**Meet the editor**

Mohammad Mokhtari was born in Nosara, Neyshabur and has obtained his BSc from Azarabadegan, MSc from Southampton and PhD from Bergen Universities. He has worked as a scientific assistant at the University of Utrecht, principal geophysicist at Norsk Hydro and NIOC. As Director of Seismology Research Center and member of Board of Directors he has established the

Iranian National Broad Band Seismic Network. He was the co-founder and Director of National Center for Earthquake Prediction and founding member of Risk Management Excellence Center, IIEES; coordinator of two International Conferences in Tehran; member of Passive Seismic Equipment Expert Panel, CTBTO; Visiting researcher, Geoscience of Australia, Indian Ocean Tsunami Hazard Assessment. He was a member of ICG/IOTWS Working Group 1 and 3; member of Editorial Board, WDR2009, report. Dr. Mokhtari supervised 18 MSc and 7 PhD students mainly in Exploration Seismology, Seismology and Tsunami. His publications consist of over 35 papers in international journals, over 85 conference presentations and four books in different earth science subjects.

Contents

**Preface IX** 

**Part 1 Advanced Measurement Methodologies 1** 

**Technologies and Its Applications 3** 

Chapter 3 **Proximal Records of Paleotsunami Runup in Barrage** 

Curt D. Peterson and Kenneth M. Cruikshank

**Creek Floodplains from Late-Holocene Great Earthquakes in the Central Cascadia Suduction Zone, Oregon, USA 35** 

**Groundwater Resources and Associated Water Supplies 87** 

Chapter 2 **Tsunami Detection by Ionospheric Sounding: New Tools for Oceanic Monitoring 19** 

**Part 2 Tsunami Effect on Infrastructures 59** 

Chapter 5 **Tsunamis as Long-Term Hazards to Coastal** 

Chapter 6 **Experimental and Numerical Modeling** 

Tze Liang Lau, Tatsuo Ohmachi,

**Restoration and Reconstruction 61** 

Karen G. Villholth and Bhanu Neupane

**of Tsunami Force on Bridge Decks 105** 

Shusaku Inoue and Panitan Lukkunaprasit

**Occurrences in the Northern Caribbean 133** 

Mario Octavio Cotilla Rodríguez and Diego Córdoba Barba

Chapter 1 **Advances for Tsunami Measurement** 

Hiroyuki Matsumoto

Giovanni Occhipinti

Chapter 4 **Post-Tsunami Lifeline** 

**Part 3 Case Studies 131** 

Chapter 7 **Comments About Tsunami** 

Yasuko Kuwata

### Contents

#### **Preface XI**


	- **Part 2 Tsunami Effect on Infrastructures 59**
	- **Part 3 Case Studies 131**

X Contents


	- **Part 4 Post-Tsunami Preparedness 199**

### Preface

The term tsunami comes from a Japanese word that means "harbor" (tsu) and "wave" (nami). In the past, the phenomenon was referred to as a tidal wave. However, in the international scientific community this word describes waves generated by sudden vertical movements of the ocean floor, triggered by large earthquakes, volcanic eruptions, or underwater explosions.

Tsunamis can be considered transition phenomena because of their impulsive origin. They are characterized by a long wave length and period. Tsunamis can travel for thousands of kilometers across the open ocean at speeds of 600–800 km per hour, and their effects can be seen hours later on shores. As a tsunami approaches the coast, it reduces the wave celerity and increases the wave height, reaching up to 20 meters with a very high destructive power.

In the recent years the world has experienced a few mega-tsunamis which have caused extensive material damage and death tolls. The most destructive ones were in December 2004 in Sumatra, causing more than 200,000 deaths, and in March 2011 in Japan, causing a nuclear accident. The 2004 catastrophe has triggered many global initiatives such as a new tsunami detection system, more detailed coastal modeling, tsunami compatible coastal developments, integrated approach for regional early warning system, an effort of educating the public, raising awareness and preparedness.

Bearing that in mind, this multi-disciplinary book intends to cover different practical aspects of pre- and post-tsunami management including: advance measurement technology as an early warning system, some important case studies and hazard assessments; lifeline, medical and psychological aspects.

