Proceedings of the International Symposium on Engineering Lessons Learned fromthe 2011 Great East Japan Earthquake, March 1-4, 2012, Tokyo, Japan

REAL-TIME DETECTION OF THE SOURCE AREA
OF AN INTENSE TSUNAMI―CASE STUDY OF
THE 2011 GREAT EAST JAPAN EARTHQUAKE
Yutaka HAYASHI1, Hiroaki TSUSHIMA2 and Yasuhiro YOSHIDA3
1

Senior Researcher, Seismology and Volcanology Research Department, Meteorological ResearchInstitute, Japan Meteorological Agency, Tsukuba, Japan, yhayashi@mri-jma.go.jp
2Researcher, Seismology and Volcanology Research Department, Meteorological Research Institute,Japan Meteorological Agency, Tsukuba, Japan, tsushima@mri-jma.go.jp
3
Senior Researcher, Seismology and Volcanology Research Department, Meteorological ResearchInstitute, Japan Meteorological Agency, Tsukuba, Japan, yyoshida@mri-jma.go.jp

ABSTRACT: Offshore tsunami waveform data from the 2011 Great East Japan
earthquake indicate that the maximum-amplitude tsunami that struck the Tohoku district
was of relatively short period. By comparing one seismic and two tsunami source models,
we concluded that the short-period large-amplitude tsunami was generated within thearea between the epicenter and the Japan Trench. We demonstrated that all three models
are capable of real-time detection of future tsunami sources of similar intensity to that of
the 2011 tsunami.
Key Words: Impulsive tsunami source, 2011 off the Pacific coast of Tohoku earthquake,strong-motion waveform inversion, tsunami back-propagation, tsunami
waveform inversion

INTRODUCTION
A massive earthquake of moment magnitude (Mw) 9.0 occurred at 14:46 JST (05:46 UTC) on 11
March 2011, off the Pacific coast of northeastern Honshu, Japan. This earthquake, named the 2011 off
the Pacific coast of Tohoku earthquake by the Japan Meteorological Agency (JMA; Hirose et al. 2011),
is hereafter referred to as the 2011 Great East Japan earthquake. The epicenter was offshore, southeastof Sendai City (38°06.2′N, 142°51.6′E; JMA 2011). Only one previous earthquake of this seismic
intensity (7, the upper limit of the JMA scale) has been observed since the JMA introduced
instrument-based seismic intensity observations in 1996 (Hoshiba et al. 2010, 2011). Seismic
intensities of 6-upper and 6-lower (JMA scale) were observed over an area of approximately 400 km ×
100 km by many onshore seismic stations in the Tohoku and Kanto districts (Hoshiba et al. 2011).
The tsunami that accompanied the earthquake was detected offshore by coastal wave gauges(Nagai et al. 2005), real-time kinematic global positioning system (RTK-GPS) buoys (Kato et al.
2005), cabled deep ocean-bottom pressure gauges (OBPG) (e.g. Fujisawa et al. 1986; Hirata et al.
2002), and Deep-ocean Assessment and Reporting of Tsunamis (DART) buoys (González et al. 2005).
In particular, OBPGs deployed off Tohoku recorded an impulsive tsunami of short period and large632amplitude (Ito et al. 2011). The periods of the initial tsunami waves observed off Hokkaido werelonger than those observed off Tohoku. Thus, the most impulsive tsunami waves propagated westward
from the source area.
In this paper, we compare tsunami source locations determined by a seismic fault model with slip
distribution estimated by waveform inversion of strong ground motion (MODEL-1; Yoshida et al.
2011), by back-propagation of tsunami arrival times at offshore stations (MODEL-2; Hayashi et al.
2011), and from a deformation field determined by tsunami waveform inversion (MODEL-3;
Tsushima et al. 2011). We then consider the real-time application of each modeling approach to thedetection of intense tsunami sources such as that of the Great East Japan earthquake.

