High-frequency rupture properties of the Mw 9.0 off the Pacific coast of Tohoku Earthquake
© The Society of Geomagnetism and Earth, Planetary and Space Sciences (SGEPSS); The Seismological Society of Japan; The Volcanological Society of Japan; The Geodetic Society of Japan; The Japanese Society for Planetary Sciences; TERRAPUB. 2011
Received: 13 April 2011
Accepted: 14 July 2011
Published: 27 September 2011
The devastating Mw 9.0 earthquake on March 11, 2011 is one of the most complex earthquakes of all the recent large events. Its source region is compact for an earthquake of this size, but it has highly variable amount of energy release from different segments. These conditions prevent conventional back-projection analysis to reveal the details of the rupture process. We incorporate a new metric to assess coherency as a part of back-projection analysis to ensure identification of these weak features. The main features obtained with this new back-projection approach are consistent with previous back-projection results, with strongest energy release downdip and close to the epicentral location. The main rupture propagation is along strike, in southwestern direction giving rise to the length extent of the earthquake. The new coherency function also allows us to investigate rupture characteristics at the beginning of the earthquake, resolving initial updip propagation from the epicentral location. Furthermore, some of very weak energy in the stacks are identified with high coherency. These additional source regions extend the area of the earthquake farther south and north than the region that has been imaged by other back-projection studies.
The giant earthquake on March 11, 2011 (off the Pacific coast of Tohoku) released high-frequency seismic energy that are strongly dependent on both space and time. This poses a challenge for back-projection analysis (e.g., Ishii et al., 2005), because high-amplitude energy may obscure weak energy release. This was an issue when the direct P waveforms are combined with depth phases for intermediate-depth earthquakes (Kiser et al., 2011). In this case, the high-amplitude P arrivals conceal depth-phase arrivals, and the authors apply a taper to suppress P-wave artifacts. This approach would be difficult to apply in the current case where the P waveform itself is rich in amplitude variation. To overcome this problem, and to capture weak, and yet potentially important, energy releases, we introduce coherency function where the coherency of individual traces are monitored as a function of location and time. In this paper, we combine the coherency function and back-projection stacks to investigate the detailed high-frequency rupture characteristics of the March 11, 2011 M 9.0 main-shock.
This gives average cross correlation value of individual seismogram and stacked trace as a function of space (grid point i) and time. The time window T over which cross correlation is calculated, should reflect how long energy may radiate from a given location as well as ensuring that there is enough waveform within this window for meaningful cross correlation calculation. This implies that the coherency function is most useful for high-frequency data, and not so much for low-frequency analysis.
The value of the coherency function is expected to be large for identifying the location and time of energy release. This is because the coherency function is not affected by the strength of various energy release, hence weak energy release can be identified as robustly as strong energy release. On the other hand, the coherency is expected to degrade as one moves away from location where the seismograms have been aligned due to changes in lateral structure. Furthermore, if the source is distributed over a diffuse area, the coherency function will be low. Despite these issues, rupture locations are determined more reliably with the coherency function, hence we use this function to estimate spatio-temporal variations in rupture using high-frequency data. We rely on the stacked traces to obtain relative amplitude information.
One important component of back-projection analysis is the station-specific correction (Ishii et al., 2005) that accounts for wavefront distortion due to lateral variations and station error (e.g., timing issue). If this correction is not properly determined, spurious features can be introduced. Because of the narrow frequency filter used and complexity of the initial few seconds of P waveform, cycle skips are commonly observed with the cross correlation procedure for the Mw 9.0 mainshock. Instead of relying upon such questionable alignment, we utilize waveforms of the Mw 7.1 aftershock from April 7, 2011. This earthquake has clean first arrival, and can be easily cross correlated and aligned. Consequently, all results presented in this paper are with respect to the hypocentral location of this aftershock, i.e., 38.253°N 141.640°E and 49.0-km depth as provided by the National Earthquake Information Centre (NEIC).
