Correlation of change in magnetotelluric parameters before and after Sikkim Himalayan earthquakes associated with stress variation

Tectonophysics 765 (2019) 205–225

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Correlation of change in magnetotelluric parameters before and after Sikkim Himalayan earthquakes associated with stress variation

T

Shankar Konda1, Prasanta K. Patro⁎ CSIR-National Geophysical Research Institute, Uppal Road, Hyderabad 500007, India

ARTICLE INFO

ABSTRACT

Keywords: Magnetotellurics Sikkim Himalaya Apparent resistivity Stress drop variation Ground motion

MT stations that were recorded before the moderate 5.3 Mw, 2006 Sikkim Himalaya earthquake had been reoccupied after 2007 (Mw 5.0) and 2011 (Mw 6.9) seismic activities. We analyzed the broadband MT data (0.001–1000 s) within the 60 km radius from epicenters location and within the focal depth range of 9–30 km. The observed change in apparent resistivity in the period range of 0.001–10 s has shown a considerable postseismic decrease of the order 10–30%. The 2D geoelectric model constructed with post seismic acquired data also reveals the feature of a decrease in resistivity (C1). The relative change (decrease) of observed apparent resistivity ((Δρ/ρobs) after 2006, 2007 and 2011 earthquakes find a correlation with combined stress drop 11–18.2 MPa values estimated by Arms-Δσ approach. The correlation is further strengthened by a similar change in relative apparent resistivities ((Δρ/ρ2D_model) derived from 2D model with stress drop. Integrated stress drop values attribute to the accumulated apparent stresses (σa = 0.23 ∆σ) that reveals the area might be subjugated to 0.6–1.2 MPa stress before 2006, 1–1.5 MPa stress before 2007 and 1.7–2.9 MPa before 2011 earthquakes occurrence. Development of micro-cracks and generation of aqueous fluids is possible due to such stressed and slow rate of relentless thrust/shear environment within the area of crust. Therefore, resistivity could have been decreased with a fluid transition from compression to dilatancy zones at different depth levels before each earthquake and further decreased upon subsequent stress drops associated with strong ground motions. We interpreted that maximum stress might be dropped in dilatational zones oblique to the maximum compression stress directions along N200E, N-S and NNW-SSE before 2006, 2007 and 2011 earthquakes and resulted in a considerable decrease in apparent resistivity along E-W (TE-mode), i.e., ρyx component.

1. Introduction

has been reported except some minors to light magnitude earthquakes (3.0 ≥ Mw ≤ 4.9). The concentration of epicenters was found to occur along the decollement plan between the surface traces of MBT and MCT within the depth of 10–25 km (Pal et al., 2008; Hazarika and Ravi Kumar, 2012; Ajaay et al., 2017). The seismo-tectonic models envisaged that the seismicity in this region is bimodal (Pradhan et al., 2013) associated with strike-slip and thrust/shear dominated mechanisms (Ansari et al., 2014). Traverse tectonics evidenced by the NNW-SSE Tista-Gangtok and NW-SE Golpara lineaments appear to have a little role in strain concentration (Hazarika et al., 2010; Dasgupta et al., 2013). Micro seismic surveys and fault plane solutions supported that the seismicity in this region resulted from stress accumulation along thrust zones and at mid-crustal ramp beneath arcuate MCT (Ni and Barazangi, 1984; Pandey and Tandukar, 1995; De and Kayal, 2004) and releasing accumulated strain through frequent minor to large earthquakes. Presence of distinct focal mechanisms, relentless near constant

Sikkim Himalaya is a segment of the 2500 km length of the arcuate Himalayan belt which was developed as a result of the Cenozoic Convergence ~55 Ma (Molnar and Tapponnier, 1975; Rowley, 1996; Aitchison et al., 2007) between the Indian craton and Eurasia. The rate of convergence was 4 cm/yr (Copley et al., 2010) at the initial stage and continued to decrease to 12 ± 1.16 mm/yr observed by GPS measurements (Bilham et al., 1997; Mukul, 2010). The main frontal thrust (MFT), main boundary thrust (MBT) and peculiar overturned main central thrust zone MCTZ bounded by MCT-1 & MCT-2 (Bhattacharyya and Mitra, 2011; Mottram et al., 2014; Dilip et al., 2017) are the prominent intra-continental thrusts developed after the NNE convergence. All these low angle thrusts join an active plane of detachment along which the Indian crust is subducting. Apart from 1965 earthquake (Mw 5.9), and 1980 earthquake (Mw 6.2), no greater earthquake (> Mw 7.0)

Corresponding author. E-mail address: [email protected] (P.K. Patro). 1 Now at:- Oil and Natural Gas Corporation Ltd., Vadodara, Gujarat, 390009, India. ⁎

https://doi.org/10.1016/j.tecto.2019.04.034 Received 21 March 2018; Received in revised form 27 April 2019; Accepted 29 April 2019 Available online 03 May 2019 0040-1951/ © 2019 Elsevier B.V. All rights reserved.

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jolted by 4.7 Mw and 5.0 Mw strike-slip mechanism earthquakes on 18th and 20th May 2007 with epicentral coordinates 27.30°N: 88.19°E and 27.30°N: 88.15°E (http://asc-india.org/seismi/seis- sikkim.htm; Joshi et al., 2010) at hypocenters depth of about 12 km and 17 km respectively. Further, the 18th September 2011 (6.9 Mw, depth 19.7 km) (http://usgs.gov.in) strong strike-slip mechanism earthquake to the northwest of Sikkim 27.73°N: 88.15°E (Thakur et al., 2012; Sajad et al., 2016) (Fig. 1) with one foreshock (Mw 4.9, depth 26 km) and four aftershocks (Mw 3.9–5.0, depth 9–30 km) devastated the region with MMI VI-VII within 60 km radius from epicenters. The rupture area was 591 km2. NNW-SSE trending younger Tista lineament was played an important role for 2007 and 2011 earthquakes in the release of the tectonics strain with a strike-slip mechanism (Mahajan et al., 2012; Baruah et al., 2016). Earthquake phenomenon is a geodynamic process and occurs in stress-enhanced areas. Even, small stress load about 0.001–0.002 MPa can trigger earthquakes (Raju et al., 2008; Nalbant and McCloskey, 2011). In case of intra-continental thrust phenomenon, the thrust fault having large surface area allows long period of stress/strain accumulation before earthquake (Bollinger et al., 2004; Dilip et al., 2017) and shows a great drop of stress after an earthquake (Liao et al., 2002; Wesson and Boyd, 2007; Yoshida et al., 2012). The release of stored elastic energy produce considerable rearrangements of stresses, hence changes in stress/strain rate in various parts of the crust even at a greater distance from rupture zone (Abdullaev et al., 2011) causes changes in physical properties of rocks. Some recent studies stated that the generation of a moderate to the large earthquake is not only a purely mechanical process but is closely related to fluid existence (Savage, 2010; Mahesh et al., 2011). Electrical resistivity is a physical property of the earth which is highly sensitive to stress/strain changes (Barsukov, 1972; Fitterman, 1976), interconnected fluids (Yamazaki, 1967; Brace, 1975), cracks and porosity (Brace and Orange, 1968) and shows its variation in low porosity crystalline rocks influenced by dilatancy and stress drop (Honkura et al., 2013). Resistivity changes in the crust nearby the focal regions under the action of stress change before and after earthquakes have been the subject of research in recent times. The magnetotelluric method is a passive electromagnetic technique that measures fluctuations in the natural electric and magnetic fields in the orthogonal direction at the surface. It provides the conductivity structure of the subsurface (few tens of meters to several hundred kilometers). Limited electromagnetic sounding methods have been carried out in seismically active regions to understand the crustal resistivity behavior prior to and post-earthquake occurrence (Honkura et al., 2000; Yang et al., 2002; Chen et al., 2007; Park et al., 2007; Yoshimura et al., 2008). Observation of resistivity changes associated with stress variation from MT data requires better than 0.1% precession (Fitterman, 1976; Johnston, 2002; Park et al., 2007) and could be possible if there are remote reference measurements (Gamble et al., 1979). Current MT processing techniques can at best provide the daily precision of 1% (Eisel and Egbert, 2001). Few MT investigations have been probed in the study area (Patro and Harinarayana, 2009; Pavan Kumar and Manglik, 2012; Manglik et al., 2013; Pavan Kumar et al., 2014), but no attempt has been made to study the apparent resistivity change prior to and postearthquake in association with stress variation. As the MT data were acquired with remote reference before 5.3 Mw, and also after 5.3 Mw, 5.0 Mw and 6.9 Mw earthquakes, we take an opportunity to observe any possible changes in MT parameters in association with stress variation from seismic activity at hypocenter nearby sites. In this paper, we deal with the analysis of change in MT parameters derived from the time series data acquired before and after 2006 Sikkim earthquake and tried to give a valid explanation for the observed apparent resistivity change in association with stress drop from these subsequent seismic events.