For some practical reasons and to increase its accessiblity, the book is divided into four sections. Section 1 provides advanced methods for tsunami measurement and modeling such as: ocean-bottom pressure sensor, kinematic GPS buoy, satellite altimetry, Paleo tsunami and Ionospheric sounding, early warning system, and scenario based numerical modeling. Section 2 presents case studies in different tsunamigeneic zones around the world such as the Northern Caribbean, Makran region and Tamil Nadu coast in India. Furthermore, classifying tsunamis into local,

#### XII Preface

regional and global, their possible impact on the region and its immediate vicinity is highlighted. Effects of tsunami hazard on the coastal environment and infrastructure (structures, lifelines, water resources, bridges, dykes, etc.) have been presented in section 3. Finally, in section 4, which deals with post-tsunami management, the need for preparedness of emergency medicine staff and the prevention of psychological consequences of the affected survivors has been discussed.

The objective of this book is to provide a collection of expert writing on different aspects of pre- and post-tsunami developments and management techniques. It is intended to be distributed within the scientific community and among the decision makers for tsunami risk reduction. The presented chapters have been thoroughly reviewed and accepted for publication.

We would like to express our gratitude to the contributing authors who are the key factor in this achievement. The Editor expresses his deep appreciation to Prof. M. G. Ashtiany for his support and encouragement. Finally, special thanks to InTech, the publisher that initiated this book and guided and helped the Editor in its completion.

> **Prof. Mohammad Mokhtari**  International Institute of Earthquake Engineering and Seismology, Dibajie Shomali, Tehran, Iran

## **Part 1**

## **Advanced Measurement Methodologies**

**1** 

*Japan* 

Hiroyuki Matsumoto

**Advances for Tsunami Measurement** 

After the Indian Ocean tsunami from the Sumatra earthquake on 26 December 2004 (Mw > 9.0), we have realized the importance of the early tsunami warning system and its necessity for mitigating the tsunami disaster. This catastrophic event was a cue for construction of the Indian Ocean early tsunami warning system, and first of all, a global tsunami forecast system was established together with the Pacific Ocean tsunami warning system operated by U.S. and Japan. The Indian Ocean tsunami motivated the international scieity to construct global tsunami warning systems, which include seismic and sea level monitoring measurements. National Centre of Geosciences in Germany, for example, would challenge to detect relative initial tsunami height distribution by GPS arrays and the seismic stations on land, and deploy GPS buoys along the Sumatra trench for establishment of an early tsunami warning system in Indonesia. And finally the German-Indonesian Tsunami Early Warning System (GITEWS)

Before the Indian Ocean tsunami in 2004, only tide gauge records are available data in the most countries surrounding the Indian Ocean (e.g., Merrifield et al., 2005; Matsumoto et al., 2009). Moreover, some of them were not transfered in real-time but were recorded and avaliable only inside the tide gauge stations. Instrumentally observed tsunami data acquired in real-time is qualitatively used for tsunami warning issue followed by its modification and cancellation. If characteristics of forthcoming tsunami would be understood in advance, it must be helpful and useful for tsunami related disaster mitigation. Tsunami height and arrival time are the most important information after the tsunamigenic earthquake occurrence, and they are often used as tsunami observation information. Tsunami observation is traditionally carried out by tide gauges at the coast. Recently, technological

become into operation in 2010 (e.g., Rudloff et al., 2009; Münch et al., 2011).

development has been promoted to estimate tsunami features as early as possible.

megathrust earthquake in the Nankai trough, SW Japan.

**2. Tsunami measurement instruments** 

This chapter reviews tsunami measurement technologies and instruments, in particularly developed in Japan and introduces an actual tsunami observation in the source area, which became possible after the offshore tsunami observation in the last decade. In the end, potential use for early tsunami detection is discussed by applying to the presumed

Tsunami measurements are usually carried out by tide gauge or bottom pressure sensor, or kinematic GPS buoy in Japan. The most traditional procedure is to measure by tide

