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Mantle Earthquakes in the Himalayan Collision Zone Vera Schulte-Pelkum1, Gaspar Monsalve2, Anne F

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Mantle Earthquakes in the Himalayan Collision Zone Vera Schulte-Pelkum1, Gaspar Monsalve2, Anne F https://doi.org/10.1130/G46378.1 Manuscript received 22 May 2018 Revised manuscript received 20 May 2019 Manuscript accepted 22 May 2019 © 2019 The Authors. Gold Open Access: This paper is published under the terms of the CC-BY license. Published online 10 June 2019 Mantle earthquakes in the Himalayan collision zone Vera Schulte-Pelkum1, Gaspar Monsalve2, Anne F. Sheehan1, Peter Shearer3, Francis Wu4, and Sudhir Rajaure5 1Cooperative Institute for Research in Environmental Sciences and Department of Geological Sciences, University of Colorado, Boulder, Colorado 80309, USA 2Facultad de Minas, Universidad Nacional de Colombia, Medellin, Colombia 3Institute for Geophysics and Planetary Physics, University of California, San Diego, La Jolla, California 92093, USA 4Department of Geology, Binghamton University, State University of New York, Binghamton, New York 13902, USA 5Department of Mines and Geology, Lainchaur, Kathmandu 44600, Nepal ABSTRACT ity located using the network (Monsalve et al., Earthquakes are known to occur beneath southern Tibet at depths up to ~95 km. Whether 2006), deep events are seen in Nepal just south these earthquakes occur within the lower crust thickened in the Himalayan collision or in of the Lesser Himalaya (Fig. 1, cluster C) and the mantle is a matter of current debate. Here we compare vertical travel paths expressed as under Tibet (Fig. 1, clusters A and B). Figure delay times between S and P arrivals for local events to delay times of P-to-S conversions from 2B shows the same set of events on a previ- the Moho in receiver functions. The method removes most of the uncertainty introduced in ous structural depth profile from receiver func- standard analysis from using velocity models for depth location and migration. We show that tions on a line crossing the collision zone in a deep seismicity in southern Tibet is unequivocally located beneath the Moho in the mantle. N18°E orientation perpendicular to the local Deep seismicity in continental lithosphere occurs under normally ductile conditions and has range strike (Schulte-Pelkum et al., 2005). The therefore garnered interest in whether its occurrence is due to particularly cold temperatures Moho is visible as a contrast in isotropic veloc- or whether other factors are causing embrittlement of ductile material. Eclogitization in the ity, and the Main Himalayan thrust as a con- subducting Indian crust has been proposed as a cause for the deep seismicity in this area. trast in anisotropy (Fig. 2B; Schulte-Pelkum Our observation of seismicity in the mantle, falling below rather than within the crustal layer et al., 2005). Even in this best-case scenario, with proposed eclogitization, requires revisiting this concept and favors other embrittlement with relocations and structural imaging done mechanisms that operate within mantle material. with the same stations and the same velocity models, the deep seismicity appears diffuse INTRODUCTION constrain than their epicenters because the sta- in depth and is difficult to clearly assign to The Himalaya-Tibet continental collision tions used to locate them are on the surface and crust or mantle. zone has featured prominently in a debate on there is a strong tradeoff between location depth Local earthquake depths are usually calcu- whether strength in the lithosphere resides in and P and S velocity structure above the events. lated from the travel-time difference between P two layers in the brittle crust and uppermost Intralithospheric interfaces such as the Moho and S arrivals from the event at each station. In mantle separated by a weak lower crust, or in a can be seen with teleseismic converted waves receiver function studies, delays between tele- single layer largely limited to the crust (Chen (receiver functions), but determining their depth seismic P and converted S are used to determine et al., 1981, 2012; Chen and Molnar, 1983; Zhu suffers from the same tradeoff with velocity depths to interfaces under the station. In both and Helmberger, 1996; Henry et al., 1997; Cat- structure above the interface as in estimating cases, the S-P delay time is mapped to depth tin and Avouac, 2000; Jackson, 2002; Chen and earthquake depths. Comparing hypocentral by assuming P and S velocity models between Yang, 2004; Beaumont et al., 2004; Monsalve depths with structural depths is hampered by the the event or converting interface and the station. et al., 2006; de la Torre et al., 2007; Bendick and uncertainties in both parameters. The compari- Teleseismic arrivals and local events close to the Flesch, 2007; Liang et al., 2008; Priestley et al., son is accurate only if the same velocity models station have similar steep incidence angles (Fig. 2008; Craig et al., 2012). The