- Email: [email protected]
Tsunamis
4.17 Tsunamis K. Satake, National Institute of Advanced Industrial Science and Technology, Tsukuba, Japan ª 2007 Elsevier B.V. All rights reserved.
4.17.1 4.17.2 4.17.3 4.17.3.1 4.17.3.1.1 4.17.3.1.2 4.17.3.1.3 4.17.3.1.4 4.17.3.1.5 4.17.3.1.6 4.17.3.1.7 4.17.3.2 4.17.3.2.1 4.17.3.2.2 4.17.3.2.3 4.17.4 4.17.4.1 4.17.4.1.1 4.17.4.1.2 4.17.4.1.3 4.17.4.1.4 4.17.4.1.5 4.17.4.2 4.17.4.2.1 4.17.4.2.2 4.17.4.3 4.17.4.4 4.17.4.5 4.17.4.5.1 4.17.4.5.2 4.17.4.5.3 4.17.4.6 4.17.5 4.17.5.1 4.17.5.1.1 4.17.5.1.2 4.17.5.1.3 4.17.5.1.4 4.17.5.2 4.17.5.2.1 4.17.5.2.2 4.17.5.2.3 4.17.5.3 4.17.5.3.1 4.17.5.3.2 4.17.5.3.3
Introduction The 2004 Indian Ocean Tsunami Other Examples Historical Tsunamis 1998 Papua New Guinea tsunami 1993 Hokkaido tsunami 1960 Chilean tsunami 1958 Lituya Bay tsunami 1946 Aleutian tsunami 1896 Sanriku tsunami 1883 Krakatoa tsunami Prehistoric Tsunamis 1700 Cascadia tsunami Hokkaido tsunami in seventeenth century Tsunami from the Storegga slide Tsunami Observations Instrumental Measurements Coastal tsunami measurements Offshore tsunami measurements Deep ocean measurements Satellite observations Definition of amplitude and wave heights Field Survey to Measure Run-Up Heights Post-tsunami surveys Definitions of inundation and run-up heights Study of Historical Documents Geological Methods Tsunami Magnitude Scales Imamura-Iida scale, m Tsunami intensity Tsunami magnitude Mt Tsunami Catalogs Tsunami Generation Earthquake Tsunamis Ocean bottom and water surface displacements Fault parameters and ocean-bottom deformation Estimation of earthquake source parameters from tsunamis Tsunamigenic and tsunami earthquakes Tsunamis from Landslides Examples Characteristics and discrimination of landslide tsunamis Modeling landslide tsunamis Volcanic Tsunamis Volcanic tsunamis due to earthquakes Volcanic tsunamis due to landslides Volcanic tsunamis from flow of volcanic ejecta
484 484 486 486 486 486 486 487 487 487 488 488 488 489 489 489 489 489 489 489 490 490 490 490 491 491 492 492 492 492 492 493 493 493 493 493 494 494 495 496 496 496 497 497 497 498
483
484 Tsunamis 4.17.5.3.4 4.17.5.3.5 4.17.6 4.17.6.1 4.17.6.2 4.17.6.3 4.17.6.4 4.17.6.5 4.17.6.6 4.17.6.7 4.17.7 4.17.7.1 4.17.7.2 4.17.7.3 4.17.7.4 4.17.8 References
Volcanic tsunamis from submarine explosion Volcanic tsunamis from caldera collapse Tsunami Propagation Weak Coupling between Elastic Earth and Ocean Ray-Theoretical Approach Coastal Amplification and Green’s Law Tsunami as a Gravity Wave Basic Equations for Tsunami Propagation Dispersion Numerical Computations Tsunami Warning and Hazard Reduction Systems Far-Field Tsunami Warning Systems Local Tsunami Warning Systems Long-Term and Probabilistic Forecast Tsunami Hazard Maps and Public Awareness Summary
4.17.1 Introduction A tsunami is an oceanic gravity wave generated by submarine or coastal geological processes such as earthquakes, landslides, or volcanic eruptions. Most tsunamis are caused by large shallow earthquakes along subduction zones. Tsunami is a Japanese word meaning ‘harbor wave’, and is sometimes called seismic sea wave or, erroneously, tidal wave. Tsunamis are usually small and barely noticed in deep ocean, but become large and cause damage when they approach coasts or harbors. The recent 2004 tsunami in the Indian Ocean demonstrated that a tsunami can travel at trans-oceanic distances and cause damage far from the source. Tsunamis have many aspects that are studied by researchers in various fields. Their generation is related to geological processes. The propagation and offshore observation are the subjects of oceanography, and basic hydrodynamics are needed to understand the characteristics. The coastal behaviors, such as run-up on beaches or resonance in bays, are mostly studied by coastal engineers. For tsunami hazard reduction, officials in local and national governments, civil defense, and emergency managements are responsible and play important roles. In this chapter, the wide range of tsunami phenomena is reviewed. Section 4.17.2 describes the tsunami caused by the 2004 Sumatra–Andaman earthquake, the most recent and the worst tsunami disaster that affected the countries around the Indian Ocean. Other examples of recent, historical and
498 498 498 498 498 499 499 499 500 500 501 502 503 504 505 506 506
prehistoric tsunamis are reviewed in Section 4.17.3. Section 4.17.4 reviews the tsunami observation systems: instrumental measurements, field surveys, and studies of historical and geological data. The tsunami magnitude scales and catalogs are also described. Section 4.17.5 reviews the tsunami generation, by earthquakes, landslides, and volcanic eruptions, with some emphasis on the studies of tsunami sources. Section 4.17.6 discusses the tsunami propagation, from its hydrodynamic background to numerical computations. Section 4.17.7 discusses the tsunami warning systems, both basin-wide and local, and hazard mitigation methods such as probabilistic hazard assessments and hazard maps.
4.17.2 The 2004 Indian Ocean Tsunami A giant earthquake (Mw ¼ 9.1–9.3) occurred off Sumatra Island on 26 December 2004 (Lay et al., 2005; Stein and Okal, 2005). This earthquake generated tsunami which devastated the shores of Indian Ocean (Figure 1); more than 200 000 people lost their lives. The number of victims, dead and missing together, is the largest in Indonesia (163 795), followed by Sri Lanka (35 399), India (16 389), and Thailand (8 345) (International Federation of Red Cross and Red Crescent Societies, 2005). Within 30 min of the earthquake, the tsunami devastated Banda Aceh and other coastal villages of Sumatra Island (Borrero, 2005a). The tsunami arrived at
Tsunamis 30° N
485
11
10
10
6
9
5
11
1
0°
2
4
6
3
7 3 5
4
4
6
8 6
7
10
11
7
6
9
12
5
5
8
30° S
7 8
30° E
60° E
7
90° E
120° E
150° E
Figure 1 Tsunami propagation in the Indian Ocean. Each curve indicates wavefront at 1 h increments. Black star and circles indicate the epicenters of mainshock and aftershocks within 1 day, respectively (data by USGS).
Thai, Sri Lankan, and Indian coasts in about 2 h after the earthquake. The tsunami propagated farther and arrived at the eastern coast of Africa several hours after the earthquake to cause 298 fatalities in Somalia. The tsunami was also observed in the Atlantic and Pacific Oceans (Titov et al., 2005a). To document the 2004 tsunami, many scientists from all over the world visited the affected coasts (e.g., Borrero, 2005b; Liu et al., 2005; Satake et al., 2006a; Tsuji et al., 2006). The measured tsunami heights in Sumatra Island, particularly in the Aceh province, were mostly larger than 20 m with the maximum above 30 m. The tsunami heights along the Andaman Sea coast were highly variable; 5–15 m in Thailand (Figure 2) but less than 3 m in Myanmar. The tsunami heights were up to 5 m on India’s Andaman Islands. In Sri Lanka, the tsunami heights were 5–15 m. The tsunami height distribution is consistent with the damage distribution and indicates that the source of the large and damaging tsunami was concentrated in the southern 700 km section of the aftershock zone. Sedimentological characteristics of the modern tsunami deposits were examined ( Jaffe et al., 2006; Moore et al., 2006), with hopes that they provide valuable information for the studies of paleo-tsunami deposits. The tsunami was instrumentally recorded on coastal tide gauges in Indian Ocean (Merrifield et al., 2005; Nagarajan et al., 2006). This tsunami was also recorded by other instruments, some were for the first time. The tsunami propagation in deep water was
Figure 2 Tsunami damage at Ban Nam Ken, about 100 km north of Phuket, Thailand.
captured by satellite altimeter (Gower, 2005). The tsunami signal was also recorded on horizontal components of broadband seismic stations near coasts (Yuan et al., 2005) or hydrophones (Hanson et al., 2007). Seismological analyses of the Sumatra–Andaman earthquake indicate that the source region extended from off Sumatra through the Nicobar to Andaman Islands with the total length of 1200–1300 km, and the total duration of rupture was about 500 s (Ishii et al., 2005; Tsai et al., 2005). The fault slip was the largest near off the northern Sumatra, followed by off Nicobar Islands (Ammon et al., 2005). Fault slip around Andaman Islands was estimated to be small from seismological analysis and speculated to be slow (Bilham, 2005; Lay et al., 2005).
486
Tsunamis
Analyses of satellite images indicate that the sea level changes were observed in Sumatra, Nicobar, and Andaman Islands, suggesting that the source length was as long as 1600 km (Meltzner et al., 2006; Tobita et al., 2006). Field investigations indicate that the largest subsidence was 3 m at Great Nicobar Island, while the small islands west off Middle Andaman uplifted about 1.5 m (Kayanne et al., 2007; Malik and Murty, 2005). Tide gauge record at Port Blair in South Andaman shows a slow subsidence with a duration of at least 15 min but possibly less than 35 min (Singh et al., 2006). The tsunami source was estimated from tsunami arrival times as 600–700 km (Lay et al., 2005), but later revised to be 900 km (Neetu et al., 2005) extending to the south of Andaman Islands. The satellite altimeter data support longer, more than 1000 km long, tsunami source (Fine et al., 2005b; Song et al., 2005; Hirata et al., 2006). The tide gauge data do not support tsunami source beneath Andaman Islands (Tanioka et al., 2006; Fujii and Satake, 2007), although Piatanesi and Lorito (2007) estimated slip on the deeper part of slab, east of Andaman Islands. Joint inversion of tide gauge and satellite data prefer the source lengh of 900 km (Fujii and Satake, 2007) (Figures 3 and 4). Seno and Hirata (2007) speculated that the fast rupture was followed by slow slip at shallower part near the trench. The source of the previous 1881 earthquake around Car Nicobar was studied from tide gauge records around Bay of Bengal (Ortiz and Bilham, 2003). The potential tsunami hazard in Indian Ocean was pointed out before the 2004 tsunami, with the computer simulation of the 1833 earthquake off Sumatra (Cummins, 2004). Future tsunami forecast models in the region were examined by Geist et al. (2007).
4.17.3 Other Examples In this section, other examples of tsunamis are briefly described in reverse-chronological order. Their sources are shown in Figure 5. We start with the recent and historical tsunamis that were instrumentally recorded, then prehistoric tsunamis that were geologically recorded. 4.17.3.1
Historical Tsunamis
4.17.3.1.1 tsunami
1998 Papua New Guinea
Another recent tsunami hazard was from an earthquake (Mw ¼ 7.1) along the northern coast of New Guinea Island on 17 July 1998. The tsunami heights
were as large as 15 m around Sissano Lagoon near the epicenter, and the total casualties were reported to be more than 2000. The large tsunami (>5 m), however, was limited to a small region (40 km) along the coast (Kawata et al., 1999) and the amplitudes on tide gauges around Japan are <10 cm (Satake and Tanioka, 2003), suggesting a local effect of tsunami generation or amplification. Post-tsunami marine surveys revealed bathymetry features that focused tsunami energy toward Sissano Lagoon and possible sources of submarine landslide (Tappin et al., 1999). Numerical simulations show that a source in addition to the earthquake fault motion, probably from submarine landslides, is needed to reproduce the locally large tsunami heights (Matsuyama et al., 1999; Geist, 2000; Titov and Gonzalez, 2001; Synolakis et al., 2002; Satake and Tanioka, 2003). 4.17.3.1.2
1993 Hokkaido tsunami The most recent tsunami disaster in Japan was that in Okushiri Island in Japan Sea, caused by the Southwest Hokkaido earthquake on 12 July 1993 (Mw ¼ 7.8). The tsunami attacked the island within a few minutes after strong ground shaking, and caused more than 200 casualties. The tsunami heights were 5–10 m around the island, while maximum tsunami run-up of more than 30 m occurred near a small valley. Numerical computations of tsunamis were made to reproduce the measured tsunami heights, recorded tsunami waveforms on tide gauges, as well as seismic or geodetic data (Shuto and Matsutomi, 1995; Takahashi et al., 1995; Tanioka et al., 1995; Titov and Synolakis, 1997). 4.17.3.1.3
1960 Chilean tsunami The largest (Mw ¼ 9.5) earthquake in the twentieth century occurred offshore southern Chile on 22 May 1960. The tsunami caused damage on the Chilean coast, including more than 1000 casualties, then propagated across the Pacific Ocean. About 15 h later, the tsunami reached the Hawaiian Islands and took 61 lives. It continued across the Pacific, reaching Japan in about 23 h, and caused 142 casualties. The large (5 m) tsunami in Japan was due to both focusing and resonance effects. Because of the sphericity of the Earth and bathymetry in the Pacific Ocean, energy was focused toward the Japanese Islands (Satake, 1988; Shuto, 1991). Some bays along the Pacific coast of Japan have a shape whose characteristic period coincided with the predominant period (about 1 h) of the trans-Pacific tsunami (Abe, 2000). This tsunami motivated to form an international tsunami warning system in the Pacific.
Tsunamis
70°
80°
90°
100°
487
110°
20°
10°
0°
–10°
–20°
Figure 3 Locations of instrumental observation for the 2004 tsunami: tide gauge stations (triangles) and satellite path (solid line). From Fujii Y and Satake K (2007) Tsunami source model of the 2004 Sumatra–Andaman earthquake inferred from tide gauge and satellite data. Bulletin of the Seismological Society of America 97: S192–S207.
4.17.3.1.4
1958 Lituya Bay tsunami The largest water run-up was recorded in Lituya Bay, Alaska. A large strike-slip earthquake (Mw ¼ 7.9) on 10 July 1958 triggered a huge rockslide (700–900 m wide, 1000 m long, and 90 m thick, or the total volume of about 3 107 m3) in this 10 km long bay. The rockslide generated water waves that surged up the opposite slope, and stripped the trees up to 520 m altitude (Miller, 1960; Lander and Lockridge, 1989). The large waves were confined in the bay; hence, this is not considered as a typical tsunami.
Lockridge, 1989). The tsunami traveled southward through the Pacific Ocean and reached the Hawaiian Islands in the morning of 1 April. Without any warning, the tsunami took 159 lives in the Hawaiian Islands. A Pacific tsunami warning system was introduced after this tsunami. Recent studies indicate that slow earthquake rupture process and large submarine landslide are responsible to far-field and near-field tsunamis, respectively (Okal et al., 2002, 2003).
4.17.3.1.6 4.17.3.1.5
1946 Aleutian tsunami The Aleutian earthquake of 1 April 1946 was very unusual; it generated huge tsunamis in the Aleutians and Hawaii, even though its surface wave magnitude Ms was only 7.4. This is a typical example of ‘tsunami earthquake’, which generates much larger tsunamis than expected from seismic waves (Kanamori, 1972). The Scotch Cap lighthouse on a 10 m cliff on Unimak Island, about 100 km from the epicenter, was totally demolished by the tsunami (Lander and
1896 Sanriku tsunami The 1896 Sanriku earthquake is another example of ‘tsunami earthquake’ (Kanamori, 1972; Tanioka and Satake, 1996b). The earthquake (Ms ¼ 7.2) was not great in size and the ground shaking was weak (2–3 on the Japanese intensity scale, corresponding to IV–V on modified Mercalli scale). However, the following tsunami was very destructive; the run-up height was >10 m for about 200 km of Sanriku coast with the maximum of 38 m, and resulted 22 000 casualties, the worst tsunami hazard in Japan in the last few centuries.
488 Tsunamis
90°
95°
4.17.3.1.7
1883 Krakatoa tsunami The 1883 eruption of Krakatoa, Indonesia, is one of the largest volcanic eruptions recorded in human history. The sound waves generated by the eruption were heard over 4000 km across the Indian Ocean. The eruption refaced the island of Krakatoa; the northen portion of the island disappeared and a caldera up to 270 m deep replaced it. The eruption also generated tsunamis that destroyed nearby coastal villages and took about 34 000 lives (Simkin and Fiske, 1983; Nomanbhoy and Satake, 1995). The maximum tsunami height along the Sunda Strait was about 15 m. Tide gauge disturbances were observed in India, New Zealand, and even at San Francisco, but those at very far places are interpreted as sea waves coupled by air waves from the eruption (Press and Harkrider, 1966; Pelinovsky et al., 2005).
100°
15°
10°
5°
4.17.3.2
Prehistoric Tsunamis
4.17.3.2.1
1700 Cascadia tsunami Geological and historical studies show the evidence of past earthquakes and tsunamis. Evidence of a great prehistoric earthquake that occurred along the Cascadia subduction zone has been recently found both near and far from the source (Atwater et al., 2005). Along the Pacific coast of North America, geological and tree-ring evidence indicates that coseismic subsidence took place about 300 years ago. In addition, tsunami deposits found throughout the Pacific Northwest indicate that this earthquake
0° 0
5
10
15 20 Slip (m)
25
30
35
Figure 4 Slip distribution on fault for the 2004 Sumatra– Andaman earthquake by an inversion of tsunami waveforms recorded on tide gauges and sea surface height measured by satellite. From Fujii Y and Satake K (2007) Tsunami source model of the 2004 Sumatra–Andaman earthquake inferred from tide gauge and satellite data. Bulletin of the Seismological Society of America 97: S192–S207.
90° N 1958 Lituya Bay
Storegga slides
60° N 1993 Hokkaido 30° N
1946 Aleutian 1896 Sanriku 1700 Cascadia
2004 Sumatra 1998 Papua New Guinea
0° 1883 Krakatoa 30° S
1960 Chile
60° S 90° S 60° E
120° E
180°
120° W
60° W
0°
60° E
Figure 5 Distribution of tsunamis in the world. Open circles are historical events with intensity 2 and validity 3 from Historical Tsunami Data Base (see Section 4.17.4.6) (Institute of Computational Mathematics and Mathematical Geophysics 2005). Solid circles are those described in Sections 4.17.2 and 4.17.3.
Tsunamis
generated a sizable tsunami. In Japan, historical records show a widespread tsunami with heights of 1–5 m struck the Pacific coast. From these records, the date and size of the earthquake was inferred to be 26 January 1700 and Mw ¼ 8.7–9.2, respectively (Satake et al., 1996, 2003). 4.17.3.2.2 Hokkaido tsunami in seventeenth century
Tsunami deposits show that the southern Kuril trench repeatedly produced tsunamis larger than those recorded in the last few centuries in Japan. Prehistoric sand sheets extend as much as 3 km inland along the 200 km of Hokkaido’s Pacific coast, where the recent tsunamis from 1952 to 2003 Tokachi-oki or 1894 to 1973 Nemuro-oki earthquakes (Mw 8) penetrated less than 1 km from the coast. The time intervals between the extensive sand sheets average about 500 years, and the most recent event occurred in the seventeenth century. These anomalous tsunamis, characterized by large inundation area and long recurrence interval, are best explained by earthquakes that rupture multiple segments, including Tokachi-oki and Nemuro-oki, of the Kuril subduction zone (Nanayama et al., 2003; Satake et al., 2005b). 4.17.3.2.3 slide
shorter period (less than 5 min), however, the system does not always follow the actual sea level change (Satake et al., 1988). More recent gauges include bubbler, acoustic, and laser gauges. The bubbler gauge is a pneumatic or gas-purged pressure gauge system. Compressed air or nitrogen gas is purged into a tube with an orifice, and its pressure is measured. As the gas slowly escapes from the orifice, the gas pressure is considered to be equal to water pressure, then converted to water level. The acoustic gauge sends ultrasonic waves in a tube that are reflected at the sea level. Sea level changes can be measured from the travel times of the acoustic wave (Figure 6). Laser gauge employs the same principle. The Global Sea Level Observation System (GLOSS), an international program, has been formed to establish high-quality global sea level network for climate change, oceanography, and coastal research (Woodworth and Player, 2003). Many GLOSS data are available in real time through the Internet. Because the main purpose of the sea level measurements is to monitor ocean tides or long-term sea level change, the typical sampling interval is not always short enough for tsunami monitoring. For tsunami measurements, high sampling rate, at least 1 min, is required.
Tsunami from the Storegga
The Storegga slide in the Norwegian Sea is the largest submarine landslide ever mapped. The Storegga slide is estimated to have taken place thrice: once at 30–50 ka and twice at 6–8 ka, with the total volume of 5.6 103 km3 (Bugge et al., 1988). The gigantic submarine landslide probably caused a tsunami. Tsunami deposits inferred from this event have been found on the coasts of Scotland, Iceland, and Norway (Smith et al., 2004; Bondevik et al., 2005).
4.17.4 Tsunami Observations 4.17.4.1
489
Instrumental Measurements
4.17.4.1.1 Coastal tsunami measurements
Coastal sea level has been measured for more than a century by tide gauges at ports and harbors throughout the world. The classic tide gauge is a mechanical type set in a tide or stilling well. Sea level change is detected by a float in the well. Because typical period of ocean tides is several hours or longer, sea level changes inside the well can be assumed to be the same as that outside the well. For tsunamis with
4.17.4.1.2 Offshore tsunami measurements
As tsunami becomes slow but large and complex near the coast, it would be ideal to detect its arrival offshore. Offshore wave gauges and GPS tsunami gauges have been developed and installed around Japan. Among about 50 acoustic wave gauges outside harbors at the water depths of tens of meters, about 10 recorded the tsunami from the 2003 Tokachi-oki earthquake (Nagai et al., 2004). A GPS tsunami gauge system has been installed about 10 km offshore Cape Muroto, Southwestern Japan, where the water depth is about 100 m (Kato et al., 2000, 2005). The gauge uses Real-Time Kinematic GPS technique to estimate the location and altitude of water surface. The sampling interval is 1 s with an accuracy of a few cm. Because it measures water surface movements, it is free from seismic waves unlike bottom pressure gauges (see the next section), and can also be used to monitor wave heights during storms. 4.17.4.1.3
Deep ocean measurements In the deep ocean, tsunami waveforms are expected to be simpler, free from coastal topographic effects, although their amplitudes are smaller. Deep ocean
490 Tsunamis
Figure 6 Types of instruments.
measurements of tsunamis have been made by using bottom pressure gauges for early detection and warnings of tsunami. A quartz crystal sensor is typically used as a pressure sensor. The sensor provides accuracy of 1 cm in sea level at a 6000 m depth. Because the pressure transducer is very sensitive to temperature, measurements and correction for temperature are essential. In addition, because the sensor is located at the ocean bottom, it also records ground motion and works as an ocean bottom seismograph. While the frequencies are often different between seismic and tsunami waves, it is necessary to separate them. Around Japan, more than 10 bottom pressure gauges are installed on six submarine cables (Okada, 1995; Iwasaki et al., 1997; Hirata et al., 2002). Tsunamis from relatively small tsunamigenic earthquakes, the 2004 Off Kii Peninsula earthquakes (M 6.9 and 7.4), were recorded on a GPS tsunami gauge and eight bottom pressure gauges, as well as coastal tide gauges located south of Honshu and Shikoku (Satake et al., 2005a). The US system, called Deepocean Assessment and Reporting of Tsunamis (DART) or simply Tsunameter, sends signals to surface buoy via acoustic telemetry in the ocean, then via satellites to land station in real time (Gonzalez et al., 2005). 4.17.4.1.4
Satellite observations Four satellites captured the tsunami propagation across the Indian Ocean on 26 December 2004 (Figure 3): Jason-1 and TOPEX/Poseidon at about 2 h after the earthquake, Envisat at 3 h 15 min, and
US Navy’s GEOSAT follow-on at 7 h (Gower, 2005). The maximum sea surface height (SSH) of the tsunami was about 0.8 m in the Indian Ocean. The accuracy of satellite altimetry is better than 5 cm and its spatial resolution is approximately 15 arc-min (27 km). Numerical simulation from the 2004 source model could reproduce the SSHs (Smith et al., 2005; Hirata et al., 2006) (see Section 4.17.2). 4.17.4.1.5 Definition of amplitude and wave heights
The tsunami ‘amplitude’ is measured from zero to peak, either positive or negative. The ‘wave height’, also called as range or double amplitude, is measured from trough to peak. Hence, the ‘tsunami height’ is measured from trough to peak on tide gauges, whereas the ‘tsunami amplitude’ is measured from estimated tidal level at the time (Figure 7). As in seismic waves, ‘initial motion’, either positive or negative, carries information on the tsunami source. For example, the initial motions of the 2004 tsunami in Indian Ocean were negative to the east of the source (e.g., Thailand) but positive to the west (e.g., Maldives) as shown in Figure 7. 4.17.4.2 Field Survey to Measure Run-Up Heights 4.17.4.2.1
Post-tsunami surveys Post-tsunami field surveys are often carried out after large tsunamis. Usually, experts from different backgrounds, such as coastal engineering, geology, or
Tsunamis
2
Male, Maldives (Max) Amplitude
491
(Max) Wave height
Amplitude (m)
0
–2 2
Estimated tide
Arrival time Phuket, Thailand (Max) Wave height
Estimated tide
0 (Max) Amplitude –2 –3
0
3
6
9
12
Time (hour from earthquake) Figure 7 Tsunami waveforms from the 2004 Sumatra–Andaman earthquake recorded at Male (top) and Phuket (bottom).
seismology, and from different countries form international tsunami survey teams that visit affected coasts to document tsunami effects. An essential element of survey is to measure the tsunami heights along the coast (Intergovernmental Oceanographic Commission, 1998). The heights, defined below, are measured from various traces such as damage on buildings or other construction, watermarks left on walls, vegetative markers such as trim lines and saltwater discoloration or die-off, or lines of debris deposited on beaches. When no physical evidence is found, the tsunami heights are estimated based on eyewitness accounts, but considered as less reliable. Locations are usually identified by portable GPS measurements and the heights are measured by traditional leveling or laser surveying. Interviews with local people also include degree of ground shaking, tsunami arrival times, the number of waves with periods, and the duration. Post-tsunami surveys also include description and documentations of building damage and tsunami deposits. Synolakis and Okal (2005) review and summarize 15 post-tsunami surveys made by international tsunami survey teams during the 11-year period between 1992 and 2002.
Figure 8 Definitions of tsunami heights on land.