**1. Introduction** 

**Technologies and Its Applications** 

*Japan Agency for Marine-Earth Science and Technology,* 

### **Advances for Tsunami Measurement Technologies and Its Applications**

Hiroyuki Matsumoto

*Japan Agency for Marine-Earth Science and Technology, Japan* 

#### **1. Introduction**

After the Indian Ocean tsunami from the Sumatra earthquake on 26 December 2004 (Mw > 9.0), we have realized the importance of the early tsunami warning system and its necessity for mitigating the tsunami disaster. This catastrophic event was a cue for construction of the Indian Ocean early tsunami warning system, and first of all, a global tsunami forecast system was established together with the Pacific Ocean tsunami warning system operated by U.S. and Japan. The Indian Ocean tsunami motivated the international scieity to construct global tsunami warning systems, which include seismic and sea level monitoring measurements. National Centre of Geosciences in Germany, for example, would challenge to detect relative initial tsunami height distribution by GPS arrays and the seismic stations on land, and deploy GPS buoys along the Sumatra trench for establishment of an early tsunami warning system in Indonesia. And finally the German-Indonesian Tsunami Early Warning System (GITEWS) become into operation in 2010 (e.g., Rudloff et al., 2009; Münch et al., 2011).

Before the Indian Ocean tsunami in 2004, only tide gauge records are available data in the most countries surrounding the Indian Ocean (e.g., Merrifield et al., 2005; Matsumoto et al., 2009). Moreover, some of them were not transfered in real-time but were recorded and avaliable only inside the tide gauge stations. Instrumentally observed tsunami data acquired in real-time is qualitatively used for tsunami warning issue followed by its modification and cancellation. If characteristics of forthcoming tsunami would be understood in advance, it must be helpful and useful for tsunami related disaster mitigation. Tsunami height and arrival time are the most important information after the tsunamigenic earthquake occurrence, and they are often used as tsunami observation information. Tsunami observation is traditionally carried out by tide gauges at the coast. Recently, technological development has been promoted to estimate tsunami features as early as possible.

This chapter reviews tsunami measurement technologies and instruments, in particularly developed in Japan and introduces an actual tsunami observation in the source area, which became possible after the offshore tsunami observation in the last decade. In the end, potential use for early tsunami detection is discussed by applying to the presumed megathrust earthquake in the Nankai trough, SW Japan.

#### **2. Tsunami measurement instruments**

Tsunami measurements are usually carried out by tide gauge or bottom pressure sensor, or kinematic GPS buoy in Japan. The most traditional procedure is to measure by tide

Advances for Tsunami Measurement Technologies and Its Applications 5

Another concern to make use of tide gauge is its response. Differences on tsunami heights between tide gauges and eyewitnesses have been pointed out in the past. Tide gauge generally uses a narrow intake pipe between the tide well and the sea as shown in Fig. 2. This is because the main purpose of tide gauge is to observe astronomical tide with its period of a few hours or much longer of a few years' sea level change caused by global climate change. Hence short period sea level changes such as surge wave or swell are structurally cut off. Tsunamis of their period less than a few ten of minutes can be recorded by tide gauges indeed, but some considerable responses were pointed out in the past. For example Okada (1985) examined the tide gauge response after the Japan Sea earthquake (Mw7.9) in 1983, and corrected tsunami waveform in terms of nonlinear response. Namegaya et al. (2009) carried out in-situ measurement of tide gauge stations and estimated liniear and nonlinear response and corrected the tsunami waveforms from the Niigataken

Chuetsu-oki, Japan eartqhauke (Mw6.6) in 2007.

**2.2 Bottom pressure sensor** 

Fig. 2. Schematic drawing of typical float type tide gauge station in Japan.

Tide gauge

Offshore tsunami measurement makes us possible to predict tsunami arrival and provide time to evacuate from tsunami. Recent deep-sea technologies enable to observe tsunamis not only offshore but also in real-time. One of the facilities composing the early tsunami warning system is the offshore observatory. National Oceanic and Atmospheric Administration (NOAA) developed Deep-ocean Assessment and Reporting of Tsunamis (DART) system that receives water pressure from the ocean bottom firstly deployed in the Pacific and Atlantic Oceans (Gonzalez et al., 2005). Now the DART system has been extended to the Indian Ocean and each observatory is owned by not only U.S. but also Australia, Chile, Indonesia Thailand, and Russia. On the other hand, other sensors such as in-lined cabled bottom pressure sensors are developed and deployed in the seismogenic

Intake pipe Tide well Float buoy

gauge, whereas the most modern is by kinematic GPS buoy actually being in operation. Distribution of tsunami measurement instruments in Japan is shown in Fig. 1. This section describes details of tide gauge, bottom pressure sensor, and kinematic GPS buoy in their basic mechanism, outstanding problems, and applications of actual tsunami observation.