Fig. 1. Regional geo-tectonic map of Sikkim Himalaya showing Magnetotelluric (MT) site locations 1-S07, 2-S08, 3-S12, 4-S14, 5-S13 before (Black) and 1SK07, 2-SK08, 3-SK12, 4-SK14, 5-SK13 after (Blue) the earthquake events with focal mechanisms (Raju et al., 2008; Paul et al., 2015). Red Star represents the 2006 epicenter location by NGRI (Raju et al., 2008), Black circles represent the 2006 aftershocks. Yellow stars indicate epicenter and aftershocks of the 2007 earthquake (Joshi et al., 2010). Maroon stars are the epicenter and its aftershocks of 2011 earthquake (Baruah et al., 2016). Small stars indicate minor earthquakes (≤ 3 Mw) (Modified after Arora et al., 2014). (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

rate of slip causes for flexural stresses, and a large amount of strain accumulation along NE dipping critically stressed thrust zones (Bilham and Ambraseys, 2005; Ponraj et al., 2010; Purnachandra Rao et al., 2015) nucleate the complex tectonic scenario and responsible for many earthquakes in tectonically active Sikkim Himalaya. MT investigation carried out in the study area primarily aimed to delineate the crustal structures and geo-electric signatures. On 14th February, 2006 (2006-2-14 00.55.25 UTC, 6.25.23 am IST) (http:// usgs.gov.in), a thrust mechanism earthquake (Mukul et al., 2014) of Mw = 5.3 (ML 5.7) (htpp://imd.gov.in) along with its aftershocks (>70) was caused severe crustal damage in the area of investigation after the data was acquired in 2005. The epicenter was reported at 27.410 N and 88.550 E, near Phodong with a focal depth of 20 ± 8 km (http://asc-india.org/lib/20060214-sikkim.htm; Raju et al., 2008) (Fig. 1). It caused the damage equivalent to III-VII MMI (Som et al., 2008). The rupture area was 15.99 km2 calculated using the relation log (A) = −3.99 + 0.98 Mw. Slip on the rupture was 0.22 m and maximum slip can be expected to 0.69 m (McGarr and Fletcher, 2003). From epicenter location and focal depth, 2006 thrust mechanism earthquake occurred on or along the plane of detachment. The study area again 206

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2. Geology and seismo-tectonics

electrical disturbances in this mountainous region, we selected sites to achieve the dipole length of about 40 m from center point towards N, S, E and W directions. Magnetic and electrical data were acquired using induction coils MFS-06 and Pb–PbCl2 non-polarizing electrodes respectively. ADU06 data acquisition system ((M/s Metronix GmbH)) was used to record time series data. At every station, we recorded time series data for about 3–4 days. Remote reference recording was carried out during the data acquisition to reduce the effects of local noise and for enhancing the data quality and precision. A long duration of recording was helped in obtaining consistent apparent resistivity and phase curves in the MT dead band (1–10 s). After the completion of MT data acquisition in 2005, the study area was experienced by three subsequent moderate to large earthquakes M5.3 (2006), M5.0 (2007), M6.9 (2011) along with their aftershocks. These raised an opportunity to study the possible changes in apparent resistivity in comparison with data acquired before the 2006 earthquake. To observe any changes in the subsurface current directions and crustal resistivity, we were reoccupied the old sites just after 6.9 Mw, 2011 earthquake. MT data were acquired during the 2011 and 2012 field seasons with the same kind of MT equipment in the same frequency range (10000–0.001 Hz). Though, we could successfully reoccupy 13 old stations as close as 50 m to their initial locations except for SK14, where the reoccupied site is 2 km away from the original site (S14). We could not re-occupy a few stations due to industrial growth and urbanization. Data were processed with the same parameter approach as that of 2005 data. In this study, we compared only five sites SK07, SK08, SK12, SK13 and SK14 with pre-earthquake sites S07, S08, S12, S13 and S14 where the epicenters of 5.3 Mw, 5.0 Mw, and Mw 6.9 earthquakes are nearby within the 60 km radius. Sites SK07, K08, SK12 and SK14 were fallen in LHD and SK13 was occupied in MCTZ.

The Sikkim Himalaya designated as IV in seismic zone map of India (BIS 1983:2002) is located in 280 07′ 48″ N to 270 04′46 ” N, and 880 00′ 58″ E to 880 55′25″ E (Fig. 1). The molasses type deposits of the Siwalik formation in the south separate from low-grade metamorphic rocks of lesser Himalayan Domain (LHD). Further north, The MCT separates the LHD from the medium to high-grade metamorphic rocks of Higher Himalayan Domain (HHD) (Pal et al., 2008). The boundary between the Siwaliks and the LHD is marked by Main Boundary Thrust (MBT), and the boundary between LHD and HHD is marked by MCT. Limited geophysical and micro-seismic investigations have been probed the area compared to the other parts of Himalaya belt. Fault plan solutions suggested that the MBT is a deep-rooted mantle reaching fault. The MCT a ductile shear zone (Catlos et al., 2004; Mottram et al., 2014) often referred to as main central thrust zone (MCTZ) bounded between MCT-1 and MCT-2. Ni and Barazangi (1984) suggested that the MBT is active and mantle reaching fault where MCT is a dormant feature now. The crustal shortening in Sikkim Himalaya has been accommodated by traverse tectonics rather than under thrusting (Hazarika and Ravi Kumar, 2012). NW-SE tending Tista/Gangtok lineaments and WNWESE of MCT betting Golpara lineaments reinforce the traverse tectonics (De and Kayal, 2004). Micro-earthquake data revealed that the earthquakes along active Golpara lineament are at a mid-crustal level (10–25 km) and shallow to lower (0–45 km) (De and Kayal, 2003; Arora et al., 2014) crustal level in other parts of Sikkim Himalayan zone (Thakur et al., 2012). Dipping of high resistive Indian crust at angle 60 with 40 km thickness under lesser Himalaya and 220 with 60 km thickness beneath HHC (Singh et al., 2007; Acton et al., 2011) with negative gradient of gravity reflect the nature of subduction (Tiwari et al., 2006; Ansari et al., 2014). The MFT, MBT, and MCT all join the north dipping decollement (MHT) which is locked to the base of the mid-crustal ramp that leads to stress build up and clustered microseismic activity around the ramp (Pandey and Tandukar, 1995; Bilham et al.,1997). Patro and Harinarayana (2009) and Pavan Kumar et al. (2014) mapped the high conductive nature of trapped sediments with fluids at a shallow level along MFT, MBT and MCT thrust zones. The prominent thrusts are critically enough stressed and accumulated enough strain to be released for future earthquakes (Raju et al., 2008). The higher stress drop in Sikkim Himalaya relative to low-stress drop in Himalayan region is attributing to the distinct feature affinity with the strike-slip source mechanism (Hazarika and Ravi Kumar, 2012). The micro-seismic study yielded that the 453 km2 rupture area for 6.9 Mw, 2011 earthquake caused 4–5 MPa stress drop with a dextral strike-slip mechanism (Paul et al., 2015). Coulomb stress analysis revealed that the main and aftershocks of 2006 and 2011 seismic events were generated in increased stress zones and change in coulomb stresses occurred at 12–25 km depth (Raju et al., 2008; Baruah et al., 2016). Sikkim Himalaya has highly complex tectonic setup between MCTZ and LHD inferred from the change in the strike directions together with the variations in the skew angle (Pavan Kumar and Manglik, 2012; Manglik et al., 2013). The variable regional strike directions, complex structures and bimodal nature of seismicity in the area of investigation demand the necessity of suitable 3-D model which could clear the traverse tectonics and nature of complex subsurface signatures. Further, no attempt has been made to study the change in magnetotelluric parameters pre to, and post-earthquake in this region attracts MT study.

