustal seismicity in the Laptev Sea is restricted to areas adjacent to

the boundary of the thick lithosphere. Two strike-slip earthquakeswith centroid depths of 22 and 24 km were reported south of theMoma Rift (Fig. 7) by Franke et al. (2000), where crustal thicknessestimates vary from ∼37 to 40 km (Mackey et al. 1998) and aregreater than those of 20–30 km near the Laptev Sea (Kogan 1974;Avetisov & Guseva 1991; Mackey et al. 1998). Both these earth-quakes are small (Mw = 4.7 and 4.6, respectively), and Franke et al.(2000) acknowledge that the principal source of error is likely to bethe misidentification of depth phases, determined from teleseismicarray data.

Another similarity between this area and the Baikal Rift is thesuggestion that high topography emplaced upon the Siberian shieldbecame rifted away from the main part of a fold belt. The OmoloiRiver Basin (Omoloiskaya, ORB in Fig. 7) appears to be the onshoreextension of the Ust Lena Graben. Though it is not currently markedby significant seismic activity, it is filled with Miocene to Quater-nary sediments and is bounded by a normal fault on the easternside (Kropotkin & Titkov 1981). According to Ivanov & Belyayev(1973), the basin is the most southwestern manifestation of Ceno-zoic rifting in this area, with sediments up to 2 km thick. Its currentseismic inactivity is likely to be related to the recent northward mi-gration of the pole of rotation that has now stopped onshore rifting(discussed above in Section 4.2.1). This now inactive rift followsthe edge of the thick lithosphere and, were it to have continued, itwould have rifted the Verkhoyansk mountains away from the rest ofthe Mesozoic fold-and-thrust belt. A large range-parallel negativefree-air gravity anomaly in the Siberian shield to the west of theVerkhoyansk Range indicates a flexural foreland basin, suggestingthat the range is emplaced upon and loading the Siberian shield (seeMcKenzie & Fairhead (1997)). The configuration thus resemblesthat in the still-active NE Rift in the Baikal rift zone, which crossesthe Vitim Embayment (Fig. 6).




4.2.4 Lomonosov ridge and Eurasian basin

Whilst the origin of the Arctic Ocean as a whole remains contro-versial, the Gakkel Ridge is recognized as an ultra-slow spreadingmid-ocean ridge at the centre of the Eurasian Basin (reviewed byCochran et al. 2003; Dick et al. 2003), delineated by seismicityalong the active ridge axis (Fig. 8).

The Eurasian basin is flanked by two continental margins, theLomonosov Ridge and the Eurasian continental shelf (Kara Sea andBarents Sea). The Lomonosov ridge is an aseismic bathymetric ridgeof continental crust which forms the margin of both the Eurasian andAmerasian Basins. Cochran et al. (2003) note that the Lomonosovridge and the edge of the Eurasian continental shelf are equidistantfrom the Gakkel ridge, forming a conjugate pair. The shape of thetwo margins is also closely matched by the Gakkel Ridge itself.These observations strongly suggest that the Lomonosov Ridgerepresents a sliver of continental material that has been rifted offthe edge of the Eurasian continental shelf. Fig. 8 shows that the edgeof the thick lithosphere coincides closely with the current edge of theEurasian continental shelf, and that the Amerasian Basin does notappear to be underlain by thick lithosphere although the presenceof thick lithosphere with horizontal extent of less than 200–400km cannot be ruled out by the methods of Priestley et al. (2006).This suggests that when the Eurasian Basin began to open at 55 Ma(Vogt et al. 1979), rifting began at the edge of the thick lithospherebeneath the Eurasia continental shelf. This behaviour represents theconclusion of the process that is in its early stages in the Baikal andLaptev regions. In all these areas, it suggests that an abrupt edge inthe thick lithosphere marks a rheological transition that controls thelocation of rifting.