MODEL-1: STRONG-MOTION WAVEFORM INVERSION
Strong-motion data observed near an earthquake epicenter (< 500 km) can provide more detailed
information about the source process of a large earthquake than can be obtained from teleseismic bodyFig. 1 Finite-source model of the 2011 Great East Japan earthquake from inversion of strong-motion
waves (Yoshida et al. 2011). (a) Moment rate function. (b) Slip distribution on the fault. The large
green star represents the epicenter of the main shock (Mw = 9.0); gray circles represent aftershocks(MJMA ≥ 5.0) within 24 h of the main shock. Crosses represent grid points on the fault plane used for
calculation of synthetic waveforms. Triangles denote seismic stations used in this analysis. Slip
distribution contour interval is 4 m. Twenty-three strong-motion seismograms were used from two
K-NET stations (IWT007 and MYG011; Kinoshita 1998) and 13 KiK-net stations (IWTH08, IWTH13,
IWTH14, IWTH21, MYGH04, MYGH12, MYGH06, MYGH08, FKSH17, FKSH19, FKSH12,
IBRH13, and IBRH15; Aoi et al. 2000) deployed by the National Research Institute for Earth Scienceand Disaster Prevention (NIED) as well as eight stations from JMA (ERIMO, HITACH, IYASAT,KATASH, ASHIKA, HANNOU, TENNOD, and NAGARA).

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wave data. Several studies have used strong-motion data to investigate the 2011 Great East Japan
earthquake (e.g. Yoshida et al. 2011, Suzuki et al. 2011). Here, we summarize the results ofstrong-motion waveform inversion by Yoshida et al. (2011).
Twenty-three near-field strong-motion seismograms were used (Fig. 1). Acceleration seismograms
were first integrated to velocity, then band-pass filtered (0.01–0.15 Hz) and decimated to 0.5 Hz. A
time window of 250 s was used, starting from 10 s before the P-wave arrivals. Strike and dip of thefault plane were assumed to be 201° and 9°, respectively. The fault was assumed to measure 475 km
along strike and 175 km in the dip direction. It was assumed that the rupture initiated at the hypocenterof the main shock determined by JMA. The fault plane was divided into sub-faults of 25 km × 25 km.
Green’s functions were calculated for each sub-fault by the discrete wavenumber method(Bouchon 1981) using reflection–transmission matrices (Kennett and Kerry 1979). The inelasticity
effect was accounted for by using a complex velocity field (Takeo 1985). The moment rate function
for each sub-fault was expressed by 20 simple triangle functions of 8 s duration and overlapping by 4 s.
The maximum rupture velocity was set at 2.5 km/s to minimize variance. The linear multiple time
window inversion method was used with constraints on smoothness of the spatiotemporal slip
distribution (e.g. Ide et al. 1996, Nakayama and Takeo 1997). Smoothness parameters(hyperparameters) were selected to minimize Akaike’s Bayesian information criterion (ABIC) (Akaike
1980, Fukahata et al. 2003). Waveforms were aligned by onset time and weighted equally for all
stations.
Fig. 1 shows the slip distribution obtained from the regional strong-motion data analysis. The total
seismic moment was 3.4 × 1022 Nm (Mw = 9.0). The slip area extends eastward from the hypocenter tothe shallowest part of the fault plane and maximum slip is 38 m. The fit between observed andsynthetic waveforms is reasonably good, when a reduction in variance of the fitted model is assumed
to be about 91%.
The sequence of released seismic moment is as follows. In the first stage of the rupture (0–40 s), therupture expands radially outward from the hypocenter. In the next stage (40–80 s), the rupture area
extends both north and south along the shallow part of the fault plane. The large amount of slip during
this stage may have generated the impulsive tsunami. The rupture velocity is very slow (about 1 km/s)
during the first and second stages. In the third stage (after 80 s), the rupture extends southward,
reaching the southern end of the fault plane at 160 s.