To investigate if the rupture of the 2011 Tohoku earthquake shows frequency dependence as observed for the 2010 Mw 8.8 Maule, Chile earthquake (Kiser and Ishii, 2011a), we filter the data to two narrow frequency bands. We refer to the two frequencies as high (0.8 to 2 Hz) and low frequency (10 to 20 second period). The low-frequency back-projection results do not have spatial or temporal resolution to identify details of the rupture process due to long period and long wavelength. We, therefore, analyze the high-frequency results to characterize the rupture, then compare main features of high-frequency results to low-frequency results. Note that the back-projection approach provides location and timing of relative energy release during an earthquake, but cannot constrain parameters such as dip of the fault or absolute slip (e.g., Ishii et al., 2005, 2007).
4.1 High-frequency back-projection results
Assuming that the epicentral location of the earthquake used to obtain the station-specific corrections Δt k is correct, we can use the coherency function to estimate the epicentral location and time of the mainshock. The epicentral location from back-projection analysis is at 142.8ΔE and 38.0ΔN, and the event begins at 05:46:28 UT. We estimate that the uncertainty in longitude and latitude to be about ±0.2°, and uncertainty in time to be about ±7 seconds. The hypocen-tral time of the earthquake from both NEIC and Japan Meteorological Agency (JMA) is 05:46:23, within uncertainty of the back-projection estimate. The two catalogues give somewhat different epicentral locations of 142.369° E 38.322°N (NEIC) and 142.86° E 38.103°N (JMA). The back-projection location is closer to the JMA determination than the NEIC solution.
The second subevent, spanning a time between 40 to 120 seconds from the beginning of the event and releasing the largest amount of energy, is the most complex of all the subevents. Much of the energy release occurs in the first 60 seconds of this subevent within a large area between about 37.5°N and 39°N (Fig. 2). This area covers much of the seismogenic zone at this location, and includes the area that ruptured during the first subevent around the epicentre. Energy release is highest within the northern portion of this area and downdip of the epicentral location. The complexity of this subevent is not obvious from Fig. 2, but if the energy release locations are monitored, they move around considerably, sometimes rupturing a region that have already released energy (e.g., first subevent area). Near the end of this subevent, starting around 100 seconds since the event initiation, the back-projection results show bilateral rupture propagation. This propagation is nearly parallel to the trench, and one front moves northeast, up to about 39.5°N, and the other front moves southwest, down to about 37.5°N (Fig. 2). The last part of the northeastern propagation is somewhat suspicious. The propagation of back-projected energy is very similar to that expected from imperfect data coverage. Even though coherency is above the background level, part of this northeastern finger may be a spurious feature, and that the actual northeastern rupture may stop at about 39°N. The southern propagation of the bilateral rupture encounters a patch that does not participate in the rupture during the entire mainshock (Fig. 2). This patch is very cleanly imaged in the coherency function, but is also visible in the stack. The rupture literally moves around this aseismic patch, and initiates rupture in a region south of the second subevent.
The third subevent, starting with the rupture that propagates around the aseismic zone, is much weaker compared to the second subevent (Fig. 2), and is a very brief episode with duration of about 25 seconds. The initial energy release south of the aseismic patch is the strongest, followed by very diffuse energy propagating slightly updip. This subevent is followed by about 15 seconds of quiescence where no significant coherency is observed within the studied area.
At about 140 seconds, there is a subevent that has the characteristics of a point source within the limitation of the current technique and data set (Fig. 2). It occurs within the rupture area of preceding subevent 3, at about 141°E and 37°N. It may be the very first aftershock of the mainshock.
The fifth subevent, and the last event involving significant area, starts about 145 seconds and lasts for about 35 seconds. This event covers an area south of subevent 3 with practically no overlap (Fig. 2), and has very weak, diffuse energy release that is difficult to identify in normal back-projection stacks. However, the coherency is above the background level, suggesting that the weak energy seen in the back-projection stacks are indeed real. The main propagation direction during this subevent is updip, toward the trench. In fact, it appears that it may even reach the trench location, although the amplitude of high-frequency energy release is almost at the background level by the time the energy arrives at the trench.