4. Data analysis Time-dependent electromagnetic data is compared at five sites S07SK07, S08-SK08, S12-SK12, S13-SK13 and S14-SK14 acquired before and after the earthquakes which are in and near to the epicentral region. These sites were relatively free from cultural electromagnetic noises and recorded with remote reference. Two components of the electric field and three components of the magnetic field were simultaneously measured at each site, including the remote reference site. Data recorded in the frequency range of 10,000–0.001 Hz with the sampling frequency (40,960 Hz, 4096 Hz, 512 Hz, 128 Hz, 64 Hz, and 2 Hz), but restrict the data analysis to the frequency range of 10,000–0.01 Hz. Single site time series data was processed with remote reference using commercial software MAPROS (M/s Metronix GmbH). Selective stacking method was used to estimate the impedance tensor for the period range 0.0001–100 s. Remote-reference processing resulted in a significant improvement in the quality of impedance (Z). The apparent resistivity and impedance phase curves for each frequency at each site, together with before and after 2006 earthquake are shown in Fig. 2a & Fig. 2c. The difference in apparent resistivity is shown in ratios (ρ2/ρ1) (Fig. 2b) and for impedance phase is in difference (Δϕ) (Fig. 2d). The apparent resistivity is decreased in both data sets with an increase in phase. Sites S07-SK07 represented with moderate resistivity (≤1000 Ω.m) at long periods and reduction towards shorter period. High resistive (≥1000-10,000 Ω.m) response curves for S13-SK13 and S14-SK14 at 0.0001s-1s are due to HHC rocks in MCTZ. We observed that at S13-SK13 and S14-SK14, the apparent resistivity of yx component data is higher than the xy component, and reverse at other sites. However, observed apparent resistivity at all the sites show a common trend of decrease in apparent resistivity towards the longer period, but the difference in apparent resistivity (ρ2/ρ1) and phase (Δϕ) are relatively observable in the range of 0.1 to 1000 Hz. There is an increase in apparent resistivity also at higher frequencies (>1000 Hz). Phase values are also got changed remarkably at all data points. The difference shown in negative for both apparent resistivity and impedance phase

3. MT data acquisition Broadband MT data were collected along an NNE-SSW profile (120 km) from Siliguri to Yumthang covered by 18 stations that cut across the major tectonic domains of the Sikkim Himalaya in 2005 (Fig. 1). The station spacing was about 5–8 km. Data were acquired in the frequency range of 10000 Hz to 0.001 Hz. As it was challenging to get relatively flat open fields of 250 × 250 m dimension away from 207

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Fig. 2. Comparison of observed apparent resistivity and phase (xy and yx components) of S07, S08, S12, S13, and S14 acquired before 2006 earthquake and of SK07, SK08, SK12, SK13, and SK14 acquired after 2006 earthquake (a & c), apparent resistivity difference in ratio (b), and Impedance phase difference (d).

values indicates a decreasing behavior at all the sites after the earthquakes. Maximum change is observed at sites SK12, SK13 and SK14 were comparatively near to the epicenters of earthquakes. From the close inspection of all the sites (Fig. 2a), the apparent resistivity in the frequency range of 0.1 Hz to 1000 Hz shows a post-seismic decrease of the order of 10–30%. Resistivity is decreased by 1–10% in its shallow part (0–2 km) and by 20–30% in its deep part (2–18 km). In the preceding, we have calculated the percentage of error in data to ascertain that the observed change in apparent resistivity and impedance phase is significant than the uncertainty in the data (Table 1). The rhoplus consistent test is used for the MT responses of each site,

and reasonable data points along the smooth curve are considered for the calculation of MT parameters (skew, phase tensor, polar diagrams, strike direction) and for further data analysis. 4.1. Strike analysis The geo-electrical strike is a direction representing the orientation of electric current flow in the subsurface and can alter due to lateral conductive in-homogeneity and anomalous crustal resistivity of the Earth. In this study, we performed multi-site multi-frequency (MSMF), and single site multi-frequency (SSMF) strike analysis of MT transfer 208

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functions using Smith code (Smith, 1997). Representation of strike direction is done using the rose diagram in bidirectional binned at every 50 window. The yielded strike direction from MSMF analysis in the period range of 0.0001–100 s for the entire five MT sites acquired before 2006 and after 2006, 2007 and 2011 earthquakes is represented in Fig. 3a. From the comparison of geo-electric strike direction for all the five sites recorded before (black color) 2006 and after (red color) three seismic activities, we observe that there is a small change in regional strike direction from each other. The trend of strike is consistent over the entire period range for pre-seismic sites S08, S12, and S14 and shown a deviation for sites S07 and S13. Sites SK07 and SK13 show a similar direction and the other three sites resulted in a different direction of the strike. The degree of regional strike direction rotated from N450W towards the south at S07 to N200E towards the north at S13. But the regional strike direction for the sites acquired after earthquakes shows a different pattern. The degree of regional strike

Table 1 Percentage of error in apparent resistivity and impedance phase for threeperiod ranges of 0.001–0.1 s, 0.1–10 s, and 10–100 s for five sites acquired before and after the 2006 earthquake. Note that for the period range 10–100 s a high error percent (marked as grey) is observed in apparent resistivity and phase for the sites S07-SK07, S13-SK13 and S14-SK14.

Fig. 3. a. Strike directions in the period range of 0.0001–100 s for five sites acquired before (black color) and after (red color) the 2006 earthquake obtained by Smith (1997) approach. b. Strike directions in the period range of 0.0001–100 s for five sites acquired before (top) and after (bottom) the 2006 earthquake obtained from phase tensor analysis. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.) 209

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Table 2 Comparison of strike directions for three-period ranges of 0.0001–100 s for five sites acquired before and after the 2006 earthquake. Period/site

S07

SK7

S08

SK8

S12

SK12

S13

SK13

S14

SK14

0.0001–0.01 s 0.01–1 s 1–100 s

−4 −30 43

25 −33 23

−38 −36 −35

−42 −12 40

−29 −31 9

−25 0 −11

−35 33 −23

36 25 −22

−37 0 −16

−34 −3 −1

direction rotated from N200E towards the south at SK07 to N400E at SK08. At site SK12, the strike direction is N180W, and at SK14 the strike is completely N-S. Further north, SK13 is showing the strike direction N220E. As the Smith approach does not resolve 900 ambiguities in MT strike determination, the orthogonal geo-electric strike directions, i.e., E–W is also possible for these strike directions. Strike directions are also calculated for three-period bands 0.0001–0.01 s, 0.01–1 s, and 1–100 s (Table 2). Strike direction for each period band is dissimilar from one another and exhibited a common trend of NW-NE.

4.4. Static shift Static or telluric shift on MT data generally cause by resistivity inhomogeneities. Galvanic and eddy current distortions are the two phenomena can be thought of arising static shifts in this environment. The electric field can be distorted by induced eddy currents and galvanic phenomena at higher frequencies, and it is subjected to galvanic distortion at very low frequencies. During the 2D inversion, we have inverted the data for static shift. The static shift values for both TE and TM modes are shown in Fig. 6. All the TE and TM mode static shift values for all the sites are confined to less than or equal to 1, which clarifies that the data has no/insignificant static shift.

4.2. Polar diagrams Analysis of impedance polar plots from MT data provides a measure for the MT data dimensionality information about the subsurface geologic structure. Principal-impedance polar diagram elongates over a resistive structure perpendicular to strike direction and elongates parallel to strike direction over the conductive body. Fig. 4 illustrates the polar diagrams for the sites recorded before and after the earthquake activities. Zxy is the principal impedance, and Zxx is a diagonal one. Zxy and Zxx are shown in black and red. Except for SK07 and S07 where the polar plots indicate 1D structure, all the other sites polar plots are 2D/ 3D in nature with peanut shape elongated in a direction either parallel or perpendicular to strike. From close observation, S08-SK08, S12SK12, S13-SK13 and S14-SK14 sites are represented a considerable difference in polar plots. SK13 towards high-frequency range (shallow part) exhibited a distinct nature of difference in polar diagrams and indicated 2D anisotropy or complex regional 3D structures.

4.5. Stress variation The present area comprised of LHD and HHC rocks bounded by critically stressed thrust zones MBT and MCT. The region with slow rate of convergence 4 mm/yr in LHD and 8 mm/yr north to LHD in HHC (Mukul, 2010) witnessed by high stress drop 5–40 MPa (Raj et al., 2009; Hazarika and Ravi Kumar, 2012; Raghukanth et al., 2012) compared to entire Himalayan region and it depends on the slip rate, area of rupture mechanism and rupture velocity (Nath et al., 2005; Paul et al., 2015). The crust was experienced with NE-SW, N-S, and NNW-SSE compression stresses before the 2006 and 2011 earthquakes (Raju et al., 2008; Baruah et al., 2016). As Sikkim Himalaya is a part of the Himalayan region, the phenomenon of maximum shear strain accumulation 0.4–0.1 μstrain/yr near MCT across Himalayan region (Ponraj et al., 2010) is observed from geodetic measurements. The observed 38.6 MPa stress drop from 2007 and 5–20 MPa stress drop from 2011 earthquake indicate the accumulation of maximum stress/strain by compressional activity before these events. In the rocks, change in resistivity is more pronounces to change in shear stress/strain than the change of elastic property (Brace and Orange, 1968; Honkura et al., 1976; Drury, 1979; Yang et al., 2002). Based on the field observation of stress changes (1 MPa) accompanying earthquakes, resistivity changes of at least 1% might be expected to accompany crustal failure (Johnston, 2002). Geophysical and lab experiments reported several percentages of change in resistivity before and after earthquakes which are related to the change in stress/strain (Park et al., 2007; Abdullaev et al., 2011; Honkura et al., 2013). The observed change in apparent resistivity, along with a change in impedance phase, polar diagram, strike direction, and phase tensor of all sites in this study also invokes that these changes are related actively or passively to the variation in stress/strain levels prior to and post-seismic activities. Hence, we strongly believe that the observed variations in MT parameters might be accompanied by the stress variation at different depth zones caused due to the crustal disturbance by rupture propagation associated with moderate to large earthquakes. One of the most important characteristics of the earthquake source process is the amount of stress released on the fault, often referred to as stress drop. Stress drop is the difference between the average stress level before and after the earthquake (Allmann and Shearer, 2009; Courboulex et al., 2016). Estimation of stress drop and its variation provide information about the mechanism of earthquake rupture. Amount of stress drop can help to describe the tectonic and frictional stress levels which can alter the resistivity before and after the earthquake. Two types of changes of electrical resistivity would be inferred