MODEL-2: BACK-PROPAGATION OF TSUNAMI ARRIVAL TIMESIn this section, we summarize the estimation of the tsunami source area by back-propagation oftsunami arrival times by Hayashi et al. (2011).
The tsunami was detected at various offshore observation stations, including coastal wave gauges(Nagai et al. 2005), real-time kinematic global positioning system (RTK-GPS) buoys (Kato et al. 2005,
cabled deep ocean-bottom pressure gauges (OBPG) (e.g. Fujisawa et al. 1986, Kanazawa andHasegawa 1997, Hirata et al. 2002) Time-series of changes of sea-level and ocean-bottom pressurerecorded following the 2011 Great East Japan earthquake were acquired from 21 observation stations(Figs. 2 and 3). Arrival times for each phase of the tsunami were manually read from waveforms (Fig.
2).
To define the edges of the tsunami source area, Huygens’ Principle was applied to back-propagate
the tsunami from each observation station. For these calculations tsunami travel-time software (TTT v.
3.0; Geoware in Hawaii, USA) was used with bathymetric data at 1 min intervals (ETOPO1; Amanteand Eakins 2009). The phase velocity of tsunami propagation was assumed to be equal to the squareroot of gravity multiplied by water depth. For very large earthquakes, the time difference between themain shock and generation of the tsunami is not negligible (Seno and Hirata 2007). Therefore, forback-propagation from tsunami arrival times, time values were modified by 1 min / 120 km,
corresponding to a distance from the epicenter to the contact point of the back-propagation curve withthe tsunami source area (Fig. 3a). This correction was applied to account for differences between thetime of the main shock and the generation of the tsunami and is equivalent to assuming an average634

apparent (i.e. projected to the seafloor) fault rupture velocity of 2 km/s.
The tsunami source area determined by back-propagation of tsunami arrivals at offshore
observation stations (Fig. 3a) was approximately 500 km long with a maximum width of about 200 km.
The eastern edge of the source area was close to and approximately parallel to the western flank of theFig. 2 Tsunami waveforms recorded offshore during and after the 2011 Great East Japan earthquake
and phase nomenclature (Hayashi et al. 2011). Waveform data were low-pass filtered with cut-off
period of 2 min. Online data transmission was interrupted at some observation stations, but data
recovered afterwards are plotted. See Fig. 3 for locations of observation stations.

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Fig. 3 Tsunami source area of the 2011 Tohoku earthquake determined by back-propagation of (a)tsunami arrivals and (b) arrival times of primary crests at offshore observation stations (Hayashi et al.
2011).

Japan Trench, and the southern edge was at about latitude 36°N. However, the aftershock area (Hirose
et al. 2011) extended to the eastern side of the trench and farther south than the tsunami source area.
Sea-level changes observed at stations TM1, TM2, GPS804, GPS802, GPS803, and GPS801 were
almost synchronous with the arrival of the seismic wave, indicating that these offshore stations were
within the tsunami source area.
Back-propagation was also applied to primary crests to identify the area of major seafloor uplift in
the tsunami source area. However, to avoid incorporating data strongly affected by nonlinear effects ordispersion, it is necessary to limit the input data to near-field tsunami observations in deep water. To
achieve this, only the primary crests observed by OBPGs and GPS buoys were used for thisback-propagation.
All of the back-propagation curves of primary crests, except those from stations GPS807 andGPS806, pass through an area centered around N38° E143.5° (indicated by the gray dotted line in Fig.
3b), several tens of kilometers east of the epicenter. If the area of seafloor uplift was confined to thissmall area, most of the primary crest arrival times observed at GPS buoys and OBPGs are reasonably636

explained. However, this is but one of several solutions that can explain the observed arrival times ofprimary crests, as discussed below.
During a great earthquake such as the 2011 Great East Japan earthquake, there may be seismic
activity and uplift at multiple locations within the source area. If this is the case, it is difficult to
identify the area of maximum uplift by back-propagation of primary crests.
Furthermore, the back-propagation analysis was based on the assumption that tsunami phasevelocity is equal to the square root of gravity multiplied by water depth. If the phase velocity is
nonlinear, wave crests travel faster than this. On the other hand, dispersion results in wave crests
moving more slowly. These effects may cause some estimation errors in areas of large uplift.
It may have been the effect of dispersion that made it difficult to explain the back-propagation
curves from stations GPS807 and GPS806 (Fig. 3); otherwise, the area of seafloor uplift may haveextended farther in the north–south direction, rather than occurring only to the east of the epicenter.