The high-frequency mainshock rupture is nearly done with the fifth subevent except for another point-source-like event at about 195 seconds. This final event occurs at the edge of the rupture areas of subevents 3 and 4, at about 37°N and 142.8°E. Again, because of its size and proximity to previously ruptured area, it may be another early aftershock.
4.2 Frequency dependence
The relative source-time function at high and low frequencies considered in this paper are, in general, consistent with one another (Fig. 3(a)). The large energy release occurs at about 80 seconds into the event, and the duration is about 210 to 220 seconds. Despite the fact that the two frequencies are not forced to have the same epicentral time, the timing at which the amplitude increases to above background noise level agrees well.
The low-frequency source-time function is characterized by initial increase in energy, the main peak which decays away rather rapidly, and a second peak at the end of the event. The main energy release around 80 seconds is better defined than at high frequency. This is a clear peak with duration of about 70 seconds whereas the high-frequency energy increase is achieved in steps with multiple peaks associated with various subevents. Based upon this observation, some subevents are more effective at releasing high-frequency energy than low-frequency energy. The opposite happens at the end of the mainshock where the relative low-frequency amplitude at about 180 seconds is larger than the relative high-frequency amplitude.
Because the spatial resolution is poor at low frequency due to larger wavelength and longer period, detailed rupture characteristics, such as that revealed at high frequency, cannot be obtained. Figure 3(b) shows the 80-seconds peak at low frequency, and the lack of lateral resolution compared to high-frequency results is immediately obvious. For this peak in energy release, the locations of high- and low-frequency results match well, although the low-frequency peak is slightly south of the high-frequency region. On the other hand, the second peak at about 180 seconds is significantly offset between the high-and low-frequency energy release locations (Fig. 3(c)). The low-frequency energy release is occurring around the trench location, most likely at the trench, while most of the high-frequency energy release does not reach the trench. This peak at low frequency follows the updip propagation of high-frequency energy that happens during subevent 5, eventually reaching the trench (Fig. 2). This suggests that once rupture propagates into the trench, the slip becomes slower, effectively releasing energy at lower frequency.
If one performs back-projection analysis at even lower frequency with additional data for improved spatial resolution (Kiser and Ishii, manuscript in preparation), the largest energy release appears updip and northeast of the epicentral location (Kiser and Ishii, manuscript in preparation). This location is consistent with seafloor GPS observations (e.g., Sato et al., 2011), suggesting that major slip is associated with such low-frequency energy release.
The high-frequency back-projection results can also be compared to source regions of past tsunamigenic earthquakes, and the general agreement is good (Fig. 4(c)). The main high-frequency energy release occurs in the area where many tsunamigenic earthquakes have repeated in the past. This highly locked zone has been producing magnitude 7.5 to 8.0 earthquakes with the recurrence time of about 37 years (The Headquarters for Earthquake Research Promotion, 2011). The three sources shown in Fig. 4(b) that lie within the northern rupture area are inferred source areas of 1915, 1936, and 1978 events (Hatori, 1987). The 1978 source region, which is the most downdip source along the coastline, are known to have ruptured in 1897, 1861, and 1835 (e.g., Hatori, 1987). To the south of this area, there have not been as many earthquakes with damaging tsunamis, but a portion of the southern rupture area imaged by the back-projection analysis overlaps with the source region of the 1938 tsunamigenic earthquake (e.g., Hatori, 1987; Fig. 4(c)). The southern part that broke during the March 11, 2011 mainshock is also not unknown to large earthquakes, and rupture regions of such earthquakes inferred from aftershock distribution have updip extent (Uchida et al., 2009) that roughly matches with the broad updip extent of subevent 5 (Fig. 2).