4.17.4.2.2 Definitions of inundation and run-up heights
Tsunami height on land is called ‘flow depth’ when measured from the ground level, and ‘inundation height’ measured relative to sea level at the time of tsunami arrival. The height at maximum inundation is called ‘run-up height’. Because the tide levels at the time of measurements and tsunami arrival are usually different, the corrections are necessary for survey data. Unless the sea surface of inundated tsunami is flat, run-up height is not necessarily the same as the inundation height or flow depth at the shore. The horizontal distance of water penetration measured from the coast is called ‘inundation distance’ (Figure 8). 4.17.4.3
Study of Historical Documents
Tsunamis and their effects documented in historical records are also used to study old earthquakes and tsunamis for which no instrumental data are available. Some countries in Europe, South America, or Asia have a long history of written records that provide accurate dates for past earthquakes and
492 Tsunamis
tsunamis. For example, tsunami heights from the 1755 Lisbon earthquake have been estimated from historical documents as 6 m in Lisbon, more than 15 m at Southwest Portugal, and more than 10 m in Spain and Morocco (Baptista et al., 1998). Estimation of tsunami heights from damage description requires many assumptions. For example, the Japanese descriptions of the January 1700 tsunami, generated by a giant earthquake in Cascadia subduction zone, include flood in farmed fields and a government warehouse and damage to coastal houses (Atwater et al., 2005). To convert these reported effects into estimated heights, Satake et al. (2003) considered various factors, such as the relationship between tsunami damage and heights, sea level change in the last 300 years, and similarity with the 1960 Chilean tsunami. Japanese tsunami data indicate that the house damage is partial when tsunami height is less than 2 m, but becomes significant when tsunami height becomes more than 4 m (Hatori, 1984).
4.17.4.4
Geological Methods
Even older tsunamis are studied from their deposits by geological means (Dawson and Shi, 2000). In Europe, inferred tsunami deposits from the Storegga slides in the North Sea have been found and studied in Scotland, Norway, and Iceland (Dawson et al., 1988). In the US, tsunami deposits from prehistoric great earthquakes along the Cascadia subduction zone have been studied (Atwater et al., 2005). In Australia, geological evidence of Late Pleistocene tsunamis has been studied (Bryant et al., 1996). In Japan, tsunami deposits in various environments such as lagoons, beaches, or marshes have been studied for modern, historic, and prehistoric tsunamis (Minoura and Nakaya, 1991; Nanayama et al., 2003). On some coasts, particularly those remote from tsunami sources, it can be difficult to distinguish between tsunami and storm deposits. The universal sedimentological criteria for identifying tsunami deposits have not been well established (Nanayama et al., 2000; Tuttle et al., 2004). For this context, studies of deposits from modern tsunamis are important and careful descriptions have been done for the recent tsunamis such as the 1993 Hokkaido tsunami (Nishimura and Miyaji, 1995), the 1998 Papua New Guinea tsunami (Gelfenbaum and Jaffe 2003), or the 2004 Indian Ocean tsunami (see Section 4.17.2). These case studies indicate that the tsunami deposits
are highly variable and depend on the coastal topography and geology.
4.17.4.5
Tsunami Magnitude Scales
4.17.4.5.1
Imamura-Iida scale, m The traditional tsunami magnitude scale is the socalled Imamura-Iida scale, m. Although the original definition was descriptive, the value is approximately equal to m ¼ log2 h
½1
where h is the maximum run-up height in meters (Iida et al., 1967). This scale is similar to the earthquake intensity scale, and is especially convenient for old tsunamis from which no instrumental records exist. Hatori (1979) extended the Imamura-Iida’s m scale to include far-field tsunami data as follows: n o pffiffiffi m ¼ 3 þ log ðh=0:5ÞðR=1000Þ1=2 =log 5
½2
where h is in meters and R is the distance from source to coast in kilometers. This indicates pffiffiffithat the tsunami height decays with distances as 1= R; which is theoretically predicted for nondispersive waves traveling a long distance (Comer, 1980). 4.17.4.5.2
Tsunami intensity Soloviev (1970) pointed out that Imamura-Iida’s m scale is more like an earthquake intensity scale rather than a magnitude. He also distinguished the maximum tsunami height h and the mean tsunami height Hav. He then defined tsunami intensity I as I ¼ log2
pffiffiffi 2Hav
½3
Comparison of [1] and [3] suggests that the mean pffiffiffi tsunami height is given as 1= 2 times the maximum height. The maximum intensity I on the coast nearest to the source is used to quantify the tsunami source. 4.17.4.5.3
Tsunami magnitude M t Another magnitude scale Mt, called tsunami magnitude, was defined and assigned for many earthquakes by Abe (1979). The definition of Mt for a transPacific tsunami is (Abe, 1979) Mt ¼ log H þ C þ 9:1
½4
and for a regional (100 km < < 3500 km) tsunami is (Abe, 1981) Mt ¼ log H þ log þ 5:8
½5
Tsunamis
where H is a maximum amplitude on tide gauges in meters, C is a distance factor depending on a combination of the source and the observation points, and is the nautical distance in kilometers. The above formulas were calibrated with the moment magnitude scale, Mw, of earthquakes. Equation [5] indicates that for the same Mt, the tsunami amplitude H decays with 1/, which is different from the assumption used for Imamura-Iida’s m scale, and may be valid only for regional distances. 4.17.4.6
Tsunami Catalogs
The location, size, run-up heights, and damage from past tsunamis have been compiled by several researchers and published as catalogs. For the Pacific tsunamis, the Iida catalog (Iida et al., 1967) contains numerous tsunami height data as far back as AD 173. Russian catalogs (Soloviev and Go, 1974a, 1974b; Soloviev et al., 1992) include Pacific tsunami data (but except for Russia) with some tide gauge records, tabulated tsunami heights, and descriptions for each tsunami. Regional catalogs including tide gauge records were also published for Alaska (Lander, 1996), the west coast of the US (Lander et al., 1993), Japan (Iida, 1984; Watanabe, 1998), and Indian Ocean (Berninghausen, 1966; Murty and Rafiq, 1991). Digital tsunami databases have been compiled, termed Historical Tsunami Data Base at the Institute of Computational Mathematics and Mathematical Geophysics, Novosibirsk (Gusiakov, 2005; Institute of Computational Mathematics and Mathematical Geophysics, 2005) and at the National Geophysical Data Center, NOAA.
4.17.5 Tsunami Generation In this section, we review tsunami generation by earthquakes, landslides, and volcanic eruptions. 4.17.5.1
Earthquake Tsunamis
4.17.5.1.1 Ocean bottom and water surface displacements
Analytical solutions for linear gravity waves generated from several types of sources have been obtained by various researchers. Kajiura (1963) generalized the solutions by introducing the tsunami Green’s functions. He showed that the initial conditions given at the ocean bottom and at the water surface differ by a factor of 1/cosh kd, where k is wave number of
493
displacement. For large wavelength, >> 2d, this factor becomes 1 because cosh kd ¼ cosh (2d/) 1. Therefore, if the wavelength of ocean-bottom deformation is much larger than the water depth, it can be assumed that the water surface displacement is the same as the bottom displacement. Kajiura (1970) examined the energy exchange between the solid Earth and ocean water for tsunamis generated by bottom deformation with a finite rise time. He showed that the efficiency of tsunami generation, defined as a ratio of dynamic tsunami energy to static energy, becomes nearly 1, if the duration is short compared to the time required for a tsunami to travel through the source area. For an earthquake with M 8 at the water depth of 5000 m, the source size is about 100 km and the tsunami travel time over that distance is about 8 min. If the duration is much shorter, for example, 1 min or less, the efficiency is more than 0.9. This means that the bottom deformation can be considered as instantaneous. If the pffiffiffiffiffiffiffiffi duration is too short, shorter than 2 d =g ; a part of the energy flux goes into acoustic wave. For the above case, this limit is 45 s. 4.17.5.1.2 Fault parameters and ocean-bottom deformation
The ocean-bottom deformation due to faulting can be calculated using the elastic theory of dislocation. The displacement in an infinite homogeneous medium due to a dislocation across a fault is given by Volterra’s theorem (Steketee, 1958) and the explicit formulas for half-space are given (e.g., Mansinha and Smylie, 1971; Okada, 1985). Usually, only the vertical component of the ocean bottom is considered for tsunami generation. When the tsunami source is on a steep ocean slope and the horizontal displacement is large, the vertical displacement of water due to the horizontal displacement of the slope must also be considered (Tanioka and Satake, 1996a). The fault parameters needed to compute surface deformation are the fault location, geometry (strike, dip, and rake), size (length L and width W ), and average slip u (or the strike-slip and dip-slip components). The seismic moment M0 is given as M0 ¼ uS ¼ uLW
½6
where S is the fault area and is rigidity. The moment magnitude is defined as (Hanks and Kanamori, 1979) Mw ¼
2 log M0 – 10:7 3
½7
494 Tsunamis
where M0 is given in dyne cm (107 N m). The effects of fault parameters on tsunami generation were examined by Yamashita and Sato (1974) and more recently by Geist (1998).
4.17.5.1.3 Estimation of earthquake source parameters from tsunamis
Tsunami data can be used to study earthquake source processes in a way similar to that in which seismic waves are used. Abe (1973) analyzed the tsunami arrival times and first-motion data from the 1968 Tokachi-oki earthquake. By drawing inverse refraction diagrams from each station (see Section 4.17.6.2), he estimated the tsunami source area, which agrees well with the aftershock area. From the first motion of tsunami waves, the initial water surface disturbance was estimated as uplift at the southeastern edge and subsidence at the northwestern edge. This pattern is very similar to the vertical bottom deformation due to the faulting, which was independently estimated from seismological analysis. Numerical simulation of tsunami has been carried out for many fault models around Japan (e.g., Aida, 1978). The best fault models were obtained by trial and error comparisons of observed and computed tsunami waveforms or coastal tsunami heights. Tsunami waveforms recorded on tide gauges were then inverted to estimate the slip distribution on the fault (Satake, 1989). In this method, the fault plane is first divided into several subfaults, and the deformation on the ocean bottom is computed for each subfault with a unit amount of slip. Using this as an initial condition, tsunami waveforms are numerically computed on actual bathymetry. The observed tsunami waveforms are expressed as a superposition of waveforms computed for each subfault as follows: Aij ðt Þ ? xj ¼ bi ðt Þ
½8
where Aij (t) is the computed waveform as a function of time t, or Green function, at the ith station from the j th subfault; xj is the amount of slip on the j th subfault; and bi(t) is the observed tsunami waveform at the ith station. The slip xj on each subfault can be estimated by a leastsquares inversion of the above set of equations. The tsunami waveform inversion has been applied to many large and great earthquakes in the Alaska-Aleutians ( Johnson, 1999), northern Japan (Satake, 1989; Tanioka et al., 1995), the Nankai trough (Satake, 1993; Tanioka and Satake, 2001a, 2001b), and the 2004 Sumatra–Andaman earthquake (Fujii and Satake, 2007). These results show that the slip distributions
are not uniform but are concentrated in small regions, which correspond to asperities (Figure 3). 4.17.5.1.4 Tsunamigenic and tsunami earthquakes
We call an earthquake that generates tsunami as ‘tsunamigenic earthquake’. A ‘tsunami earthquake’, on the contrary, is defined as an earthquake which excites much larger tsunamis than expected from its seismic waves (Kanamori, 1972). The discrepancy can be quantified in terms of the surface wave magnitude (Ms) and tsunami magnitude (Mt) scales; when Mt is larger than Ms by more than 0.5 units, it is called a tsunami earthquake (Abe, 1989). Newman and Okal (1998) introduced a new parameter, called slowness parameter Y ¼ log10 ðE E =M0 Þ, a ratio of radiated seismic energy EE to seismic moment M0. They show that Y is about –5 for regular earthquakes but smaller by 1–1.4 for tsunami earthquakes, and they proposed to use this parameter for discrimination of tsunami earthquakes. Several proposals have been made for the generating mechanisms of tsunami earthquakes. Slow and long rupture processes (Kanamori, 1972) reconcile the large discrepancy between seismic and tsunami waves. High-angle secondary faulting in the accretionary wedge can explain the relatively large tsunami excitation (Fukao, 1979). For a given slip, the seismic moment is smaller in the shallower part of subduction zones because of the smaller rigidity, and the tsunami becomes larger (Okal, 1988). Tsunami heights may not be anomalously large if compared to Mw rather than Ms (Pelayo and Wiens, 1992); the large discrepancy between the tsunami and Ms is due to a saturation of Ms at around 7.3 for those earthquakes that occurred in the accretionary wedge. Recent tsunami waveform modeling shows that the fault parameters and locations of tsunami earthquakes share a common feature: narrow and shallow faulting near the trench axis (Figure 9). The 1992 Nicaragua earthquake (Ms ¼ 7.2, Mt ¼ 7.9) was the first tsunami earthquake recorded on modern broadband seismic instruments, and seismological studies showed that the duration was very long for its size, about 100 s (Kanamori and Kikuchi, 1993). Comparison of numerically computed tsunami waveforms with tide gauge records (Satake, 1994) showed that a narrow (40 km) and shallow (extending only to the upper 10 km of the ocean bottom) fault is responsible for the tsunami generation. Among the tsunami earthquakes, the 1896 Sanriku earthquake (Ms ¼ 7.2, Mt ¼ 8.6) and the 1946 Aleutian
Tsunamis
495
Figure 9 Vertical deformation of seafloor (top) due to faulting (bottom) for typical interplate earthquake (left) and tsunami earthquakes (right). Modified from Satake K and Tanioka Y (1999) Sources of tsunami and tsunamigenic earthquakes in subduction zones. Pure and Applied Geophysics 154: 467–483.
earthquake (Ms ¼ 7.3, Mt ¼ 9.3) were very anomalous (Kanamori, 1972; Kanamori and Kikuchi, 1993). While the tsunami modeling showed a similar character in fault parameters to the Nicaragua earthquake, that is, narrow fault near trench axis (Tanioka and Satake, 1996b; Johnson and Satake, 1997; Tanioka and Seno, 2001a, 2001b), Kanamori and Kikuchi (1993) suggested that submarine slumping was involved in the unusually large tsunami generation. These tsunami earthquakes occurred in trenches with large amounts of sediments and an accretionary prism where slumping occasional occur. Other tsunami earthquakes such as the 1992 Nicaragua earthquake occurred in subduction zones with little sediment, where a slow rupture is responsible for large tsunamis (Kanamori and Kikuchi, 1993). Tanioka et al. (1997) showed that the 1896 tsunami earthquake occurred in a region where the oceanbottom topography is rough, characterized by a welldeveloped horst and graben structure. Polet and Kanamori (2000) extended this model to global subduction zones, based on the examination of source spectra of large (M > 7) earthquakes in the 1990s. Okal and Newman (2001), however, failed to find any correlation between the slowness parameter Y and either depth, focal mechanism, or seismic moment for central America, Java, and Peru,
suggesting an absence of regional signal for the occurrence of tsunami earthquakes. Bilek and Lay (2002), from analysis of source time functions of tsunamigenic and tsunami earthquakes, found that shallower earthquakes have longer duration and claimed that tsunami earthquakes are simply larger earthquakes at very shallow depth. Seno (2002) proposed that increase in fluid pressure causes temporal change in frictional property near trench axis and responsible for slow rupture; hence, the tsunami earthquakes are transient phenomena. Seno and Hirata (2007) showed that the 2004 earthquake had some component of ‘tsunami earthquake’, slow duration, and updip rupture of seismogenic zone. They inferred that the seismogenic rupture propagated toward north at the rupture velocity of 2.5 km s1, but was followed by slower rupture (0.7 km s1) at shallower part, which has about a third of the total seismic moment.
4.17.5.2
Tsunamis from Landslides
Tsunamis from submarine landslides have been a hot research topic since the 1990s. Abnormally large and concentrated tsunami heights from the 1992 Flores Island, Indonesia, and 1998 Sissano, Papua New Guinea, earthquakes and the nonseismic 1994
496 Tsunamis
Skagway tsunami, Alaska, provided case studies. Marine surveys revealed detailed bathymetric features of landslides or slumps. Physical experiments in tanks and numerical computations of landslidegenerated tsunamis have been developed. Such multidisciplinary studies have been summarized in several special issues (Keating et al., 2000; Bardet et al., 2003; Tappin, 2004). 4.17.5.2.1
Examples Post-tsunami field surveys after large tsunamigenic earthquakes sometimes show abnormally large tsunami heights concentrated in a small portion of coasts. While offshore bathymetry or coastal topography (such as V-shaped bays) can cause such focusing of incoming tsunamis, locally large tsunami may be caused by a local landslide, either subaerial or submarine, triggered by ground shaking. Large water waves caused by a rockslide in Lituya Bay at the time of 1958 earthquake are described in Section 4.17.3.1.4. Concurrent with the 1964 Alaska earthquake (Mw 9.2), in addition to Pacific-wide tsunamis, local tsunamis from landslides, were reported at more than 20 locations, with the largest run-up height of 50 m (Plafker et al., 1969; Lander, 1996). More recently, the tsunami heights from the 1992 Flore Island earthquake were mostly 4–5 m on most of the northern coast of the island, but 26 m at Riangkroko (Yeh et al., 1993). The tsunami heights from the 1998 Papua New Guinea tsunami, also described in Section 4.17.3.1.1, show local peak of 15 m (Figure 10). Larger-scale submarine slumping and turbidity currents were triggered by the 1929 Grand Banks earthquake (Heezen and Ewing, 1952; Fine et al., 2005a). The maximum tsunami run-up height of 13 m and 28 casualties made it the worst tsunami disaster in Canada. The turbidity current also caused damage on the trans-Atlantic submarine cables at 12 locations over a range of 600 km. An example of nonseismic landslide tsunami is the 1994 Skagway tsunami (Kulikov et al., 1996). A submarine landslide, without seismic trigger, occurred in Skagway Harbor in Alaska and generated local tsunami of 5–10 m high. The tide gauge records show a dispersive character, because of small size of the source (see the next section). 4.17.5.2.2 Characteristics and discrimination of landslide tsunamis
Scale length of submarine mass failure is very different from that of submarine faulting. While vertical movement can be hundreds of meters, the horizontal
scale is mostly less than 10 km for submarine mass failure. Seafloor deformation due to faulting, on the contrary, has at most several meters of vertical movement but the spatial extent can be several hundreds of kilometers. As will be discussed in next section, an earthquake source almost always satisfies the longwave (shallow water) condition, while the submarine landslide may not. Hence, the water waves from submarine landslides show a dispersive character and attenuate quickly without reaching a long distance. Okal and Synolakis (2004) fit the run-up height distribution (y) along a beach (in the y-direction) by the following function: ðyÞ ¼
b ½ðy – cÞ=a2 þ 1
½9
where parameters a, b, and c are estimated from the observed data (Figure 10). They show that I2 ¼ b/a, the aspect ratio of maximum run-up heights to extent of run-up along beach, can be used as a discriminant of the nature of tsunami source, either dislocation (earthquake fault) or landslide source. Seismic dislocation sources have I2 smaller than 104, whereas a landslide source has a larger value.
4.17.5.2.3
Modeling landslide tsunamis Submarine landslide sources of tsunamis have been also modeled in many ways. These can be grouped into three categories: dynamic fluid flow, kinematic rigid-body motion, and static initial profile. In the dynamic fluid flow approach, landslide is treated as fluid and its evolution is calculated simultaneously with tsunami generation. Jiang and Leblond (1994) treated a landslide as incompressible viscous flow and seawater as inviscid flow, and formulated the multiphase fluid dynamics. Fine et al. (2005a) applied the multiphase flow method to the 1929 Grand Banks landslide tsunami; Imamura et al. (2001), Heinrich et al. (2001), and Titov and Gonzalez (2001) applied a similar method to the submarine landslide of the 1998 Papua New Guinea tsunami. In the rigid-body approach, temporal change of water height due to progress of submarine landslide is prescribed and used to compute tsunami propagation. Harbitz (1992) assumed a box-shaped slide for the Storegga slide, Ward (2001) used a chain of rigid boxed to model the Storegga and other submarine slides. Tinti et al. (1997) introduced a block model of landslide and simulated the landslide motion, then
Tsunamis
Run-up, ζ (m)
15
Aspect ratio I 2 = 4.74 × 10–4
1998 PNG
c = –6 km
5
0
4
a = 19 km
b=9m
10
6
497
0
–100 2002 PNG
100
Aspect ratio I2 = 1.17 × 10–5
b = 1.5 m c = –25 km
2 0
a = 131 km
–100
100 0 Distance y along beach (km)
Figure 10 Comparison of tsunami run-up heights for the 1998 and 2002 Papua New Guinea tsunamis. Estimated parameters of [9] are also shown. Modified from Okal EA and Synolakis CE (2004) Source discriminants for near-field tsunamis. Geophysical Journal International 158: 899–912.
used the results for the tsunami computations. Satake et al. (2002) introduced a simple kinematic landslide model on the basis of bathymetry change. In the static approach, an initial profile due to landslide is assumed. Watts et al. (2003), based on the results of physical and numerical experiments, introduced a parametric initial profile for submarine landslide, and applied such initial profiles for several landslide-generated tsunamis including the 1946 Aleutian, 1994 Skagway, and 1998 Papua New Guinea tsunamis. Synolakis et al. (2002) and Satake and Tanioka (2003) used the same approach to model the 1998 Papua New Guinea tsunami. For water waves generated from subaerial landslides, various studies such as theoretical, experimental, numerical and field approaches are summarized in Slingerland and Voight (1979).
4.17.5.3
Volcanic Tsunamis
Tsunamis are one of the most important hazards associated with volcanic activity. About a quarter of human casualties directly associated with volcanic eruption is due to tsunamis (Tilling, 1989). Volcanic tsunamis can be roughly categorized into five groups according to their origin: (1) earthquakes; (2) landslides and avalanches of cold rock and lahars; (3) flow of volcanic ejecta such as pyroclastic flow, base surge, hot rock avalanches, or lava; (4) submarine explosion;
and (5) caldera collapse or subsidence. From case studies of 106 volcanic tsunamis, Latter (1981) classified 69 of them into the 10 groups. In addition, he also classified tsunamis generated from air waves of explosions (4.5%) with possible examples in Bezymianny, Kamchatka in 1955–1956, as well as some far-field tsunamis from the 1883 Krakatoa eruption. 4.17.5.3.1 Volcanic tsunamis due to earthquakes
According to Latter (1981), some 22% of volcanic tsunamis are caused by earthquakes accompanying eruptions. Examples include the 1914 Sakurajima, Japan, 1930 Stromboli, Italy, and 1933 Severgin, Kuril Islands, tsunamis. 4.17.5.3.2 Volcanic tsunamis due to landslides
Latter (1981) separated the tsunamis generated from landslides and avalanches of cold rock (7%) and lahars (5%). A typical example of landslide is from Shimabara, Japan. In 1792, Mayuyama, a small mountain near Unzen volcano, collapsed and the landslide entered Ariake-kai, a shallow bay with typical depth of 10–20 m. Because the horizontal extent of the slide is about 3000 m, much longer than the water depth, this satisfied the long-wave approximation, hence causing large (10–20 m) tsunami height and
498 Tsunamis
significant damage (5000 casualties) on the other side of the bay. The total casualties was 15 000.
4.17.6 Tsunami Propagation
4.17.5.3.3 Volcanic tsunamis from flow of volcanic ejecta
In the Earth-ocean-coupled system, an earthquake source would generate waves both in the elastic Earth and in the ocean, and the propagation of a tsunami in the ocean layer would excite waves in the elastic Earth. In contrast, the traditional decoupled approach first calculates the elastic deformation on the ocean bottom generated by an earthquake, then uses this as an initial condition for tsunami wave propagation such that the tsunami wave in the ocean would have no influence on the elastic Earth. Theoretical treatment of tsunamis viewed as coupled system is similar to that of Rayleigh waves or spheroidal oscillations. Ward (1980) and Okal (1988) extended the free oscillation approach to solve the generation and propagation of tsunamis, while Comer (1984a) extended the Rayleigh wave approach to solve the tsunami problem. The calculated dispersion curve shows that, except at very low frequencies, the phase velocity is very similar to that of the gravity wave. Both approaches give almost identical solutions, because of the small compressibility of water (Comer 1984b). In other words, the coupling between the elastic Earth and the ocean is extremely weak, which explains why very large earthquakes are necessary to generate tsunamis.
About a third of volcanic tsunamis can be classified in this group. Latter (1981) further separated them into pyroclastic flows (ignimbrites and nuees ardentes; 20%), base surges (7%), avalanches of hot rock (8%), or lava (1%). Latter (1981) considered the 1883 Krakatoa tsunami as due to pyroclastic flow, although this is controversial (see Section 4.17.5.3.5). Other examples of pyroclastic flow include St. Vincent and Mt. Pelee in lesser Antilles in 1902 and St. Augustine (Alaska) in 1883. Other types of ejecta are base surge in Taal volcano in the Philippines, hot rock avalanche at Stromboli, Italy, and lava avalanche at Matavanu volcano, Samoa in 1906–1907.
4.17.5.3.4 Volcanic tsunamis from submarine explosion
Submarine explosion occupies about 19% of volcanic tsunamis according to Latter (1981). Submarine volcano, Bayonnaize Rocks (along Izu-Bonin Islands about 300 km south of Honshu), erupted in 1952. Tsunamis from this eruption, recorded at Hachijo Island about 130 km north, show very short period (1 min) wave with a dispersive character. Unoki and Nakano (1953) applied a Cauchy–Poisson wave (deep water waves from initial water surface disturbance) to the observed tsunami record and estimate the source size (2 km) and water surface uplift or impulse force.
4.17.5.3.5 collapse
Volcanic tsunamis from caldera
About 9% of volcanic tsunamis were from caldera collapse or subsidence. Large-scale caldera collapse was recorded at the times of the 1883 Krakatoa and the 1883 Ritter Island (Papua New Guinea) eruptions. Ward and Day (2003) modeled the Ritter Island tsunami by using the landslide model. The generation mechanism of the 1883 Krakatoa tsunami had been long debated. A numerical computation of this tsunami and comparison with tide gauge record and run-up heights indicated that the tsunami was generated by submarine eruption, rather than caldera collapse, pyroclastic flow, or falling ejecta (Nomanbhoy and Satake, 1995).
4.17.6.1 Weak Coupling between Elastic Earth and Ocean
4.17.6.2
Ray-Theoretical Approach
If the tsunami wavelength is much smaller than the scale of heterogeneity in propagation velocity, or the change in water depth, then we can apply the geometrical ray theory of optics. The wavefronts of propagating tsunami can be drawn on the basis of Huygens’ principle. Such a diagram is called a ‘refraction diagram’. Refraction diagrams can be prepared for major tsunami sources and used for tsunami warning; as soon as the epicenter is known, the tsunami arrival times can be readily estimated. Figure 1 shows the refraction diagram from the 2004 Andaman–Sumatra earthquake with wavefronts at each hour. It is shown that the tsunami was expected to arrive at Thai and Indian coasts in 2 h and at Africa in 7 hs. The refraction diagrams can also be drawn backwards from coasts. Such a diagram is called an ‘inverse refraction diagram’ and used to estimate the tsunami source area. In this case, the wavefronts or
Tsunamis
rays are traced backwards for the corresponding travel times (Miyabe, 1934). The traced wavefronts from each observed station bound the tsunami source area.
4.17.6.3 Law
Coastal Amplification and Green’s
From the conservation of energy along the rays, it can be shown that 1=2
1=2
b0 d0 h02 ¼ b1 d1 h12
½10
where d is the water depth, b is the distance between the rays, h is the tsunami amplitude, and the subscripts 0 and 1 indicate two different locations. If the tsunami amplitude at location 0 (e.g., deep water) is known, the tsunami amplitude after propagation can be estimated as h1 ¼
1=2 1=4 b0 d0 h0 b1 d1
½11
This is known as Green’s law. The ratio b1/b0 represents the spreading of rays, which can be graphically obtained from refraction diagrams. Note that this applies only to direct waves and does not account for of the reflected waves or edge waves.