Fig. 1. Locations of tide gauge stations (open circles) and offshore tsunami observatories (kinematic GPS buoys: triangles; bottom pressure sensors: squares) in Japan.

#### **2.1 Tide gauge**

Tide gauges are deployed in order for measurement of usual sea level, i.e., astronomical tide level not only in Japan but also all over the world. Several types of tide gauges are being operated in Japan. The most typical tide gauge is to use a tide well which records vertical motion of a float buoy in a well connecting by an intake pipe to the open sea (Fig. 2). The first tide gauge was established in Japan is the in the early 1890s, for which Kelvin type tide gauge produced in England was employed (GSI, available at online). This type of tide gauge had used the analogue paper chart until the 1990s, and more recently digital decoding instrument is equipped on the tide gauge. The tide gauge using a paper chart requires replacement of recording paper at some intervals. Other types of tide gauges are as follows, e.g., a pressure type which measure hydrostatic pressure equivalent to the sea level at the station, and an acoustic type which measure distance between the sea surface and the acoustic receiver at the bottom. Generally, tide gauge stations are located inside the port or the harbour. This is why tsunami height based on tide gauge means tendency value where the tide gauge station is located. In fact, tsunami heights vary depending on both the local land and subsea topographies.

gauge, whereas the most modern is by kinematic GPS buoy actually being in operation. Distribution of tsunami measurement instruments in Japan is shown in Fig. 1. This section describes details of tide gauge, bottom pressure sensor, and kinematic GPS buoy in their basic mechanism, outstanding problems, and applications of actual tsunami observation.

Fig. 1. Locations of tide gauge stations (open circles) and offshore tsunami observatories

Tide gauges are deployed in order for measurement of usual sea level, i.e., astronomical tide level not only in Japan but also all over the world. Several types of tide gauges are being operated in Japan. The most typical tide gauge is to use a tide well which records vertical motion of a float buoy in a well connecting by an intake pipe to the open sea (Fig. 2). The first tide gauge was established in Japan is the in the early 1890s, for which Kelvin type tide gauge produced in England was employed (GSI, available at online). This type of tide gauge had used the analogue paper chart until the 1990s, and more recently digital decoding instrument is equipped on the tide gauge. The tide gauge using a paper chart requires replacement of recording paper at some intervals. Other types of tide gauges are as follows, e.g., a pressure type which measure hydrostatic pressure equivalent to the sea level at the station, and an acoustic type which measure distance between the sea surface and the acoustic receiver at the bottom. Generally, tide gauge stations are located inside the port or the harbour. This is why tsunami height based on tide gauge means tendency value where the tide gauge station is located. In fact, tsunami heights vary depending on both the local

(kinematic GPS buoys: triangles; bottom pressure sensors: squares) in Japan.

**2.1 Tide gauge** 

land and subsea topographies.

Another concern to make use of tide gauge is its response. Differences on tsunami heights between tide gauges and eyewitnesses have been pointed out in the past. Tide gauge generally uses a narrow intake pipe between the tide well and the sea as shown in Fig. 2. This is because the main purpose of tide gauge is to observe astronomical tide with its period of a few hours or much longer of a few years' sea level change caused by global climate change. Hence short period sea level changes such as surge wave or swell are structurally cut off. Tsunamis of their period less than a few ten of minutes can be recorded by tide gauges indeed, but some considerable responses were pointed out in the past. For example Okada (1985) examined the tide gauge response after the Japan Sea earthquake (Mw7.9) in 1983, and corrected tsunami waveform in terms of nonlinear response. Namegaya et al. (2009) carried out in-situ measurement of tide gauge stations and estimated liniear and nonlinear response and corrected the tsunami waveforms from the Niigataken Chuetsu-oki, Japan eartqhauke (Mw6.6) in 2007.

Fig. 2. Schematic drawing of typical float type tide gauge station in Japan.