MODEL-3: TSUNAMI WAVEFORM INVERSION
Tsunami waveform inversion is a technique that applies a least-squares approach to observed tsunami
waveforms to identify the tsunami source from the spatial distribution of either fault slip or initial
sea-surface displacement. This technique has been applied in many studies of past tsunami events (e.g.
Satake 1989; Tanioka et al. 2006; Fujii and Satake 2008; Saito et al. 2010), including the 2011 Great
East Japan earthquake (Fujii et al. 2011). Some studies have proposed a tsunami forecasting algorithm
based on inversion of offshore tsunami waveforms (Titov et al. 2005; Wei et al. 2008; Tsushima et al.
2009). Here, we summarize the results of tsunami waveform inversion for the 2011 Great East Japan
earthquake by Tsushima et al. (2011).
The tsunami was recorded at several offshore tsunami stations (Fig. 2). Waveforms recorded at
offshore stations provide a tsunami source signature without distortions from complex coastal
topography (González et al. 2005). To take advantage of these undistorted signals, Tsushima et al.
(2011) chose to use tsunami waveform data recorded at only four OBPGs (Hirata et al. 2002,
Kanazawa and Hasegawa 1997) and five GPS buoys (Kato et al. 2005). Their inversion included theimpulsive tsunami waveforms that were recorded at OBPG stations TM1 and TM2 and GPS buoys
802 and 804 (Figs. 4a and 4b). The model derived from those data explains a source of impulsive
tsunami.
The modeled distribution of initial sea-surface displacement indicates the tsunami source area. The
inversion method applied by Tsushima et al. (2011) is part of a tsunami forecasting algorithm
developed by Tsushima et al. (2009). In this inversion, two constraints are imposed to stabilize thesolution of the inverse problem: a smoothing constraint and a damping constraint. The damping
constraint is based on the presumption that the initial sea-surface displacement due to a tsunamigenic
earthquake will be zero at locations distant from the epicenter. Green’s functions were computed byfinite-difference approximation of linear long-wave equations (Satake 1995).
The observed and modeled tsunami waveforms matched well (Figs. 4a and 4b), and the modeled
uplift occurred in a roughly circular area between the epicenter and the trench (Fig. 4c). Determinationof a more accurate location and extent of the source area of the impulsive tsunami could not be
achieved because of insufficient azimuthal coverage of the offshore stations included in the inversion(Tsushima et al. 2011); nonetheless, the inversion results indicate that the intense tsunami was causedby considerable sea-surface elevation near the trench.

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Fig. 4 Result of tsunami waveform inversion by Tsushima et al. (2011) using selected offshore tsunami
waveforms from the 2011 Great East Japan earthquake. Comparison of observed (black lines) andmodeled (red lines) waveforms at (a) four OBPGs and (b) five GPS buoys. Only the waveforms in
unshaded areas of each panel were included in the inversion. (c) Distribution of modeled initial
sea-surface elevation. Star indicates epicenter used as the damping constraint in the inversion. Areas
shaded gray are outside the influence area. Contour interval is 1 m.