4.17.6.4
Tsunami as a Gravity Wave
In this section, we take a two-dimensional Cartesian coordinate system with z-axis vertical upward (the origin is on the undisturbed water level) and consider propagation of a wave (Figure 11). When gravity is the restoring force, it is called a gravity wave. Euler’s equation of motion can be written as qV 1 þ V ? rV ¼ – g – rp qt
z
½12
d
u
Figure 11 Definition of coordinate system and the parameters for gravity waves.
where V is the velocity vector (whose x- and z-components are u and w, respectively), g is the gravitational acceleration, and p are water density and pressure, respectively. The phase velocity c of gravity waves is given as c¼
g k
tanh kd
1=2
¼
g 2d 1=2 tanh 2
½13
where k is the wave number, is the wave length, and d is the water depth. When the horizontal scale of motion, or the wavelength , is much larger than the water depth (>>d), the vertical acceleration of water is negligible compared to the gravity. This means that the horizontal motion of water mass is almost uniform from bottom to surface. Such a wave is called a shallow-water wave or long wave. For tsunamis, typical ocean depth is about 4–5 km and a large earthquake has a source area of several tens to hundreds of kilometers. Hence, the long-wave approximation is appropriate for most cases. When the amplitude is small compared to the water depth (h << d), the equation of motion, or conservation of momentum, can be written as qu qh ¼ –g qt qx
½14
and the equation of continuity, or conservation of mass, as qh q ¼ – ðduÞ qt qx
½15
Such a wave is called a small-amplitude, or linear long wave. The assumptions are valid for most tsunami propagation paths except for near-shore regions. From [14] and [15], by assuming that water depth d is constant, we obtain the wave equation 2 pffiffiffiffiffi q2 h 2q h ¼ c ; c ¼ gd qt 2 qx 2
½16
in which the phase velocity is determined by the water depth only. Note that the above phase velocity can be also obtained from [13] if >>d.
λ h
499
x
4.17.6.5 Basic Equations for Tsunami Propagation Including the bottom friction and the Coriolis force, the momentum equation for shallow-water (long)
500 Tsunamis
waves can be written for a three-dimensional case as follows: pffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi qU qU qU qh U U 2 þV 2 þU þV ¼ – f V – g – Cf d þh qt qx qy qx ½17 pffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi qV qV qV qh V U2þV 2 þU þV ¼ f U –g – Cf qt qx qy qy d þh
These equations, including frequency dispersion, are known as the linear Boussinesq equation. If we relax the small-amplitude (h << d) assumption, the nonlinear, nondispersive equations of motion and continuity are
and the mass equation is qh q q þ fU ðh þ d Þgþ fV ðh þ d Þg ¼ 0 qt qx qy
½18
where the coordinate system is x ¼ east and y ¼ south, f is the Coriolis parameter, Cf is a nondimensional bottom frictional coefficient, and U and V are the depth-averaged velocities in x- and y-directions, respectively. The velocities are averaged from the bottom to the surface: Z h 1 Qx u dz ¼ h þ d –d hþd Z h Qy 1 V ¼ v dz ¼ h þ d –d hþd
qu qu qh þ u ¼ –g qt qx qx
½23
qh q ¼ – fuðh þ d Þg qt qx
½24
These equations are for the finite-amplitude shallowwater waves and account for amplitude dispersion. For the linear case (ignoring the advection term), the phase pffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi velocity is given by c ¼ gðd þ hÞ. This means that the larger the amplitude, the faster the wave speed. As a consequence, peaks of a wave catch up with troughs in front of them, and the forward facing portion of the wave continues to steepen and it will eventually break.
U¼
½19
where Qx and Qy are flow rates in x- and y-directions. A typical value of a non-dimensional frictional coefficient Cf is 103 for coastal water and 102 for run-up on land (Satake, 1995).
4.17.6.6
Dispersion
Here we will briefly examine general gravity waves deviating from the small-amplitude, linear shallowwater (long) wave approximations (Peregrine, 1972; Mei, 1989). We have already seen that the phase velocity of a general gravity wave is given by [13]. If we expand the hyperbolic tangent function into its Taylor series and include the second term, [13] becomes c¼
1=2 pffiffiffiffiffi pffiffiffiffiffi 1 1 gd 1 – ðkd Þ2 gd 1 – ðkd Þ2 3 6
½20
The corresponding equation of motion is qu qh 1 q3 u ¼ –g þ d 2 2 qt qx 3 qx qt
½21
and the equation of continuity is qh qu ¼ d qt qx
½22
4.17.6.7
Numerical Computations
The tsunami propagation depends on the water depth, which is not uniform but the variation is relatively well known compared to velocity distribution of seismic waves. Hence, numerical computations on realistic or actual bathymetry have been popularly made. The historical review of tsunami numerical simulation is presented by Shuto (1991). Let us consider a simple one-dimensional propagation (on constant depth) of linear long waves. Denoting the average velocity (x-component only) as U, the equations of motion and continuity are given as follows: qU qh ¼ –g qt qx
½25
qh qU ¼ –d qt qx
½26
Using the staggered (leap-frog) grid system for U and h, and denoting time t ¼ l t as superscripts and space x ¼ m x as subscripts where l and m are integers, these become h i lþð1=2Þ
l l hmþ Um – Uml – ð1=2Þ ð1=2Þ – hm – ð1=2Þ ¼ –g x t h i h i lþð1=2Þ lþ1 l hmþ Umþ1 – Umlþð1=2Þ ð1=2Þ – hmþð1=2Þ ¼ –d t x
½27
½28
Tsunamis
It can be shown that the error will grow with time unless the following stability condition is met: pffiffiffiffiffi t x 1 or t pffiffiffiffiffi gd x gd
x t pffiffiffiffiffiffiffi 2gd
U ði,it þ 1Þ ¼ U ði,itÞ – g
hði,it þ 1Þ ¼hði,itÞ t –d ½U ðiþ1,it þ 1Þ – U ði,it þ 1Þ x
The algorithm thus is as follows: set up parameters read bathymetry data read initial condition h (i, 1) while (it
h (i )
½31
½32
h (i + 1)
t + Δt
t x U(i – 1)
U(i )
U(i + 1)
Figure 12 Staggered leap-frog system.
compute U (i, itþ1) from U (i, it) and h (i, it) set U (i, itþ1) ¼ 0 for land boundary compute h (i, itþ1) from U (i, itþ1) and h (i, it) adjust h (i, itþ1) for open boundary exchange variables h (i, itþ1)!h (i, it) U (i, itþ1)!U (i, it) itþ¼ 1
½30
There are two types of boundaries involved in the tsunami computation. The first one is land–ocean boundary. The simplest assumption is a total reflection of energy on the coast. For the computation of tsunami inundation on land, however, a moving land–ocean boundary is needed; the boundary moves depending on the relative heights of water height and land. The other kind of boundary is the open ocean boundary along the outside of the computational region. A typical assumption is the radiation condition, which means that waves go out of the computational region keeping their slopes. The computer program for one-dimensional finite-difference computation is very simple. Since there is no overlap between the two (height and velocity) grids (Figure 12) in the linear case, in the program the spatial grid points m and m þ (1/2) can be treated by the same index (e.g., i) and the time l – 1/2 and l can be treated as the same time (e.g., it). Writing U and h using the two-dimensional array as U (i, it) and h(i, it), [27] and [28] become t ½hði,itÞ x – hði – 1,itÞ
t
½29
which is known as Courant–Friedrichs–Lewy (CFL) condition. The physical meaning of the stability condition is that the time step t must be equal to or smaller than the time required for the disturbance to travel the spatial grid size x. If t is larger, the values other than (l, m – l), (l – 1, m) and (l, m þ 1) will influence the value at (l þ1, m). For the two-dimensional case, the CFL condition becomes
501
end Figure 13 shows the snapshots of tsunami propagation from the 2004 Sumatra–Andaman earthquake.
4.17.7 Tsunami Warning and Hazard Reduction Systems The December 2004 tsunami in the Indian Ocean caused the worst tsunami disaster in the history. In addition to the gigantic size of the earthquake, lack of preparation for tsunamis on the coast of Indian Ocean was also responsible for the tragedy. Preparation for tsunami includes a tsunami warning system based on seismic and sea level monitoring, hazard assessment such as making hazard maps, and education and awareness to the coastal residents. Such systems on tsunami hazard reduction rely on the scientific knowledge of earthquakes and tsunamis. The US National Tsunami Hazard Mitigation Program (Bernard, 2005) is a comprehensive program based on three components: warning guidance, hazard assessment, and mitigation (Figure 14). Tsunami warning systems can be grouped into a basin-wide (such as Pacific or Indian Ocean), regional, and local systems. Hazard assessment includes making inundation maps or probabilistic maps. Mitigation includes
502 Tsunamis
Phuket Thailand
20° subsided 10° uplifted 0°
Sri Lanka Sumatra
–10°
0h
70°
1h 80°
90°
100°
2h
110°
Wa rn
on
ing
ati
tig
gu
Mi
ida
nc
e
Figure 13 Snapshots of simulation for the 2004 tsunami in Indian Ocean.
Hilo Bay
Hazard assessment Figure 14 Three key elements for tsunami hazard reduction: warning guidance, hazard assessment, and mitigation (Bernard, 2005).
public awareness such as implementing tsunami signs or educating coastal residents. 4.17.7.1 Far-Field Tsunami Warning Systems Tsunamis travel very fast as an ocean wave, at about 710 km h1, or 0.2 km s 1, for a water depth of 4000 m, yet much slower than seismic waves, which travel at 5–10 km s1. This velocity difference between the seismic and tsunami waves makes it possible to issue a tsunami warning by detecting seismic waves before the actual tsunami arrival. For distant tsunamis that propagate across the Pacific or Indian Ocean, there are a few hours for accurate and reliable warning. The Richard H. Hagemeyer Pacific Tsunami Warning Center (PTWC), located in Hawaii, is the operational center for the international tsunami warning system, and issues warnings for regional and distant tsunamis
in the Pacific Basin to almost all the countries and states around the Pacific Ocean (McCreery, 2005). The West Coast/Alaska Tsunami Warning Center, located in Alaska, is responsible for tsunami warning along the US and Canadian coasts except for Hawaii. At these centers, seismic waveform data from several tens of stations are continuously monitored, and once a large earthquake is detected, the location, depth, and magnitudes are assessed using a combination of automatic and interactive procedures by watchstanders (Figure 15). Recent deployment of advanced seismological parameters such as Mm (Okal and Talandier, 1989), Mwp(Tsuboi et al., 1995), or Y (Newman and Okal, 1998) makes it possible to quickly assess the earthquake size and tsunami potential, and to issue the tsunami warnings in less than 25 min. To avoid false alarm, it is very important to confirm that a tsunami was actually generated; hence, sea levels are monitored with both costal tide gauges and tsunameters (DART buoys). If a tsunami is observed, then the tsunami warning is updated (Titov et al., 2005b). If the tsunami cannot be confirmed, then the warning is canceled. At the time of 26 December 2004 tsunami, PTWC issued the first information bulletin at 1:14 GMT, only 15 min after the earthquake. An earthquake was located off the west coast of Northern Sumatra, but the magnitude (Mwp) was estimated as 8.0. The second bulletin was issued at 2:08 GMT, 69 min after the earthquake yet before the tsunami arrivals at Thai, Sri Lankan, or Indian coast. The earthquake size (Mm) was updated to 8.5 and a possibility of local tsunami was mentioned in the bulletin. The French Polynesia Tsunami Warning Center developed and adopted TREMORS (Tsunami Risk Evaluation through seismic Moment in a Real time System) in 1987 (Reymond et al., 1991). TREMORS
Tsunamis
503
90° N 60° N 30° N PTWC
0° 30° S Seismic Sea level DART
60° S 90° S 0°
60° E
120° E
180°
120° W
60° W
0°
Figure 15 Distribution of seismic stations and sea level monitoring, coastal gauges, and tsunameter (DART) stations that send the data to PTWC in real time.
features automatic detection and location of seismic events using a single three-component long-period seismic station, and estimation of seismic moment through variable-period mantle magnitude Mm (Okal and Talandier, 1989). The advantage of using long-period seismic waves for tsunami warning is the ability to detect ‘tsunami earthquakes’. The disadvantage is that longer time is needed before the seismic waves are recorded. TREMORS is a hybrid system; it updates the moment estimate as more seismic data are available. The Pacific-wide tsunami warning system was established after the 1946 Aleutian tsunami that affected the Hawaiian Islands (Section 4.17.3.1.5). After the 1960 Chilean tsunami (Section 4.17.3.1.3), the Pacific tsunami warning system became international, and an International Coordination Group for The Tsunami Warning system in The Pacific (ICG/ ITSU) was formed under the auspices of UNESCO Intergovernmental Oceanographic Commission. After the 2004 tsunami in Indian Ocean, it was renamed as ICG/PTWS (ICG for the Pacific Tsunami Warning and Mitigation System), and similar groups were formed for the other ocean basins: ICG/IOTWS for Indian Ocean, ICG/NEAMTWS for the Northeastern Atlantic and the Mediterranean Sea, and ICG/CARIBE-EWS for the Caribbean Sea.
4.17.7.2
Local Tsunami Warning Systems
For local tsunamis, time is more critical. Japan Meteorological Agency ( JMA) is responsible for issuing tsunami warnings in Japan. JMA uses seismic data from a few hundreds of stations to detect tsunamigenic earthquakes, and sea level data to confirm
the tsunami generation. There are six regional tsunami warning centers and each of them is staffed with multiple seismologists on duty for 24 h a day and 7 days a week. When an earthquake is detected on the seismic network, the location and size (magnitude) are automatically estimated within a few minutes. Likelihood of tsunami generation is judged from the location and magnitude, and the staff seismologists issue tsunami warning or advisory. The warning or advisory, as well as other related information such as the earthquake location, size, and the expected tsunami arrival times, is disseminated by using the Geostationary Meteorological Satellite. In the 1950s when JMA first introduced the tsunami warning system, it took about 20 min to issue tsunami warnings (Figure 16). In 1983, when a large earthquake occurred in Japan Sea, the warning was issued in 12 min after the earthquake, but the tsunami arrived in 7 min after the earthquake and 100 people died. JMA had improved the warning system, and in 1993 the tsunami warning was issued in 5 min after the South West Hokkaido earthquake. Even though, the tsunami arrived in less than 5 min and the 230 people died (Section 4.17.3.1.2). JMA further improved the system and now issues tsunami warning within 3–5 min after a large earthquake. It is also very important to improve the accuracy of tsunami forecast and to reduce false alarm. JMA introduced a numerical simulation technique in April 1999 (Tatehata, 1997); tsunami generation and propagation for 100 000 different cases were calculated in advance, and the results have been stored as a database. When a large earthquake occurs, the most appropriate case for an actual location and a magnitude of the earthquake is retrieved from the database.
504 Tsunamis
30
Elapsed time (min)
25 20 15 10 5 0 1952 1956 1960 1964 1968 1972 1976 1980 1984 1988 1992 1996 2000 2004 Year Figure 16 Improvements in elapsed time between earthquakes and tsunami warning issued by Japan Meteorological Agency.
A similar system based on database has been developed for the Pacific Ocean (Titov et al., 2005b). After a tsunami warning is issued, the information must be disseminated to coastal residents. In Japan, coastal residents receive warning through media, TV and radio, emergency broadcast systems, or patrol cars from police or fire stations. Periodic practice and drills are very important to keep the system functional. In Japan, annual national drills are conducted on 1 September, national disaster prevention day. In May 2006, The ICG/PTWS planned and conducted the first Pacific-wide exercise on issuing, disseminating, and responding to tsunami warning messages.
4.17.7.3 Long-Term and Probabilistic Forecast The tsunami hazard probabilities can be estimated from historical data. If we know that a particular coast has experienced N tsunamis over a period of time T, assuming a Poisson process, the probability P of having at least one tsunami for future time period is given as P ¼ 1 – exp ð – N =T Þ
½33
For example, if two tsunamis were recorded in the last 100 years at a particular location, the probability of having another tsunami in the next 10 years is 18%. The 10-year probability is 63% if 10 tsunamis occurred in the last 100 years. If the water heights from past tsunamis are known, the probabilities for each range of tsunami heights can be similarly computed. This method is useful for assessing tsunamis from various origins such as trans-Pacific tsunamis.
Great earthquakes at the plate boundaries may recur more regularly than Poisson process, although it is controversial. When the cyclic nature of past earthquakes is known, it can be used to calculate the probability of earthquake occurrence. The time interval between earthquakes can be modeled using a probability density function such as Weibull or log–normal distribution (Rikitake, 1976; Utsu, 1984). The Brownian passagetime (BPT) distribution has been recently used, because it models steady tectonic loading plus Brownian perturbation (Matthews et al., 2002). Then the probability of next earthquake in the specific time window can be computed if the average recurrence period and the date of the most recent event, as well as aperiodicity parameter or variation coefficient, are known. Probabilistic earthquake forecast had some success in Japan. In March 2003, the Japanese government made a long-term forecast for great (M 8) earthquakes along the Kuril trench. The committee estimated that the probability in the next 30 years (starting March 2003) is 60% in Tokachi-oki and 20–30% in Nemuro-oki. This forecast was based on the mean inter-event time of great earthquakes for the southern Kuril region in the nineteenth and twentieth centuries (Earthquake Research Committee, 2004; Satake et al., 2005b). Six months later, on 26 September 2003, a great Tokachi-oki earthquake (M 8.0) actually occurred. When the fault parameters of the characteristic tsunamigenic earthquake are known, tsunami heights can be estimated by numerical computations, and the probabilistic tsunami heights can be calculated. Such a probabilistic estimate of tsunami heights was made in the 1980s for the Japanese coasts (Rikitake and
Tsunamis
P (≥ 0.5 m) P (≥ 1 m)
P (≥ 2 m)
505
P (≥ 5 m) P (≥ 7 m) P (≥ 10 m)
15 7
7 27
Kushiro
7 15
15
7
7
9 9
Pt. Erimo
Hokkaido 22
Tomakomai
9 9
Muroran
100% Figure 17 Probabilities for the Pacific coast of Hokkaido hit by a tsunami with wave height exceeding 0.5, 1, 2, 5, 7, and 10 m during a period from 2000 to 2010. The upper half of the coast was hit by a tsunami with 2–4 m in 2003. From Rikitake T and Aida I (1988) Tsunami hazard probability in Japan. Bulletin of Seismological Society of America 78: 1268–1278.
Aida, 1988). For the Pacific coast of central Japan where a large earthquake (Tokai earthquake) is expected, the probability of having 5 m tsunami for the period between 2000 and 2010 is as high as 41%. For the Pacific coast of Hokkaido along the Kuril Trench, where the 2003 Tokachi-oki earthquake actually generated the tsunami, the probability of 2 m tsunami was estimated as 7–15% and that of 5 m tsunami was 7% (Figure 17). The actual tsunami heights in 2003 were up to 4 m (Tanioka et al., 2004), but the coastal run-up and the tsunami source were different from those of the previous Tokachi-oki earthquake in 1952 (Satake et al., 2006b). Similar probabilistic estimates have also been made for the US coast (Geist and Parsons, 2006). 4.17.7.4 Tsunami Hazard Maps and Public Awareness Once the coastal residents receive a tsunami warning message, they need to know what it means and where to evacuate. An effective tool for the public is a map showing area to evacuate for a future tsunami. Possible inundation zones are estimated either from historical data of past tsunamis or by tsunami numerical simulation based on maximum credible earthquakes. Safe places such as tsunami shelters to evacuate can also be shown in the tsunami hazard maps. Tsunami hazard map help coastal communities prepare for tsunami hazards.
On the basis of tsunami hazard maps, evacuation plans can be established by responsible officials such as civil defense officials (Gonzalez et al., 2005). One of the lessons from the 2004 Asian tsunami is that not only coastal residents but also foreign tourists need to be informed on tsunami hazards. For the Hawaiian Islands, tsunami hazard maps are prepared for most of the occupied coasts and published in the local phone book that are available at every hotel room (Figure 18). The maps show, in addition to the inundation area and places to evacuate, evacuation instructions. For example, those in high-rise buildings are instructed to move vertically to the third or higher floors, rather than moving out of the area. Tsunami hazard assessment has been carried out on the Japanese coasts. Numerical simulations were used to estimate the time and heights of tsunami and possible damage including casualties. For example, for the future Tonankai-Nankai earthquake off the western Japan, inundation of 40 000 houses and 2200–8600 casualties are estimated (Central Disaster Management Council, 2006). For public awareness, booklets and video on tsunamis have been prepared and used to educate coastal residents. In Japan, a famous story of a village chief who evacuated the entire residents, by burning his harvest at the time of the 1854 Nankai earthquake tsunami (known as ‘‘Inamura-no-Hi’’), has been popularly used for hazard education.
506 Tsunamis
Figure 18 Tsunami evacuation maps for Waikiki beach of Honolulu, Hawaii. The maps were prepared by University of Hawaii in cooperation with Hawaii State Civil Defense Agency and published in local phone books.
4.17.8 Summary Study of tsunami has interdisciplinary natures as summarized in the chapter. The tsunami sources include giant earthquake such as the 2004 Sumatra– Andaman earthquake or unusual ‘tsunami earthquakes’ for which more seismological studies are needed. Compared to seismic sources, generation mechanism of tsunamis from landslides or volcanic eruptions is less understood. While the tsunami propagation in deep ocean has been treated by shallowwater theory and numerical computations have been made popularly, the actual propagation in deep ocean was captured by the satellite for the first time at the 2004 Indian Ocean tsunami. Similar deep-sea observations in the future would provide chance to examine the propagation characteristics such as dispersion effects. The coastal behavior of tsunamis will need more studies on theory, field-data, and physical experiments. Geological studies of past tsunamis are also important, as tsunami is one of infrequent phenomena. The 2004 tsunami in the Indian Ocean demonstrated that the effect of tsunami can be global. The tsunami damage was extended through the Indian Ocean, and the tsunami was recorded throughout the world. In addition, the victims of the 2004 tsunami were not only limited to the coastal residents in Asian countries but also included tourists from Europe. For the reduction of future tsunami
damage, it is essential to educate the global citizens about tsunamis. After the 1960 Chilean tsunami, our scientific knowledge on tsunami has significantly improved and we can now detect the occurrence of earthquake and assess the possibility of tsunamis in real time. Such scientific and technological development can be and should be used to save human lives in the world.
References Abe K (1973) Tsunami and mechanism of great earthquakes. Physics of the Earth and Planetary Interiors 7: 143–153. Abe K (1979) Size of great earthquakes of 1873–1974 inferred from tsunami data. Journal of Geophysical Research 84: 1561–1568. Abe K (1981) Physical size of tsunamigenic earthquakes of the northwestern Pacific. Physics of the Earth and Planetary Interiors 27: 194–205. Abe K (1989) Quantification of tsunamigenic earthquakes by the Mt scale. Tectonophysics 166: 27–34. Abe K (2000) Predominance of long periods in large Pacific tsunamis. Science of Tsunami Hazards 18: 15–34. Aida I (1978) Reliability of a tsunami source model derived from fault parameters. Journal of Physics of the Earth 26: 57–73. Ammon CJ, Ji C, Thio HK, et al. (2005) Rupture process of the 2004 Sumatra–Andaman earthquake. Science 308: 1133–1139. Atwater BF, Musumi-Rokkaku S, Satake K, Tsuji Y, Ueda K, and Yamaguchi DK (2005) The Orphan Tsunami of 1700. USGS Professional Paper 1707: 133 pp. Baptista MA, Heitor S, Miranda JM, Miranda P, and Victor LM (1998) The 1755 Lisbon tsunami; Evaluation of the tsunami parameters. Journal of Geodynamics 25: 143–157.
Tsunamis Bardet JP, Synolakis CE, Davies HL, Imamura F, and Okal EA (2003) Landslide tsunamis: Recent findings and research directions. Pure and Applied Geophysics 160: 1793–1809. Bernard EN (2005) The US National Tsunami Hazard Mitigation Program: A successful State–Federal partnership. Natural Hazards 35: 5–24. Berninghausen WH (1966) Tsunamis and seismic seiches reported from regions adjacent to the Indian Oceans. Bulletin of the Seismological Society of America 56: 69–74. Bilek SL and Lay T (2002) Tsunami earthquakes possibly widespread manifestations of frictional conditional stability. Geophysical Research Letters 29: doi:10.1029/2002GL015215. Bilham R (2005) A flying start, then a slow slip. Science 308: 1126–1127. Bondevik S, Lovholt F, Harbitz C, Mangerud J, Dawson A, and Svendsen JI (2005) The Storegga Slide tsunami – comparing field observations with numerical simulations. Marine and Petroleum Geology 22: 195–208. Borrero JC (2005a) Field survey of Northern Sumatra and Banda Aceh, Indonesia after the tsunami and earthquake of 26 December 2004. Seismological Research Letters 76: 312–320. Borrero JC (2005b) Field data and satellite imagery of tsunami effects in Banda Aceh. Science 308: 1596. Bryant EA, Young RW, and Price DM (1996) Tsunami as a major control on coastal evolution, southeastern Australia. Journal of Coastal Research 12: 831–840. Bugge T, Belderson RH, and Kenyon NH (1988) The Storegga slide. Philosophical Transactions of the Royal Society of London Series A 325: 357–388. Central Disaster Management Council (2006) Hazard Estimation for Earthquakes Along Japan and Kuril Trenches (in Japanese). Tokyo: Cabinets Office. Comer RP (1980) Tsunami height and earthquake magnitude: theoretical basis of an empirical relation. Geophysical Research Letters 7: 445–448. Comer RP (1984a) The tsunami mode of a flat earth and its excitation by earthquake sources. Geophysical Journal of the Royal Astronomical Society 77: 1–27. Comer RP (1984b) Tsunami generation: a comparison of traditional and normal mode approaches. Geophysical Journal of the Royal Astronomical Society 77: 29–41. Cummins P (2004) Small threat, but warning sounded for tsunami research. AusGeo News 75: 4–7. Dawson AG, Long D, and Smith DE (1988) The Storegga Slides – Evidence from eastern Scotland for a possible tsunami. Marine Geology 82: 271–276. Dawson AG and Shi SZ (2000) Tsunami deposits. Pure and Applied Geophysics 157: 875–897. Earthquake Research Committee (2004) Long-term evaluation of seismicity along the Kuril Trench (in Japanese). Publications of Earthquake Research Committee, pp. 1–74 . Tokyo: Earthquake Research Committee. Fine, IV, Rabinovich AB, Bornhold BD, Thomson RE, and Kulikov EA (2005a) The Grand Banks landslide-generated tsunami of November 18, 1929: preliminary analysis and numerical modeling. Marine Geology 215: 45–57. Fine, IV, Rabinovich AB, and Thomson RE (2005b) The dual source region for the 2004 Sumatra tsunami. Geophysical Research Letters 32: L16602, doi: 10.1029/2005GL023521. Fujii Y and Satake K (2007) Tsunami source model of the 2004 Sumatra–Andaman earthquake inferred from tide gauge and satellite data. Bulletin of the Seismological Society of America 97: S192–S207. Fukao Y (1979) Tsunami earthquakes and subduction processes near deep-sea trenches. Journal of Geophysical Research 84: 2303–2314. Geist E (1998) Local tsunamis and earthquake source parameters. Advances in Geophysics 39: 117–209.