4.3. Phase tensor Caldwell et al. (2004) formulated phase tensor (PT) Ф = X−1Y, where X and Y are real and imaginary components of the impedance tensor. Calculation of the PT requires no assumption about the dimensionality of the underlying conductivity distribution and is applicable where both the heterogeneity and regional structure are 3-D. Calculated skew angle (β), a measure of phase tensor's asymmetry and dimensionality parameter for all sites from phase tensor analysis equation β = ½ tan−1(Φ12 - Φ21) / (Φ11+ Φ22), where Φ12, Φ21, Φ11, Φ22 are the components of the phase tensor Φ (Fig. 5). The color of ellipses represents the skew angle (β, in degrees). Results reveal that phase tensor asymmetry and skew angles at each site are relatively difference for different frequency and having a change in strike direction. Strike analysis, and Polar plots suggested that the region comprised with variable strikes with 2D anisotropy or complex 3D structure, hence we analyzed the phase tensor orientation direction (γ = α β) for the entire frequency band to verify the obtained regional strike direction from Smith approach (Fig. 3b). The variation in the orientation of normalized ellipses with major axis fairly consistent with computed strike direction at all sites. Distortion of phase tensor ellipses (1100 s) strongly suggests the presence of the complex nature of subsurface structures. The skew angle for sites S07 and SK07 are 1D for the frequency range 0.001–0.01 s and possibly 2-D nature towards longer periods. Sites S12, SK12, and S08, SK08 at 1–100 s indicate a 3-D structure with high skew values. A high skew value observed for the SK13 site at the near-surface provides information about the complex 3D structure lying within the MCTZ. 210

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Fig. 4. Polar diagrams at frequencies indicated left of each plot for sites recorded before (S07, S08, S12, S13, and S14) and after (SK07, SK08, SK12, SK13, and SK14) the 2006 earthquake. The Blue indicates the principal impedance, Zxy. The red shows the diagonal impedance, Zxx. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

in association with stress/stress drop variation. In the first, resistivity may increase with shear stress, reaches a minimum together with a sudden release of shear stress and returns to a higher value immediately afterward. In the second, resistivity again may decrease with increasing stress but, in contrast to the first type of changes, it may decrease further upon the stress drop (Brace and Orange, 1966; Fitterman, 1976; Byerlee and Wyss, 1978). In recent years, people try to explore related earthquake anisotropic resistivity changes in association with the stress level status in/nearby the earthquake focal regions. Here, we attempted to estimate apparent stress and stress drop variations from source parameters of 2006, 2007 and 2011 earthquakes and correlate with the observed changes in MT parameters in this paper.

4.6. Estimation of stress drop Global dependence on stress drop plays a pivotal role in changing the earth physical properties due to crustal deformation processes. Stress drop might be static and dynamic. There is a difference between static stress drop and dynamic stress drop. Dynamic stress drop is the stress available effectively to drive fault motion through the ground, whereas static stress drop (source size or final displacement) is the difference between the stress level before and after the earthquake (Allmann and Shearer, 2009). Several methods have been adopted to estimate static stress drop, apparent stress from dynamic source parameters like radiated energy, corner frequency, source radius, moment magnitude using various empirical relationships (Brune, 1970; Kanamori and Anderson, 1975; Hanks and Kanamori, 1979; Baltay 211

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Fig. 5. Phase tensor ellipse plots for sites recorded before (S07, S08, S12, S13, and S14) and after (SK07, SK08, SK12, SK13, and SK14) the 2006 earthquakes. Ellipse orientation represents the strike direction and color indicates the skew angle values in degree. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

Fig. 6. Static shift values for all the five sites recorded before (S07, S08, S12, S13, and S14) and after (SK07, SK08, SK12, SK13, and SK14) the 2006 earthquake.

et al., 2011; Kumar et al., 2016; Kanamori and Heaton, 2000; KeilesBorok, 1959, see Eqs. 1 and 2) in seismically active areas. Brune's static stress drop (∆σ) which is more commonly applied is having the interrelationship with earthquake source parameters (Moment (M0), Corner frequency (fc), Shear wave velocity (Vs), Source radius (r), Rupture area (A) and Rupture width (D).

= 2 Es /Mo

(1)

= 7M0 /16r3

(2)

data was acquired for calculation of static stress drop. Seismic moment (M0) is calculated from the area of rupture (A) and width of fault (D) i.e. M0 = μAD (Brune, 1970). The obtained seismic moment is ranging from 1.0 × 1023–2.8 × 1026 dyne.cm. The value of corner frequency (fc) 0.1–1.53 Hz and source radii (r) 0.87–13.42 km are computed from formulae suggested by Hazarika and Ravi Kumar (2012) and Kumar et al. (2016) respectively. Area of rupture for the 5.3 Mw, 2006 earthquake is of about 15.99 km2 with 5.65 km rupture length and 2.82 km rupture width and slipped by 0.22 m. Ruptured the area of about 8.12 km2 during the 2007 earthquake and 530 km2 during the 2011 earthquake and caused severe crustal damage within the radius of 60 km. The value of source parameters obtained for the 2011

From the possible empirical green's relations, dynamic source parameterization has been done for the earthquakes occurred after the 212

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earthquake-induced ground motions (Baltay et al., 2013; Lior and Ziv, 2016) i.e.

Table 3 Regional velocity and density model for stress drop analysis. Depth (km)

Density (kg/m3)

Vs (km/s)

μ(Ν/m2)

0–2 2–16 16–39 39–59 >59

2670 2710 2900 3100 3300

3200 3470 3610 4100 4410

2.73E+10 3.26E+10 3.78E+10 5.21E+10 6.42E+10

Arms = Arms

fc (4)

fmax

where Arms is the root mean square acceleration, ρ is the density, RθΦ is the radiation pattern, fc is the corner frequency, fmax is the observational upper-frequency limit, and R is the hypocentral distance (geometrical spreading). Simply speaking, the energy released during an earthquake travels as a seismic wave. Earthquakes with higher radiated energy and rupture velocity will have more intense ground motions. The variation of earthquake-induced ground motions depends on crustal conditions and source to site distances are thus related to the variability of stress drop (Causse and Song, 2015; Courboulex et al., 2016). The attenuation of ground motion relation is important to model the stress drop variation (Oth et al., 2017). Arms method is a better estimator for stress drop variation which is less applied and having less uncertainty because of the square root of rather than cubed fc. 10% in fc uncertainty will only lead to 5% relative error in stress drop. Hence, we have examined the variability of stress drop based on the root mean square acceleration Arms - stress drop (Δσ) method to understand the observed change in apparent resistivity after 2006, 2007 and 2011 earthquakes and aftershocks. Initially, the range of frequency/period at which impedance responses recorded at each MT site is transformed to depth (h) using Nibblett and Sayn-Wittgenstein (1960) code. The estimated depths were constrained by constructing a 2D model. The 2D modeling was carried out using WinGLink software package. Firstly, 2D inversion was performed with both data sets acquired before 2006 and after 2006, 2007 and 2011 earthquakes. Reasonable data points from the Rhoplus consistent test for each site were considered for 2D inversion. Rotated the data by −100 (N100W) (Patro and Harinarayana, 2009) and considered xy component as TM mode (i.e., N-S electric field). The model grid consists of 66 rows and 75 columns with irregular grid spacing. 100 Ω.m was considered as homogeneous half space for the starting model. An error floor of 20% was assigned for TE and TMmode apparent resistivity data and 5% for impedance phases. The regularization parameter (τ) was kept as 10 to generate a smooth model. Influence of static shift reduces on data in the 2D inversion with assigned larger error floor value for apparent resistivity than the impedance phases error floor value. The final models (rms = 2.12; rms = 1.967) obtained after 90 iterations are presented in Fig. 7a & b. RMS error for each site of both data sets is shown in Fig. 8. In order to demarcate the depth extent of the model, the sensitivity of the model was examined. We start with the base 2D models (Fig. 7a & b) and the resistivity values are replaced with 500 Ω.m from the depth of 10 km and below (see Fig. 7a1 & b1). These models were again inverted with the same parameters as that of base model. We notice in the final models that the region below 10 km is sensitive to the data (see Fig. 7a2 & b2). Modeling was performed again with 500 Ω.m half space

earthquake consistent with the values estimated by Paul et al. (2015). Regional velocity and density model (Table 3) prepared in the depth range inferred from gravity and seismic studies (Tiwari et al., 2006; Singh et al., 2010; Acton et al., 2011; Ansari et al., 2014) for the area of investigation to utilize in estimation of the stress drop and its variation for these earthquakes. Assuming the constant rupture velocity (Vs) at the source region, determined the Brune static stress drop (∆σ) for 2006, 2007 and 2011 events (Table 4) from ∆σ-M0-fc relation (Brune, 1970; Aitchison et al., 2007, see Eq. 3) i.e. (3)