DISCUSSION
The strong impulsive tsunami waves of the 2011 Great East Japan earthquake are represented by
the primary crests recorded at offshore observation stations off the Pacific coast of the Tohoku district(Fig. 2). Tsunami waveform inversion (MODEL-3) of selected data (that were not distorted by coastal
topography) from offshore stations showed considerable sea-surface elevation in the region between
the epicenter and the Japan Trench (Fig. 4c). The arrival times of primary crests at most offshore
stations are consistent with an area of maximum uplift between the epicenter and the trench (Fig. 3b).
Furthermore, the area of large slip obtained by inversion of seismic strong-motion waves (Fig. 1b) isalmost coincident with the source area determined by back-propagation (Fig. 3b) and the area of
sea-level uplift determined by tsunami waveform inversion (Fig. 4b). Therefore, we conclude that the638

source of the initial impulsive tsunami was probably within the area between the epicenter and theJapan Trench.
Data used in MODEL-1 were regional strong-motion data recorded on land over several minutes
starting from the first seismic activity of the 2011 Great East Japan earthquake. Data used in
MODEL-3 were tsunami waveform data recorded at offshore observation stations during the 35-min
period following the main shock. Application of these modeling methods to data from the 2011 Great
East Japan earthquake suggests that both have the potential to provide forewarning of an intense
tsunami generated by a near-field large earthquake, provided that there is sufficient processing andanalysis time.
Tsunami back-propagation using selected near-field offshore data showed that the method of
MODEL-2 may be able to provide forewarning of an intense tsunami generated by a large near-field
earthquake, provided that real-time data transmission is available. For MODEL-2, individual
back-propagations were initially performed based on initial tsunami arrival times (Fig. 3a) and on
primary crest arrival times (Fig. 3b) at offshore observation stations; both inversions included

Fig. 5 Tests of rapid determination based on initial 30 min of data at offshore observation stations: (a)tsunami source area determined by back-propagation of tsunami arrivals and (b) intense tsunami
source determined from occurrence times of primary tsunami crests.

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Fig. 6 Tests of rapid determination based on initial 20 min of data at offshore observation stations: (a)tsunami source area determined by back-propagation of tsunami arrivals and (b) intense tsunami
source determined from occurrence times of primary tsunami crests.

waveform data recorded more than 30 min after the main shock. However, when input data werelimited to those recorded within 30 min of the main shock (Fig. 5a and 5b), the source area of theintense tsunami was only slightly east of that determined from longer wave trains (Fig. 5b). The data
recorded within 20 min of the main shock (Fig. 6) were insufficient to determine the location of thesource of the intense tsunami.
However, if the back-propagation analyses of Figs 5a and 6a had been applied in real time, thetsunami arrival-time data recorded within 20–30 min after the main shock may have provided
important information on the extent of the greate tsunami source area. Back-propagation analysis
using the 30-min dataset (Fig. 5a) estimated a tsunami source area that was both wider than 180 km
and longer than 480 km. The boundaries of the tsunami source area so defined by back-propagation of
DART buoy 21418, GPS buoys along the coastline, and OBPGs KPG2 and BOSO3. And, theepicenter and the stations where sea-level changes were almost synchronous with the arrival of seismic
wave (TM1, TM2, GPS801, GPS802, GPS803, and GPS 804) must be within the tsunami source area.
Similarly, back-propagation analysis using the 20 min dataset indicated a wider than 130 km and640

longer than 480 km. From tsunami data obtained in the Pacific Ocean around Japan, Abe (1975)
empirically derived the following relationships:

M 0 = 1.23 × 106 S 3 2 , S = 0.8St ,(1)

where M0, S, and St are seismic moment (Nm), rectangular area (m2) of the seismic fault, andrectangular area (m2) of the tsunami source, respectively.
Kanamori (1977) used worldwide earthquake data to establish the following relationship between
seismic moment M0 and moment magnitude Mw:

log10 M 0 = 1.5M W + 9.1 .

(2)

Applying Eqs. (1) and (2) to the 20- and 30-min datasets gives Mw values of ≥8.7 and ≥8.8,
respectively, slightly lower than the final moment magnitude provided by the JMA earthquake
catalogue (Mw 9.0).