Finally, recent studies (e.g., Minoura, 2008) have scientifically confirmed the earthquake of 869 A.D. (Jogan) and associated tsunami that devastated similar region as that affected by the event in 2011. Identification of old tsunami deposits has shown extensive inundation in areas around the Sendai plain from the Jogan earthquake (e.g., Sawai et al., 2008). Tsunami simulations that attempt to produce similar inundation have shown that the earthquake source is unlikely to be at the outer rise or shallowest part of the plate interface (e.g., Satake et al., 2008). The preferred source is within the seismogenic zone, with depth ranging somewhere between 15 and 50 km (Satake et al., 2008). The southern part of the area that broke during subevent 2 in 2011 (Fig. 2) is within the proposed source region for the 869 A.D. Jogan earthquake, and may explain the level of tsunami damage in the Sendai area.
The back-projection results suggest that many segments of the plate interface that have broken independently in the past have moved during the mainshock of March 11, 2011. The rupture area shows involvement of at least three large patches with minimal overlap. This multi-segment process, as well as nearly complete breakage of the seismo-genic zone, are similar to other recent great earthquakes (e.g., 2010 Maule, Chile; Kiser and Ishii, 2011a, b). If we exclude the outer-rise activity, the distribution of aftershocks is also mostly consistent with inferred rupture area (Fig. 2), although the match is not as good as those observed for other great earthquakes (Kiser and Ishii, 2011b). One feature that has been observed in some finite-fault models (e.g., Ammon et al., 2011; Hayes, 2011; Shao et al., 2011) and not in the back-projection results (e.g., Hutko, 2011; Kiser, 2011; Wang and Mori, 2011; Yao et al., 2011) is the large slip updip of the epicentre, extending almost to the trench. Even using the coherency function, no sign of high-frequency energy release from this region is observed, i.e., the gap in energy release close to the trench in Fig. 2 is real. In fact, this updip gap in high-frequency energy release seems to be where the finite-fault models indicate highest slip. This may mean that the slip mechanism in this region is very different from any of the recent great earthquakes, i.e., deformation or slip without any high-frequency energy release. In contrast, high-frequency energy release is observed up to the trench at the southern end of the main-shock rupture area. This area also releases energy at lower frequency, suggestive of possible near-trench deformation.
The author would like to thank the Geological Survey of Canada and Incorporated Research Institutions for Seismology for making digital data of the Canadian National Seismograph Network and the data from the United States readily available via the internet. Comments from Haruko Sekiguchi and an anonymous reviewer have helped to improve this manuscript. Some of the figures were generated using the Generic Mapping Tools (Wessel and Smith, 1991).
- Ammon, C. J., T. Lay, and H. Kanamori, Seismicity animations, fault rupture model, etc. of the great 2011 Tohoku-Chiho Taiheiyo-Oki earthquake sequence, http://eqseis.geosc.psu.edu/~cammon/Japan2011EQ/, 2011.
- Hashimoto, C., A. Noda, T. Sagiya, and M. Matsuura, Interplate seismo-genic zones along the Kuril-Japan trench inferred from GPS data inversion, Nature Geosci.,2, 141–144, 2009.View ArticleGoogle Scholar
- Hatori, T., Distribution of seismic intensity and tsunami of the 1793 Miyagi-oki earthquake, northeastern Japan, Bull. Earthq. Res. Inst.,62, 297–309, 1987.Google Scholar
- Hayes, G., Finite fault model, http://earthquake.usgs.gov/earthquakes/eqinthenews/2011/usc0001xgp/finitefault.php, 2011.
- Hutko, A., Japan March 11, 2011 05:46:23 Mw8.9, http://es.ucsc.edu/~ahutko/JapanMarch2011/, 2011.