distribution of are used for depth determination for both data 3C). Direct comparison of such S-P delay times seismicity with depth in strained regions serves sets. More typically, results are compared be- between the two data sets therefore avoids the as a proxy for the lithospheric strength profile tween different studies, and a lack of consistency introduction of bias from using velocity models because the occurrence of earthquakes implies between the velocity models used exacerbates for depth migration and from unaccounted-for sufficient strength to sustain brittle failure, due the uncertainties. We propose a simple work- lateral variations, because the ray paths sample to either cold conditions (e.g., Chen and Mol- around to the depth determination uncertainty similar volumes. The two comparison data sets nar, 1983; Sloan and Jackson, 2012; Blanchette by comparing seismicity and structure directly are constructed as follows. et al., 2018) or local embrittlement processes as delay times between shear and compressional (e.g., Incel et al., 2017, 2019; Prieto et al., 2017). waves (“S minus P”, or S-P), rather than convert- Local Seismicity The task is therefore to map where exactly local ing both to depth first. For the local seismicity used in the S-P delay earthquakes occur with respect to layers within comparison, we used a previously relocated and the crustal and uppermost mantle column. A DATA AND METHODS published local event set of ~500 earthquakes difficulty is posed by uncertainty in estimated We used data from the CE 2001–2003 (Monsalve et al., 2006; Fig. 1), each located depths for both local seismicity and the Moho. HIMNT (Himalayan Nepal Tibet) seismic with at least 10 P and S arrivals picked with Crustal earthquake depths are more difficult to broadband experiment (Fig. 1). Of the seismic- the HIMNT network; pick statistics are shown CITATION: Schulte-Pelkum, V., et al., 2019, Mantle earthquakes in the Himalayan collision zone: Geology, v. 47, p. 815–819, https:// doi .org /10 .1130 /G46378.1 Geological Society of America | GEOLOGY | Volume 47 | Number 9 | www.gsapubs.org 815 Downloaded from http://pubs.geoscienceworld.org/gsa/geology/article-pdf/47/9/815/4831709/815.pdf by guest on 02 October 2021 Figure 1. Map of study area, Himalaya-Tibet collision zone. Inset map shows regional geographical context, with gray shading 85° 86° 87° 88° showing topography for orientation. Blue circles in inset are 12 events from the compilation of Chen and Yang (2004) with depths Elevation RC14 of 80 km or more. Magenta open circles in inset map are our deep (m) SAGA SSAN ONRN events 1–6 (Table DR1 [see footnote 1]). In large map, background 8000 LAZE color is elevation shaded by topography. Small circles are best- 7000 relocated local seismicity from 2001 to 2003 (Monsalve et al., 29° 29° 6000 SAJA 2006, 2008, 2009). Large circles are the three deep events from the XIXI MNBU 5000 4 1986 DINX compilation of Chen and Yang (2004, their events T8, T9, and T12; NAIL mb5.5 MAZA displayed here with event year and body wave magnitude) that 4000 5 3000 1 fall within the map area. Events are color coded by depth below 2 3 sea level. Small circles have thin white outline for events in the 2000 RBSH A 1000 6 1992 initial catalog (Monsalve et al., 2009; white circles in Fig. 2B) and m 4.9 28° BUNG Everest b 28° thicker black outline for events used in delay-time analysis after 0 B Kathmandu selection for station distance, location error, and pick uncertainty SUKT NAMC (blue circles in Fig. 2B). Black triangles are HIMNT (Hima layan 80° 90° Kanchenjunga THAK JIRI Nepal Tibet) stations used in this study. Red line marks the profile PHAP shown in Figures 2 and 3. White squares show locations of Kath- Tibetan RUMJ SIND TUML mandu, Mount Everest, and Kanchenjunga. Black labeled ellipses Plateau PHID roughly outline clusters of deep seismicity A–C discussed in the 30° 30° HILE 27° ILAM 27°N text; while the three larger events from Chen and Yang (2004) are C GAIG JANA located near them, the clusters’ distribution in space, time, and 20° 20° 1988 m 6.4 magnitude do not fit aftershock characteristics (see the Data Re- b −10 20 45 60 75 95Depth BIRA 50 km pository). Dashed rectangle shows map area of Figure 4A, and 80° 90° pale orange squares show receiver function piercing points at (km) 75 km depth for stations in Figure 4. 85° 86° 87° 88°E in Figure DR1 in the GSA Data Repository1. ples in Fig. 4B). A group of stations in southern Figure 3 demonstrates that the error introduced For each event in this catalog located within the Tibet shows a positive amplitude arrival pre- by a possible bias in velocity model is of second network footprint, we selected stations with P ceding the Moho peak (Fig. 4B), which marks order because of the small epicentral distances and S picks at epicentral distances of <35 km to the top of a layer with elevated lower-crustal involved, whereas a velocity model bias leads to sample a near-station volume.
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