507
Geist E (2000) Origin of the 17 July 1998 Papua New Guinea tsunami; earthquake or landslide? Seismological Research Letters 71: 344–351. Geist E and Parsons T (2006) Probabilistic analysis of tsunami hazards. Natural Hazards 37: 277–314. Geist E, Titov VV, Arcas D, Pollitz FF, and Bilek SL (2007) Implications of the December 26, 2004 Sumatra–Andaman Earthquake on Tsunami Forecast and Assessment Models for Great Subduction Zone Earthquakes. Bulletin of the Seismological Society of America 97: S249–S270. Gelfenbaum G and Jaffe B (2003) Erosion and sedimentation from the 17 July, 1998 Papua New Guinea tsunami. Pure and Applied Geophysics 160: 1969–1999. Gonzalez FI, Bernard EN, Meinig C, Eble MC, Mofjeld HO, and Stalin S (2005) The NTHMP tsunameter network. Natural Hazards 35: 25–39. Gower J (2005) Jason 1 detects the 26 December 2004 tsunami. Eos, Transactions, American Geophysical Union 86: 37–38. Gusiakov VK (2005) Tsunami generation potential of different tsunamigenic regions in the Pacific. Marine Geology 215: 3–9. Hanks T and Kanamori H (1979) A moment magnitude scale. Journal of Geophysical Research 84: 2348–2350. Hanson JF, Reasoner C, and Bowman JR (2007) High frequency tsunami signals of the great Indonesian earthquakes of 26 December and 28 March 2005. Bulletin of the Seismological Society of America 97: S232–S248. Harbitz CB (1992) Model simulations of tsunamis generated by the Storegga slides. Marine Geology 105: 1–21. Hatori T (1979) Relation between tsunami magnitude and wave energy. Bulletin of the Earthquake Research Institute (University of Tokyo) 54: 531–541. Hatori T (1984) On the damage to houses due to tsunamis. Bulletin of the Earthquake Research Institute (University of Tokyo) 59: 433–439. Heezen BC and Ewing M (1952) Turbidity currents and submarine slumps, and the 1929 Grand Banks earthquake. American Journal of Science 250: 849–873. Heinrich P, Piatanesi A, and Hebert H (2001) Numerical modelling of tsunami generation and propagation from submarine slumps: the 1998 Papua New Guinea event. Geophysical Journal International 145: 97–111. Hirata K, Aoyagi M, Mikada H, et al. (2002) Real-time geophysical measurements on the deep seafloor using submarine cable in the southern Kurile subduction zone. IEEE Journal of Oceanic Engineering 27: 170–181. Hirata K, Satake K, Tanioka Y, et al. (2006) The 2004 Indian Ocean tsunami: Tsunami source model from satellite altimetry. Earth, Planets and Space 58: 195–201. Iida K (1984) Catalog of Tsunamis in Japan and Its Neighboring Countries, 52pp. Toyato: Aichi Institute of Technology. Iida K, Cox DC, and Pararas-Carayannis G (1967) Preliminary catalog of tsunamis occurring in the Pacific Ocean. Honolulu: University of Hawaii, 274pp. Imamura F, Hashi K, and Imteaz MMA (2001) Modeling for tsunamis generated by landsliding and debris flow. In: Hebenstreit GT (ed.) Tsunami Research at the End of a Critical Decade,, pp. 209–228. Dordrecht: Kluwer Academic. Institute of Computational Mathematics and Mathematical Geophysics (2005) Integrated Tsunami Database for the World Ocean, Novosibirsk, CD-ROM. Intergovernmental Oceanographic Commission (1998) Manuals and Guides 37: Post-tsunami Survey Field Guide. Paris: UNESCO. International Federation of Red Cross and Red Crescent Societies (2005) World Disasters Report, 251pp. Geneva: International Federation of Red Cross and Red Crescent Societies. Ishii M, Shearer PM, Houston H, and Vidale JE (2005) Extent, duration and speed of the 2004 Sumatra–Andaman
508 Tsunamis earthquake imaged by the Hi-Net array. Nature 435: 933–936. Iwasaki SI, Eguchi T, Fujinawa Y, et al. (1997) Precise tsunami observation system in the deep ocean by an ocean bottom cable network for the prediction of earthquakes and tsunamis. In: Hebenstreit G (ed.) Perspectives on Tsunami Hazard Reduction, pp. 47–66. Dordrecht: Kluwer Academic. Jaffe BE, Borrero JC, Prasetya GS, et al. (2006) Northwest Sumatra and offshore islands field survey after the December 2004 Indian Ocean tsunami. Earthquake Spectra 22: S105–S135. Jiang L and Leblond PH (1994) 3-Dimensional modeling of tsunami generation due to a submarine mudslide. Journal of Physical Oceanography 24: 559–572. Johnson JM (1999) Heterogeneous coupling along AlaskaAleutians as inferred from tsunami, seismic and geodetic inversions. Advances in Geophysics 39: 1–116. Johnson JM and Satake K (1997) Estimation of seismic moment and slip distribution of the April 1, 1946, Aleutian tsunami earthquake. Journal of Geophysial Research 102: 11765–11774. Kajiura K (1963) The leading wave of a tsunami. Bulletin of the Earthquake Research Institute (University of Tokyo) 41: 535–571. Kajiura K (1970) Tsunami source, energy and the directivity of wave radiation. Bulletin of the Earthquake Research Institute (University of Tokyo) 48: 835–869. Kanamori H (1972) Mechanism of tsunami earthquakes. Physics of the Earth and Planetary Interiors 6: 246–259. Kanamori H and Kikuchi M (1993) The 1992 Nicaragua earthquake: A slow tsunami earthquake associated with subducted sediments. Nature 361: 714–716. Kato T, Terada Y, Ito K, et al. (2005) Tsunami due to the 2004 September 5th off Kii peninsula earthquake, Japan, recorded by a new GPS buoy. Earth, Planets and Space 57: 297–301. Kato T, Terada Y, Kinoshita M, et al. (2000) Real-time observation of tsunami by RTK-GPS. Earth, Planets and Space 52: 841–845. Kawata Y, Benson B, Borrero JC, et al. (1999) Tsunami in Papua New Guinea was as intense as first thought. Eos, Transactions, American Geophysical Union 80: 101–105. Kayanne H, Ikeda Y, Echigo T, et al. (2007) Coseismic and postseismic creep in Andaman Islands associated with the 2004 Sumatra–Andaman earthquake. Geophysical Research Letters 34: L01310, doi: 10.129/2006GL028200. Keating BH, Waythomas CF, and Dawson AG (2000) Special issue: Landslides and Tsunamis–Introduction. Pure and Applied Geophysics 157: 871–873. Kulikov EA, Rabinovich AB, Thomson RE, and Bornhold BD (1996) The landslide tsunami of November 3, 1994, Skagway Harbor, Alaska. Journal of Geophysical Research 101: 6609–6615. Lander JF (1996) Tsunamis Affecting Alaska: 1737–1996. NGDC Key to Geophysical Research Documentation No. 29, Boulder, CO. Lander JF and Lockridge PA (1989) United States Tsunamis. Boulder, CO: National Geophysical Data Center. Lander JF, Lockridge PA, and Kozuch MJ (1993) Tsunamis Affecting the West Coast of the United States 1806–1992. Boulder, CO: National Geophysical Data Center. Latter JH (1981) Tsunamis of volcanic origin: summary of causes, with particular reference to Krakatoa, 1883. Bulletin of Volcanology 44: 467–490. Lay T, Kanamori H, Ammon CJ, et al. (2005) The great SumatraAndaman earthquake of 26 December 2004. Science 308: 1127–1133. Liu PLF, Lynett P, Fernando H, et al. (2005) Observations by the International Tsunami Survey Team in Sri Lanka. Science 308: 1595–1595.
Malik JN and Murty CVR (2005) Landscape changes in Andaman and Nicobar Islands (India) due to Mw 9.3 tsunamigenic Sumatra earthquake of 26 December 2004. Current Science 88: 1384–1386. Mansinha L and Smylie DE (1971) The displacement fields of inclined faults. Bulletin of the Seismological Society of America 61: 1433–1440. Matsuyama M, Walsh JP, and Yeh H (1999) The effect of bathymetry on tsunami characteristics at Sisano Lagoon, Papua New Guinea. Geophysical Research Letters 26: 3513–3516. Matthews MV, Ellsworth WL, and Reasenberg PA (2002) A Brownian model for recurrent earthquakes. Bulletin of the Seismological Society of America 92: 2233–2250. McCreery CS (2005) Impact of the National Tsunami Hazard Mitigation Program on operations of the Richard H. Hagemeyer Pacific Tsunami Warning Center. Natural Hazards 35: 73–88. Mei CC (1989) The Applied Dynamics of Ocean Surface Waves. Advances Series on Ocean Engineering, volume 1. Singapore: World Scientific. Meltzner AJ, Sieh K, Abrams M, et al. (2006) Uplift and subsidence associated with the great Aceh–Andaman earthquake of 2004. Journal of Geophysical Research 111: B02407 (doi: 10.1029/2005JB003891). Merrifield MA, Firing YL, Aarup T, et al. (2005) Tide gauge observations of the Indian Ocean tsunami, December 26, 2004. Geophysical Research Letters 32, doi: 10.1029/ 2005GL022610. Miller DJ (1960) Giant waves in Lituya Bay, Alaska. USGS Professional Paper 354-C: 51–86. Minoura K and Nakaya S (1991) Traces of tsunami preserved in inter-tidal lacustrine and marsh deposits - Some examples from northeast Japan. The Journal of Geology 99: 265–287. Miyabe N (1934) An investigation of the Sanriku tunami based on mareogram data. Bulletin of the Earthquake Research Institute (University of Tokyo) 1: 112–126. Moore A, Nishimura Y, Gelfenbaum G, Kamataki T, and Triyono R (2006) Sedimentary deposits of the 26 December 2004 tsunami on the northwest. Earth, Planets and Space 58: 253–258. Murty TS and Rafiq M (1991) A tentative list of tsnamis in the marginal seas of the North Indian Ocean. Natural Hazards 4: 81–83. Nagai T, Ogawa H, Nukada K, and Kudaka M (2004) Characteristics of the observed 2003 Tokachi-oki earthquake tsunami profile. Coastal Engineering Journal 46: 315–327. Nagarajan B, Suresh I, Sundar D, et al. (2006) The Great Tsunami of 26 December 2004: A description based on tidegauge data from the Indian subcontinent and surrounding areas. Earth, Planets and Space 58: 211–215. Nanayama F, Satake K, Furukawa R, et al. (2003) Unusually large earthquakes inferred from tsunami deposits along the Kuril trench. Nature 424: 660–663. Nanayama F, Shigeno K, Satake K, et al. (2000) Sedimentary differences between the 1993 Hokkaido-nansei-oki tsunami and the 1959 Miyakojima typhoon at Taisei, southwestern Hokkaido, northern Japan. Sedimentary Geology 135: 255–264. Neetu S, Suresh I, Shankar R, et al. (2005) Comment on ‘‘The Great Sumatra-Andaman Earthquake of 26 December 2004’’. Science 310: 1431. Newman AV and Okal EA (1998) Teleseismic estimates of radiated seismic energy: The E/M-0 discriminant for tsunami earthquakes. Journal of Geophysical Research 103: 26885–26898. Nishimura Y and Miyaji N (1995) Tsunami deposits from the 1993 Southwest Hokkaido earthquake and the 1640
Tsunamis Hokkaido-Komagatake eruption, northern Japan. Pure and Applied Geophysics 144: 719–733. Nomanbhoy N and Satake K (1995) Generation mechanism of tsunamis from the 1883 Krakatau eruption. Geophysical Research Letters 22: 509–512. Okada M (1995) Tsunami observation by ocean bottom pressure gauge. In: Tsuchiya Y, Shuto N (eds.) Tsunami: Progress in Prediction, Disaster Prevention and Warning, pp. 287–303. Dordrecht: Kluwer Academic. Okada Y (1985) Surface deformation due to shear and tensile faults in a half-space. Bulletin of the Seismological Society of America 75: 1135–1154. Okal EA (1988) Seismic parameters controlling far-field tsunami amplitudes: A review. Natural Hazards 1: 67–96. Okal EA and Newman AV (2001) Tsunami earthquakes: The quest for a regional signal. Physics of the Earth and Planetary Interiors 124: 45–70. Okal EA, Plafker G, Synolakis CE, and Borrero JC (2003) Nearfield survey of the 1946 Aleutian Tsunami on Unimak and Sanak Islands. Bulletin of the Seismological Society of America 93: 1226–1234. Okal EA and Synolakis CE (2004) Source discriminants for nearfield tsunamis. Geophysical Journal International 158: 899–912. Okal EA, Synolakis CE, Fryer GJ, et al. (2002) A field survey of the 1946 Aleutian tsunami in the far field. Seismological Research Letters 73: 490–503. Okal EA and Talandier J (1989) Mm: A variable-period mantle magnitude. Journal of Geophysical Research 94: 4169–4193. Ortiz M and Bilham R (2003) Source area and rupture parameters of the 31 December 1881 M-w ¼ 7.9 Car Nicobar earthquake estimated from tsunamis recorded in the Bay of Bengal. Journal of Geophysical Research 108: (doi: 10.129/ 2002JB001941). Pelayo AM and Wiens DA (1992) Tsunami earthquakes: slow thrust-faulting events in the accretionary wedge. Journal of Geophysical Research 97: 15321–15337. Pelinovsky E, Choi BH, Stromkov A, Didenkulova I, and Kim H-S (2005) Analysis of tide-gauge records of the 1883 Krakatau tsunami. In: Satake K (ed.) Tsunamis: Case Studies and Recent Developments,, pp. 57–78. Dordrecht: Springer. Peregrine H (1972) Equations for water waves and the approximations behind them. In: Meyer RE (ed.) Waves on Beaches and Resulting Sediment Transport, pp. 95–121. Academic Press. Piatanesi A and Lorito S (2007) Rupture process of the 2004 Sumatra–Andaman earthquake from tsunami waveform inversion. Bulletin of the Seismological Society of America 97: S223–S231. Plafker G, Kachadoorian R, Eckel EB, and Mayo LR (1969) Effects of the earthquake of March 27, 1964 on various communities. USGS Professional Paper 542-G: 1–50. Polet J and Kanamori H (2000) Shallow subduction zone earthquakes and their tsunamigenic potential. Geophysical Journal International 142: 684–702. Press F and Harkrider D (1966) Air-sea waves from the explosion of Krakatoa. Science 154: 1325–1327. Reymond D, Hyvernaud O, and Talandier J (1991) Automatic detection, location and quantification of earthquakes: application to tsunami warning. Pure and Applied Geophysics 135: 361–382. Rikitake T (1976) Recurrence of great earthquakes at subduction zones. Tectonophysics 35: 335–362. Rikitake T and Aida I (1988) Tsunami hazard probability in Japan. Bulletin of the Seismological Society of America 78: 1268–1278. Satake K (1988) Effects of Bathymetry on Tsunami Propagation - Application of Ray Tracing to Tsunamis. Pure and Applied Geophysics 126: 27–36.
509
Satake K (1989) Inversion of tsunami waveforms for the estimation of heterogeneous fault motion of large submarine earthquakes – the 1968 Tokachi-Oki and 1983 Japan Sea earthquakes. Journal of Geophysical Research 94: 5627–5636. Satake K (1993) Depth distribution of coseismic slip along the Nankai Trough, Japan, from joint inversion of geodetic and tsunami data. Journal of Geophysical Research 98: 4553–4565. Satake K (1994) Mechanism of the 1992 Nicaragua tsunami earthquake. Geophysical Research Letters 21: 2519–2522. Satake K (1995) Linear and nonlinear computations of the 1992 Nicaragua earthquake tsunami. Pure and Applied Geophysics 144: 455–470. Satake K, Aung TT, Sawai Y, et al. (2006a) Tsunami heights and damage along the Myanmar coast from the December 2004 Sumatra-Andaman earthquake. Earth, Planets and Space 58: 243–252. Satake K, Baba T, Hirata K, et al. (2005a) Tsunami source of the 2004 off the Kii peninsula earthquakes inferred from offshore tsunami and coastal tide gauges. Earth, Planets and Space 57: 173–178. Satake K, Hirata K, Yamaki S, and Tanioka Y (2006b) Re-estimation of tsunami source of the 1952 Tokachi-oki earthquake. Earth, Planets and Space 58: 535–542. Satake K, Nanayama F, Yamaki S, Tanioka Y, and Hirata K (2005b) Variability among tsunami sources in the 17th–21st centuries along the southern Kuril trench. In: Satake K (ed.) Tsunamis: Case Studies and Recent Developments, pp. 157–170. Dordrecht: Springer. Satake K, Okada M, and Abe K (1988) Tide gauge response to tsunamis: Measurements at 40 tide gauge stations in Japan. Journal of Marine Research 46: 557–571. Satake K, Shimazaki K, Tsuji Y, and Ueda K (1996) Time and site of a giant earthquake in Cascadia inferred from Japanese tsunami records of January 1700. Nature 379: 246–249. Satake K, Smith JR, and Shinozaki K (2002) Three-dimensional reconstruction and tsunami model of the Nuuanu and Wailau landslides, Hawaii. In: Takahashi E, Lipman P, Garcia M, Naka J, and Aramaki S (eds.) Hawaiian Volcanoes: Deep Underwater Perspectives. 128, pp. 333–346. Washington D.C: American Geophysical Union. Satake K and Tanioka Y (1999) Sources of tsunami and tsunamigenic earthquakes in subduction zones. Pure and Applied Geophysics 154: 467–483. Satake K and Tanioka Y (2003) The July 1998 Papua New Guinea earthquake: Mechanism and quantification of unusual tsunami generation. Pure and Applied Geophysics 160: 2087–2118. Satake K, Wang KL, and Atwater BF (2003) Fault slip and seismic moment of the 1700 Cascadia earthquake inferred from Japanese tsunami descriptions. Journal of Geophysical Research 108: (doi: 10.1029/2003JB002521). Seno T (2002) Tsunami earthquake as transient phenomena. Geophysical Research Letters 29: (doi: 10.029/ 2002GL014868). Seno T and Hirata K (2007) Did the 2004 Sumatra–Andaman earthquake involve a component of tsunami earthquakes? Bulletin of the Seismological Society of America 97: S296–S306. Shuto N (1991) Numerical simulation of tsunamis – Its present and near future. Natural Hazards 4: 171–191. Shuto N and Matsutomi H (1995) Field survey of the (1993) Hokkaido-Nansei-Oki earthquake tsunami. Pure and Applied Geophysics 144: 649–663. Simkin T and Fiske R (1983) Krakatau, 1883 – the volcanic eruption and its effects. Washington, D.C: Smithsonian Institution Press. Singh SK, Ortiz M, Gupta HK, and Ramadass DGA (2006) Slow slip below Port Blair, Andaman, during the great Sumatra–
510 Tsunamis Andaman earthquake of 26 December 2004. Geophysical Research Letters 33: (doi: 10.1029/2005GL025025). Slingerland R and Voight B (1979) Occurrence, properties, and predictive models of landslide-generated water waves. In: Voight B (ed.) Rockslides and Avalanches, 1, pp. 317–397. Amsterdam: Elsevier. Smith DE, Shi S, Cullingford RA, et al. (2004) The Holocene Storegga slide tsunami in the United Kingdom. Quaternary Science Reviews 23: 2291–2321. Smith WHF, Scharroo R, Titov VV, Arcas D, and Arbic BK (2005) Satellite altimeters measure tsunami, early model estimates confirmed. Oceanography 18: 10–12. Soloviev SL (1970) Recurrence of tsunamis in the Pacific. In: Adams W M (ed.) Tsunamis in the Pacific Ocean, pp. 149–164. Honolulu: East-West Center Press. Soloviev SL and Go CN (1974a) Catalogue of Tsunamis on the Western Shore of the Pacific Ocean (173–1968). Moscow: Nauka Publishing House. Soloviev SL and Go CN (1974b) Catalogue of Tsunamis on the Eastern shore of the Pacific Ocean (1513–1968). Moscow: Nauka Publishing House. Soloviev SL, Go CN, and Kim KS (1992) Catalog of Tsunamis in the Pacific 1969–1982. Moscow: Academy of Science of USSR. Song YT, Ji C, Fu L-L, et al. (2005) The 26 December 2004 tsunami source estimated from satellite radar altimetry and seismic waves. Geophysical Research Letters 32, doi: 10.1029/2005GL023683. Stein S and Okal EA (2005) Speed and size of the Sumatra earthquake. Nature 434: 581–582. Steketee JA (1958) On Volterra’s dislocations in a semi-infinite elastic medium. Canadian Journal of Physics 36: 192–205. Synolakis CE, Bardet JP, Borrero JC, et al. (2002) The slump origin of the 1998 Papua New Guinea tsunami. Proceedings of the Royal society of London Series A 458: 763–789. Synolakis CE and Okal EA (2005) 1992–2002: Perspective on a decade of post-tsunami surveys. In: Satake K (ed.) Tsunamis: Case Studies and Recent Developments, pp. 1–29. Dordrecht: Springer. Takahashi T, Shuto N, Imamura F, and Ortiz M (1995) Source models for the 1993 Hokkaido-Nansei-Oki earthquake tsunami. Pure and Applied Geophysics 144: 747–767. Tanioka Y, Nishimura Y, Hirakawa K, et al. (2004) Tsunami runup heights of the 2003 Tokachi-oki earthquake. Earth, Planets and Space 56: 359–365. Tanioka Y, Ruff L, and Satake K (1997) What controls the lateral variation of large earthquake occurrence along the Japan trench? Island Arc 6: 261–266. Tanioka Y and Satake K (1996a) Tsunami generation by horizontal displacement of ocean bottom. Geophysical Research Letters 23: 861–864. Tanioka Y and Satake K (1996b) Fault parameters of the 1896 Sanriku tsunami earthquake estimated from tsunami numerical modeling. Geophysical Research Letters 23: 1549–1552. Tanioka Y and Sataka K (2001a) Coseismic slip distribution of the 1946 Nankai earthquake and aseismic slips caused by the earthquake. Earth, Planets and Space 53: 235–241. Tanioka Y and Satake K (2001b) Detailed coseismic slip distribution of the 1944 Tonankai earthquake estimated from tsunami waveforms. Geophysical Research Letters 28: 1075–1078. Tanioka Y, Satake K, and Ruff L (1995) Total analysis of the 1993 Hokkaido-Nansei-Oki earthquake using seismic-wave, tsunami, and geodetic data. Geophysical Research Letters 22: 9–12. Tanioka Y and Seno T (2001a) Sediment effect on tsunami generation of the 1896 Sanriku tsunami earthquake. Geophysical Research Letters 28: 3389–3392.
Tanioka Y and Seno T (2001b) Detailed analysis of tsunami waveforms generated by the 1946 Aleutian tsunami earthquake. Natural Hazards and Earth System Sciences 1: 171–175. Tanioka Y, Yudhicara, Kusunose T, et al. (2006) Rupture process of the 2004 great Sumatra-Andaman earthquake estimated from tsunami waveforms. Earth, Planets and Space 58: 203–209. Tappin DR (2004) Special issue: Submarine-slump-generated tsunamis – Preface. Marine Geology 203: 199–200. Tappin DR, Matsumoto P, Watts P, et al. (1999) Sediment slump likely caused 1998 Papua New Guinea tsunami. Eos, Transactions, American Geophysical Union 80: 329, 334, 340. Tatehata H (1997) The new tsunami warning system of the Japan Meteorological Agency. In: Hebenstreit G (ed.) Perspectives on Tsunami Hazards Reduction, pp. 175–188. Dordrecht: Kluwer. Tilling RI (1989) Volcanic Hazards and Their Mitigation - Progress and Problems. Reviews of Geophysics 27: 237–269. Tinti S, Bortolucci E, and Vannini C (1997) A block-based theoretical model suited to gravitational sliding. Natural Hazards 16: 1–28. Titov V, Rabinovich AB, Mofjeld HO, Thomson RE, and Gonzalez FI (2005a) The global reach of the 26 December 2004 Sumatra tsunami. Science 309: 2045–2048. Titov VV and Gonzalez FI (2001) Numerical study of the source of the July 17, 1998 PNG tsunami. In: Hebensreit GT (ed.) Tsunami Research at the End of a Critical Decade, pp. 197–207. Dordrecht: Kluwer Academic. Titov VV, Gonzalez FI, Bernard EN, et al. (2005b) Real-time tsunami forecasting: Challenges and solutions. Natural Hazards 35: 41–58. Titov VV and Synolakis CE (1997) Extreme inundation flows during the Hokkaido-Nansei-Oki tsunami. Geophysical Research Letters 24: 1315–1318. Tobita M, Suito H, Imakiire T, Kato M, Fujiwara S, and Murakami M (2006) Outline of vertical displacement of the 2004 and 2005 Sumatra earthquakes revealed by satellite radar imagery. Earth, Planets and Space 58: e1–e4. Tsai VC, Nettles M, Ekstrom G, and Dziewonski AM (2005) Multiple CMT source analysis of the 2004 Sumatra earthqauke. Geophysical Research Letters 32: (doi: 10.1029/ 2005GL023813). Tsuboi S, Abe K, Takano K, Yamanaka Y (1995) RapidDetermination of Mw from Broad-Band P-Wave-Forms. Bulletin of the Seismological Society of America 85: 606–613. Tsuji Y, Namegaya Y, Matsumoto H, et al. (2006) The 2004 Indian tsunami in Thailand: Surveyed runup heights and tide gauge records. Earth, Planets and Space 58: 223–232. Tuttle MP, Ruffman A, Anderson T, and Jeter H (2004) Distinguishing tsunami from storm deposits in eastern North America: The 1929 grand banks tsunami versus the 1991 Halloween storm. Seismological Research Letters 75: 117–131. Unoki S and Nakano M (1953) On the Cauchy-Poisson waves caused by the eruption of a submarine volcano (III). Oceanographical Magazine 4: 139–150. Utsu T (1984) Estimation of parameters for recurrence models of earthquakes. Bulletin of the Earthquake Research Institute (University of Tokyo) 59: 53–66. Ward S (1980) Relationships of tsunami generation and an earthquake source. Journal of Physics of the Earth 28: 441–474. Ward SN (2001) Landslide tsunami. Journal of Geophysical Research 106: 11201–11215. Ward SN and Day S (2003) Ritter Island Volcano – lateral collapse and the tsunami of 1888. Geophysical Journal International 154: 891–902.
Tsunamis Watanabe H (1998) Comprehensive List of Tsunamis to Hit the Japanese Islands. Tokyo: University of Tokyo Press. Watts P, Grilli ST, Kirby JT, Fryer GJ, and Tappin DR (2003) Landslide tsunami case studies using a Boussinesq model and a fully nonlinear tsunami generation model. Natural Hazards Earth System Sciences 3: 391–402. Woodworth PL and Player R (2003) The permanent service for mean sea level: An update to the 21st century. Journal of Coastal Research 19: 287–295. Yamashita T and Sato R (1974) Generation of tsunami by a fault model. Journal of Physics of the Earth 22: 415–440. Yeh H, Imamura F, Synolakis C, Tsuji Y, Liu P, and Shi S (1993) The Flores Island tsunamis. Eos, Transactions, American Geophysical Union 74: 369, 371–373.
511
Yuan XH, Kind R, and Pedersen HA (2005) Seismic monitoring of the Indian Ocean tsunami. Geophysical Research Letters 32: L15308, doi: 10.1029/2005GL023464.