#### **2.2 Bottom pressure sensor**

Offshore tsunami measurement makes us possible to predict tsunami arrival and provide time to evacuate from tsunami. Recent deep-sea technologies enable to observe tsunamis not only offshore but also in real-time. One of the facilities composing the early tsunami warning system is the offshore observatory. National Oceanic and Atmospheric Administration (NOAA) developed Deep-ocean Assessment and Reporting of Tsunamis (DART) system that receives water pressure from the ocean bottom firstly deployed in the Pacific and Atlantic Oceans (Gonzalez et al., 2005). Now the DART system has been extended to the Indian Ocean and each observatory is owned by not only U.S. but also Australia, Chile, Indonesia Thailand, and Russia. On the other hand, other sensors such as in-lined cabled bottom pressure sensors are developed and deployed in the seismogenic

Advances for Tsunami Measurement Technologies and Its Applications 7

placed on the top of offshore buoy and the other is placed on land-based station. After practical operation period, about 10 kinematic GPS buoys have been deployed 10-20 km offshore from the coast in Japan (Figs. 1 and 4). Although GPS buoy cannot be deployed over several kilometres further offshore because of the limitation of communication distance between the GPS buoy and the base station, it has demonstrated an advantage for early tsunami detection. The tsunami from the off Kii peninsula earthquake was recorded by the GPS buoy for the first time 8 min before its arrival at the nearest tide gauge station (Kato et al., 2005). Tsunami from the off Kii peninsula earthquake was detected by both the offshore pressure sensors and the kinematic GPS buoys in which tsunami heights were recorded to be ca. 10 cm and ca. 20 cm in peak-to-peak amplitude, respectively, whereas the tsunami height recorded by the tide gauge was to be 50 to 100 cm. This is attributed to the shoreing

Fig. 4. Kinematic GPS buoy deployed offshore of NE Japan (photo by Port and Airport

Offshore tsunami observation has an advantage for far-field tsunami as mentioned above. However, for the near-field tsunamis that are generated near the tsunami measurement sensors it have not been experienced and discussed about usage of acquired data. This section describes an example of actual tsunami observations in particular in the nearsource area by the bottom pressure sensors by the cabled observatory system, and discuss their unique phenomena during the tsunami generation process for use of tsunami early detections. Offshore tsunami observations have been done in the past as reviewed in the previous section. At the beginning of the offshore observation of tsunami, pressure fluctuation caused by the seismic wave apparently much intense than that by the tsunami wave (e.g., Filloux, 1982). This is why mathematical low-pass filtering is necessary to detect tsunami signals. In fact, low-pass filtering was applied in most cases of tsunamigenic earthquakes in order to identify tsunami signals afterwards for scientific

effect.

Research Institute).

**3. Tsunami measurement applications** 

zone in Japan. Figure 1 represents the current bottom pressure sensors locations being operated in Japan either. The first offshore observatory in Japan has been deployed in 1978 off Omaezaki, central Japan, where the probability of the presumed Tokai earthquake is expected to be 87 % by the Earthquake Research Committee of the Headquarters for Earthquake Research Promotion, i.e., the Japanese Government. Then this types of tsunami measurement have followed until now, and eight observatories in total have been deployed in Japan.

Fig. 3. In-lined cabled bottom pressure sensor deployed in the ocean.

Offshore tsunami detected by bottom pressure sensor is given by Filloux (1982) for the first time. Eble and Gonzalez (1986) performed the long-term observation on bottom pressure sensors and reported detection of offshore tsunami signals from three different earthquakes during their observational period. Hino et al. (2001) and Hirata et al. (2003), for example, used less than a few centimetres tsunamis from the moderate-to-large earthquakes occurred in the Japan trench and the Kuril trench, respectively, that could be detected by Japanese cabled bottom tsunami sensors. Matsumoto and Mikada (2005) and Satake et al. (2005) used offshore tsunami recorded by bottom pressure sensors in order to constrain fault models of the off Kii peninsula earthquake (Mw 7.4) in Japan, and demonstrated advances offshore observation for tsunami. Tsunami from the off Kii peninsula earthquake was also observed at the tide gauge stations along the coast nearby. Bottom pressure sensors could detect tsunami signals about 20 min before its arrival at the nearest tide gauge stations. Thus it shows that offshore tsunami observation has an advantage of the tsunami detection for farfield tsunamis.