CONCLUSIONS
Based on our comparison of models derived independently by (1) strong ground motion waveform
inversion, (2) back-propagation of tsunami arrival times at offshore observation stations, and (3)tsunami waveform inversion, we concluded that the origin of the destructive impulsive tsunami thatstruck the Pacific coast of Tohoku after the 2011 Great East Japan earthquake probably formed
between the epicenter and the Japan Trench.
Each of the above methods can be applied in real time, and each would have detected the source of
the intense tsunami from either seismic or tsunami waveform data within about 35 min of the main
shock. Furthermore, seismic waveform inversion or tsunami waveform inversion would havedetermined a reasonably accurate seismic magnitude of this mega-earthquake within about 20 min of
the main shock.
We thus conclude that individual real-time application of each of the above methods could detect
intense tsunami sources and approximate seismic magnitudes of future large earthquakes in the Japanregion. Furthermore, we expect that these models used in combination would improve our ability toprovide timely warnings of impending tsunamis.

ACKNOWLEDGEMENTS
This work was partly supported by the Science and Technology Research Partnership for Sustainable
Development (SATREPS) Research Project on Enhancement of Technology to Develop
Tsunami-resilient Community, promoted by the Japan Science and Technology Agency (JST) andJapan International Cooperation Agency (JICA). Tsunami waveform data were acquired from the Portand Harbour Bureau of the Ministry of Land, Infrastructure, Transport and Tourism (MLIT); the Japan
Agency for Marine-Earth Science and Technology (JAMSTEC); the Earthquake Research Institute(ERI) at Tokyo University; the Japan Meteorological Agency (JMA); the Russian Far Eastern Regional
Hydrometeorological Research Institute (RFERHRI); and the United States National Oceanic andAtmospheric Administration (NOAA). We used Generic Mapping Tools software (Wessel and Smith
1998) to prepare our figures.

REFERENCES
Abe, K. (1975). “Reliable estimation of the seismic moment of large earthquakes.” Journal of Physics
of the Earth, Vol. 23, 381-390.
Akaike, H. (1980). “Likelihood and the Bayes procedure, in Bayesian statics.” edited by J. M.