= M0 f c 3 (2 /2.34 x Vs)3

106 R 2R (2 )2

6.64 MPa static stress drop is obtained for 2006 earthquake, 6.15–8.55 MPa for 2007 earthquake, and 4.13–7.43 MPa for 2011 earthquake, whereas 1.42–1.98 MPa accumulated apparent stress (σa) before 2007 earthquake and 1.08–1.72 MPa before the 2011 earthquake estimated from scaling relationship σa = ∆σ x 0.23 (Savage and Wood, 1971). Paul et al. (2015) estimated 4.4 MPa stress drop for the 2011 earthquake is comparable with the value obtained in this study. One basic factor that may cause the stress drop to vary is changing physical properties of the crust, particularly shear wave velocity and density with depth around rupture area (Goebel et al., 2015). We found the Brune static stress drop in this area is varied with the moment, and cubed corner frequency, but independent on the proportional relationship, i.e., M0fc3 for 2006, 2007 and 2011 earthquakes. Besides the corner frequency and seismic moment, the variability of stress drop is also noticed with variation in shear wave velocity at the rupture and source duration. The Different values of shear-wave velocities (3.47 km/s and 4.2 km/s) of the medium at a hypocentral depth of earthquakes in the study area exhibited a difference in stress drop. Five MT sites acquired before 2006 earthquake are situated at different epicentral and hypocentral distances with different level of depth of penetration. Variation of Brune static stress drop only concerning with hypocentral depth could not help to understand the observed change in resistivity variation at the acquired data sites. Practically, most of the studies using seismograms to determine stress drop are based on the fc determination. But, corner frequency is very difficult to estimate accurately and errors in the corner frequency will strongly influence the variability of Brune stress drop due to the term fc3. To avoid this strong dependence on corner frequency, Hanks (1979) derived Arms for dynamic stress drop relation by utilizing source parameters and

Table 4 Details of source parameters and static stress drop for 2006, 2007 and 2011 earthquakes. Year

Mw

FD (km)

A (km2)

M0 (N.m)

fc(Hz)

r (km)

Δσ (MPa)

σa (MPa)

2011 2011 2011 2011 2011 2007 2007 2006

6.9 4.5 3.9 4.2 5.0 5.0 4.7 5.3

19.7 9 29 28 16 17 12 12

591.6 2.6 0.7 1.3 8.1 8.1 4.1 16.0

3.07E+19 7.87E+15 1.19E+15 3.30E+15 4.95E+16 4.95E+16 1.54E+16 1.18E+17

0.09 1.67 3.22 2.26 0.88 0.88 1.32 0.65

14.10 0.77 0.42 0.59 1.51 1.51 0.97 1.96

4.72 7.49 7.20 6.90 6.16 6.16 7.29 6.69

1.08 1.72 1.66 1.59 1.42 1.42 1.68 1.54

(Mw = Moment magnitude; FD = Focal depth; A = Area of rupture; M0 = Seismic moment; fc = Corner frequency; r = Source radius; Δσ = Static stress drop; σa Apparent stress) 213

=

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Fig. 7. The geoelectric structure model from 2D joint inversion of TE and TM mode data (a & b) for all five MT sites acquired before and after 2006 seismic activity (a & b). 2D base model with 500 Ω.m half space below 10 km depth (a1, b1). 2D inversion result for a1, b1 (a2, b2). 2D base model with 500 Ω.m half space below 15 km depth (a3, b3). 2D inversion result for a3, b3 (a4, b4). 2D base model with 500 Ω.m half space below 20 km depth (a5, b5). 2D inversion result for a5, b5 (a6, b6). Data points with red circles are at deeper depth than the ascertained depth boundary. MBT- Main boundary thrust, MCT-1, MCT-2 -Main central thrust. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

214

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Fig. 8. RMS error for sites recorded before (S07, S08, S12, S13, and S14) and after (SK07, SK08, SK12, SK13, and SK14) the 2006 earthquake.

Fig. 9. Low-frequency level Ground displacement vs.hypocentral distances for xy and yx components of five MT sites acquired after 2006 earthquake.

given on the base models from the depth of 15 km and below (Fig. 7a3 & b3) to examine the data sensitivity beyond 15 km. The final model after inversion, suggest that the data sensitivity is restricted within the 10 km below S07-SK07 and S08-SK08 sites (Fig. 7a4 & b4). But the data sensitivity noticed to be percolated below S12-SK12, S13-SK13, and S14SK14 from the development of conductivity/resistivity features after 15 km depth (Fig. 7a4 & b4). Continuing the modeling efforts further, 500 Ω.m given half space on the base models from the depth of 20 km and below (Fig. 7a5 & b5) for further examination of data sensitivity to determine maximum depth penetration boundary. Final models after inverted shows a decrease in the sensitivity of data and restricted beyond the depth of 28 km for model 7a (Fig. 7a6) and 26 km for model 7b (Fig. 7b6). Based on the number of runs and models, the maximum depth of penetration of data for each site was ascertained and compared the final ascertained depth boundary values with the maximum depth calculated from of frequency to depth transformation. The ascertained depth boundary along with the calculated depth of each site is shown on the 2D models (Fig. 7a6 & b6). Except for the depth of few data points of the longer period shown with red circles in Fig. 7a6 and b6, depth of all the data points at all the data sites are confined within the ascertained depth boundary. The focal depth ranges of the main and aftershocks of 2006, 2007 and 2011 earthquakes in the area of study were estimated of about 9–30 km. Therefore, from the demarcation of depth boundary, we kept the depth boundary for maximum depth of penetration as 30 km for all the data sites and considered as final depth boundary for the data analysis in this paper. The calculated depth for all the data points beyond the final depth boundary is not taken into consideration. After finalization of estimated depth values of each data points from 2D models, calculated the hypocentral distance (R) for each data point of each site from the hypocenter depth of each earthquake. The depth of hypocenter is constant, but the hypocentral distance (R) for each MT

site may vary depending on its the epicentral distance from its epicenter position which is vertically above the hypocenter on the surface (see Fig. 13 for hypocenter and hypocentral distance). Hypocentral distance (R) is computed from the relation of epicentral distance and depth of penetration of each data point for each MT site as R = √∆2 + h2. (∆ = epicentral distance, h = depth of penetration). Then, Arms values are estimated from (Lior and Ziv, 2016) equation:

Arms = (2 ) 2

0

(

fc 2

kT 1 +

kfc

1.50.25

)

2

(5)

where, Ω0 the low-frequency plateau level of the displacement spectrum, fc is the corner frequency, T is the data interval (R/Vs + 1/fc), and k is the kappa (local site attenuation parameter). The spectral parameters Ω0 and fc hold fundamental information regarding the physical attributes of the earthquake source. The spectral plateau of the low-frequency ground displacement spectrum Ω0depends on seismic moment, shear wave velocity and crustal density distribution of the earthquake affected region. The behavior of the low-frequency level of displacement spectrum along the path is obtained towards each site from the functional relation (Brune, 1970) 0

=

M0 R FG (R) 4

3

(6)

M0 (seismic moment), Rθϕ is the radiation pattern (Rθϕ = 0.65), G(R) is the geometrical spreading factor (G(R) = R−1 for R < 70 km), R is hypocentral distance, ρ is the density, and F is the free-surface amplification (F = 2). The ground displacement with hypocentral distance (Fig. 9) shows that the maximum ground displacement is observed towards the site occupied with shorter hypocentral distance from each considered earthquake hypocenter and attenuated with an increase in distance. 215

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Fig. 10. Comparison of Arms values vs. hypocentral distance plot obtained from two different approaches for xy and yx components of five MT sites acquired after 2006 earthquake.

The difference in ground displacement values for xy and yx component as well as for each site is due to the difference in depth of penetration of xy and yx component data points and hypocentral distance of each site. Fmax depends on the Brune static stress drop and seismic moment M0. Fmax can directly be computed from Fmax = 7.31x103M0–0.12 (Faccioli, 1986) and Δσ = −2.91logfmax + 5.8 (Satoh et al., 2000). Here, fmax is computed from the relation Δσ = −2.91logfmax + 5.8 using Brune static stress drop at source. fmax also depends on the local site attenuation parameter (kappa) k. The k value within 4 km depth range at all sites is found to be 0.002–0.006 s estimated from k = 0.145–0.12ln (Vuc) (1.65 km/s ≤ Vs ≤ 3.35 km/s (Campbell, 2009; Allen, 2012). Considering local site attenuation parameter, variation in fmax is computed for the velocity ≤ 3.35 km/s at each site using equation fmax = k-0.0258/−0.0005 (Kumar et al., 2014), then Arms values are measured as per the Eq. (5). Peak ground acceleration (PGA) is one of the most commonly used ground motion parameters. Earthquake stress drop is having a strong relation to the PGA, i.e., ∆σα PGA1.25 (Hanks and Johnson, 1976). Cotton et al. (2013) also examined the link between variability of stress drop to the PGA as σLn (∆σ) = (6/5) σLn (PGA) (Sigma (σ) = Standard Deviation). These relations suggest that the variability of PGA is directly connected to the variability of stress drop. Arms having an established relation with Peak ground acceleration (PGA) (McGuire and Hanks, 1980).