#### **2.3 Kinematic GPS buoy**

Kinematic GPS buoy is a new technological system developed in the late 1990s to observe tsunami at the offshore sea surface (Kato et al., 2000). GPS, i.e. Global Positioning System technology widely used on land is to be applied to the sea surface. The current kinematic GPS buoy monitors a moving platform in real-time with an accuracy of a few centimetres by relative positioning. It requires two GPS receivers to measure the relative position, one is

zone in Japan. Figure 1 represents the current bottom pressure sensors locations being operated in Japan either. The first offshore observatory in Japan has been deployed in 1978 off Omaezaki, central Japan, where the probability of the presumed Tokai earthquake is expected to be 87 % by the Earthquake Research Committee of the Headquarters for Earthquake Research Promotion, i.e., the Japanese Government. Then this types of tsunami measurement have followed until now, and eight observatories in total have been deployed

Fig. 3. In-lined cabled bottom pressure sensor deployed in the ocean.

Offshore tsunami detected by bottom pressure sensor is given by Filloux (1982) for the first time. Eble and Gonzalez (1986) performed the long-term observation on bottom pressure sensors and reported detection of offshore tsunami signals from three different earthquakes during their observational period. Hino et al. (2001) and Hirata et al. (2003), for example, used less than a few centimetres tsunamis from the moderate-to-large earthquakes occurred in the Japan trench and the Kuril trench, respectively, that could be detected by Japanese cabled bottom tsunami sensors. Matsumoto and Mikada (2005) and Satake et al. (2005) used offshore tsunami recorded by bottom pressure sensors in order to constrain fault models of the off Kii peninsula earthquake (Mw 7.4) in Japan, and demonstrated advances offshore observation for tsunami. Tsunami from the off Kii peninsula earthquake was also observed at the tide gauge stations along the coast nearby. Bottom pressure sensors could detect tsunami signals about 20 min before its arrival at the nearest tide gauge stations. Thus it shows that offshore tsunami observation has an advantage of the tsunami detection for far-

Kinematic GPS buoy is a new technological system developed in the late 1990s to observe tsunami at the offshore sea surface (Kato et al., 2000). GPS, i.e. Global Positioning System technology widely used on land is to be applied to the sea surface. The current kinematic GPS buoy monitors a moving platform in real-time with an accuracy of a few centimetres by relative positioning. It requires two GPS receivers to measure the relative position, one is

in Japan.

field tsunamis.

**2.3 Kinematic GPS buoy** 

placed on the top of offshore buoy and the other is placed on land-based station. After practical operation period, about 10 kinematic GPS buoys have been deployed 10-20 km offshore from the coast in Japan (Figs. 1 and 4). Although GPS buoy cannot be deployed over several kilometres further offshore because of the limitation of communication distance between the GPS buoy and the base station, it has demonstrated an advantage for early tsunami detection. The tsunami from the off Kii peninsula earthquake was recorded by the GPS buoy for the first time 8 min before its arrival at the nearest tide gauge station (Kato et al., 2005). Tsunami from the off Kii peninsula earthquake was detected by both the offshore pressure sensors and the kinematic GPS buoys in which tsunami heights were recorded to be ca. 10 cm and ca. 20 cm in peak-to-peak amplitude, respectively, whereas the tsunami height recorded by the tide gauge was to be 50 to 100 cm. This is attributed to the shoreing effect.

Fig. 4. Kinematic GPS buoy deployed offshore of NE Japan (photo by Port and Airport Research Institute).

### **3. Tsunami measurement applications**

Offshore tsunami observation has an advantage for far-field tsunami as mentioned above. However, for the near-field tsunamis that are generated near the tsunami measurement sensors it have not been experienced and discussed about usage of acquired data. This section describes an example of actual tsunami observations in particular in the nearsource area by the bottom pressure sensors by the cabled observatory system, and discuss their unique phenomena during the tsunami generation process for use of tsunami early detections. Offshore tsunami observations have been done in the past as reviewed in the previous section. At the beginning of the offshore observation of tsunami, pressure fluctuation caused by the seismic wave apparently much intense than that by the tsunami wave (e.g., Filloux, 1982). This is why mathematical low-pass filtering is necessary to detect tsunami signals. In fact, low-pass filtering was applied in most cases of tsunamigenic earthquakes in order to identify tsunami signals afterwards for scientific