Relevant Websites http://tsun.sscc.ru – Institute of Computational Mathematics and Mathematical Geophysics, Novosibirsk. http://www.ngdc.noaa.gov – National Geophysical Data Center, NOAA.
4.17.1 4.17.2 4.17.3 4.17.3.1 4.17.3.1.1 4.17.3.1.2 4.17.3.1.3 4.17.3.1.4 4.17.3.1.5 4.17.3.1.6 4.17.3.1.7 4.17.3.2 4.17.3.2.1 4.17.3.2.2 4.17.3.2.3 4.17.4 4.17.4.1 4.17.4.1.1 4.17.4.1.2 4.17.4.1.3 4.17.4.1.4 4.17.4.1.5 4.17.4.2 4.17.4.2.1 4.17.4.2.2 4.17.4.3 4.17.4.4 4.17.4.5 4.17.4.5.1 4.17.4.5.2 4.17.4.5.3 4.17.4.6 4.17.5 4.17.5.1 4.17.5.1.1 4.17.5.1.2 4.17.5.1.3 4.17.5.1.4 4.17.5.2 4.17.5.2.1 4.17.5.2.2 4.17.5.2.3 4.17.5.3 4.17.5.3.1 4.17.5.3.2 4.17.5.3.3
Introduction The 2004 Indian Ocean Tsunami Other Examples Historical Tsunamis 1998 Papua New Guinea tsunami 1993 Hokkaido tsunami 1960 Chilean tsunami 1958 Lituya Bay tsunami 1946 Aleutian tsunami 1896 Sanriku tsunami 1883 Krakatoa tsunami Prehistoric Tsunamis 1700 Cascadia tsunami Hokkaido tsunami in seventeenth century Tsunami from the Storegga slide Tsunami Observations Instrumental Measurements Coastal tsunami measurements Offshore tsunami measurements Deep ocean measurements Satellite observations Definition of amplitude and wave heights Field Survey to Measure Run-Up Heights Post-tsunami surveys Definitions of inundation and run-up heights Study of Historical Documents Geological Methods Tsunami Magnitude Scales Imamura-Iida scale, m Tsunami intensity Tsunami magnitude Mt Tsunami Catalogs Tsunami Generation Earthquake Tsunamis Ocean bottom and water surface displacements Fault parameters and ocean-bottom deformation Estimation of earthquake source parameters from tsunamis Tsunamigenic and tsunami earthquakes Tsunamis from Landslides Examples Characteristics and discrimination of landslide tsunamis Modeling landslide tsunamis Volcanic Tsunamis Volcanic tsunamis due to earthquakes Volcanic tsunamis due to landslides Volcanic tsunamis from flow of volcanic ejecta
484 484 486 486 486 486 486 487 487 487 488 488 488 489 489 489 489 489 489 489 490 490 490 490 491 491 492 492 492 492 492 493 493 493 493 493 494 494 495 496 496 496 497 497 497 498
483
484 Tsunamis 4.17.5.3.4 4.17.5.3.5 4.17.6 4.17.6.1 4.17.6.2 4.17.6.3 4.17.6.4 4.17.6.5 4.17.6.6 4.17.6.7 4.17.7 4.17.7.1 4.17.7.2 4.17.7.3 4.17.7.4 4.17.8 References
Volcanic tsunamis from submarine explosion Volcanic tsunamis from caldera collapse Tsunami Propagation Weak Coupling between Elastic Earth and Ocean Ray-Theoretical Approach Coastal Amplification and Green’s Law Tsunami as a Gravity Wave Basic Equations for Tsunami Propagation Dispersion Numerical Computations Tsunami Warning and Hazard Reduction Systems Far-Field Tsunami Warning Systems Local Tsunami Warning Systems Long-Term and Probabilistic Forecast Tsunami Hazard Maps and Public Awareness Summary
4.17.1 Introduction A tsunami is an oceanic gravity wave generated by submarine or coastal geological processes such as earthquakes, landslides, or volcanic eruptions. Most tsunamis are caused by large shallow earthquakes along subduction zones. Tsunami is a Japanese word meaning ‘harbor wave’, and is sometimes called seismic sea wave or, erroneously, tidal wave. Tsunamis are usually small and barely noticed in deep ocean, but become large and cause damage when they approach coasts or harbors. The recent 2004 tsunami in the Indian Ocean demonstrated that a tsunami can travel at trans-oceanic distances and cause damage far from the source. Tsunamis have many aspects that are studied by researchers in various fields. Their generation is related to geological processes. The propagation and offshore observation are the subjects of oceanography, and basic hydrodynamics are needed to understand the characteristics. The coastal behaviors, such as run-up on beaches or resonance in bays, are mostly studied by coastal engineers. For tsunami hazard reduction, officials in local and national governments, civil defense, and emergency managements are responsible and play important roles. In this chapter, the wide range of tsunami phenomena is reviewed. Section 4.17.2 describes the tsunami caused by the 2004 Sumatra–Andaman earthquake, the most recent and the worst tsunami disaster that affected the countries around the Indian Ocean. Other examples of recent, historical and
498 498 498 498 498 499 499 499 500 500 501 502 503 504 505 506 506
prehistoric tsunamis are reviewed in Section 4.17.3. Section 4.17.4 reviews the tsunami observation systems: instrumental measurements, field surveys, and studies of historical and geological data. The tsunami magnitude scales and catalogs are also described. Section 4.17.5 reviews the tsunami generation, by earthquakes, landslides, and volcanic eruptions, with some emphasis on the studies of tsunami sources. Section 4.17.6 discusses the tsunami propagation, from its hydrodynamic background to numerical computations. Section 4.17.7 discusses the tsunami warning systems, both basin-wide and local, and hazard mitigation methods such as probabilistic hazard assessments and hazard maps.
4.17.2 The 2004 Indian Ocean Tsunami A giant earthquake (Mw ¼ 9.1–9.3) occurred off Sumatra Island on 26 December 2004 (Lay et al., 2005; Stein and Okal, 2005). This earthquake generated tsunami which devastated the shores of Indian Ocean (Figure 1); more than 200 000 people lost their lives. The number of victims, dead and missing together, is the largest in Indonesia (163 795), followed by Sri Lanka (35 399), India (16 389), and Thailand (8 345) (International Federation of Red Cross and Red Crescent Societies, 2005). Within 30 min of the earthquake, the tsunami devastated Banda Aceh and other coastal villages of Sumatra Island (Borrero, 2005a). The tsunami arrived at
Tsunamis 30° N
485
11
10
10
6
9
5
11
1
0°
2
4
6
3
7 3 5
4
4
6
8 6
7
10
11
7
6
9
12
5
5
8
30° S
7 8
30° E
60° E
7
90° E
120° E
150° E
Figure 1 Tsunami propagation in the Indian Ocean. Each curve indicates wavefront at 1 h increments. Black star and circles indicate the epicenters of mainshock and aftershocks within 1 day, respectively (data by USGS).
Thai, Sri Lankan, and Indian coasts in about 2 h after the earthquake. The tsunami propagated farther and arrived at the eastern coast of Africa several hours after the earthquake to cause 298 fatalities in Somalia. The tsunami was also observed in the Atlantic and Pacific Oceans (Titov et al., 2005a). To document the 2004 tsunami, many scientists from all over the world visited the affected coasts (e.g., Borrero, 2005b; Liu et al., 2005; Satake et al., 2006a; Tsuji et al., 2006). The measured tsunami heights in Sumatra Island, particularly in the Aceh province, were mostly larger than 20 m with the maximum above 30 m. The tsunami heights along the Andaman Sea coast were highly variable; 5–15 m in Thailand (Figure 2) but less than 3 m in Myanmar. The tsunami heights were up to 5 m on India’s Andaman Islands. In Sri Lanka, the tsunami heights were 5–15 m. The tsunami height distribution is consistent with the damage distribution and indicates that the source of the large and damaging tsunami was concentrated in the southern 700 km section of the aftershock zone. Sedimentological characteristics of the modern tsunami deposits were examined ( Jaffe et al., 2006; Moore et al., 2006), with hopes that they provide valuable information for the studies of paleo-tsunami deposits. The tsunami was instrumentally recorded on coastal tide gauges in Indian Ocean (Merrifield et al., 2005; Nagarajan et al., 2006). This tsunami was also recorded by other instruments, some were for the first time. The tsunami propagation in deep water was
Figure 2 Tsunami damage at Ban Nam Ken, about 100 km north of Phuket, Thailand.
captured by satellite altimeter (Gower, 2005). The tsunami signal was also recorded on horizontal components of broadband seismic stations near coasts (Yuan et al., 2005) or hydrophones (Hanson et al., 2007). Seismological analyses of the Sumatra–Andaman earthquake indicate that the source region extended from off Sumatra through the Nicobar to Andaman Islands with the total length of 1200–1300 km, and the total duration of rupture was about 500 s (Ishii et al., 2005; Tsai et al., 2005). The fault slip was the largest near off the northern Sumatra, followed by off Nicobar Islands (Ammon et al., 2005). Fault slip around Andaman Islands was estimated to be small from seismological analysis and speculated to be slow (Bilham, 2005; Lay et al., 2005).
486
Tsunamis
Analyses of satellite images indicate that the sea level changes were observed in Sumatra, Nicobar, and Andaman Islands, suggesting that the source length was as long as 1600 km (Meltzner et al., 2006; Tobita et al., 2006). Field investigations indicate that the largest subsidence was 3 m at Great Nicobar Island, while the small islands west off Middle Andaman uplifted about 1.5 m (Kayanne et al., 2007; Malik and Murty, 2005). Tide gauge record at Port Blair in South Andaman shows a slow subsidence with a duration of at least 15 min but possibly less than 35 min (Singh et al., 2006). The tsunami source was estimated from tsunami arrival times as 600–700 km (Lay et al., 2005), but later revised to be 900 km (Neetu et al., 2005) extending to the south of Andaman Islands. The satellite altimeter data support longer, more than 1000 km long, tsunami source (Fine et al., 2005b; Song et al., 2005; Hirata et al., 2006). The tide gauge data do not support tsunami source beneath Andaman Islands (Tanioka et al., 2006; Fujii and Satake, 2007), although Piatanesi and Lorito (2007) estimated slip on the deeper part of slab, east of Andaman Islands. Joint inversion of tide gauge and satellite data prefer the source lengh of 900 km (Fujii and Satake, 2007) (Figures 3 and 4). Seno and Hirata (2007) speculated that the fast rupture was followed by slow slip at shallower part near the trench. The source of the previous 1881 earthquake around Car Nicobar was studied from tide gauge records around Bay of Bengal (Ortiz and Bilham, 2003). The potential tsunami hazard in Indian Ocean was pointed out before the 2004 tsunami, with the computer simulation of the 1833 earthquake off Sumatra (Cummins, 2004). Future tsunami forecast models in the region were examined by Geist et al. (2007).
4.17.3 Other Examples In this section, other examples of tsunamis are briefly described in reverse-chronological order. Their sources are shown in Figure 5. We start with the recent and historical tsunamis that were instrumentally recorded, then prehistoric tsunamis that were geologically recorded. 4.17.3.1
Historical Tsunamis
4.17.3.1.1 tsunami
1998 Papua New Guinea
Another recent tsunami hazard was from an earthquake (Mw ¼ 7.1) along the northern coast of New Guinea Island on 17 July 1998. The tsunami heights
were as large as 15 m around Sissano Lagoon near the epicenter, and the total casualties were reported to be more than 2000. The large tsunami (>5 m), however, was limited to a small region (40 km) along the coast (Kawata et al., 1999) and the amplitudes on tide gauges around Japan are <10 cm (Satake and Tanioka, 2003), suggesting a local effect of tsunami generation or amplification. Post-tsunami marine surveys revealed bathymetry features that focused tsunami energy toward Sissano Lagoon and possible sources of submarine landslide (Tappin et al., 1999). Numerical simulations show that a source in addition to the earthquake fault motion, probably from submarine landslides, is needed to reproduce the locally large tsunami heights (Matsuyama et al., 1999; Geist, 2000; Titov and Gonzalez, 2001; Synolakis et al., 2002; Satake and Tanioka, 2003). 4.17.3.1.2
1993 Hokkaido tsunami The most recent tsunami disaster in Japan was that in Okushiri Island in Japan Sea, caused by the Southwest Hokkaido earthquake on 12 July 1993 (Mw ¼ 7.8). The tsunami attacked the island within a few minutes after strong ground shaking, and caused more than 200 casualties. The tsunami heights were 5–10 m around the island, while maximum tsunami run-up of more than 30 m occurred near a small valley. Numerical computations of tsunamis were made to reproduce the measured tsunami heights, recorded tsunami waveforms on tide gauges, as well as seismic or geodetic data (Shuto and Matsutomi, 1995; Takahashi et al., 1995; Tanioka et al., 1995; Titov and Synolakis, 1997). 4.17.3.1.3
1960 Chilean tsunami The largest (Mw ¼ 9.5) earthquake in the twentieth century occurred offshore southern Chile on 22 May 1960. The tsunami caused damage on the Chilean coast, including more than 1000 casualties, then propagated across the Pacific Ocean. About 15 h later, the tsunami reached the Hawaiian Islands and took 61 lives. It continued across the Pacific, reaching Japan in about 23 h, and caused 142 casualties. The large (5 m) tsunami in Japan was due to both focusing and resonance effects. Because of the sphericity of the Earth and bathymetry in the Pacific Ocean, energy was focused toward the Japanese Islands (Satake, 1988; Shuto, 1991). Some bays along the Pacific coast of Japan have a shape whose characteristic period coincided with the predominant period (about 1 h) of the trans-Pacific tsunami (Abe, 2000). This tsunami motivated to form an international tsunami warning system in the Pacific.
Tsunamis
70°
80°
90°
100°
487
110°
20°
10°
0°
–10°
–20°
Figure 3 Locations of instrumental observation for the 2004 tsunami: tide gauge stations (triangles) and satellite path (solid line). From Fujii Y and Satake K (2007) Tsunami source model of the 2004 Sumatra–Andaman earthquake inferred from tide gauge and satellite data. Bulletin of the Seismological Society of America 97: S192–S207.
4.17.3.1.4
1958 Lituya Bay tsunami The largest water run-up was recorded in Lituya Bay, Alaska. A large strike-slip earthquake (Mw ¼ 7.9) on 10 July 1958 triggered a huge rockslide (700–900 m wide, 1000 m long, and 90 m thick, or the total volume of about 3 107 m3) in this 10 km long bay. The rockslide generated water waves that surged up the opposite slope, and stripped the trees up to 520 m altitude (Miller, 1960; Lander and Lockridge, 1989). The large waves were confined in the bay; hence, this is not considered as a typical tsunami.
Lockridge, 1989). The tsunami traveled southward through the Pacific Ocean and reached the Hawaiian Islands in the morning of 1 April. Without any warning, the tsunami took 159 lives in the Hawaiian Islands. A Pacific tsunami warning system was introduced after this tsunami. Recent studies indicate that slow earthquake rupture process and large submarine landslide are responsible to far-field and near-field tsunamis, respectively (Okal et al., 2002, 2003).
4.17.3.1.6 4.17.3.1.5
1946 Aleutian tsunami The Aleutian earthquake of 1 April 1946 was very unusual; it generated huge tsunamis in the Aleutians and Hawaii, even though its surface wave magnitude Ms was only 7.4. This is a typical example of ‘tsunami earthquake’, which generates much larger tsunamis than expected from seismic waves (Kanamori, 1972). The Scotch Cap lighthouse on a 10 m cliff on Unimak Island, about 100 km from the epicenter, was totally demolished by the tsunami (Lander and
1896 Sanriku tsunami The 1896 Sanriku earthquake is another example of ‘tsunami earthquake’ (Kanamori, 1972; Tanioka and Satake, 1996b). The earthquake (Ms ¼ 7.2) was not great in size and the ground shaking was weak (2–3 on the Japanese intensity scale, corresponding to IV–V on modified Mercalli scale). However, the following tsunami was very destructive; the run-up height was >10 m for about 200 km of Sanriku coast with the maximum of 38 m, and resulted 22 000 casualties, the worst tsunami hazard in Japan in the last few centuries.
488 Tsunamis
90°
95°
4.17.3.1.7
1883 Krakatoa tsunami The 1883 eruption of Krakatoa, Indonesia, is one of the largest volcanic eruptions recorded in human history. The sound waves generated by the eruption were heard over 4000 km across the Indian Ocean. The eruption refaced the island of Krakatoa; the northen portion of the island disappeared and a caldera up to 270 m deep replaced it. The eruption also generated tsunamis that destroyed nearby coastal villages and took about 34 000 lives (Simkin and Fiske, 1983; Nomanbhoy and Satake, 1995). The maximum tsunami height along the Sunda Strait was about 15 m. Tide gauge disturbances were observed in India, New Zealand, and even at San Francisco, but those at very far places are interpreted as sea waves coupled by air waves from the eruption (Press and Harkrider, 1966; Pelinovsky et al., 2005).
100°
15°
10°
5°
4.17.3.2
Prehistoric Tsunamis
4.17.3.2.1
1700 Cascadia tsunami Geological and historical studies show the evidence of past earthquakes and tsunamis. Evidence of a great prehistoric earthquake that occurred along the Cascadia subduction zone has been recently found both near and far from the source (Atwater et al., 2005). Along the Pacific coast of North America, geological and tree-ring evidence indicates that coseismic subsidence took place about 300 years ago. In addition, tsunami deposits found throughout the Pacific Northwest indicate that this earthquake
0° 0
5
10
15 20 Slip (m)
25
30
35
Figure 4 Slip distribution on fault for the 2004 Sumatra– Andaman earthquake by an inversion of tsunami waveforms recorded on tide gauges and sea surface height measured by satellite. From Fujii Y and Satake K (2007) Tsunami source model of the 2004 Sumatra–Andaman earthquake inferred from tide gauge and satellite data. Bulletin of the Seismological Society of America 97: S192–S207.
90° N 1958 Lituya Bay
Storegga slides
60° N 1993 Hokkaido 30° N
1946 Aleutian 1896 Sanriku 1700 Cascadia
2004 Sumatra 1998 Papua New Guinea
0° 1883 Krakatoa 30° S
1960 Chile
60° S 90° S 60° E
120° E
180°
120° W
60° W
0°
60° E
Figure 5 Distribution of tsunamis in the world. Open circles are historical events with intensity 2 and validity 3 from Historical Tsunami Data Base (see Section 4.17.4.6) (Institute of Computational Mathematics and Mathematical Geophysics 2005). Solid circles are those described in Sections 4.17.2 and 4.17.3.
Tsunamis
generated a sizable tsunami. In Japan, historical records show a widespread tsunami with heights of 1–5 m struck the Pacific coast. From these records, the date and size of the earthquake was inferred to be 26 January 1700 and Mw ¼ 8.7–9.2, respectively (Satake et al., 1996, 2003). 4.17.3.2.2 Hokkaido tsunami in seventeenth century
Tsunami deposits show that the southern Kuril trench repeatedly produced tsunamis larger than those recorded in the last few centuries in Japan. Prehistoric sand sheets extend as much as 3 km inland along the 200 km of Hokkaido’s Pacific coast, where the recent tsunamis from 1952 to 2003 Tokachi-oki or 1894 to 1973 Nemuro-oki earthquakes (Mw 8) penetrated less than 1 km from the coast. The time intervals between the extensive sand sheets average about 500 years, and the most recent event occurred in the seventeenth century. These anomalous tsunamis, characterized by large inundation area and long recurrence interval, are best explained by earthquakes that rupture multiple segments, including Tokachi-oki and Nemuro-oki, of the Kuril subduction zone (Nanayama et al., 2003; Satake et al., 2005b). 4.17.3.2.3 slide
shorter period (less than 5 min), however, the system does not always follow the actual sea level change (Satake et al., 1988). More recent gauges include bubbler, acoustic, and laser gauges. The bubbler gauge is a pneumatic or gas-purged pressure gauge system. Compressed air or nitrogen gas is purged into a tube with an orifice, and its pressure is measured. As the gas slowly escapes from the orifice, the gas pressure is considered to be equal to water pressure, then converted to water level. The acoustic gauge sends ultrasonic waves in a tube that are reflected at the sea level. Sea level changes can be measured from the travel times of the acoustic wave (Figure 6). Laser gauge employs the same principle. The Global Sea Level Observation System (GLOSS), an international program, has been formed to establish high-quality global sea level network for climate change, oceanography, and coastal research (Woodworth and Player, 2003). Many GLOSS data are available in real time through the Internet. Because the main purpose of the sea level measurements is to monitor ocean tides or long-term sea level change, the typical sampling interval is not always short enough for tsunami monitoring. For tsunami measurements, high sampling rate, at least 1 min, is required.
Tsunami from the Storegga
The Storegga slide in the Norwegian Sea is the largest submarine landslide ever mapped. The Storegga slide is estimated to have taken place thrice: once at 30–50 ka and twice at 6–8 ka, with the total volume of 5.6 103 km3 (Bugge et al., 1988). The gigantic submarine landslide probably caused a tsunami. Tsunami deposits inferred from this event have been found on the coasts of Scotland, Iceland, and Norway (Smith et al., 2004; Bondevik et al., 2005).
4.17.4 Tsunami Observations 4.17.4.1
489
Instrumental Measurements
4.17.4.1.1 Coastal tsunami measurements
Coastal sea level has been measured for more than a century by tide gauges at ports and harbors throughout the world. The classic tide gauge is a mechanical type set in a tide or stilling well. Sea level change is detected by a float in the well. Because typical period of ocean tides is several hours or longer, sea level changes inside the well can be assumed to be the same as that outside the well. For tsunamis with
4.17.4.1.2 Offshore tsunami measurements
As tsunami becomes slow but large and complex near the coast, it would be ideal to detect its arrival offshore. Offshore wave gauges and GPS tsunami gauges have been developed and installed around Japan. Among about 50 acoustic wave gauges outside harbors at the water depths of tens of meters, about 10 recorded the tsunami from the 2003 Tokachi-oki earthquake (Nagai et al., 2004). A GPS tsunami gauge system has been installed about 10 km offshore Cape Muroto, Southwestern Japan, where the water depth is about 100 m (Kato et al., 2000, 2005). The gauge uses Real-Time Kinematic GPS technique to estimate the location and altitude of water surface. The sampling interval is 1 s with an accuracy of a few cm. Because it measures water surface movements, it is free from seismic waves unlike bottom pressure gauges (see the next section), and can also be used to monitor wave heights during storms. 4.17.4.1.3
Deep ocean measurements In the deep ocean, tsunami waveforms are expected to be simpler, free from coastal topographic effects, although their amplitudes are smaller. Deep ocean
490 Tsunamis
Figure 6 Types of instruments.
measurements of tsunamis have been made by using bottom pressure gauges for early detection and warnings of tsunami. A quartz crystal sensor is typically used as a pressure sensor. The sensor provides accuracy of 1 cm in sea level at a 6000 m depth. Because the pressure transducer is very sensitive to temperature, measurements and correction for temperature are essential. In addition, because the sensor is located at the ocean bottom, it also records ground motion and works as an ocean bottom seismograph. While the frequencies are often different between seismic and tsunami waves, it is necessary to separate them. Around Japan, more than 10 bottom pressure gauges are installed on six submarine cables (Okada, 1995; Iwasaki et al., 1997; Hirata et al., 2002). Tsunamis from relatively small tsunamigenic earthquakes, the 2004 Off Kii Peninsula earthquakes (M 6.9 and 7.4), were recorded on a GPS tsunami gauge and eight bottom pressure gauges, as well as coastal tide gauges located south of Honshu and Shikoku (Satake et al., 2005a). The US system, called Deepocean Assessment and Reporting of Tsunamis (DART) or simply Tsunameter, sends signals to surface buoy via acoustic telemetry in the ocean, then via satellites to land station in real time (Gonzalez et al., 2005). 4.17.4.1.4
Satellite observations Four satellites captured the tsunami propagation across the Indian Ocean on 26 December 2004 (Figure 3): Jason-1 and TOPEX/Poseidon at about 2 h after the earthquake, Envisat at 3 h 15 min, and
US Navy’s GEOSAT follow-on at 7 h (Gower, 2005). The maximum sea surface height (SSH) of the tsunami was about 0.8 m in the Indian Ocean. The accuracy of satellite altimetry is better than 5 cm and its spatial resolution is approximately 15 arc-min (27 km). Numerical simulation from the 2004 source model could reproduce the SSHs (Smith et al., 2005; Hirata et al., 2006) (see Section 4.17.2). 4.17.4.1.5 Definition of amplitude and wave heights
The tsunami ‘amplitude’ is measured from zero to peak, either positive or negative. The ‘wave height’, also called as range or double amplitude, is measured from trough to peak. Hence, the ‘tsunami height’ is measured from trough to peak on tide gauges, whereas the ‘tsunami amplitude’ is measured from estimated tidal level at the time (Figure 7). As in seismic waves, ‘initial motion’, either positive or negative, carries information on the tsunami source. For example, the initial motions of the 2004 tsunami in Indian Ocean were negative to the east of the source (e.g., Thailand) but positive to the west (e.g., Maldives) as shown in Figure 7. 4.17.4.2 Field Survey to Measure Run-Up Heights 4.17.4.2.1
Post-tsunami surveys Post-tsunami field surveys are often carried out after large tsunamis. Usually, experts from different backgrounds, such as coastal engineering, geology, or
Tsunamis
2
Male, Maldives (Max) Amplitude
491
(Max) Wave height
Amplitude (m)
0
–2 2
Estimated tide
Arrival time Phuket, Thailand (Max) Wave height
Estimated tide
0 (Max) Amplitude –2 –3
0
3
6
9
12
Time (hour from earthquake) Figure 7 Tsunami waveforms from the 2004 Sumatra–Andaman earthquake recorded at Male (top) and Phuket (bottom).
seismology, and from different countries form international tsunami survey teams that visit affected coasts to document tsunami effects. An essential element of survey is to measure the tsunami heights along the coast (Intergovernmental Oceanographic Commission, 1998). The heights, defined below, are measured from various traces such as damage on buildings or other construction, watermarks left on walls, vegetative markers such as trim lines and saltwater discoloration or die-off, or lines of debris deposited on beaches. When no physical evidence is found, the tsunami heights are estimated based on eyewitness accounts, but considered as less reliable. Locations are usually identified by portable GPS measurements and the heights are measured by traditional leveling or laser surveying. Interviews with local people also include degree of ground shaking, tsunami arrival times, the number of waves with periods, and the duration. Post-tsunami surveys also include description and documentations of building damage and tsunami deposits. Synolakis and Okal (2005) review and summarize 15 post-tsunami surveys made by international tsunami survey teams during the 11-year period between 1992 and 2002.
Figure 8 Definitions of tsunami heights on land.