Arms =

motion attenuation relationship derived by Nath et al. (2005) for Sikkim Himalaya at each site (see Fig. 11) without considering local site conditions.

Ln (PGA) =

( ) 2fmax fc

1.08 ln R + 0.007R

(8)

(R < 100 km) (M = moment magnitude, R = hypocentral distance) PGA values are varied with hypocentral distances in the range of 0.02–0.38 g for 2006, 0.04–0.09 g and 0.1–0.26 g for 2011 seismic events. The surface PGA values 0.12 g at SK08 (Rangpo), and 0.16 g at SK14 (Dikchu) obtained for 2011 earthquake in this study are closer to the values 0.15 g (Rangpo), and 0.24 g (Dikchu) estimated from ground motion studies for 2011 earthquake (Raghukanth et al., 2012). Calculated Arms values from Arms - PGA relation (7) to validate the Arms values obtained from the Eq. (5). Both results found a good consistent for all earthquakes at each MT site (Fig. 10). We confirm that the stress drop variation can be predictable from PGA values which are closely matched with the Arms values obtained from Eq. (5). Therefore, The PGA attenuation models presented here (Fig. 11) to compare Arms based stress drop variation for each site. Finally, stress drop variation values deduced from the ∆σ-Arms (equation 4) method for each site and for each depth level after applying the correction of geometric spreading and site effects. Stress drop variation observed during the 2006 earthquake is ranging from 2.5 to 4.9 MPa (Fig. 12a). The 2007 earthquakes dropped the stress level in the range of 4.5–6.5 MPa (Fig. 12b) and 7.6–12.4 MPa for 2011 earthquake along with their aftershocks (Fig. 12c). Observed 7.2–12.8 MPa range of stress drop in this study well agree with the estimated stress drop ranging from 5 to 20 MPa for 2011earthquake (Raghukanth et al., 2012; Paul et al., 2015). We noticed that stress drop depends on the focal mechanism and 2–3 times higher stress drops for strike-slip

PGA 2 ln

3.6 + 0.72M

(7)

In the absence of strong ground motion data, we estimated PGA values (geometric mean of the horizontal component) using ground 216

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Fig. 11. Contour map of PGA values for the five MT sites ((a) - xy component; (b) - yx component) acquired after the earthquakes 2006 (Red Star), 2007 (Yellow star) and 2011 (Blue star). (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

earthquakes than thrust type. Stress drop variation values derived from the ∆σ-Arms method are much lower than those derived from Brune static stress drop values. The reason may be due to variation in stress drop depends on attenuation of ground accelerations with hypocentral distance for each site. The shorter hypocentral distance of data points affected higher stress drop and decreased with increasing hypocentral distance. The Source duration or Duration of rupture propagation is represented as Td = T0 + bR, where T0 (=1/ fc), is the source duration at the origin and b is the coefficient that controls the increase of duration with distance (Atkinson and Boore, 1995; Allen, 2012). We

examined that the stress drop is also varied with source duration Td of an earthquake with hypocentral distance R (Fig. 13). A hypothesis of stress drop variation increases with hypocentral depth, because of an increase of crustal stresses at greater depths and lower the stress drop due to slower rupture velocities at shallow part (Allmann and Shearer, 2009) is ascertained from the stress drop variation results. Given all above, the region where the pre-seismic apparent resistive data was acquired may be experienced different stress drops at different depth levels and varied based on attenuation of strong ground motions with hypocentral distances from the 2006, 2007 and 2011 source regions. 217

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Fig. 12. a. Variation of Stress drop of the earthquake event 2006, 5.3 Mw and its aftershocks (Black circles) for five MT sites ((a) - xy component; (b) - yx component) with hypocentral distance. b. Variation of Stress drop of combined earthquake events 2007 (5.0 Mw, 4.7 Mw, 3.5Mw) for five MT sites ((a) - xy component; (b) - yx component) with hypocentral distance. c. Variation of Stress drop of the combined earthquake events 2011 (6.9 Mw, 4.5 Mw, 3.9 Mw, 4.2 Mw, 5.0 Mw) for five MT sites ((a) - xy component; (b) - yx component) with hypocentral distance.

4.7. Observed and calculated apparent resistivity data

penetration of ρxy component data is also different from the ρyx component at each site and even shallower than the ρyx. Therefore, we executed the 2D models for TM-mode (ρxy component) and TE-mode (ρyx component) separately from the data acquired before (Fig. 15a & c) and after (Fig. 15b & d) earthquake activity. Geo-electrical features from these models observed to be different from each other. The observed difference in apparent resistivity of xy and yx components for each site also plotted with their hypocentral distance and depth of penetration of data points. ∆ρxy plot represent xy component data (Fig. 16a) and ∆ρyx plot represent yx component data (Fig. 16b). It is noticed from Figs. 14, 15 and 16 that, the depth of penetration of the data for five sites was confined within the estimated focal depth range

Observed apparent resistivity are two individual components (ρxy, ρxy) recorded in MT data. The response of each component is observed to be different. Firstly, we constructed 2D models from joint inversion of ρxy and ρyx (TM = xy; TE = yx) component data for both the data sets recorded before and after 2006 earthquake (Fig. 14a & b). The depth of penetration of xy and yx components of each site are averaged for each period and represented with color circles. Considering the response of ρxy and ρyx components before and after earthquakes separately, the observed change in apparent resistivity for the ρxy component is appeared to be different from ρyx component data. The depth of 218

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Fig. 13. Source duration Vs. Hypocentral distance plot for 2006 (left) and 2011 (right) earthquakes at five MT sites.

Fig. 14. The 2-D geoelectric model derived from joint inversion of TE and TM mode data for all five MT sites acquired before 2006 and after the 2006 seismic activity. Average depth of penetration of xy and yx components data shown in color circles for each period with color scale. MBT-Main boundary thrust, MCT-1, MCT-2 -Main central thrust.

of 9–30 km. Interpretation from 2D inversion models (Fig. 14a & b); Geoelectric conductive/resistivie features are seems to be different from each other. Model 14b characterized with the conductive zone (C1) beneath SK13 and SK14 sites compared to model 14a and reveal that the resistivity of the subsurface is decreased after the earthquake activities. The observed difference in apparent resistivity for yx component is larger than the xy components and indicate a decrease in resistivity (values with negative sign) below sites SK14 and SK13 (see Fig. 16b). The feature of decrease in observed apparent resistivity of yx component coincide with the increase in conductivity zone delineated from 2D model demonstrated with TE (= yx) mode data (Fig. 15d). Earthquakes generally occur in stress-enhanced areas, and resistivity is highly sensitive to stress/strain changes. Changes in resistivity of rocks dependence on stress variation (Brace, 1975; Barsukov, 1972) and stress/resistivity relation (Johnston, 2002) formulated as:

/ = Kr

stresses that had been developed before the occurrence of 2007 and 2011 earthquake are also taken into consideration and subtracted from stress drop values for individual earthquakes. Then, final stress drop values from each earthquake are integrated to find a correlation with observed change in apparent resistivity. Fig. 17 shows the integrated stress drop variation values for xy (∆σxy) and yx (∆σxy) components of five MT sites. Estimated stress drop is observed to be greater near SK13 and SK14 and decreased towards SK08 and SK07 which were hypocentrically far from the earthquake source. The observed difference in apparent resistivity is also maximum at SK13 and SK14 and reduced near SK08 and SK07 (Fig. 16a & b). The distribution of stress drop variation is unequal in ∆σxy and ∆σxy and the range of stress drop variation 11–18.2 MPa for yx component is relatively higher than xy component (11.5–17.9 MPa). The higher range of associated stress drop values for yx component at SK13 and SK14 strongly supports the behavior of decrease in observed apparent resistivity (Fig. 15d & 16b). Further, relative change in apparent resistivity (Δρ/ρobs) for xy and yx components is derived from observed data. Computed predicted values of Δρ/ρpre for xy and yx components using Kr and Δσ from the relation (see Eq. (9)) (Atkinson and Boore, 1995). Both Δρ/ρobs and Δρ/ρpre values are plotted in same graph against the estimated stress drop (Δσ) (Fig. 18a). If Δσ changes then, ρ2 also changes and thus ρ2 ≠ ρ1. In this case, the Δρ (ρ2-ρ1) gives the difference. The difference (Δρ) must be positive and negative. If ρ2 increases than the ρ1, then the Δρ gives positive value and if ρ2 decreases than the ρ1, then the Δρ gives negative. The relative change Δρ/ρ can also gives positive and negative values depending on the Δρ. From the concept of decrease in resistivity (ρ2 < ρ1, Δρ = −ve) with an increase in stress drop, the predicted