Advances for Tsunami Measurement Technologies and Its Applications 9

magnitude over 6.0 by means of signal-to-noise ratio. Conditioning these criteria, 16 earthquakes were selected listed in Table 2. Among those 16 earthquakes, three earthquakes on 26 September 2003, on 29 November 2004, and 11 September 2009 generated the tsunamis which were observed at the tide gauge stations at the coast. Locations of the selected earthquakes and the PGs are compared in Figure 5. Both the 2003 and 2009 tsunamigenic earthquakes' epicenters were located beneath PG1, on the other hand, that of the 2004

Latitude (○N) Longitude (○E) Water depth (m)

Longitude (○E)

1 2003.09.26 04:05:07.42 41.779 144.079 45.1 8.0 observed

10 2004.11.29 03:32:14.53 42.946 145.276 48.2 7.1 observed

15 2008.09.11 09:20:51.35 41.776 144.151 30.9 7.1 observed

Generally, bottom pressure sensors measure the vibration regarding pressure and temperature, from which physical value is processed compensating the temperature collections. For the principal of the pressure sensors, narrow sample rate gives low resolution response. 10 Hz sampling is the minimum sample rate for the reliable value. We have analyzed the obtained PGs dataset to make spectrograms. Numerical technique to

Table 2. Significant earthquakes occurred near the pressure sensors off Hokkaido, Japan.

Depth (km)

Magnitude Tsunami

PG1 41.7040 144.4375 2218 PG2 42.2365 144.8454 2210

Table 1. Location of bottom pressure sensors off Hokkaido, Japan.

Latitude (○N)

2 2003.09.26 06:08:01.84 41.710 143.692 21.4 7.1 3 2003.09.26 15:26:58.10 42.189 144.776 27.4 6.1 4 2003.09.27 05:38:22.31 42.026 144.728 34.4 6.0 5 2003.09.29 11:36:55.06 42.360 144.553 42.5 6.5 6 2003.10.08 18:06:56.79 42.565 144.670 51.4 6.4 7 2003.10.11 09:08:48.15 41.864 144.440 27.8 6.1 8 2003.12.29 10:30:55.40 42.419 144.756 38.9 6.0 9 2004.11.11 19:02:46.17 42.083 144.486 38.6 6.3

11 2004.11.29 03:36:41.19 42.884 145.236 45.6 6.0 12 2004.12.06 23:15:11.81 42.848 145.343 48.8 6.9 13 2005.01.18 23:09:06.65 42.876 145.007 49.8 6.4 14 2007.02.17 09:02:56.63 41.732 143.723 40.1 6.2

16 2009.06.05 12:30:33.80 41.812 143.620 31.3 6.4

earthquake was located out of PGs.

Date Time

**3.2 Data processing procedure** 

analyze 10 Hz time-series PGs dataset is as follows;

(JST)

purpose. In Japan, the Japan Meteorological Agency (JMA) is responsible for tsunami warning issue, and offshore measurement data are processed by using 1-2 min moving averaging technique. If a large earthquake would take place offshore and accompany a tsunami, i.e., a far-field tsunami, it would not be so difficult to notify tsunami signals as done by the present procedure. Most pressure sensors have been deployed in the tsunami source area. For near-field tsunami, however, there has not been established that data processing methods prepared so far. We urgently need data processing procedure for the near-field tsunamis.

#### **3.1 Bottom pressure sensor off Hokkaido, Japan**

Japan Agency for Marine-Earth Science and Technology (JAMSTEC) is operating four offshore observatories in the seismogenic zone in Japan; off Muroto cape and off Kumano in the Nankai trough, SW Japan, off Hatsushima Island in the Sagami trough, central Japan, and off Hokkaido in the Kuril trench, northern Japan. The present study introduces the offshore observatory off Hokkaido deployed in 1999 (Hirata et al., 2002). Figure 5 shows that the location of bottom pressure sensors connecting by the submarine cable. The cabled observatory has two bottom pressure sensors, and those data is telemetered to JAMSTEC in real-time. Two bottom pressure sensors as referred by PG1 and PG2 hereafter are deployed at the water depths of 2218 m and 2210 m, respectively, and their locations are listed in Table 1.