641

Bernardo, M. H. DeGroot, D. V. Lindley, and A. F. M. Smith, University Press, Valencia, Spain.
Amante, C. and B. W. Eakins (2009). “ETOPO1 1 arc-minute global relief model: Procedures, data
source and analysis.” NOAA Technical Memorandum, NESDIS NGDC-24, 19p.
Aoi, S., K. Obara, S. Hori, K. Kasahara, and Y. Okada (2000). “New strong-motion observation
network: Kik-net.” EOS, Transactions, American Geophysical Union, Vol. 81, F863.
Bouchon, M. (1981). “A simple method to calculate Green’s functions for elastic layered media.”
Bulletin of the Seismological Society of America, Vol. 71, 959-971.
Fujii, Y., and K. Satake (2008). “Tsunami sources of the November 2006 and January 2007 great Kuril
earthquakes.” Bulletin of the Seismological Society of America, Vol. 98, 1559-1571, doi:10.1785/
0120070221.
Fujii, Y., K. Satake, S. Sakai, M. Shinohara, and T. Kanazawa (2011). “Tsunami source of the 2011 off
the Pacific coast of Tohoku Earthquake.” Earth, Planets and Space, Vol. 63, 815-820.
Fujisawa, I., S. Tateyama, and J. Fujisaki (1986). “Permanent ocean-bottom earthquake and tsunami
observation system off the Boso Peninsula”, Weather Service Bulletin, Vol. 53, 127-166. (inJapanese)
Fukahata, Y., Y. Yagi, and M. Matsu’ura (2003). “Waveform inversion for seismic source processes
using ABIC with two sorts of prior constraints: comparison between proper and improper
formulations.” Geophysical Research Letters, Vol. 30, 1305, doi:10.1029/2002GL016293.
González, F. I., E. N. Bernard, C. Meinig, M. C. Ebel, H. O. Mofjeld, and S. Stalin (2005). “TheNTHMP tsunameter network.” Natural Hazards, Vol. 35, 25-39, doi:10.1007/s11069-004-2402-4.
Hayashi, Y., H. Tsushima, K. Hirata, K. Kimura, and K. Maeda (2011). “Tsunami source area of the2011 off the Pacific Coast of Tohoku Earthquake determined from tsunami arrival times at offshore
observation station.” Earth, Planets and Space, Vol. 63, 809-813.
Hirata, K., M. Aoyagi, H. Mikada, K. Kawaguchi, Y. Kaiho, R. Iwase, S. Morita, I. Fujisawa, H.
Sugioka, K. Mitsuzawa, K. Suyehiro, H. Kinoshita, and N. Fujiwara (2002). “Real-time
geophysical measurements on the deep seafloor using submarine cable in the southern Kurile
subduction zone.” IEEE Journal of Oceanic Engineering, Vol. 27, 170-181.
Hirose, F., K. Miyaoka, N. Hashimoto, T. Yamazaki, and M. Nakamura (2011). “Outline of the 2011
off the Pacific coast of Tohoku Earthquake (Mw 9.0)—Seismicity: foreshocks, mainshock,aftershocks, and induced activity—.” Earth Planets, and Space, Vol. 63, 513-518.
Hoshiba, M., K. Ohtake, K. Iwakiri, T. Aketagawa, H. Nakamura, and S. Yamamoto (2010). “Howprecisely can we anticipate seismic intensities? A study of uncertainty of anticipated seismic
intensities for the Earthquake Early Warning method in Japan.” Earth Planets, and Space, Vol. 62,
611-620.
Hoshiba, M., K. Ohtake, K. Iwakiri, T. Aketagawa, H. Nakamura, and S. Yamamoto (2011). “Outlineof the 2011 off the Pacific coast of Tohoku earthquake (Mw9.0) – earthquake early warning andobserved seismic intensity –.” Earth Planets, and Space, Vol. 63, 547-551.
Ide, S., M. Takeo, and Y. Yoshida (1996). “Source process of the 1995 Kobe earthquake:
Determination of spatio-temporal slip distribution by Bayesian modeling.” Bulletin of the
Seismological Society of America, Vol. 86, 547-566.
Ito, Y., T. Tsuji, Y. Osada, M. Kido, D. Inazu, Y. Hayashi, H. Tsushima, R. Hino, and H. Fujimoto(2011). “Frontal wedge deformation near the source region of the 2011 Tohoku-Oki earthquake.”
Geophysical Research Letters, Vol.38, L00G05, doi:10.1029/2011GL048355.
Japan Meteorological Agency (2011). “Monthly Report on Earthquakes and Volcanoes in Japan
March 2011.” pp.57-148. (in Japanese)
Kanamori, H. (1977). “The energy release in great earthquakes.” Journal of Geophysical Research,
Vol. 82, 2981-2987.
Kanazawa, T., and H. Hasegawa (1997). Ocean-bottom observatory for earthquakes and tsunami off
Sanriku, north-east Japan using submarine cable, International Workshop on Scientific Use ofSubmarine Cables, Comm. for Sci. Use of Submarine Cables, Okinawa, Japan, 208-209.
Kato, T., Y. Terada, K. Ito, R. Hattori, T. Abe, T. Miyake, S. Koshimura, and T. Nagai (2005).
“Tsunami due to the 2004 September 5th off the Kii Peninsula earthquake, Japan, recorded by a642