(9)

where, ρ = Initial resistivity, ∆ρ = Change in resistivity, Kr = constant (3 × 10−3 MPa−1), Δσ = Change in stress (stress drop). Therefore, the change in resistivity is a linear relation with ∆σ. From ∆σ-Arms method, different ranges of stress drop values are estimated from each earthquake. Stress drop values also attribute towards the amount of accumulated stresses σa = ∆σ x 0.23 (Savage and Wood, 1971). The concept of change in resistivity with stress variation suggests that the resistivity may decrease initially with increasing stress and decrease further upon the stress drop (Brace and Orange, 1966; Fitterman, 1976; Byerlee and Wyss, 1978). In this study, we examine the correlation of observed change in apparent resistivity with stress drop after seismic events. Hence, additional accumulated apparent 219

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Fig. 15. The geoelectric model from 2D inversion for TE and TM mode data for all five MT sites acquired before and after 2006 seismic activity. (a) 2D model from TM mode data acquired before 2006, (b) 2D model from TE mode data acquired before 2006, (c) 2D model from TM mode data acquired after 2006, (d) 2D model from TE mode data acquired after 2006 seismic activity. Depth of penetration of xy and yx components data shown in color circles for each period with color scale. MBT-Main boundary thrust, MCT-1, MCT-2 -Main central thrust.

Fig. 16. Observed difference in apparent resistivity of xy and yx component data Δρxy (a) and Δρyx (b) at five MT sites acquired before and after 2006 seismic activity. Stars represent the seismic events with focal depth in the study area. (2006, 5.3 Mw (Red), 2007, 5.0 Mw, 4.7 Mw, 3.5 Mw (yellow) and, 2011, 6.9 Mw, 4.5 Mw, 3.9 Mw, 4.2 Mw, 5.0 Mw (Blue), black circles represent aftershocks of 2006 event). (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

values of Δρ/ρpre are shown in negative to compare the relative change in observed apparent resistivity (Δρ/ρobs). In Fig. 18a; the Δρ/ρobs and Δρ/ρpre are confined within the ±1 except for few data points of Δρ/ρobs at site SK08. The predicted change in Δρ/ρpre is lesser than the Δρ/ρobs (see Fig. 18a). To see the clear variation, the Δρ/ρpre values are plotted in zooming mode with ±0.2 scale. Red and black colored lines indicate xy componenet and yx componenet data of Δρ/ρpre. The predicted change in both components of Δρ/ρpre for all five sites showed a straight line with Kr. The slope of five straight lines is limited to −0.2 to 0.2 and varied with Δσ. The negative values of Δρ/ρobs for SK07 fall below the straight line. The Δρ/ρobs values for SK12 fall above and below the straight line with positive and negative. But, for SK13 and SK14, most of Δρ/ρobs values fall below the straight line with negative values against Δσ. The negative values of Δρ/ρobs indicate a decrease in ρ2 (post seismic acquired data). It is noticed from Fig. 18a that, the

decrease in Δρ/ρobs related to increase in Δσ. Despite of small scale variation (within ±0.2), the Δρ/ρpre also observed the relation of decrease apparent resistivity with an increase in stress drop (see Fig. 18a). The observed decrease in Δρ/ρobs is minimum at SK07, SK08 due to limited range of stress drop. Limited or shallow depth of penetration of SK07, and SK08 were attracted less stress drop concerning hypocentral distance and failed to fallow the linear relation. The decrease in ρ2 with negative Δρ/ρobs for SK12 related to an increase in stress drop but followed a nonlinear relation. In the case of SK13 and SK14, we can find a linear relation between Δρ/ρobs and Δσ. For all Δσ values, the Δρ/ρobs is attributed a decreasing trend represented with negative values. At SK13 and SK14, both xy and yx components data exhibited a maximum decrease in Δρ/ρobs (negative values) with an increase in Δσ. The linear relation is valid to these sites which were nearby the hypocenter with maximum data points with depth of penetration confined within the 220

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Fig. 17. Estimated Stress drop variation model for ∆σxy and ∆σyx of five MT sites. Stars represent the seismic events with focal depth in the study area. (2006, 5.3 Mw (Red), 2007, 5.0 Mw, 4.7 Mw, 3.5 Mw (yellow) and, 2011, 6.9 Mw, 4.5 Mw, 3.9 Mw, 4.2 Mw, 5.0 Mw (Blue), block circles represent aftershocks of 2006 event). (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

higher stress drop zone of the earthquakes. A feature of decreased resistivity zone (C1) (Fig. 14b) below post-earthquake acquired data sites SK13 and SK14 with higher stress drop values (Fig. 17b) complemented the linear relation. Examined higher decrease in Δρ/ρpre at sites SK13 and SK14 than the other sites also justify the linear relation. To complement the linear relation between Δρ/ρobs and Δσ further, synthetic data were computed from 2D models (Fig. 14a & b) through forward modeling. Relative change in apparent resistivies from model (Δρ/ ρ2D_model) are derived and represented against to stress drop variation values (Fig. 18b). The similarity in the behavior of decrease of Δρ/ ρ2D_model values for SK13 and SK14 with stress drop strengthen the linear relation between Δρ/ρobs and Δσ (see Fig. 18a & b). Therefore, it is inferred that the observed decrease at SK13 and SK14 for both ρxy and ρyx data is having a fair correlation with increase in stress drop values at these sites. The correlation corroborate the observed decrease in apparent resistivity that might be occurred in association with higher stress drop near the epicenter sites after earthquakes within the focal depth range of 4–18 km.

stress levels, and strike-slip focal mechanisms of 2007 and 2011 were caused great variation in stress drop 5–38 MPa in this area (Hazarika and Ravi Kumar, 2012; Raghukanth et al., 2012). Crustal resistivity is thought to change from the deformation of the crust with the diffusion of crustal fluids under the action of variation in stress levels. Johnston (2002) stated that at least 1% of resistivity changes might be expected at 1 MPa of stress change by accompanying crustal failure during an earthquake. The 11–18.2 MPa range of stress drop would be able to reduce 11–18% in observed apparent resistivity at post seismic sites. With this intension, a possible change in apparent resistivity is calculated from stress/resistivity relation using estimated stress drop values for each site. The calculated 11–18% decrease in apparent resistivity is lesser than the observed 10–30% percentage of decrease, but decrease trend of both ρxy and ρyx components data are well comparable. The 10–30% of the observed change in apparent resistivity with greater than the uncertainty <0.1–5% in the data quality for 0.001 s - 10s period range reinforce its significant change. Post seismic restoration of apparent resistivity to pre-seismic level could not be able to observe after 2006 and 2007 by probing MT survey. However, we interpreted that presence of critically sheared/stressed zones in association with relentless compression activity (Bollinger et al., 2004; Raju et al., 2008; Ponraj et al., 2010; Baruah et al., 2016) before and un-restored crustal failure with strong ground motions induced by subsequent micromoderate earthquakes could have been attributed post seismic decrease with higher than estimated stress drop 11–18.2 MPa. Development of micro-cracks in low-high grade metamorphic rocks along active stressed and thrust/shear zones and generation of crustal fluids by metamorphic dehydration may be possible in and around pre-rupture regions. In this condition, diffusion of crustal fluids between compression-dilatants regions (increased stress level) before the earthquake to rupture initiation zones could have been occurred and attributed to the apparent resistivity decrease and further decreased with maximum stress drop. SK07 and SK08 data points were hypocentrally far from the earthquake source region. Apparent resistivity at S07-SK07 and S08-SK08 sites represented a decreasing behavior (1000 Ω.m – 1 Ω.m) towards 0.001–100 s period corresponds to ~ 4-12 km (Fig. 2 & Fig. 14). The 12 km thickness of low-grade meta sediments with aqueous fluids generated in the shearing stress environment beneath LHD attributes to the low resistivity nature (Fig. 14). Observed low S - wave velocity due to fluid filled sediment zone within the depth range of 2-10 km (Singh et al., 2010) also supported the conductive feature delineated from MT study. Little compression in low porosity meta-sediments would raise one to two orders of magnitude in the increase of resistivity. The low porous meta-sediments with aqueous fluids (3–4% porosity) might be stressed little enough and responded to the estimated stress drop during