4.17.4.2.2 Definitions of inundation and run-up heights
Tsunami height on land is called ‘flow depth’ when measured from the ground level, and ‘inundation height’ measured relative to sea level at the time of tsunami arrival. The height at maximum inundation is called ‘run-up height’. Because the tide levels at the time of measurements and tsunami arrival are usually different, the corrections are necessary for survey data. Unless the sea surface of inundated tsunami is flat, run-up height is not necessarily the same as the inundation height or flow depth at the shore. The horizontal distance of water penetration measured from the coast is called ‘inundation distance’ (Figure 8). 4.17.4.3
Study of Historical Documents
Tsunamis and their effects documented in historical records are also used to study old earthquakes and tsunamis for which no instrumental data are available. Some countries in Europe, South America, or Asia have a long history of written records that provide accurate dates for past earthquakes and
492 Tsunamis
tsunamis. For example, tsunami heights from the 1755 Lisbon earthquake have been estimated from historical documents as 6 m in Lisbon, more than 15 m at Southwest Portugal, and more than 10 m in Spain and Morocco (Baptista et al., 1998). Estimation of tsunami heights from damage description requires many assumptions. For example, the Japanese descriptions of the January 1700 tsunami, generated by a giant earthquake in Cascadia subduction zone, include flood in farmed fields and a government warehouse and damage to coastal houses (Atwater et al., 2005). To convert these reported effects into estimated heights, Satake et al. (2003) considered various factors, such as the relationship between tsunami damage and heights, sea level change in the last 300 years, and similarity with the 1960 Chilean tsunami. Japanese tsunami data indicate that the house damage is partial when tsunami height is less than 2 m, but becomes significant when tsunami height becomes more than 4 m (Hatori, 1984).
4.17.4.4
Geological Methods
Even older tsunamis are studied from their deposits by geological means (Dawson and Shi, 2000). In Europe, inferred tsunami deposits from the Storegga slides in the North Sea have been found and studied in Scotland, Norway, and Iceland (Dawson et al., 1988). In the US, tsunami deposits from prehistoric great earthquakes along the Cascadia subduction zone have been studied (Atwater et al., 2005). In Australia, geological evidence of Late Pleistocene tsunamis has been studied (Bryant et al., 1996). In Japan, tsunami deposits in various environments such as lagoons, beaches, or marshes have been studied for modern, historic, and prehistoric tsunamis (Minoura and Nakaya, 1991; Nanayama et al., 2003). On some coasts, particularly those remote from tsunami sources, it can be difficult to distinguish between tsunami and storm deposits. The universal sedimentological criteria for identifying tsunami deposits have not been well established (Nanayama et al., 2000; Tuttle et al., 2004). For this context, studies of deposits from modern tsunamis are important and careful descriptions have been done for the recent tsunamis such as the 1993 Hokkaido tsunami (Nishimura and Miyaji, 1995), the 1998 Papua New Guinea tsunami (Gelfenbaum and Jaffe 2003), or the 2004 Indian Ocean tsunami (see Section 4.17.2). These case studies indicate that the tsunami deposits
are highly variable and depend on the coastal topography and geology.
4.17.4.5
Tsunami Magnitude Scales
4.17.4.5.1
Imamura-Iida scale, m The traditional tsunami magnitude scale is the socalled Imamura-Iida scale, m. Although the original definition was descriptive, the value is approximately equal to m ¼ log2 h
½1
where h is the maximum run-up height in meters (Iida et al., 1967). This scale is similar to the earthquake intensity scale, and is especially convenient for old tsunamis from which no instrumental records exist. Hatori (1979) extended the Imamura-Iida’s m scale to include far-field tsunami data as follows: n o pffiffiffi m ¼ 3 þ log ðh=0:5ÞðR=1000Þ1=2 =log 5
½2
where h is in meters and R is the distance from source to coast in kilometers. This indicates pffiffiffithat the tsunami height decays with distances as 1= R; which is theoretically predicted for nondispersive waves traveling a long distance (Comer, 1980). 4.17.4.5.2
Tsunami intensity Soloviev (1970) pointed out that Imamura-Iida’s m scale is more like an earthquake intensity scale rather than a magnitude. He also distinguished the maximum tsunami height h and the mean tsunami height Hav. He then defined tsunami intensity I as I ¼ log2
pffiffiffi 2Hav
½3
Comparison of [1] and [3] suggests that the mean pffiffiffi tsunami height is given as 1= 2 times the maximum height. The maximum intensity I on the coast nearest to the source is used to quantify the tsunami source. 4.17.4.5.3
Tsunami magnitude M t Another magnitude scale Mt, called tsunami magnitude, was defined and assigned for many earthquakes by Abe (1979). The definition of Mt for a transPacific tsunami is (Abe, 1979) Mt ¼ log H þ C þ 9:1
½4
and for a regional (100 km < < 3500 km) tsunami is (Abe, 1981) Mt ¼ log H þ log þ 5:8
½5
Tsunamis
where H is a maximum amplitude on tide gauges in meters, C is a distance factor depending on a combination of the source and the observation points, and is the nautical distance in kilometers. The above formulas were calibrated with the moment magnitude scale, Mw, of earthquakes. Equation [5] indicates that for the same Mt, the tsunami amplitude H decays with 1/, which is different from the assumption used for Imamura-Iida’s m scale, and may be valid only for regional distances. 4.17.4.6
Tsunami Catalogs
The location, size, run-up heights, and damage from past tsunamis have been compiled by several researchers and published as catalogs. For the Pacific tsunamis, the Iida catalog (Iida et al., 1967) contains numerous tsunami height data as far back as AD 173. Russian catalogs (Soloviev and Go, 1974a, 1974b; Soloviev et al., 1992) include Pacific tsunami data (but except for Russia) with some tide gauge records, tabulated tsunami heights, and descriptions for each tsunami. Regional catalogs including tide gauge records were also published for Alaska (Lander, 1996), the west coast of the US (Lander et al., 1993), Japan (Iida, 1984; Watanabe, 1998), and Indian Ocean (Berninghausen, 1966; Murty and Rafiq, 1991). Digital tsunami databases have been compiled, termed Historical Tsunami Data Base at the Institute of Computational Mathematics and Mathematical Geophysics, Novosibirsk (Gusiakov, 2005; Institute of Computational Mathematics and Mathematical Geophysics, 2005) and at the National Geophysical Data Center, NOAA.
4.17.5 Tsunami Generation In this section, we review tsunami generation by earthquakes, landslides, and volcanic eruptions. 4.17.5.1
Earthquake Tsunamis
4.17.5.1.1 Ocean bottom and water surface displacements
Analytical solutions for linear gravity waves generated from several types of sources have been obtained by various researchers. Kajiura (1963) generalized the solutions by introducing the tsunami Green’s functions. He showed that the initial conditions given at the ocean bottom and at the water surface differ by a factor of 1/cosh kd, where k is wave number of
493
displacement. For large wavelength, >> 2d, this factor becomes 1 because cosh kd ¼ cosh (2d/) 1. Therefore, if the wavelength of ocean-bottom deformation is much larger than the water depth, it can be assumed that the water surface displacement is the same as the bottom displacement. Kajiura (1970) examined the energy exchange between the solid Earth and ocean water for tsunamis generated by bottom deformation with a finite rise time. He showed that the efficiency of tsunami generation, defined as a ratio of dynamic tsunami energy to static energy, becomes nearly 1, if the duration is short compared to the time required for a tsunami to travel through the source area. For an earthquake with M 8 at the water depth of 5000 m, the source size is about 100 km and the tsunami travel time over that distance is about 8 min. If the duration is much shorter, for example, 1 min or less, the efficiency is more than 0.9. This means that the bottom deformation can be considered as instantaneous. If the pffiffiffiffiffiffiffiffi duration is too short, shorter than 2 d =g ; a part of the energy flux goes into acoustic wave. For the above case, this limit is 45 s. 4.17.5.1.2 Fault parameters and ocean-bottom deformation
The ocean-bottom deformation due to faulting can be calculated using the elastic theory of dislocation. The displacement in an infinite homogeneous medium due to a dislocation across a fault is given by Volterra’s theorem (Steketee, 1958) and the explicit formulas for half-space are given (e.g., Mansinha and Smylie, 1971; Okada, 1985). Usually, only the vertical component of the ocean bottom is considered for tsunami generation. When the tsunami source is on a steep ocean slope and the horizontal displacement is large, the vertical displacement of water due to the horizontal displacement of the slope must also be considered (Tanioka and Satake, 1996a). The fault parameters needed to compute surface deformation are the fault location, geometry (strike, dip, and rake), size (length L and width W ), and average slip u (or the strike-slip and dip-slip components). The seismic moment M0 is given as M0 ¼ uS ¼ uLW
½6
where S is the fault area and is rigidity. The moment magnitude is defined as (Hanks and Kanamori, 1979) Mw ¼
2 log M0 – 10:7 3
½7
494 Tsunamis
where M0 is given in dyne cm (107 N m). The effects of fault parameters on tsunami generation were examined by Yamashita and Sato (1974) and more recently by Geist (1998).
4.17.5.1.3 Estimation of earthquake source parameters from tsunamis
Tsunami data can be used to study earthquake source processes in a way similar to that in which seismic waves are used. Abe (1973) analyzed the tsunami arrival times and first-motion data from the 1968 Tokachi-oki earthquake. By drawing inverse refraction diagrams from each station (see Section 4.17.6.2), he estimated the tsunami source area, which agrees well with the aftershock area. From the first motion of tsunami waves, the initial water surface disturbance was estimated as uplift at the southeastern edge and subsidence at the northwestern edge. This pattern is very similar to the vertical bottom deformation due to the faulting, which was independently estimated from seismological analysis. Numerical simulation of tsunami has been carried out for many fault models around Japan (e.g., Aida, 1978). The best fault models were obtained by trial and error comparisons of observed and computed tsunami waveforms or coastal tsunami heights. Tsunami waveforms recorded on tide gauges were then inverted to estimate the slip distribution on the fault (Satake, 1989). In this method, the fault plane is first divided into several subfaults, and the deformation on the ocean bottom is computed for each subfault with a unit amount of slip. Using this as an initial condition, tsunami waveforms are numerically computed on actual bathymetry. The observed tsunami waveforms are expressed as a superposition of waveforms computed for each subfault as follows: Aij ðt Þ ? xj ¼ bi ðt Þ
½8
where Aij (t) is the computed waveform as a function of time t, or Green function, at the ith station from the j th subfault; xj is the amount of slip on the j th subfault; and bi(t) is the observed tsunami waveform at the ith station. The slip xj on each subfault can be estimated by a leastsquares inversion of the above set of equations. The tsunami waveform inversion has been applied to many large and great earthquakes in the Alaska-Aleutians ( Johnson, 1999), northern Japan (Satake, 1989; Tanioka et al., 1995), the Nankai trough (Satake, 1993; Tanioka and Satake, 2001a, 2001b), and the 2004 Sumatra–Andaman earthquake (Fujii and Satake, 2007). These results show that the slip distributions
are not uniform but are concentrated in small regions, which correspond to asperities (Figure 3). 4.17.5.1.4 Tsunamigenic and tsunami earthquakes
We call an earthquake that generates tsunami as ‘tsunamigenic earthquake’. A ‘tsunami earthquake’, on the contrary, is defined as an earthquake which excites much larger tsunamis than expected from its seismic waves (Kanamori, 1972). The discrepancy can be quantified in terms of the surface wave magnitude (Ms) and tsunami magnitude (Mt) scales; when Mt is larger than Ms by more than 0.5 units, it is called a tsunami earthquake (Abe, 1989). Newman and Okal (1998) introduced a new parameter, called slowness parameter Y ¼ log10 ðE E =M0 Þ, a ratio of radiated seismic energy EE to seismic moment M0. They show that Y is about –5 for regular earthquakes but smaller by 1–1.4 for tsunami earthquakes, and they proposed to use this parameter for discrimination of tsunami earthquakes. Several proposals have been made for the generating mechanisms of tsunami earthquakes. Slow and long rupture processes (Kanamori, 1972) reconcile the large discrepancy between seismic and tsunami waves. High-angle secondary faulting in the accretionary wedge can explain the relatively large tsunami excitation (Fukao, 1979). For a given slip, the seismic moment is smaller in the shallower part of subduction zones because of the smaller rigidity, and the tsunami becomes larger (Okal, 1988). Tsunami heights may not be anomalously large if compared to Mw rather than Ms (Pelayo and Wiens, 1992); the large discrepancy between the tsunami and Ms is due to a saturation of Ms at around 7.3 for those earthquakes that occurred in the accretionary wedge. Recent tsunami waveform modeling shows that the fault parameters and locations of tsunami earthquakes share a common feature: narrow and shallow faulting near the trench axis (Figure 9). The 1992 Nicaragua earthquake (Ms ¼ 7.2, Mt ¼ 7.9) was the first tsunami earthquake recorded on modern broadband seismic instruments, and seismological studies showed that the duration was very long for its size, about 100 s (Kanamori and Kikuchi, 1993). Comparison of numerically computed tsunami waveforms with tide gauge records (Satake, 1994) showed that a narrow (40 km) and shallow (extending only to the upper 10 km of the ocean bottom) fault is responsible for the tsunami generation. Among the tsunami earthquakes, the 1896 Sanriku earthquake (Ms ¼ 7.2, Mt ¼ 8.6) and the 1946 Aleutian
Tsunamis
495
Figure 9 Vertical deformation of seafloor (top) due to faulting (bottom) for typical interplate earthquake (left) and tsunami earthquakes (right). Modified from Satake K and Tanioka Y (1999) Sources of tsunami and tsunamigenic earthquakes in subduction zones. Pure and Applied Geophysics 154: 467–483.
earthquake (Ms ¼ 7.3, Mt ¼ 9.3) were very anomalous (Kanamori, 1972; Kanamori and Kikuchi, 1993). While the tsunami modeling showed a similar character in fault parameters to the Nicaragua earthquake, that is, narrow fault near trench axis (Tanioka and Satake, 1996b; Johnson and Satake, 1997; Tanioka and Seno, 2001a, 2001b), Kanamori and Kikuchi (1993) suggested that submarine slumping was involved in the unusually large tsunami generation. These tsunami earthquakes occurred in trenches with large amounts of sediments and an accretionary prism where slumping occasional occur. Other tsunami earthquakes such as the 1992 Nicaragua earthquake occurred in subduction zones with little sediment, where a slow rupture is responsible for large tsunamis (Kanamori and Kikuchi, 1993). Tanioka et al. (1997) showed that the 1896 tsunami earthquake occurred in a region where the oceanbottom topography is rough, characterized by a welldeveloped horst and graben structure. Polet and Kanamori (2000) extended this model to global subduction zones, based on the examination of source spectra of large (M > 7) earthquakes in the 1990s. Okal and Newman (2001), however, failed to find any correlation between the slowness parameter Y and either depth, focal mechanism, or seismic moment for central America, Java, and Peru,
suggesting an absence of regional signal for the occurrence of tsunami earthquakes. Bilek and Lay (2002), from analysis of source time functions of tsunamigenic and tsunami earthquakes, found that shallower earthquakes have longer duration and claimed that tsunami earthquakes are simply larger earthquakes at very shallow depth. Seno (2002) proposed that increase in fluid pressure causes temporal change in frictional property near trench axis and responsible for slow rupture; hence, the tsunami earthquakes are transient phenomena. Seno and Hirata (2007) showed that the 2004 earthquake had some component of ‘tsunami earthquake’, slow duration, and updip rupture of seismogenic zone. They inferred that the seismogenic rupture propagated toward north at the rupture velocity of 2.5 km s1, but was followed by slower rupture (0.7 km s1) at shallower part, which has about a third of the total seismic moment.
4.17.5.2
Tsunamis from Landslides
Tsunamis from submarine landslides have been a hot research topic since the 1990s. Abnormally large and concentrated tsunami heights from the 1992 Flores Island, Indonesia, and 1998 Sissano, Papua New Guinea, earthquakes and the nonseismic 1994
496 Tsunamis
Skagway tsunami, Alaska, provided case studies. Marine surveys revealed detailed bathymetric features of landslides or slumps. Physical experiments in tanks and numerical computations of landslidegenerated tsunamis have been developed. Such multidisciplinary studies have been summarized in several special issues (Keating et al., 2000; Bardet et al., 2003; Tappin, 2004). 4.17.5.2.1
Examples Post-tsunami field surveys after large tsunamigenic earthquakes sometimes show abnormally large tsunami heights concentrated in a small portion of coasts. While offshore bathymetry or coastal topography (such as V-shaped bays) can cause such focusing of incoming tsunamis, locally large tsunami may be caused by a local landslide, either subaerial or submarine, triggered by ground shaking. Large water waves caused by a rockslide in Lituya Bay at the time of 1958 earthquake are described in Section 4.17.3.1.4. Concurrent with the 1964 Alaska earthquake (Mw 9.2), in addition to Pacific-wide tsunamis, local tsunamis from landslides, were reported at more than 20 locations, with the largest run-up height of 50 m (Plafker et al., 1969; Lander, 1996). More recently, the tsunami heights from the 1992 Flore Island earthquake were mostly 4–5 m on most of the northern coast of the island, but 26 m at Riangkroko (Yeh et al., 1993). The tsunami heights from the 1998 Papua New Guinea tsunami, also described in Section 4.17.3.1.1, show local peak of 15 m (Figure 10). Larger-scale submarine slumping and turbidity currents were triggered by the 1929 Grand Banks earthquake (Heezen and Ewing, 1952; Fine et al., 2005a). The maximum tsunami run-up height of 13 m and 28 casualties made it the worst tsunami disaster in Canada. The turbidity current also caused damage on the trans-Atlantic submarine cables at 12 locations over a range of 600 km. An example of nonseismic landslide tsunami is the 1994 Skagway tsunami (Kulikov et al., 1996). A submarine landslide, without seismic trigger, occurred in Skagway Harbor in Alaska and generated local tsunami of 5–10 m high. The tide gauge records show a dispersive character, because of small size of the source (see the next section). 4.17.5.2.2 Characteristics and discrimination of landslide tsunamis
Scale length of submarine mass failure is very different from that of submarine faulting. While vertical movement can be hundreds of meters, the horizontal
scale is mostly less than 10 km for submarine mass failure. Seafloor deformation due to faulting, on the contrary, has at most several meters of vertical movement but the spatial extent can be several hundreds of kilometers. As will be discussed in next section, an earthquake source almost always satisfies the longwave (shallow water) condition, while the submarine landslide may not. Hence, the water waves from submarine landslides show a dispersive character and attenuate quickly without reaching a long distance. Okal and Synolakis (2004) fit the run-up height distribution (y) along a beach (in the y-direction) by the following function: ðyÞ ¼
b ½ðy – cÞ=a2 þ 1
½9
where parameters a, b, and c are estimated from the observed data (Figure 10). They show that I2 ¼ b/a, the aspect ratio of maximum run-up heights to extent of run-up along beach, can be used as a discriminant of the nature of tsunami source, either dislocation (earthquake fault) or landslide source. Seismic dislocation sources have I2 smaller than 104, whereas a landslide source has a larger value.
4.17.5.2.3
Modeling landslide tsunamis Submarine landslide sources of tsunamis have been also modeled in many ways. These can be grouped into three categories: dynamic fluid flow, kinematic rigid-body motion, and static initial profile. In the dynamic fluid flow approach, landslide is treated as fluid and its evolution is calculated simultaneously with tsunami generation. Jiang and Leblond (1994) treated a landslide as incompressible viscous flow and seawater as inviscid flow, and formulated the multiphase fluid dynamics. Fine et al. (2005a) applied the multiphase flow method to the 1929 Grand Banks landslide tsunami; Imamura et al. (2001), Heinrich et al. (2001), and Titov and Gonzalez (2001) applied a similar method to the submarine landslide of the 1998 Papua New Guinea tsunami. In the rigid-body approach, temporal change of water height due to progress of submarine landslide is prescribed and used to compute tsunami propagation. Harbitz (1992) assumed a box-shaped slide for the Storegga slide, Ward (2001) used a chain of rigid boxed to model the Storegga and other submarine slides. Tinti et al. (1997) introduced a block model of landslide and simulated the landslide motion, then
Tsunamis
Run-up, ζ (m)
15
Aspect ratio I 2 = 4.74 × 10–4
1998 PNG
c = –6 km
5
0
4
a = 19 km
b=9m
10
6
497
0
–100 2002 PNG
100
Aspect ratio I2 = 1.17 × 10–5
b = 1.5 m c = –25 km
2 0
a = 131 km
–100
100 0 Distance y along beach (km)
Figure 10 Comparison of tsunami run-up heights for the 1998 and 2002 Papua New Guinea tsunamis. Estimated parameters of [9] are also shown. Modified from Okal EA and Synolakis CE (2004) Source discriminants for near-field tsunamis. Geophysical Journal International 158: 899–912.
used the results for the tsunami computations. Satake et al. (2002) introduced a simple kinematic landslide model on the basis of bathymetry change. In the static approach, an initial profile due to landslide is assumed. Watts et al. (2003), based on the results of physical and numerical experiments, introduced a parametric initial profile for submarine landslide, and applied such initial profiles for several landslide-generated tsunamis including the 1946 Aleutian, 1994 Skagway, and 1998 Papua New Guinea tsunamis. Synolakis et al. (2002) and Satake and Tanioka (2003) used the same approach to model the 1998 Papua New Guinea tsunami. For water waves generated from subaerial landslides, various studies such as theoretical, experimental, numerical and field approaches are summarized in Slingerland and Voight (1979).
4.17.5.3
Volcanic Tsunamis
Tsunamis are one of the most important hazards associated with volcanic activity. About a quarter of human casualties directly associated with volcanic eruption is due to tsunamis (Tilling, 1989). Volcanic tsunamis can be roughly categorized into five groups according to their origin: (1) earthquakes; (2) landslides and avalanches of cold rock and lahars; (3) flow of volcanic ejecta such as pyroclastic flow, base surge, hot rock avalanches, or lava; (4) submarine explosion;
and (5) caldera collapse or subsidence. From case studies of 106 volcanic tsunamis, Latter (1981) classified 69 of them into the 10 groups. In addition, he also classified tsunamis generated from air waves of explosions (4.5%) with possible examples in Bezymianny, Kamchatka in 1955–1956, as well as some far-field tsunamis from the 1883 Krakatoa eruption. 4.17.5.3.1 Volcanic tsunamis due to earthquakes
According to Latter (1981), some 22% of volcanic tsunamis are caused by earthquakes accompanying eruptions. Examples include the 1914 Sakurajima, Japan, 1930 Stromboli, Italy, and 1933 Severgin, Kuril Islands, tsunamis. 4.17.5.3.2 Volcanic tsunamis due to landslides
Latter (1981) separated the tsunamis generated from landslides and avalanches of cold rock (7%) and lahars (5%). A typical example of landslide is from Shimabara, Japan. In 1792, Mayuyama, a small mountain near Unzen volcano, collapsed and the landslide entered Ariake-kai, a shallow bay with typical depth of 10–20 m. Because the horizontal extent of the slide is about 3000 m, much longer than the water depth, this satisfied the long-wave approximation, hence causing large (10–20 m) tsunami height and
498 Tsunamis
significant damage (5000 casualties) on the other side of the bay. The total casualties was 15 000.
4.17.6 Tsunami Propagation
4.17.5.3.3 Volcanic tsunamis from flow of volcanic ejecta
In the Earth-ocean-coupled system, an earthquake source would generate waves both in the elastic Earth and in the ocean, and the propagation of a tsunami in the ocean layer would excite waves in the elastic Earth. In contrast, the traditional decoupled approach first calculates the elastic deformation on the ocean bottom generated by an earthquake, then uses this as an initial condition for tsunami wave propagation such that the tsunami wave in the ocean would have no influence on the elastic Earth. Theoretical treatment of tsunamis viewed as coupled system is similar to that of Rayleigh waves or spheroidal oscillations. Ward (1980) and Okal (1988) extended the free oscillation approach to solve the generation and propagation of tsunamis, while Comer (1984a) extended the Rayleigh wave approach to solve the tsunami problem. The calculated dispersion curve shows that, except at very low frequencies, the phase velocity is very similar to that of the gravity wave. Both approaches give almost identical solutions, because of the small compressibility of water (Comer 1984b). In other words, the coupling between the elastic Earth and the ocean is extremely weak, which explains why very large earthquakes are necessary to generate tsunamis.
About a third of volcanic tsunamis can be classified in this group. Latter (1981) further separated them into pyroclastic flows (ignimbrites and nuees ardentes; 20%), base surges (7%), avalanches of hot rock (8%), or lava (1%). Latter (1981) considered the 1883 Krakatoa tsunami as due to pyroclastic flow, although this is controversial (see Section 4.17.5.3.5). Other examples of pyroclastic flow include St. Vincent and Mt. Pelee in lesser Antilles in 1902 and St. Augustine (Alaska) in 1883. Other types of ejecta are base surge in Taal volcano in the Philippines, hot rock avalanche at Stromboli, Italy, and lava avalanche at Matavanu volcano, Samoa in 1906–1907.
4.17.5.3.4 Volcanic tsunamis from submarine explosion
Submarine explosion occupies about 19% of volcanic tsunamis according to Latter (1981). Submarine volcano, Bayonnaize Rocks (along Izu-Bonin Islands about 300 km south of Honshu), erupted in 1952. Tsunamis from this eruption, recorded at Hachijo Island about 130 km north, show very short period (1 min) wave with a dispersive character. Unoki and Nakano (1953) applied a Cauchy–Poisson wave (deep water waves from initial water surface disturbance) to the observed tsunami record and estimate the source size (2 km) and water surface uplift or impulse force.
4.17.5.3.5 collapse
Volcanic tsunamis from caldera
About 9% of volcanic tsunamis were from caldera collapse or subsidence. Large-scale caldera collapse was recorded at the times of the 1883 Krakatoa and the 1883 Ritter Island (Papua New Guinea) eruptions. Ward and Day (2003) modeled the Ritter Island tsunami by using the landslide model. The generation mechanism of the 1883 Krakatoa tsunami had been long debated. A numerical computation of this tsunami and comparison with tide gauge record and run-up heights indicated that the tsunami was generated by submarine eruption, rather than caldera collapse, pyroclastic flow, or falling ejecta (Nomanbhoy and Satake, 1995).
4.17.6.1 Weak Coupling between Elastic Earth and Ocean
4.17.6.2
Ray-Theoretical Approach
If the tsunami wavelength is much smaller than the scale of heterogeneity in propagation velocity, or the change in water depth, then we can apply the geometrical ray theory of optics. The wavefronts of propagating tsunami can be drawn on the basis of Huygens’ principle. Such a diagram is called a ‘refraction diagram’. Refraction diagrams can be prepared for major tsunami sources and used for tsunami warning; as soon as the epicenter is known, the tsunami arrival times can be readily estimated. Figure 1 shows the refraction diagram from the 2004 Andaman–Sumatra earthquake with wavefronts at each hour. It is shown that the tsunami was expected to arrive at Thai and Indian coasts in 2 h and at Africa in 7 hs. The refraction diagrams can also be drawn backwards from coasts. Such a diagram is called an ‘inverse refraction diagram’ and used to estimate the tsunami source area. In this case, the wavefronts or
Tsunamis
rays are traced backwards for the corresponding travel times (Miyabe, 1934). The traced wavefronts from each observed station bound the tsunami source area.
4.17.6.3 Law
Coastal Amplification and Green’s
From the conservation of energy along the rays, it can be shown that 1=2
1=2
b0 d0 h02 ¼ b1 d1 h12
½10
where d is the water depth, b is the distance between the rays, h is the tsunami amplitude, and the subscripts 0 and 1 indicate two different locations. If the tsunami amplitude at location 0 (e.g., deep water) is known, the tsunami amplitude after propagation can be estimated as h1 ¼
1=2 1=4 b0 d0 h0 b1 d1
½11
This is known as Green’s law. The ratio b1/b0 represents the spreading of rays, which can be graphically obtained from refraction diagrams. Note that this applies only to direct waves and does not account for of the reflected waves or edge waves.