4.3 Patterns: seismicity and lithosphere thicknessin extending regions

Two main patterns stand out in the rifting discussed in the Baikal,Laptev Sea and Eurasian basins. The first is a tendency for theactive extension to follow the edge of the thick lithosphere mappedby Priestley et al. (2006). In some places it is localised along theedge, as in the NE Baikal rift (Fig. 6) and during the formation ofthe Eurasian Basin (Fig. 8), whereas in others the edge of the thicklithosphere limits a broader region of rifting, as in the Laptev Sea(Fig. 7). In the Lake Baikal region itself, the rift approaches theedge of the thick lithosphere at a steep angle, before turning andfollowing it (Fig. 6).

The second pattern is that in both the Baikal rift zone and theLaptev Sea teleseismically confirmed lower crustal earthquakes(shown in red on Figs 7 and 8) are restricted to locations wherethe rifting closely follows or abuts the edge of the thick lithosphere.Where rift systems are located away from the thick lithosphere(e.g. Lake Baikal itself and the E Laptev Sea) the deepest tele-seimically confirmed earthquake are in the range 15–20 km typicalof most regions of continental extension (Chen & Molnar 1983;Jackson & White 1989; Maggi et al. 2000b). To the extent thatthe quality of the locations can be assessed, the locally recordedseismicity appears roughly to conform to this pattern.

Emmerson et al. (2006) used the method of McKenzie et al.(2005) to calculate geotherms in the Siberian shield and the Palaeo-zoic Sayan-Baikal fold belt from mantle nodule data. At Vitim andTariat in the Sayan-Baikal fold belt they estimated Moho tempera-tures of 750 and 860 ◦C , whilst at Udachnaya and Obnazhennayain the Siberian shield they obtained lower temperatures of 580 and550 ◦C. Both Emmerson et al. (2006) and Radziminovich (2010)

suggested that lower crustal seismicity in the NE Rift of Baikalmay be due to a cooler crustal temperature structure in that area,related to its situation adjacent to the colder Siberian shield. Such acooling effect may produce a geotherm in the NE Rift intermediatebetween those determined by Emmerson et al. (2006). However itcannot produce a Moho temperature as low as ∼350 ◦C, the usualcut-off temperature for crustal seismicity (Chen & Molnar 1983;Scholz 2002), as the Moho temperatures in both areas are signif-icantly higher than this, even away from active extension. Mohotemperatures in the Laptev Sea are not constrained by nearby stud-ies on mantle nodule data, but, beneath the actively rifting area,estimates of Moho depth (20–30 km, from Kogan 1974; Avetisov &Guseva 1991; Mackey et al. 1998) are shallower than in the NE Riftand lithospheric thickness variations appear to be similar (Priestleyet al. 2006, Figs 5a and 6a) so the Moho temperatures are unlikelyto be significantly cooler. Thus in both the NE Rift and Laptev Sea athermal effect cannot solely explain the occurrence of lower crustalseismicity and there must be an additional compositional effect,probably related to dehydration (Jackson et al. 2008).

5 D I S C U S S I O N

5.1 Temperatures

It seems inescapable that the lower crustal seismicity close to theMoho in all the areas described here, as well as in other areasdiscussed by Maggi et al. (2000b), is taking place at temperatureswell above the ∼350 ◦C limit that normally represents the mid-crustal cut-off, and possibly as high as 500–600 ◦C. The most likelyexplanation is a loss of water from the lower crust, leaving a residualmaterial with a higher melting temperature (Jackson et al. 2004,2008). A probable candidate for that material, based on the studiesof Austrheim & Boundy (1994), is an anhydrous granulite-faciesmetamorphic assemblage, formed through partial melting and theremoval of the hydrous granitic melt. The granulitic residue canremain in a friction-dominated brittle regime to temperatures ofat least 600 ◦C. An important feature of dry granulite is that itcan remain metastable, and mechanically strong, at such elevatedtemperatures, even at pressures corresponding to depths of 50 kmor more, where the equilibrium assemblage is eclogite (Austrheim& Boundy 1994; Bjornerud et al. 2002; Lund et al. 2004). Thuswe suggest that much of the lower-crustal near-Moho earthquakeactivity on the continents is in rocks of a dry granulite facies, evenif such rocks are not exposed at the surface. Such rocks will bein the friction-dominated regime and mechanically strong. This isan important link between the depth distribution of seismicity andrheology.