- Ishii, M., P. M. Shearer, H. Houston, and J. E. Vidale, Extent, duration and speed of the 2004 Sumatra-Andaman earthquake imaged by the Hi-net array, Nature,435, 933–936, 2005.Google Scholar
- Ishii, M., P. M. Shearer, H. Houston, and J. E. Vidale, Teleseismic P wave imaging of the 26 December 2004 Sumatra-Andaman and 28 March 2005 Sumatra earthquake ruptures using the Hi-Net array, J. Geophys. Res.,112, B11307, doi:10.1029/2006JB004700, 2007.View ArticleGoogle Scholar
- Kiser, E., Preliminary rupture modelling of the March 11, 2011 Tohoku-Chiho Taiheiyo-Oki earthquake and sequence of events using the USArray Transportable array, http://seismology.harvard.edu/researchjapan.html, 2011.
- Kiser, E. and M. Ishii, The 2010 Mw 8.8 Chile earthquake: Triggering on multiple segments and frequency-dependent rupture behavior, Geophys. Res. Lett.,38, doi:10.1029/2011GL047140, 2011a.Google Scholar
- Kiser, E. and M. Ishii, Rupture details of great earthquakes between 2004 and 2010 imaged by the back-projection technique, 2011b (in preparation).Google Scholar
- Kiser, E., M. Ishii, C. H. Langmuir, P. M. Shearer, and H. Hirose, Insights into the mechanism of intermediate-depth earthquakes from source properties as imaged by back-projection of multiple seismic phases, J. Geophys. Res.,116, B06310, 2011.Google Scholar
- Loveless, J. P. and B. J. Meade, Geodetic imaging of plate motions, slip rates, and partitioning of deformation in Japan, J. Geophys. Res.,115, B02410, 2010.Google Scholar
- Minoura, K., Quantitative analysis of tsunami by the Jogan earthquake, Report of Research Project, Grant-in-Aid for Scientific Research (B) (1), No. 17310103, 2008.Google Scholar
- Satake, K., Y. Namegaya, and S. Yamaki, Numerical simulation of the AD 869 Jogan tsunami in Ishinomaki and Sendai plains, Res. Rep. Active Faults Hist. Earthq.,8, 71–89, 2008.Google Scholar
- Sato, M., T. Ishikawa, N. Ujihara, S. Yoshida, M. Fujita, M. Mochizuki, and A. Asada, Displacement above the hypocenter of the 2011 Tohoku-Oki earthquake, Science,332, 1395, doi:10.1126/science.1207401, 2011.View ArticleGoogle Scholar
- Sawai, Y., Y. Fujii, O. Fujiwara, T. Kamataki, J. Komatsubara, Y. Okamura, K. Satake, and M. Shishikura, Marine incursions of the past 1500 years and evidence of tsunamis at Suijin-numa, a coastal lake facing the Japan Trench, The Holocene,18, 517–528, 2008.View ArticleGoogle Scholar
- Shao, G., X. Li, C. Ji, and T. Maeda, Preliminary result of the Mar 11, 2011 Mw 9.2 Honshu earthquake, http://www.geol.ucsb.edu/faculty/ji/bigearthquakes/2011/03/0311/Honshumain.html, 2011.
- The Headquarters for Earthquake Research Promotion, Summary of long-term seismic probability for subduction zone earthquakes, http://www.jishin.go.jp/main/choukihyoka/kaikou.htm, 2011.
- Uchida, N., J. Nakajima, A. Hasegawa, and T. Matsuzawa, What controls interplate coupling?: Evidence for abrupt change in coupling across a border between two overlying plates in the NE Japan subduction zone. Earth Planet. Sci. Lett.,283, 111–121, 2009.View ArticleGoogle Scholar
- Wang, D. and J. Mori, Back projection of the rupture for the 2011 Tohoku earthquake, http://www.eqh.dpri.kyoto-u.ac.jp/~mori/Sendai/sendai.htm, 2011.
- Wessel, P. and W. H. F. Smith, Free software helps map and display data. EOS Trans. AGU,72, 441, 1991.View ArticleGoogle Scholar
- Yao, H., P. M. Shearer, and T. Uchide, Iterative back projection of the rupture and spatial-temporal distribution of subevents of the 2011 Tohoku, Japan earthquake, http://igppweb.ucsd.edu/~huyao/TohokuEarthquake.html, 2011.