4.17.6.4
Tsunami as a Gravity Wave
In this section, we take a two-dimensional Cartesian coordinate system with z-axis vertical upward (the origin is on the undisturbed water level) and consider propagation of a wave (Figure 11). When gravity is the restoring force, it is called a gravity wave. Euler’s equation of motion can be written as qV 1 þ V ? rV ¼ – g – rp qt
z
½12
d
u
Figure 11 Definition of coordinate system and the parameters for gravity waves.
where V is the velocity vector (whose x- and z-components are u and w, respectively), g is the gravitational acceleration, and p are water density and pressure, respectively. The phase velocity c of gravity waves is given as c¼
g k
tanh kd
1=2
¼
g 2d 1=2 tanh 2
½13
where k is the wave number, is the wave length, and d is the water depth. When the horizontal scale of motion, or the wavelength , is much larger than the water depth (>>d), the vertical acceleration of water is negligible compared to the gravity. This means that the horizontal motion of water mass is almost uniform from bottom to surface. Such a wave is called a shallow-water wave or long wave. For tsunamis, typical ocean depth is about 4–5 km and a large earthquake has a source area of several tens to hundreds of kilometers. Hence, the long-wave approximation is appropriate for most cases. When the amplitude is small compared to the water depth (h << d), the equation of motion, or conservation of momentum, can be written as qu qh ¼ –g qt qx
½14
and the equation of continuity, or conservation of mass, as qh q ¼ – ðduÞ qt qx
½15
Such a wave is called a small-amplitude, or linear long wave. The assumptions are valid for most tsunami propagation paths except for near-shore regions. From [14] and [15], by assuming that water depth d is constant, we obtain the wave equation 2 pffiffiffiffiffi q2 h 2q h ¼ c ; c ¼ gd qt 2 qx 2
½16
in which the phase velocity is determined by the water depth only. Note that the above phase velocity can be also obtained from [13] if >>d.
λ h
499
x
4.17.6.5 Basic Equations for Tsunami Propagation Including the bottom friction and the Coriolis force, the momentum equation for shallow-water (long)
500 Tsunamis
waves can be written for a three-dimensional case as follows: pffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi qU qU qU qh U U 2 þV 2 þU þV ¼ – f V – g – Cf d þh qt qx qy qx ½17 pffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi qV qV qV qh V U2þV 2 þU þV ¼ f U –g – Cf qt qx qy qy d þh
These equations, including frequency dispersion, are known as the linear Boussinesq equation. If we relax the small-amplitude (h << d) assumption, the nonlinear, nondispersive equations of motion and continuity are
and the mass equation is qh q q þ fU ðh þ d Þgþ fV ðh þ d Þg ¼ 0 qt qx qy
½18
where the coordinate system is x ¼ east and y ¼ south, f is the Coriolis parameter, Cf is a nondimensional bottom frictional coefficient, and U and V are the depth-averaged velocities in x- and y-directions, respectively. The velocities are averaged from the bottom to the surface: Z h 1 Qx u dz ¼ h þ d –d hþd Z h Qy 1 V ¼ v dz ¼ h þ d –d hþd
qu qu qh þ u ¼ –g qt qx qx
½23
qh q ¼ – fuðh þ d Þg qt qx
½24
These equations are for the finite-amplitude shallowwater waves and account for amplitude dispersion. For the linear case (ignoring the advection term), the phase pffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi velocity is given by c ¼ gðd þ hÞ. This means that the larger the amplitude, the faster the wave speed. As a consequence, peaks of a wave catch up with troughs in front of them, and the forward facing portion of the wave continues to steepen and it will eventually break.
U¼
½19
where Qx and Qy are flow rates in x- and y-directions. A typical value of a non-dimensional frictional coefficient Cf is 103 for coastal water and 102 for run-up on land (Satake, 1995).
4.17.6.6
Dispersion
Here we will briefly examine general gravity waves deviating from the small-amplitude, linear shallowwater (long) wave approximations (Peregrine, 1972; Mei, 1989). We have already seen that the phase velocity of a general gravity wave is given by [13]. If we expand the hyperbolic tangent function into its Taylor series and include the second term, [13] becomes c¼
1=2 pffiffiffiffiffi pffiffiffiffiffi 1 1 gd 1 – ðkd Þ2 gd 1 – ðkd Þ2 3 6
½20
The corresponding equation of motion is qu qh 1 q3 u ¼ –g þ d 2 2 qt qx 3 qx qt
½21
and the equation of continuity is qh qu ¼ d qt qx
½22
4.17.6.7
Numerical Computations
The tsunami propagation depends on the water depth, which is not uniform but the variation is relatively well known compared to velocity distribution of seismic waves. Hence, numerical computations on realistic or actual bathymetry have been popularly made. The historical review of tsunami numerical simulation is presented by Shuto (1991). Let us consider a simple one-dimensional propagation (on constant depth) of linear long waves. Denoting the average velocity (x-component only) as U, the equations of motion and continuity are given as follows: qU qh ¼ –g qt qx
½25
qh qU ¼ –d qt qx
½26
Using the staggered (leap-frog) grid system for U and h, and denoting time t ¼ l t as superscripts and space x ¼ m x as subscripts where l and m are integers, these become h i lþð1=2Þ
l l hmþ Um – Uml – ð1=2Þ ð1=2Þ – hm – ð1=2Þ ¼ –g x t h i h i lþð1=2Þ lþ1 l hmþ Umþ1 – Umlþð1=2Þ ð1=2Þ – hmþð1=2Þ ¼ –d t x
½27
½28
Tsunamis
It can be shown that the error will grow with time unless the following stability condition is met: pffiffiffiffiffi t x 1 or t pffiffiffiffiffi gd x gd
x t pffiffiffiffiffiffiffi 2gd
U ði,it þ 1Þ ¼ U ði,itÞ – g
hði,it þ 1Þ ¼hði,itÞ t –d ½U ðiþ1,it þ 1Þ – U ði,it þ 1Þ x
The algorithm thus is as follows: set up parameters read bathymetry data read initial condition h (i, 1) while (it
h (i )
½31
½32
h (i + 1)
t + Δt
t x U(i – 1)
U(i )
U(i + 1)
Figure 12 Staggered leap-frog system.
compute U (i, itþ1) from U (i, it) and h (i, it) set U (i, itþ1) ¼ 0 for land boundary compute h (i, itþ1) from U (i, itþ1) and h (i, it) adjust h (i, itþ1) for open boundary exchange variables h (i, itþ1)!h (i, it) U (i, itþ1)!U (i, it) itþ¼ 1
½30
There are two types of boundaries involved in the tsunami computation. The first one is land–ocean boundary. The simplest assumption is a total reflection of energy on the coast. For the computation of tsunami inundation on land, however, a moving land–ocean boundary is needed; the boundary moves depending on the relative heights of water height and land. The other kind of boundary is the open ocean boundary along the outside of the computational region. A typical assumption is the radiation condition, which means that waves go out of the computational region keeping their slopes. The computer program for one-dimensional finite-difference computation is very simple. Since there is no overlap between the two (height and velocity) grids (Figure 12) in the linear case, in the program the spatial grid points m and m þ (1/2) can be treated by the same index (e.g., i) and the time l – 1/2 and l can be treated as the same time (e.g., it). Writing U and h using the two-dimensional array as U (i, it) and h(i, it), [27] and [28] become t ½hði,itÞ x – hði – 1,itÞ
t
½29
which is known as Courant–Friedrichs–Lewy (CFL) condition. The physical meaning of the stability condition is that the time step t must be equal to or smaller than the time required for the disturbance to travel the spatial grid size x. If t is larger, the values other than (l, m – l), (l – 1, m) and (l, m þ 1) will influence the value at (l þ1, m). For the two-dimensional case, the CFL condition becomes
501
end Figure 13 shows the snapshots of tsunami propagation from the 2004 Sumatra–Andaman earthquake.
4.17.7 Tsunami Warning and Hazard Reduction Systems The December 2004 tsunami in the Indian Ocean caused the worst tsunami disaster in the history. In addition to the gigantic size of the earthquake, lack of preparation for tsunamis on the coast of Indian Ocean was also responsible for the tragedy. Preparation for tsunami includes a tsunami warning system based on seismic and sea level monitoring, hazard assessment such as making hazard maps, and education and awareness to the coastal residents. Such systems on tsunami hazard reduction rely on the scientific knowledge of earthquakes and tsunamis. The US National Tsunami Hazard Mitigation Program (Bernard, 2005) is a comprehensive program based on three components: warning guidance, hazard assessment, and mitigation (Figure 14). Tsunami warning systems can be grouped into a basin-wide (such as Pacific or Indian Ocean), regional, and local systems. Hazard assessment includes making inundation maps or probabilistic maps. Mitigation includes
502 Tsunamis
Phuket Thailand
20° subsided 10° uplifted 0°
Sri Lanka Sumatra
–10°
0h
70°
1h 80°
90°
100°
2h
110°
Wa rn
on
ing
ati
tig
gu
Mi
ida
nc
e
Figure 13 Snapshots of simulation for the 2004 tsunami in Indian Ocean.
Hilo Bay
Hazard assessment Figure 14 Three key elements for tsunami hazard reduction: warning guidance, hazard assessment, and mitigation (Bernard, 2005).
public awareness such as implementing tsunami signs or educating coastal residents. 4.17.7.1 Far-Field Tsunami Warning Systems Tsunamis travel very fast as an ocean wave, at about 710 km h1, or 0.2 km s 1, for a water depth of 4000 m, yet much slower than seismic waves, which travel at 5–10 km s1. This velocity difference between the seismic and tsunami waves makes it possible to issue a tsunami warning by detecting seismic waves before the actual tsunami arrival. For distant tsunamis that propagate across the Pacific or Indian Ocean, there are a few hours for accurate and reliable warning. The Richard H. Hagemeyer Pacific Tsunami Warning Center (PTWC), located in Hawaii, is the operational center for the international tsunami warning system, and issues warnings for regional and distant tsunamis
in the Pacific Basin to almost all the countries and states around the Pacific Ocean (McCreery, 2005). The West Coast/Alaska Tsunami Warning Center, located in Alaska, is responsible for tsunami warning along the US and Canadian coasts except for Hawaii. At these centers, seismic waveform data from several tens of stations are continuously monitored, and once a large earthquake is detected, the location, depth, and magnitudes are assessed using a combination of automatic and interactive procedures by watchstanders (Figure 15). Recent deployment of advanced seismological parameters such as Mm (Okal and Talandier, 1989), Mwp(Tsuboi et al., 1995), or Y (Newman and Okal, 1998) makes it possible to quickly assess the earthquake size and tsunami potential, and to issue the tsunami warnings in less than 25 min. To avoid false alarm, it is very important to confirm that a tsunami was actually generated; hence, sea levels are monitored with both costal tide gauges and tsunameters (DART buoys). If a tsunami is observed, then the tsunami warning is updated (Titov et al., 2005b). If the tsunami cannot be confirmed, then the warning is canceled. At the time of 26 December 2004 tsunami, PTWC issued the first information bulletin at 1:14 GMT, only 15 min after the earthquake. An earthquake was located off the west coast of Northern Sumatra, but the magnitude (Mwp) was estimated as 8.0. The second bulletin was issued at 2:08 GMT, 69 min after the earthquake yet before the tsunami arrivals at Thai, Sri Lankan, or Indian coast. The earthquake size (Mm) was updated to 8.5 and a possibility of local tsunami was mentioned in the bulletin. The French Polynesia Tsunami Warning Center developed and adopted TREMORS (Tsunami Risk Evaluation through seismic Moment in a Real time System) in 1987 (Reymond et al., 1991). TREMORS
Tsunamis
503
90° N 60° N 30° N PTWC
0° 30° S Seismic Sea level DART
60° S 90° S 0°
60° E
120° E
180°
120° W
60° W
0°
Figure 15 Distribution of seismic stations and sea level monitoring, coastal gauges, and tsunameter (DART) stations that send the data to PTWC in real time.
features automatic detection and location of seismic events using a single three-component long-period seismic station, and estimation of seismic moment through variable-period mantle magnitude Mm (Okal and Talandier, 1989). The advantage of using long-period seismic waves for tsunami warning is the ability to detect ‘tsunami earthquakes’. The disadvantage is that longer time is needed before the seismic waves are recorded. TREMORS is a hybrid system; it updates the moment estimate as more seismic data are available. The Pacific-wide tsunami warning system was established after the 1946 Aleutian tsunami that affected the Hawaiian Islands (Section 4.17.3.1.5). After the 1960 Chilean tsunami (Section 4.17.3.1.3), the Pacific tsunami warning system became international, and an International Coordination Group for The Tsunami Warning system in The Pacific (ICG/ ITSU) was formed under the auspices of UNESCO Intergovernmental Oceanographic Commission. After the 2004 tsunami in Indian Ocean, it was renamed as ICG/PTWS (ICG for the Pacific Tsunami Warning and Mitigation System), and similar groups were formed for the other ocean basins: ICG/IOTWS for Indian Ocean, ICG/NEAMTWS for the Northeastern Atlantic and the Mediterranean Sea, and ICG/CARIBE-EWS for the Caribbean Sea.
4.17.7.2
Local Tsunami Warning Systems
For local tsunamis, time is more critical. Japan Meteorological Agency ( JMA) is responsible for issuing tsunami warnings in Japan. JMA uses seismic data from a few hundreds of stations to detect tsunamigenic earthquakes, and sea level data to confirm
the tsunami generation. There are six regional tsunami warning centers and each of them is staffed with multiple seismologists on duty for 24 h a day and 7 days a week. When an earthquake is detected on the seismic network, the location and size (magnitude) are automatically estimated within a few minutes. Likelihood of tsunami generation is judged from the location and magnitude, and the staff seismologists issue tsunami warning or advisory. The warning or advisory, as well as other related information such as the earthquake location, size, and the expected tsunami arrival times, is disseminated by using the Geostationary Meteorological Satellite. In the 1950s when JMA first introduced the tsunami warning system, it took about 20 min to issue tsunami warnings (Figure 16). In 1983, when a large earthquake occurred in Japan Sea, the warning was issued in 12 min after the earthquake, but the tsunami arrived in 7 min after the earthquake and 100 people died. JMA had improved the warning system, and in 1993 the tsunami warning was issued in 5 min after the South West Hokkaido earthquake. Even though, the tsunami arrived in less than 5 min and the 230 people died (Section 4.17.3.1.2). JMA further improved the system and now issues tsunami warning within 3–5 min after a large earthquake. It is also very important to improve the accuracy of tsunami forecast and to reduce false alarm. JMA introduced a numerical simulation technique in April 1999 (Tatehata, 1997); tsunami generation and propagation for 100 000 different cases were calculated in advance, and the results have been stored as a database. When a large earthquake occurs, the most appropriate case for an actual location and a magnitude of the earthquake is retrieved from the database.
504 Tsunamis
30
Elapsed time (min)
25 20 15 10 5 0 1952 1956 1960 1964 1968 1972 1976 1980 1984 1988 1992 1996 2000 2004 Year Figure 16 Improvements in elapsed time between earthquakes and tsunami warning issued by Japan Meteorological Agency.
A similar system based on database has been developed for the Pacific Ocean (Titov et al., 2005b). After a tsunami warning is issued, the information must be disseminated to coastal residents. In Japan, coastal residents receive warning through media, TV and radio, emergency broadcast systems, or patrol cars from police or fire stations. Periodic practice and drills are very important to keep the system functional. In Japan, annual national drills are conducted on 1 September, national disaster prevention day. In May 2006, The ICG/PTWS planned and conducted the first Pacific-wide exercise on issuing, disseminating, and responding to tsunami warning messages.
4.17.7.3 Long-Term and Probabilistic Forecast The tsunami hazard probabilities can be estimated from historical data. If we know that a particular coast has experienced N tsunamis over a period of time T, assuming a Poisson process, the probability P of having at least one tsunami for future time period is given as P ¼ 1 – exp ð – N =T Þ
½33
For example, if two tsunamis were recorded in the last 100 years at a particular location, the probability of having another tsunami in the next 10 years is 18%. The 10-year probability is 63% if 10 tsunamis occurred in the last 100 years. If the water heights from past tsunamis are known, the probabilities for each range of tsunami heights can be similarly computed. This method is useful for assessing tsunamis from various origins such as trans-Pacific tsunamis.
Great earthquakes at the plate boundaries may recur more regularly than Poisson process, although it is controversial. When the cyclic nature of past earthquakes is known, it can be used to calculate the probability of earthquake occurrence. The time interval between earthquakes can be modeled using a probability density function such as Weibull or log–normal distribution (Rikitake, 1976; Utsu, 1984). The Brownian passagetime (BPT) distribution has been recently used, because it models steady tectonic loading plus Brownian perturbation (Matthews et al., 2002). Then the probability of next earthquake in the specific time window can be computed if the average recurrence period and the date of the most recent event, as well as aperiodicity parameter or variation coefficient, are known. Probabilistic earthquake forecast had some success in Japan. In March 2003, the Japanese government made a long-term forecast for great (M 8) earthquakes along the Kuril trench. The committee estimated that the probability in the next 30 years (starting March 2003) is 60% in Tokachi-oki and 20–30% in Nemuro-oki. This forecast was based on the mean inter-event time of great earthquakes for the southern Kuril region in the nineteenth and twentieth centuries (Earthquake Research Committee, 2004; Satake et al., 2005b). Six months later, on 26 September 2003, a great Tokachi-oki earthquake (M 8.0) actually occurred. When the fault parameters of the characteristic tsunamigenic earthquake are known, tsunami heights can be estimated by numerical computations, and the probabilistic tsunami heights can be calculated. Such a probabilistic estimate of tsunami heights was made in the 1980s for the Japanese coasts (Rikitake and
Tsunamis
P (≥ 0.5 m) P (≥ 1 m)
P (≥ 2 m)
505
P (≥ 5 m) P (≥ 7 m) P (≥ 10 m)
15 7
7 27
Kushiro
7 15
15
7
7
9 9
Pt. Erimo
Hokkaido 22
Tomakomai
9 9
Muroran
100% Figure 17 Probabilities for the Pacific coast of Hokkaido hit by a tsunami with wave height exceeding 0.5, 1, 2, 5, 7, and 10 m during a period from 2000 to 2010. The upper half of the coast was hit by a tsunami with 2–4 m in 2003. From Rikitake T and Aida I (1988) Tsunami hazard probability in Japan. Bulletin of Seismological Society of America 78: 1268–1278.
Aida, 1988). For the Pacific coast of central Japan where a large earthquake (Tokai earthquake) is expected, the probability of having 5 m tsunami for the period between 2000 and 2010 is as high as 41%. For the Pacific coast of Hokkaido along the Kuril Trench, where the 2003 Tokachi-oki earthquake actually generated the tsunami, the probability of 2 m tsunami was estimated as 7–15% and that of 5 m tsunami was 7% (Figure 17). The actual tsunami heights in 2003 were up to 4 m (Tanioka et al., 2004), but the coastal run-up and the tsunami source were different from those of the previous Tokachi-oki earthquake in 1952 (Satake et al., 2006b). Similar probabilistic estimates have also been made for the US coast (Geist and Parsons, 2006). 4.17.7.4 Tsunami Hazard Maps and Public Awareness Once the coastal residents receive a tsunami warning message, they need to know what it means and where to evacuate. An effective tool for the public is a map showing area to evacuate for a future tsunami. Possible inundation zones are estimated either from historical data of past tsunamis or by tsunami numerical simulation based on maximum credible earthquakes. Safe places such as tsunami shelters to evacuate can also be shown in the tsunami hazard maps. Tsunami hazard map help coastal communities prepare for tsunami hazards.
On the basis of tsunami hazard maps, evacuation plans can be established by responsible officials such as civil defense officials (Gonzalez et al., 2005). One of the lessons from the 2004 Asian tsunami is that not only coastal residents but also foreign tourists need to be informed on tsunami hazards. For the Hawaiian Islands, tsunami hazard maps are prepared for most of the occupied coasts and published in the local phone book that are available at every hotel room (Figure 18). The maps show, in addition to the inundation area and places to evacuate, evacuation instructions. For example, those in high-rise buildings are instructed to move vertically to the third or higher floors, rather than moving out of the area. Tsunami hazard assessment has been carried out on the Japanese coasts. Numerical simulations were used to estimate the time and heights of tsunami and possible damage including casualties. For example, for the future Tonankai-Nankai earthquake off the western Japan, inundation of 40 000 houses and 2200–8600 casualties are estimated (Central Disaster Management Council, 2006). For public awareness, booklets and video on tsunamis have been prepared and used to educate coastal residents. In Japan, a famous story of a village chief who evacuated the entire residents, by burning his harvest at the time of the 1854 Nankai earthquake tsunami (known as ‘‘Inamura-no-Hi’’), has been popularly used for hazard education.
506 Tsunamis
Figure 18 Tsunami evacuation maps for Waikiki beach of Honolulu, Hawaii. The maps were prepared by University of Hawaii in cooperation with Hawaii State Civil Defense Agency and published in local phone books.
4.17.8 Summary Study of tsunami has interdisciplinary natures as summarized in the chapter. The tsunami sources include giant earthquake such as the 2004 Sumatra– Andaman earthquake or unusual ‘tsunami earthquakes’ for which more seismological studies are needed. Compared to seismic sources, generation mechanism of tsunamis from landslides or volcanic eruptions is less understood. While the tsunami propagation in deep ocean has been treated by shallowwater theory and numerical computations have been made popularly, the actual propagation in deep ocean was captured by the satellite for the first time at the 2004 Indian Ocean tsunami. Similar deep-sea observations in the future would provide chance to examine the propagation characteristics such as dispersion effects. The coastal behavior of tsunamis will need more studies on theory, field-data, and physical experiments. Geological studies of past tsunamis are also important, as tsunami is one of infrequent phenomena. The 2004 tsunami in the Indian Ocean demonstrated that the effect of tsunami can be global. The tsunami damage was extended through the Indian Ocean, and the tsunami was recorded throughout the world. In addition, the victims of the 2004 tsunami were not only limited to the coastal residents in Asian countries but also included tourists from Europe. For the reduction of future tsunami
damage, it is essential to educate the global citizens about tsunamis. After the 1960 Chilean tsunami, our scientific knowledge on tsunami has significantly improved and we can now detect the occurrence of earthquake and assess the possibility of tsunamis in real time. Such scientific and technological development can be and should be used to save human lives in the world.
References Abe K (1973) Tsunami and mechanism of great earthquakes. Physics of the Earth and Planetary Interiors 7: 143–153. Abe K (1979) Size of great earthquakes of 1873–1974 inferred from tsunami data. Journal of Geophysical Research 84: 1561–1568. Abe K (1981) Physical size of tsunamigenic earthquakes of the northwestern Pacific. Physics of the Earth and Planetary Interiors 27: 194–205. Abe K (1989) Quantification of tsunamigenic earthquakes by the Mt scale. Tectonophysics 166: 27–34. Abe K (2000) Predominance of long periods in large Pacific tsunamis. Science of Tsunami Hazards 18: 15–34. Aida I (1978) Reliability of a tsunami source model derived from fault parameters. Journal of Physics of the Earth 26: 57–73. Ammon CJ, Ji C, Thio HK, et al. (2005) Rupture process of the 2004 Sumatra–Andaman earthquake. Science 308: 1133–1139. Atwater BF, Musumi-Rokkaku S, Satake K, Tsuji Y, Ueda K, and Yamaguchi DK (2005) The Orphan Tsunami of 1700. USGS Professional Paper 1707: 133 pp. Baptista MA, Heitor S, Miranda JM, Miranda P, and Victor LM (1998) The 1755 Lisbon tsunami; Evaluation of the tsunami parameters. Journal of Geodynamics 25: 143–157.
Tsunamis Bardet JP, Synolakis CE, Davies HL, Imamura F, and Okal EA (2003) Landslide tsunamis: Recent findings and research directions. Pure and Applied Geophysics 160: 1793–1809. Bernard EN (2005) The US National Tsunami Hazard Mitigation Program: A successful State–Federal partnership. Natural Hazards 35: 5–24. Berninghausen WH (1966) Tsunamis and seismic seiches reported from regions adjacent to the Indian Oceans. Bulletin of the Seismological Society of America 56: 69–74. Bilek SL and Lay T (2002) Tsunami earthquakes possibly widespread manifestations of frictional conditional stability. Geophysical Research Letters 29: doi:10.1029/2002GL015215. Bilham R (2005) A flying start, then a slow slip. Science 308: 1126–1127. Bondevik S, Lovholt F, Harbitz C, Mangerud J, Dawson A, and Svendsen JI (2005) The Storegga Slide tsunami – comparing field observations with numerical simulations. Marine and Petroleum Geology 22: 195–208. Borrero JC (2005a) Field survey of Northern Sumatra and Banda Aceh, Indonesia after the tsunami and earthquake of 26 December 2004. Seismological Research Letters 76: 312–320. Borrero JC (2005b) Field data and satellite imagery of tsunami effects in Banda Aceh. Science 308: 1596. Bryant EA, Young RW, and Price DM (1996) Tsunami as a major control on coastal evolution, southeastern Australia. Journal of Coastal Research 12: 831–840. Bugge T, Belderson RH, and Kenyon NH (1988) The Storegga slide. Philosophical Transactions of the Royal Society of London Series A 325: 357–388. Central Disaster Management Council (2006) Hazard Estimation for Earthquakes Along Japan and Kuril Trenches (in Japanese). Tokyo: Cabinets Office. Comer RP (1980) Tsunami height and earthquake magnitude: theoretical basis of an empirical relation. Geophysical Research Letters 7: 445–448. Comer RP (1984a) The tsunami mode of a flat earth and its excitation by earthquake sources. Geophysical Journal of the Royal Astronomical Society 77: 1–27. Comer RP (1984b) Tsunami generation: a comparison of traditional and normal mode approaches. Geophysical Journal of the Royal Astronomical Society 77: 29–41. Cummins P (2004) Small threat, but warning sounded for tsunami research. AusGeo News 75: 4–7. Dawson AG, Long D, and Smith DE (1988) The Storegga Slides – Evidence from eastern Scotland for a possible tsunami. Marine Geology 82: 271–276. Dawson AG and Shi SZ (2000) Tsunami deposits. Pure and Applied Geophysics 157: 875–897. Earthquake Research Committee (2004) Long-term evaluation of seismicity along the Kuril Trench (in Japanese). Publications of Earthquake Research Committee, pp. 1–74 . Tokyo: Earthquake Research Committee. Fine, IV, Rabinovich AB, Bornhold BD, Thomson RE, and Kulikov EA (2005a) The Grand Banks landslide-generated tsunami of November 18, 1929: preliminary analysis and numerical modeling. Marine Geology 215: 45–57. Fine, IV, Rabinovich AB, and Thomson RE (2005b) The dual source region for the 2004 Sumatra tsunami. Geophysical Research Letters 32: L16602, doi: 10.1029/2005GL023521. Fujii Y and Satake K (2007) Tsunami source model of the 2004 Sumatra–Andaman earthquake inferred from tide gauge and satellite data. Bulletin of the Seismological Society of America 97: S192–S207. Fukao Y (1979) Tsunami earthquakes and subduction processes near deep-sea trenches. Journal of Geophysical Research 84: 2303–2314. Geist E (1998) Local tsunamis and earthquake source parameters. Advances in Geophysics 39: 117–209.