5.2 Lithosphere thickness

There are two types of area underlain by thick lithosphere in theregions discussed in this paper. The first is within mountain beltsthat have experienced late Cenozoic shortening, such as the TibetanPlateau and, we believe, the Tien Shan. The narrow width of theTien Shan means that there may be lithospheric structure that isnot resolved by the methods of Priestley et al. (2006) and some re-gional studies have suggested the central Tien Shan is underlain byrelatively thin (∼100 km) lithosphere, however as discussed in Sec-tion 4.3 this remains controversial. Places where the lithosphere hasbeen thickened by late Cenozoic shortening are deforming rapidlyand are not expected to be in thermal equilibrium. Earthquakes




are generally restricted to the upper crust, except beneath the topo-graphic fronts, where it is reasonable to assign the deeper seismicityto foreland material thrust beneath the range. The thick lithosphericroot is a plausibly the result of recent rapid shortening. Dependingon the rate and geometry of its thickening process, the crust mayretain or develop a colder geotherm during the shortening itself, per-haps accounting for the relatively deep seismicity cut-off of about30 km within the Tien Shan (Fig. 3d), but thereafter the middle crustwill heat up on the timescale of tens of million years ago throughinternal heat generation. If the crust is thick enough, as it is in Tibettoday, this process can eventually lead to granulite facies conditionsin the middle crust, and may have been responsible for widespreadcratonization in the Precambrian (McKenzie & Priestley 2008).

Other areas with thick lithosphere, such as the Siberian and northIndian shields, are not associated with Cenozoic shortening. Inthese, and some other shield regions such as East Africa, the under-lying lithosphere is Archaean or Proterozoic in age and likely to bein thermal equilibrium. The thick lithosphere will allow a relativelycool geotherm, depressing the 600 ◦C isotherm to the Moho andallowing seismicity throughout the crust, but only if the lower crustis anhydrous.

In the oceans, the maximum lithosphere thickness of about105 km (McKenzie et al. 2005) is limited by a convective insta-bility that causes the old, cold lithosphere to detach (Parsons &McKenzie 1978). Where the lithosphere on the continents is sub-stantially thicker than this limit, it is likely to have been made so byshortening, and its mantle part stabilized against delamination bychemical depletion by melting, which will have removed iron andgarnet. Shortening of the lithosphere may also have been responsi-ble for crustal thickening and granulite generation in the crust but,as McKenzie & Priestley (2008) point out, geochemical and otherarguments require that depletion and melting must have occurredbefore the thickening event, not after. Nonetheless, it is at least pos-sible to see why strong seismogenic lower crust is often associatedwith thick lithosphere.

5.3 Shortening and rifting

In regions of shortening, the most important pattern revealed hereis that lower crustal earthquakes are restricted to the forelands andthe range fronts they underthrust, and do not occur beneath thehighest central parts of mountain belts. This pattern is likely tobe general. In Switzerland, microseismicity is limited to 15–20 kmdepth beneath the Alps, but increases to 30 km in the northern Alpineforeland, where it involves the whole crust (Deichmann et al. 2000).Diehl et al. (2009) found that seismicity occurred throughout thecrust in both the northern (Molasse Basin) and southern (Po Basin)forelands of the Alps. In South America, Assumpcao & Araujo(1993) report reliable locations of lower crustal earthquakes in theeastern forelands of the Andes in NW Argentina. Meigs & Nabelek(2010) conclude that planar thrust faults penetrate the lower crustin this area. For reasons explained above and in Section 3.3, theselower crustal earthquakes are a proxy for crustal strength, and anindicator that the rigid lower crust of the forelands supports therange fronts, rather than the mantle beneath them.