new GPS buoy”, Earth Planets, and Space, Vol. 57, 279-301.
Kennett, L. N. and N. J. Kerry (1979). “Seismic waves in a stratified half space.” Geophysical Journalof the Royal Astronomical Society, Vol. 57, 557-583.
Kinoshita, S. (1998). “Kyoshin net (K-NET).” Seismological Research Letters, Vol. 69, 209-332.
Nagai, T., S. Satomi, Y. Terada, T. Kato, K. Nukada, and M. Kudaka (2005). “GPS buoy and 4 seabedinstalled wave gauge application to offshore tsunami observation”, Proceedings of the 15th
International Offshore and Polar Engineering Conference, Vol. III, 292-299.
Nakayama, W. and M. Takeo (1997). “Slip history of the 1994 Sanriku-haruka-oki, Japan, earthquake
deduced from strong-motion data.” Bulletin of the Seismological Society of America, Vol. 87,
918-931.
Saito, T., K. Satake, and T. Furumura (2010). “Tsunami waveform inversion including dispersive
waves: the 2004 earthquake off Kii Peninsula, Japan.” Journal of Geophysical Research, Vol. 115,
B06303, doi:10.1029/2009JB006884.
Satake, K. (1989). “Inversion of tsunami waveforms for the estimation of heterogeneous fault motionof large submarine earthquakes: The 1968 Tokachi-oki and 1983 Japan Sea earthquakes.” Journalof Geophysical Research, Vol. 94, 5627-5636, doi:10.1029/JB094iB05p05627.
Satake K. (1995). “Linear and nonlinear computations of the 1992 Nicaragua earthquake tsunami.”
Pure and Applied Geophysics, Vol. 144, 455-470, doi:10.1007/BF00874378.
Seno, T. and K. Hirata (2007). “Did the 2004 Sumatra-Andaman earthquake involve a component oftsunami earthquakes?” Bulletin of the Seismological Society of America, Vol. 97, S296–S306.
Suzuki, W., S. Aoi, H. Sekiguchi, and T. Kunugi (2011). “Rupture process of the 2011 Tohoku-Oki
mega-thrust earthquake (M9.0) inverted from strong-motion data.” Geophysical Research Letters,Vol. 38, L00G16, doi:10.1029/2011GL049136.
Takeo, M. (1985). “Near-field synthetic seismograms taking into account the effect of an elasticity.”
Papers in Meteorology and Geophysics, Vol. 36, 245-257. (in Japanese)
Tanioka, Y., Yudhicara, T. Kususose, S. Kathiroll, Y. Nishimura, S. Iwasaki, and K. Satake (2006).
“Rupture process of the 2004 great Sumatra-Andaman earthquake estimated from tsunami
waveforms.” Earth, Planets and Space, Vol. 58, 203-209.
Titov, V. V., F. I. González, E. N. Bernard, M. C. Eble, H. O. Mojfeld, J. C. Newman, and A. J.
Venturato (2005). “Real-time tsunami forecasting: Challenges and solutions.” Natural Hazards, Vol.
35, 35-41, doi:10.1007/s11069-004-2403-3.
Tsushima, H., R. Hino, H. Fujimoto, Y. Tanioka, and F. Imamura (2009). “Near-field tsunami
forecasting from cabled ocean bottom pressure data.” Journal of Geophysical Research, Vol. 114,
B06309, doi:10.1029/2008JB005988.
Tsushima, H., K. Hirata, Y. Hayashi, Y. Tanioka, K. Kimura, S. Sakai, M. Shinohara, T. Kanazawa, R.
Hino, and K. Maeda (2011). “Near-field tsunami forecasting using offshore tsunami data from the2011 off the Pacific coast of Tohoku Earthquake.” Earth, Planets and Space, Vol. 63, 821-826.
Wei, Y., E. N. Bernard, L. Tang, R. Weiss, V. V. Titov, C. Moore, M. Spillane, M. Hopkins, and U.
Kanoglu (2008). “Real-time experimental forecast of the Peruvian tsunami of August 2007 for U.S.
coastlines.” Geophysical Research Letters, Vol. 35, L04609, doi:10.1029/2007GL032250.
Wessel, P., and W. H. F. Smith (1998). “New improved version of Generic Mapping Tools released.”
EOS, Transactions, American Geophysical Union, Vol. 79, 579.
Yoshida, Y., H. Ueno, D. Muto, and S. Aoki (2011). “Source process of the 2011 Off the Pacific Coastof Tohoku Earthquake with the combination of teleseismic and strong motion data.” Earth, Planets
and Space, Vol. 63, 565-569.

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