507
Geist E (2000) Origin of the 17 July 1998 Papua New Guinea tsunami; earthquake or landslide? Seismological Research Letters 71: 344–351. Geist E and Parsons T (2006) Probabilistic analysis of tsunami hazards. Natural Hazards 37: 277–314. Geist E, Titov VV, Arcas D, Pollitz FF, and Bilek SL (2007) Implications of the December 26, 2004 Sumatra–Andaman Earthquake on Tsunami Forecast and Assessment Models for Great Subduction Zone Earthquakes. Bulletin of the Seismological Society of America 97: S249–S270. Gelfenbaum G and Jaffe B (2003) Erosion and sedimentation from the 17 July, 1998 Papua New Guinea tsunami. Pure and Applied Geophysics 160: 1969–1999. Gonzalez FI, Bernard EN, Meinig C, Eble MC, Mofjeld HO, and Stalin S (2005) The NTHMP tsunameter network. Natural Hazards 35: 25–39. Gower J (2005) Jason 1 detects the 26 December 2004 tsunami. Eos, Transactions, American Geophysical Union 86: 37–38. Gusiakov VK (2005) Tsunami generation potential of different tsunamigenic regions in the Pacific. Marine Geology 215: 3–9. Hanks T and Kanamori H (1979) A moment magnitude scale. Journal of Geophysical Research 84: 2348–2350. Hanson JF, Reasoner C, and Bowman JR (2007) High frequency tsunami signals of the great Indonesian earthquakes of 26 December and 28 March 2005. Bulletin of the Seismological Society of America 97: S232–S248. Harbitz CB (1992) Model simulations of tsunamis generated by the Storegga slides. Marine Geology 105: 1–21. Hatori T (1979) Relation between tsunami magnitude and wave energy. Bulletin of the Earthquake Research Institute (University of Tokyo) 54: 531–541. Hatori T (1984) On the damage to houses due to tsunamis. Bulletin of the Earthquake Research Institute (University of Tokyo) 59: 433–439. Heezen BC and Ewing M (1952) Turbidity currents and submarine slumps, and the 1929 Grand Banks earthquake. American Journal of Science 250: 849–873. Heinrich P, Piatanesi A, and Hebert H (2001) Numerical modelling of tsunami generation and propagation from submarine slumps: the 1998 Papua New Guinea event. Geophysical Journal International 145: 97–111. Hirata K, Aoyagi M, Mikada H, et al. (2002) Real-time geophysical measurements on the deep seafloor using submarine cable in the southern Kurile subduction zone. IEEE Journal of Oceanic Engineering 27: 170–181. Hirata K, Satake K, Tanioka Y, et al. (2006) The 2004 Indian Ocean tsunami: Tsunami source model from satellite altimetry. Earth, Planets and Space 58: 195–201. Iida K (1984) Catalog of Tsunamis in Japan and Its Neighboring Countries, 52pp. Toyato: Aichi Institute of Technology. Iida K, Cox DC, and Pararas-Carayannis G (1967) Preliminary catalog of tsunamis occurring in the Pacific Ocean. Honolulu: University of Hawaii, 274pp. Imamura F, Hashi K, and Imteaz MMA (2001) Modeling for tsunamis generated by landsliding and debris flow. In: Hebenstreit GT (ed.) Tsunami Research at the End of a Critical Decade,, pp. 209–228. Dordrecht: Kluwer Academic. Institute of Computational Mathematics and Mathematical Geophysics (2005) Integrated Tsunami Database for the World Ocean, Novosibirsk, CD-ROM. Intergovernmental Oceanographic Commission (1998) Manuals and Guides 37: Post-tsunami Survey Field Guide. Paris: UNESCO. International Federation of Red Cross and Red Crescent Societies (2005) World Disasters Report, 251pp. Geneva: International Federation of Red Cross and Red Crescent Societies. Ishii M, Shearer PM, Houston H, and Vidale JE (2005) Extent, duration and speed of the 2004 Sumatra–Andaman
508 Tsunamis earthquake imaged by the Hi-Net array. Nature 435: 933–936. Iwasaki SI, Eguchi T, Fujinawa Y, et al. (1997) Precise tsunami observation system in the deep ocean by an ocean bottom cable network for the prediction of earthquakes and tsunamis. In: Hebenstreit G (ed.) Perspectives on Tsunami Hazard Reduction, pp. 47–66. Dordrecht: Kluwer Academic. Jaffe BE, Borrero JC, Prasetya GS, et al. (2006) Northwest Sumatra and offshore islands field survey after the December 2004 Indian Ocean tsunami. Earthquake Spectra 22: S105–S135. Jiang L and Leblond PH (1994) 3-Dimensional modeling of tsunami generation due to a submarine mudslide. Journal of Physical Oceanography 24: 559–572. Johnson JM (1999) Heterogeneous coupling along AlaskaAleutians as inferred from tsunami, seismic and geodetic inversions. Advances in Geophysics 39: 1–116. Johnson JM and Satake K (1997) Estimation of seismic moment and slip distribution of the April 1, 1946, Aleutian tsunami earthquake. Journal of Geophysial Research 102: 11765–11774. Kajiura K (1963) The leading wave of a tsunami. Bulletin of the Earthquake Research Institute (University of Tokyo) 41: 535–571. Kajiura K (1970) Tsunami source, energy and the directivity of wave radiation. Bulletin of the Earthquake Research Institute (University of Tokyo) 48: 835–869. Kanamori H (1972) Mechanism of tsunami earthquakes. Physics of the Earth and Planetary Interiors 6: 246–259. Kanamori H and Kikuchi M (1993) The 1992 Nicaragua earthquake: A slow tsunami earthquake associated with subducted sediments. Nature 361: 714–716. Kato T, Terada Y, Ito K, et al. (2005) Tsunami due to the 2004 September 5th off Kii peninsula earthquake, Japan, recorded by a new GPS buoy. Earth, Planets and Space 57: 297–301. Kato T, Terada Y, Kinoshita M, et al. (2000) Real-time observation of tsunami by RTK-GPS. Earth, Planets and Space 52: 841–845. Kawata Y, Benson B, Borrero JC, et al. (1999) Tsunami in Papua New Guinea was as intense as first thought. Eos, Transactions, American Geophysical Union 80: 101–105. Kayanne H, Ikeda Y, Echigo T, et al. (2007) Coseismic and postseismic creep in Andaman Islands associated with the 2004 Sumatra–Andaman earthquake. Geophysical Research Letters 34: L01310, doi: 10.129/2006GL028200. Keating BH, Waythomas CF, and Dawson AG (2000) Special issue: Landslides and Tsunamis–Introduction. Pure and Applied Geophysics 157: 871–873. Kulikov EA, Rabinovich AB, Thomson RE, and Bornhold BD (1996) The landslide tsunami of November 3, 1994, Skagway Harbor, Alaska. Journal of Geophysical Research 101: 6609–6615. Lander JF (1996) Tsunamis Affecting Alaska: 1737–1996. NGDC Key to Geophysical Research Documentation No. 29, Boulder, CO. Lander JF and Lockridge PA (1989) United States Tsunamis. Boulder, CO: National Geophysical Data Center. Lander JF, Lockridge PA, and Kozuch MJ (1993) Tsunamis Affecting the West Coast of the United States 1806–1992. Boulder, CO: National Geophysical Data Center. Latter JH (1981) Tsunamis of volcanic origin: summary of causes, with particular reference to Krakatoa, 1883. Bulletin of Volcanology 44: 467–490. Lay T, Kanamori H, Ammon CJ, et al. (2005) The great SumatraAndaman earthquake of 26 December 2004. Science 308: 1127–1133. Liu PLF, Lynett P, Fernando H, et al. (2005) Observations by the International Tsunami Survey Team in Sri Lanka. Science 308: 1595–1595.
Malik JN and Murty CVR (2005) Landscape changes in Andaman and Nicobar Islands (India) due to Mw 9.3 tsunamigenic Sumatra earthquake of 26 December 2004. Current Science 88: 1384–1386. Mansinha L and Smylie DE (1971) The displacement fields of inclined faults. Bulletin of the Seismological Society of America 61: 1433–1440. Matsuyama M, Walsh JP, and Yeh H (1999) The effect of bathymetry on tsunami characteristics at Sisano Lagoon, Papua New Guinea. Geophysical Research Letters 26: 3513–3516. Matthews MV, Ellsworth WL, and Reasenberg PA (2002) A Brownian model for recurrent earthquakes. Bulletin of the Seismological Society of America 92: 2233–2250. McCreery CS (2005) Impact of the National Tsunami Hazard Mitigation Program on operations of the Richard H. Hagemeyer Pacific Tsunami Warning Center. Natural Hazards 35: 73–88. Mei CC (1989) The Applied Dynamics of Ocean Surface Waves. Advances Series on Ocean Engineering, volume 1. Singapore: World Scientific. Meltzner AJ, Sieh K, Abrams M, et al. (2006) Uplift and subsidence associated with the great Aceh–Andaman earthquake of 2004. Journal of Geophysical Research 111: B02407 (doi: 10.1029/2005JB003891). Merrifield MA, Firing YL, Aarup T, et al. (2005) Tide gauge observations of the Indian Ocean tsunami, December 26, 2004. Geophysical Research Letters 32, doi: 10.1029/ 2005GL022610. Miller DJ (1960) Giant waves in Lituya Bay, Alaska. USGS Professional Paper 354-C: 51–86. Minoura K and Nakaya S (1991) Traces of tsunami preserved in inter-tidal lacustrine and marsh deposits - Some examples from northeast Japan. The Journal of Geology 99: 265–287. Miyabe N (1934) An investigation of the Sanriku tunami based on mareogram data. Bulletin of the Earthquake Research Institute (University of Tokyo) 1: 112–126. Moore A, Nishimura Y, Gelfenbaum G, Kamataki T, and Triyono R (2006) Sedimentary deposits of the 26 December 2004 tsunami on the northwest. Earth, Planets and Space 58: 253–258. Murty TS and Rafiq M (1991) A tentative list of tsnamis in the marginal seas of the North Indian Ocean. Natural Hazards 4: 81–83. Nagai T, Ogawa H, Nukada K, and Kudaka M (2004) Characteristics of the observed 2003 Tokachi-oki earthquake tsunami profile. Coastal Engineering Journal 46: 315–327. Nagarajan B, Suresh I, Sundar D, et al. (2006) The Great Tsunami of 26 December 2004: A description based on tidegauge data from the Indian subcontinent and surrounding areas. Earth, Planets and Space 58: 211–215. Nanayama F, Satake K, Furukawa R, et al. (2003) Unusually large earthquakes inferred from tsunami deposits along the Kuril trench. Nature 424: 660–663. Nanayama F, Shigeno K, Satake K, et al. (2000) Sedimentary differences between the 1993 Hokkaido-nansei-oki tsunami and the 1959 Miyakojima typhoon at Taisei, southwestern Hokkaido, northern Japan. Sedimentary Geology 135: 255–264. Neetu S, Suresh I, Shankar R, et al. (2005) Comment on ‘‘The Great Sumatra-Andaman Earthquake of 26 December 2004’’. Science 310: 1431. Newman AV and Okal EA (1998) Teleseismic estimates of radiated seismic energy: The E/M-0 discriminant for tsunami earthquakes. Journal of Geophysical Research 103: 26885–26898. Nishimura Y and Miyaji N (1995) Tsunami deposits from the 1993 Southwest Hokkaido earthquake and the 1640
Tsunamis Hokkaido-Komagatake eruption, northern Japan. Pure and Applied Geophysics 144: 719–733. Nomanbhoy N and Satake K (1995) Generation mechanism of tsunamis from the 1883 Krakatau eruption. Geophysical Research Letters 22: 509–512. Okada M (1995) Tsunami observation by ocean bottom pressure gauge. In: Tsuchiya Y, Shuto N (eds.) Tsunami: Progress in Prediction, Disaster Prevention and Warning, pp. 287–303. Dordrecht: Kluwer Academic. Okada Y (1985) Surface deformation due to shear and tensile faults in a half-space. Bulletin of the Seismological Society of America 75: 1135–1154. Okal EA (1988) Seismic parameters controlling far-field tsunami amplitudes: A review. Natural Hazards 1: 67–96. Okal EA and Newman AV (2001) Tsunami earthquakes: The quest for a regional signal. Physics of the Earth and Planetary Interiors 124: 45–70. Okal EA, Plafker G, Synolakis CE, and Borrero JC (2003) Nearfield survey of the 1946 Aleutian Tsunami on Unimak and Sanak Islands. Bulletin of the Seismological Society of America 93: 1226–1234. Okal EA and Synolakis CE (2004) Source discriminants for nearfield tsunamis. Geophysical Journal International 158: 899–912. Okal EA, Synolakis CE, Fryer GJ, et al. (2002) A field survey of the 1946 Aleutian tsunami in the far field. Seismological Research Letters 73: 490–503. Okal EA and Talandier J (1989) Mm: A variable-period mantle magnitude. Journal of Geophysical Research 94: 4169–4193. Ortiz M and Bilham R (2003) Source area and rupture parameters of the 31 December 1881 M-w ¼ 7.9 Car Nicobar earthquake estimated from tsunamis recorded in the Bay of Bengal. Journal of Geophysical Research 108: (doi: 10.129/ 2002JB001941). Pelayo AM and Wiens DA (1992) Tsunami earthquakes: slow thrust-faulting events in the accretionary wedge. Journal of Geophysical Research 97: 15321–15337. Pelinovsky E, Choi BH, Stromkov A, Didenkulova I, and Kim H-S (2005) Analysis of tide-gauge records of the 1883 Krakatau tsunami. In: Satake K (ed.) Tsunamis: Case Studies and Recent Developments,, pp. 57–78. Dordrecht: Springer. Peregrine H (1972) Equations for water waves and the approximations behind them. In: Meyer RE (ed.) Waves on Beaches and Resulting Sediment Transport, pp. 95–121. Academic Press. Piatanesi A and Lorito S (2007) Rupture process of the 2004 Sumatra–Andaman earthquake from tsunami waveform inversion. Bulletin of the Seismological Society of America 97: S223–S231. Plafker G, Kachadoorian R, Eckel EB, and Mayo LR (1969) Effects of the earthquake of March 27, 1964 on various communities. USGS Professional Paper 542-G: 1–50. Polet J and Kanamori H (2000) Shallow subduction zone earthquakes and their tsunamigenic potential. Geophysical Journal International 142: 684–702. Press F and Harkrider D (1966) Air-sea waves from the explosion of Krakatoa. Science 154: 1325–1327. Reymond D, Hyvernaud O, and Talandier J (1991) Automatic detection, location and quantification of earthquakes: application to tsunami warning. Pure and Applied Geophysics 135: 361–382. Rikitake T (1976) Recurrence of great earthquakes at subduction zones. Tectonophysics 35: 335–362. Rikitake T and Aida I (1988) Tsunami hazard probability in Japan. Bulletin of the Seismological Society of America 78: 1268–1278. Satake K (1988) Effects of Bathymetry on Tsunami Propagation - Application of Ray Tracing to Tsunamis. Pure and Applied Geophysics 126: 27–36.
509
Satake K (1989) Inversion of tsunami waveforms for the estimation of heterogeneous fault motion of large submarine earthquakes – the 1968 Tokachi-Oki and 1983 Japan Sea earthquakes. Journal of Geophysical Research 94: 5627–5636. Satake K (1993) Depth distribution of coseismic slip along the Nankai Trough, Japan, from joint inversion of geodetic and tsunami data. Journal of Geophysical Research 98: 4553–4565. Satake K (1994) Mechanism of the 1992 Nicaragua tsunami earthquake. Geophysical Research Letters 21: 2519–2522. Satake K (1995) Linear and nonlinear computations of the 1992 Nicaragua earthquake tsunami. Pure and Applied Geophysics 144: 455–470. Satake K, Aung TT, Sawai Y, et al. (2006a) Tsunami heights and damage along the Myanmar coast from the December 2004 Sumatra-Andaman earthquake. Earth, Planets and Space 58: 243–252. Satake K, Baba T, Hirata K, et al. (2005a) Tsunami source of the 2004 off the Kii peninsula earthquakes inferred from offshore tsunami and coastal tide gauges. Earth, Planets and Space 57: 173–178. Satake K, Hirata K, Yamaki S, and Tanioka Y (2006b) Re-estimation of tsunami source of the 1952 Tokachi-oki earthquake. Earth, Planets and Space 58: 535–542. Satake K, Nanayama F, Yamaki S, Tanioka Y, and Hirata K (2005b) Variability among tsunami sources in the 17th–21st centuries along the southern Kuril trench. In: Satake K (ed.) Tsunamis: Case Studies and Recent Developments, pp. 157–170. Dordrecht: Springer. Satake K, Okada M, and Abe K (1988) Tide gauge response to tsunamis: Measurements at 40 tide gauge stations in Japan. Journal of Marine Research 46: 557–571. Satake K, Shimazaki K, Tsuji Y, and Ueda K (1996) Time and site of a giant earthquake in Cascadia inferred from Japanese tsunami records of January 1700. Nature 379: 246–249. Satake K, Smith JR, and Shinozaki K (2002) Three-dimensional reconstruction and tsunami model of the Nuuanu and Wailau landslides, Hawaii. In: Takahashi E, Lipman P, Garcia M, Naka J, and Aramaki S (eds.) Hawaiian Volcanoes: Deep Underwater Perspectives. 128, pp. 333–346. Washington D.C: American Geophysical Union. Satake K and Tanioka Y (1999) Sources of tsunami and tsunamigenic earthquakes in subduction zones. Pure and Applied Geophysics 154: 467–483. Satake K and Tanioka Y (2003) The July 1998 Papua New Guinea earthquake: Mechanism and quantification of unusual tsunami generation. Pure and Applied Geophysics 160: 2087–2118. Satake K, Wang KL, and Atwater BF (2003) Fault slip and seismic moment of the 1700 Cascadia earthquake inferred from Japanese tsunami descriptions. Journal of Geophysical Research 108: (doi: 10.1029/2003JB002521). Seno T (2002) Tsunami earthquake as transient phenomena. Geophysical Research Letters 29: (doi: 10.029/ 2002GL014868). Seno T and Hirata K (2007) Did the 2004 Sumatra–Andaman earthquake involve a component of tsunami earthquakes? Bulletin of the Seismological Society of America 97: S296–S306. Shuto N (1991) Numerical simulation of tsunamis – Its present and near future. Natural Hazards 4: 171–191. Shuto N and Matsutomi H (1995) Field survey of the (1993) Hokkaido-Nansei-Oki earthquake tsunami. Pure and Applied Geophysics 144: 649–663. Simkin T and Fiske R (1983) Krakatau, 1883 – the volcanic eruption and its effects. Washington, D.C: Smithsonian Institution Press. Singh SK, Ortiz M, Gupta HK, and Ramadass DGA (2006) Slow slip below Port Blair, Andaman, during the great Sumatra–
510 Tsunamis Andaman earthquake of 26 December 2004. Geophysical Research Letters 33: (doi: 10.1029/2005GL025025). Slingerland R and Voight B (1979) Occurrence, properties, and predictive models of landslide-generated water waves. In: Voight B (ed.) Rockslides and Avalanches, 1, pp. 317–397. Amsterdam: Elsevier. Smith DE, Shi S, Cullingford RA, et al. (2004) The Holocene Storegga slide tsunami in the United Kingdom. Quaternary Science Reviews 23: 2291–2321. Smith WHF, Scharroo R, Titov VV, Arcas D, and Arbic BK (2005) Satellite altimeters measure tsunami, early model estimates confirmed. Oceanography 18: 10–12. Soloviev SL (1970) Recurrence of tsunamis in the Pacific. In: Adams W M (ed.) Tsunamis in the Pacific Ocean, pp. 149–164. Honolulu: East-West Center Press. Soloviev SL and Go CN (1974a) Catalogue of Tsunamis on the Western Shore of the Pacific Ocean (173–1968). Moscow: Nauka Publishing House. Soloviev SL and Go CN (1974b) Catalogue of Tsunamis on the Eastern shore of the Pacific Ocean (1513–1968). Moscow: Nauka Publishing House. Soloviev SL, Go CN, and Kim KS (1992) Catalog of Tsunamis in the Pacific 1969–1982. Moscow: Academy of Science of USSR. Song YT, Ji C, Fu L-L, et al. (2005) The 26 December 2004 tsunami source estimated from satellite radar altimetry and seismic waves. Geophysical Research Letters 32, doi: 10.1029/2005GL023683. Stein S and Okal EA (2005) Speed and size of the Sumatra earthquake. Nature 434: 581–582. Steketee JA (1958) On Volterra’s dislocations in a semi-infinite elastic medium. Canadian Journal of Physics 36: 192–205. Synolakis CE, Bardet JP, Borrero JC, et al. (2002) The slump origin of the 1998 Papua New Guinea tsunami. Proceedings of the Royal society of London Series A 458: 763–789. Synolakis CE and Okal EA (2005) 1992–2002: Perspective on a decade of post-tsunami surveys. In: Satake K (ed.) Tsunamis: Case Studies and Recent Developments, pp. 1–29. Dordrecht: Springer. Takahashi T, Shuto N, Imamura F, and Ortiz M (1995) Source models for the 1993 Hokkaido-Nansei-Oki earthquake tsunami. Pure and Applied Geophysics 144: 747–767. Tanioka Y, Nishimura Y, Hirakawa K, et al. (2004) Tsunami runup heights of the 2003 Tokachi-oki earthquake. Earth, Planets and Space 56: 359–365. Tanioka Y, Ruff L, and Satake K (1997) What controls the lateral variation of large earthquake occurrence along the Japan trench? Island Arc 6: 261–266. Tanioka Y and Satake K (1996a) Tsunami generation by horizontal displacement of ocean bottom. Geophysical Research Letters 23: 861–864. Tanioka Y and Satake K (1996b) Fault parameters of the 1896 Sanriku tsunami earthquake estimated from tsunami numerical modeling. Geophysical Research Letters 23: 1549–1552. Tanioka Y and Sataka K (2001a) Coseismic slip distribution of the 1946 Nankai earthquake and aseismic slips caused by the earthquake. Earth, Planets and Space 53: 235–241. Tanioka Y and Satake K (2001b) Detailed coseismic slip distribution of the 1944 Tonankai earthquake estimated from tsunami waveforms. Geophysical Research Letters 28: 1075–1078. Tanioka Y, Satake K, and Ruff L (1995) Total analysis of the 1993 Hokkaido-Nansei-Oki earthquake using seismic-wave, tsunami, and geodetic data. Geophysical Research Letters 22: 9–12. Tanioka Y and Seno T (2001a) Sediment effect on tsunami generation of the 1896 Sanriku tsunami earthquake. Geophysical Research Letters 28: 3389–3392.
Tanioka Y and Seno T (2001b) Detailed analysis of tsunami waveforms generated by the 1946 Aleutian tsunami earthquake. Natural Hazards and Earth System Sciences 1: 171–175. Tanioka Y, Yudhicara, Kusunose T, et al. (2006) Rupture process of the 2004 great Sumatra-Andaman earthquake estimated from tsunami waveforms. Earth, Planets and Space 58: 203–209. Tappin DR (2004) Special issue: Submarine-slump-generated tsunamis – Preface. Marine Geology 203: 199–200. Tappin DR, Matsumoto P, Watts P, et al. (1999) Sediment slump likely caused 1998 Papua New Guinea tsunami. Eos, Transactions, American Geophysical Union 80: 329, 334, 340. Tatehata H (1997) The new tsunami warning system of the Japan Meteorological Agency. In: Hebenstreit G (ed.) Perspectives on Tsunami Hazards Reduction, pp. 175–188. Dordrecht: Kluwer. Tilling RI (1989) Volcanic Hazards and Their Mitigation - Progress and Problems. Reviews of Geophysics 27: 237–269. Tinti S, Bortolucci E, and Vannini C (1997) A block-based theoretical model suited to gravitational sliding. Natural Hazards 16: 1–28. Titov V, Rabinovich AB, Mofjeld HO, Thomson RE, and Gonzalez FI (2005a) The global reach of the 26 December 2004 Sumatra tsunami. Science 309: 2045–2048. Titov VV and Gonzalez FI (2001) Numerical study of the source of the July 17, 1998 PNG tsunami. In: Hebensreit GT (ed.) Tsunami Research at the End of a Critical Decade, pp. 197–207. Dordrecht: Kluwer Academic. Titov VV, Gonzalez FI, Bernard EN, et al. (2005b) Real-time tsunami forecasting: Challenges and solutions. Natural Hazards 35: 41–58. Titov VV and Synolakis CE (1997) Extreme inundation flows during the Hokkaido-Nansei-Oki tsunami. Geophysical Research Letters 24: 1315–1318. Tobita M, Suito H, Imakiire T, Kato M, Fujiwara S, and Murakami M (2006) Outline of vertical displacement of the 2004 and 2005 Sumatra earthquakes revealed by satellite radar imagery. Earth, Planets and Space 58: e1–e4. Tsai VC, Nettles M, Ekstrom G, and Dziewonski AM (2005) Multiple CMT source analysis of the 2004 Sumatra earthqauke. Geophysical Research Letters 32: (doi: 10.1029/ 2005GL023813). Tsuboi S, Abe K, Takano K, Yamanaka Y (1995) RapidDetermination of Mw from Broad-Band P-Wave-Forms. Bulletin of the Seismological Society of America 85: 606–613. Tsuji Y, Namegaya Y, Matsumoto H, et al. (2006) The 2004 Indian tsunami in Thailand: Surveyed runup heights and tide gauge records. Earth, Planets and Space 58: 223–232. Tuttle MP, Ruffman A, Anderson T, and Jeter H (2004) Distinguishing tsunami from storm deposits in eastern North America: The 1929 grand banks tsunami versus the 1991 Halloween storm. Seismological Research Letters 75: 117–131. Unoki S and Nakano M (1953) On the Cauchy-Poisson waves caused by the eruption of a submarine volcano (III). Oceanographical Magazine 4: 139–150. Utsu T (1984) Estimation of parameters for recurrence models of earthquakes. Bulletin of the Earthquake Research Institute (University of Tokyo) 59: 53–66. Ward S (1980) Relationships of tsunami generation and an earthquake source. Journal of Physics of the Earth 28: 441–474. Ward SN (2001) Landslide tsunami. Journal of Geophysical Research 106: 11201–11215. Ward SN and Day S (2003) Ritter Island Volcano – lateral collapse and the tsunami of 1888. Geophysical Journal International 154: 891–902.
Tsunamis Watanabe H (1998) Comprehensive List of Tsunamis to Hit the Japanese Islands. Tokyo: University of Tokyo Press. Watts P, Grilli ST, Kirby JT, Fryer GJ, and Tappin DR (2003) Landslide tsunami case studies using a Boussinesq model and a fully nonlinear tsunami generation model. Natural Hazards Earth System Sciences 3: 391–402. Woodworth PL and Player R (2003) The permanent service for mean sea level: An update to the 21st century. Journal of Coastal Research 19: 287–295. Yamashita T and Sato R (1974) Generation of tsunami by a fault model. Journal of Physics of the Earth 22: 415–440. Yeh H, Imamura F, Synolakis C, Tsuji Y, Liu P, and Shi S (1993) The Flores Island tsunamis. Eos, Transactions, American Geophysical Union 74: 369, 371–373.
511
Yuan XH, Kind R, and Pedersen HA (2005) Seismic monitoring of the Indian Ocean tsunami. Geophysical Research Letters 32: L15308, doi: 10.1029/2005GL023464.
Relevant Websites http://tsun.sscc.ru – Institute of Computational Mathematics and Mathematical Geophysics, Novosibirsk. http://www.ngdc.noaa.gov – National Geophysical Data Center, NOAA.