The increased strength of the lower crust is likely to be dueto an anhydrous, granulite-facies composition. This conclusion issupported in places, by evidence from seismic refraction and wide-angle reflection profiles (Wang et al. 2003; Zhao et al. 2006), whichshow that the Junggar and Tarim Basins have crustal structurestypical of stable continental platforms and lower-crustal veloci-ties that are consistent with a mafic granulite-facies composition

(Christensen & Mooney 1995). In contrast, the Qaidam Basin(where lower crustal earthquakes are not observed) lacks the high-velocity lower-crustal layer that indicates mafic granulite-faciescompositions (Zhao et al. 2006).

A consequence is that range fronts may be thrust several hundredkm onto the strong foreland, whose limit beneath the mountainsextends at least as far as the lower crustal earthquakes while theshortening is active (Fig. 9a) and, if the strong crust of the forelandis associated with thick lithosphere, then at the end of the shorteningevent, the limit of underthrusting will also be indicated by thegradient of lithosphere thickness. We suspect this is the significanceof the edge of the thick lithosphere east of the Verkhoyansk mountainfront (Fig. 7) and south of the Vitim Embayment (Fig. 6). A similarsituation probably exists in North America where the Rockies andMackenzie mountain fronts, emplaced onto the North Americanshield (Bally et al. 1966; Hyndman et al. 2005) lie several hundredkilometres east of the edge of the thick lithosphere (McKenzie &Priestley 2008).

If the mountain ranges are subsequently extended, it seems, fromthe evidence in the Baikal rift zone, Laptev Sea, and Eurasian basinmargin (Figs 6 and 7) very likely that the edge of the thick litho-sphere might again become reactivated (Fig. 9b), possibly becausethe original rheological contrast between the foreland and the rangeinterior is inherited or preserved. Within the 200–400 km resolutionof Priestley et al. (2006)’s maps, the rifting coincides very well withthe boundaries of the thick lithosphere. If the rifting currently onalong the edge of the thick lithosphere in the NE Rift of Baikalwere to continue, it would eventually separate the range front to thenorth from the interior of the Sayan-Baikal fold belt. Indeed, if rift-ing were to progress to become seafloor spreading, it would createthe configuration seen in the Appalachians, where the thin-skinnedrange front emplaced onto the North American shield (Cook et al.1979; Cook & Vasudevan 2006), and lying above thick lithosphere(McKenzie & Priestley 2008), is separated by an ocean from therest of the mountain range in Europe.

5.4 Cautions

The regions whose seismicity were examined in this paper revealpatterns that, through plausible links between seismicity, temper-ature, lithosphere thickness and crustal rheology, help understandsome major features on continental geology, topography and evolu-tion. However, two notes of caution are justified.

First, while the association of rifting with the edge of thicklithosphere is obvious in the NE Baikal rift, the Laptev Sea andEurasian basin, in other areas it appears possible for rifting to pen-etrate thick lithosphere. For instance the southern part of the EastAfrican rift system appears to penetrate lithosphere over 200 kmthick (McKenzie & Priestley 2008; Priestley et al. 2008b). Wherethis rift system enters the Precambrian shield system, it is also as-sociated with earthquakes throughout the crust (Foster & Jackson1998). However the shield system is far from homogeneous, and inplaces the rift system follows mobile belts, such as the ProterozoicUbendian Belt that separates the Archean Zambian and Tanzaniancratons (e.g. Ring 1994) and presumably represents some inheritedcrustal weakness. Whether such structures, which are only about100 km in width, correlate with structural features in the mantlelithosphere is uncertain, as with existing capability, features of thatsize cannot be resolved by surface wave tomography.