5. Discussion The acquired MT sites S07-SK07, S08-SK08, S12-SK12 and S14SK14 before and after earthquakes were fallen in the deformed region of LHD domain north to MBT except S13-SK13 which is in MCTZ. The post seismic acquired MT data in this region exhibited a change in apparent resistivity as well as impedance phase at each site in comparison with the data acquired before 2006 (Fig. 2). Observed change (decrease) in apparent resistivity from both MT data sites is in the range of 10–30%. The 2D geoelectric model constructed with post seismic acquired data also exhibit decrease in apparent resistivity (C1) (Fig. 14b) when compared with the 2D model developed from the data acquired before 2006. The overall stress drop variation driven by subsequent seismic slips in 2006, 2007 and 2011 is in the range of 11–18.2 MPa. The active decollement plane which is causing 11–17 MPa shear stresses (Lamb, 2006) overlain by meta-sediments and upper crust bounded by MBT and MCT (Fig. 1). Micro seismic studies reveal that the diffused micro-to-moderate seismicity concentrated primarily between the MBT and MCT. These thrusts were critically enough stressed to be released before the occurrence of the 2006 earthquake (Raju et al., 2008). Examination of source zone and stress regime from 2006, 2007 and 2011 events (Raju et al., 2008; Joshi et al., 2010, and Baruah et al., 2016), it is reported that Sikkim Himalaya and its vicinity, experienced by the direction of horizontal compressive stresses and the main /aftershocks occurred in the increased stress zones. Differences in rupture properties of 2006, 2007 and 2011 earthquakes may also be attributed to differences of dynamic friction or 221

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Fig. 18. a. Relation between relative change in observed apparent resistivity (Δρ/ρobs) before and after 2006, 2007 and 2011 earthquakes and estimated stress drop variation (Δσ) for all five sites. Negative values indicate a decrease in apparent resistivity after seismic activities. Predicted relative change in apparent resistivity (Δρ/ ρpre) obtained from the Eq. (9) (Δρ/ρ = KrΔσ) shown as straight lines with slope Kr = 0.003 MPa (xy-componenet = red color; yx-component = black color) b. Relation of change in apparent resistivity (Δρ/ρ2D_model) before and after 2006, 2007 and 2011 earthquakes derived from 2D models (Fig. 14a & 14b) with estimated stress drop variation (Δσ) for all five sites. Negative values indicate a decrease in apparent resistivity after seismic activities. Predicted relative change in apparent resistivity (Δρ/ρpre) obtained from the Eq. (9) (Δρ/ρ = KrΔσ) shown as straight lines with slope Kr = 0.003 MPa (xy-componenet = red color; yx-component = black color). (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

the ground motion leads to further decrease in resistivity. If it is valid, we can interpret that the variation of apparent resistivity is observed under the action of stress and its variation before and after an earthquake may confirm the zone with fluid presence. The observed moderate resistive (≥ 1000 Ωm) subsurface beneath SK12 site (Fig. 14) was well compare from another MT study (Pavan Kumar et al., 2014). SK12 site was situated close to the main and aftershocks of 2007 (Fig. 1 & Fig. 12b). 4–6 MPa stress drop values obtained from 2007 event near to SK12 comparatively less than the maximum stress drop 38.6 MPa observed during 2007 earthquake (Hazarika and Ravi Kumar, 2012) and caused to decease resistivity somewhat higher than SK07 and SK08. The difference of change in apparent resistivity is decreased towards sites SK12, SK08, and SK07. The low-stress drop values (Fig. 17) towards these sites due to attenuation of ground acceleration concerning epicentral and hypocentral distances (Fig. 9) support the difference in apparent resistivity change. SK13 and SK14 sites were reoccupied at 3–8 km of epicentral distance from 2006 main/aftershocks and 35–40 km from 2011 main/ aftershocks. SK13 and SK14 sites are exhibited a comparatively maximum post-seismic decrease in apparent resistivity (Fig. 14, 16b & 18). These sites are near to MCT, where the maximum shear stress/strain 0.4–0.1 μstrain/yr (Ponraj et al., 2010) is being accommodated. The data at SK13 and SK14 have strongly affected from nearby 16km2 rupture area with 22 cm slip during 2006 and 591 km2 rupture area with 137 cm slip during 2011event. Nearly 3.1–5.2 MPa stress drop observed from 2006 seismic event and 9.2–12.8 MPa stress drop observed from the 2011 earthquake. Higher stress drop values observed towards SK13 and SK14 sites (Fig. 17 & 18). Higher PGA values (0.2–0.38 g) and with shorter duration of rupture propagation (Figs. 11 & 13) support for the maximum stress drop at these sites. High resistive shallow formation (>1000–10,000 Ωm) at sites SK13 and SK14 (Fig. 14 & Fig. 15) replicate the HHC which might be allowed maximum apparent stress to accumulate. The observed maximum apparent resistivity decrease in the depth range of 4–18 km at both SK13 and SK14 (Fig. 14b, 16b & 18) found a fair correlation with the maximum stress drop values (Fig. 17 & 18a). Prominent apparent resistivity change was observed in ρyx (E-W) component data (Fig. 2, 15d & 16b). Fault plan solutions showed that the region was dominated by N200E directive sub-horizontal stress (Raju et al., 2006) during 2006 earthquake and NNW-SSE horizontal compression stress (Baruah et al., 2016) during 2011 seismic event. Joshi et al. (2010) revealed that the rupture initiated in 2007 under the influence of the northward directed tectonic stresses. Shear failure with adequate stress drop oblique to the direction of regional compressional stresses could give a valid explanation for the maximum decrease in ρyx by subsequent seismic activities 2006, 2007 and 2011. Resistivity in the maximum ruptured stress direction might be increased prior to the earthquake as cracks close at the initial stage and might be reduced by enhancing the interconnectivity with deepcrustal fluids within the opened fractures/cracks. The resistivity could have further been reduced with sudden drop of compression/shear stresses at different depth levels in response to strong ground motions. The depth of penetration of ρyx component data was deeper than ρxy. Therefore, more data points with the higher depth of penetration of ρyx component were fallen in the high-stress drop region might also be a possible reason for more decreased behavior than ρxy. Sites for short period data (> 0.001 s) exhibited an increased apparent resistivity phenomenon. Local crustal disturbances and anisotropic attenuation of

ground acceleration might be caused for rearrangement stresses in the shallow crust support the observed results. The site SK13 which is in MCTZ reported a maximum change in MT parameters compared to S13. Regional strike direction calculated in the period range of 0.0001–0.01 s at SK13 and S13 both showed a distinct pattern. SK13 exhibited elongated and complex nature of polar diagrams and phase tensors with high skew values which are different from the pre-earthquake acquired S13 site (Fig. 3, 4 & 5). Structural and geological disturbance with strong ground motions within the hypocentral distance range also might be influenced the dimensionality parameters. Strike direction obtained from SSMF analysis at sites falling in LHD domain showed NNW-SSE direction, but the strike direction for the site SK13 in MCTZ zone is NNE-SSW indicates complex structural features. Distortion of phase tensors with large skew values and rotation strike angles replicate the anisotropy and heterogeneity in MCTZ. All these suggest that the tectonic setup of the Sikkim Himalaya within LHD and MCTZ being significantly different regarding tectonic setup, geological strikes, and structural complexities. 6. Conclusion The apparent resistivity of MT data sites acquired after 2006, 2007 and 2011 earthquakes attributed an observed change (decrease) is in the range of 10–30%. Strike direction, Polar diagrams, and Phase tensor ellipses support the observed impedance change. The feature of low resistivity zone recognized from geoelectric model with joint inversion of TE and TM mode data acquired after these earthquakes. The observed resistivity change is related to combined stress drops at different depth levels in association with strong ground motions induced from 2006, 2007 and 2011 events. The correlation between relative change in observed apparent resistivity and estimated stress drop corrobarate the relation. Alternatively, strong ground motions which caused for greater stress drop level could have been migrated the crustal fluids from compressional to dilatation zones and altered the mechanical and electrical properties resulted in change in resistivity. Hence, the conclusion assumes that the study region near the main rupture zone/ aftershocks might be critically more stressed compared to other regions before 2006 earthquake and resistivity could have been decreased by stress drop and gradually recovering its pre-seismic values. The observed change in apparent resistivity is different from site to a site related to the variation in stress drop values depending on attenuation of Arms as well as source duration with hypocentral distances. The results described in this study implies that more examples with much better accuracy will be required to observe such changes and may also need to be investigated using long-term observations. Strike, polar plot and phase tensor analysis results suggest that the study area is under the situation of three-dimensional. From this point, we will expand our study to a three-dimensional analysis including the complete data set to reveal the detailed, complex structural features and geo-electrical signatures in and around the focal regions of the area. Acknowledgment We are grateful to the Director NGRI for constant encouragement and support. Sincere thanks to the Department of Science and Technology and Ministry of Earth Sciences, Government of India, for funding the study. Thanks also to K. Chinna Reddy and Narendra Babu 223

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for their participation in the tough field campaign. Permission to access the field sites and logistic support extended by the Sikkim government during the field campaign is highly acknowledged. PKP's research is supported through MLP-6404-28 (BPK) at CSIR-NGRI. We are grateful to both the anonymous reviewers and the editor Rob Govers for their constructive suggestions which have improved the clarity in the manuscript.

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