Second, while thick lithosphere in forelands is often associatedwith strong, seismogenic, and probably granulitic, lower crust, theconverse may not always be true. For example, the Molasse Basin in




Figure 9. Cartoon showing the idealised evolution of the tectonic settings under discussion. Upper-crustal earthquakes are shown as white circles, and lower-crustal earthquakes are shown as red circles. Dark green material represents relatively strong crustal material, with an anhydrous lower crust, characterized bylower-crustal seismicity. Light green material represents relatively weak crustal material where seismicity is restricted to the upper crust. (a) A compressionalsetting (similar to Tibet or the Tien Shan). A relatively weak orogenic plateau which flows out over a strong foreland as a gravity current. (b) An extensionalsetting (similar to the NE Baikal rift or the Laptev Sea). Rifting is concentrated along the edge of thick lithosphere. This causes some high topography, whichhas been thrust over the edge of the shield in a previous episode of thin-skinned compressional deformation, to be rifted away from the greater mountainousregion.

the northern foreland of the Alps which is seismogenic throughoutthe crust, and although it does not appear to be above abnormallythick lithosphere, it is too small in areal extent to resolve with thecurrent resolution of surface wave tomography. Nor is the asso-ciation of thick lithosphere with shields, although common, com-pletely understood. Priestley & McKenzie (2006) and McKenzie &Priestley (2008) are careful to point out that not all thick lithosphereis associated with shields (e.g. Tibet), nor do all shields now havethick lithosphere (e.g. NE China), and seismology is quite unableto determine the age of any thick lithosphere. For these reasons,Priestley & McKenzie (2006) and McKenzie & Priestley (2008) re-fer to the regions of thick lithosphere as ‘cores’; a word that denotesno particular origin or age.

6 C O N C LU S I O N S

We have shown that there is a striking correlation between tec-tonic setting, apparent long-term strength and depth distributionof seismicity throughout central–northeastern Asia and that this islikely to be a common feature across many continental areas. Incentral–northeastern Asia lower crustal earthquakes (those below15–20 km) are associated with the rigid forelands in regions un-der compression, and areas where rifting abuts old, strong cratonicmaterial (such as the Siberian Shield) in extensional regions. Thetemperatures at these depths (which may be as high as ∼600◦C) aremuch hotter than the normal temperature range of seismicity in thecontinents (up to ∼350◦C) suggesting that these earthquakes occurin anhydrous granulite grade metamorphic material. This wouldalso explain their association with long-term geological strength.

The correlation observed above can therefore be explained in termsof the links between rheology and composition that have recentlybeen proposed by Jackson et al. (2008) and the references therein.

This variation in crustal rheology appears to exert a first-ordercontrol on the morphology of the continents and can explain manyof the features observed. In compressional areas these range fromthe support of mountain belts by the surrounding foreland crust tothe formation of steep range fronts through gravity currents flowingover a rigid base. In extensional areas it appears that it may affectthe localization and orientation of rifting. In combination theseobservations also explain the preservation of ancient mountain belts(such as the Appalachians) following the reversal of the tectonicstresses that formed them. Some of these processes, such as theformation of the Himalayan range front through a gravity current,have been previously discussed (e.g. Copley & McKenzie 2007),however this study allows us to extend this model to other areaswhere traditional gravity-based studies of long term strength areinappropriate (e.g. in the Tarim Basin). Others, such as the formationand preservation of orphan mountain ranges rely on a combinationof these processes operating over the course of their geologicalhistory and so only become clear when, as in this study, a range ofmodern day settings are considered in combination.

A C K N OW L E D G M E N T

We would like to thank S. Roecker for providing us with theFKCMTINV software and two anonymous reviewers for theirthoughtful comments.




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S U P P O RT I N G I N F O R M AT I O N

Additional Supporting Information may be found in the online ver-sion of this article:

Appendix A. Waveform inversion for all source parameters.Appendix B. Forward modelling for depth.Appendix C. Inversion for all source parameters using regionalseismograms.

Please note: Wiley-Blackwell are not responsible for the content orfunctionality of any supporting materials supplied by the authors.Any queries (other than missing material) should be directed to thecorresponding author for the article.



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