A megathrust M8.0 earthquake occurred in 2003 in this region (Watanabe et al., 2004), and then the seismic activities including aftershocks have become relatively high. A number of earthquakes over their magnitude 6.0 took place after 2003. In the present study, we focus on the near-field earthquakes in order to understand the observed fluctuation of water pressure during the tsunamigenic earthquake. Referring the earthquake database complied by JMA, earthquakes occurred inside ca. 100 km from the observatories are selected. Because a bottom pressure sensor is very sensitive, we focus on large earthquakes with their

Fig. 5. Offshore observatory off Hokkaido, Japan with locations of significant earthquakes' epicenters. Red indicates a tsunamigenic earthquake.

purpose. In Japan, the Japan Meteorological Agency (JMA) is responsible for tsunami warning issue, and offshore measurement data are processed by using 1-2 min moving averaging technique. If a large earthquake would take place offshore and accompany a tsunami, i.e., a far-field tsunami, it would not be so difficult to notify tsunami signals as done by the present procedure. Most pressure sensors have been deployed in the tsunami source area. For near-field tsunami, however, there has not been established that data processing methods prepared so far. We urgently need data processing procedure for the

Japan Agency for Marine-Earth Science and Technology (JAMSTEC) is operating four offshore observatories in the seismogenic zone in Japan; off Muroto cape and off Kumano in the Nankai trough, SW Japan, off Hatsushima Island in the Sagami trough, central Japan, and off Hokkaido in the Kuril trench, northern Japan. The present study introduces the offshore observatory off Hokkaido deployed in 1999 (Hirata et al., 2002). Figure 5 shows that the location of bottom pressure sensors connecting by the submarine cable. The cabled observatory has two bottom pressure sensors, and those data is telemetered to JAMSTEC in real-time. Two bottom pressure sensors as referred by PG1 and PG2 hereafter are deployed at the water depths of 2218 m and 2210 m, respectively, and their locations are listed in

A megathrust M8.0 earthquake occurred in 2003 in this region (Watanabe et al., 2004), and then the seismic activities including aftershocks have become relatively high. A number of earthquakes over their magnitude 6.0 took place after 2003. In the present study, we focus on the near-field earthquakes in order to understand the observed fluctuation of water pressure during the tsunamigenic earthquake. Referring the earthquake database complied by JMA, earthquakes occurred inside ca. 100 km from the observatories are selected. Because a bottom pressure sensor is very sensitive, we focus on large earthquakes with their

Fig. 5. Offshore observatory off Hokkaido, Japan with locations of significant earthquakes'

epicenters. Red indicates a tsunamigenic earthquake.

near-field tsunamis.

Table 1.

**3.1 Bottom pressure sensor off Hokkaido, Japan** 

magnitude over 6.0 by means of signal-to-noise ratio. Conditioning these criteria, 16 earthquakes were selected listed in Table 2. Among those 16 earthquakes, three earthquakes on 26 September 2003, on 29 November 2004, and 11 September 2009 generated the tsunamis which were observed at the tide gauge stations at the coast. Locations of the selected earthquakes and the PGs are compared in Figure 5. Both the 2003 and 2009 tsunamigenic earthquakes' epicenters were located beneath PG1, on the other hand, that of the 2004 earthquake was located out of PGs.



Table 1. Location of bottom pressure sensors off Hokkaido, Japan.

Table 2. Significant earthquakes occurred near the pressure sensors off Hokkaido, Japan.

#### **3.2 Data processing procedure**

Generally, bottom pressure sensors measure the vibration regarding pressure and temperature, from which physical value is processed compensating the temperature collections. For the principal of the pressure sensors, narrow sample rate gives low resolution response. 10 Hz sampling is the minimum sample rate for the reliable value. We have analyzed the obtained PGs dataset to make spectrograms. Numerical technique to analyze 10 Hz time-series PGs dataset is as follows;

Advances for Tsunami Measurement Technologies and Its Applications 11

PG1 PG2

 2003.9.26 (M8.0) 2008.9.11 (M7.1) 2004.11.29 (M7.1) 2003.9.29 (M6.5)

(M6.5) is also displayed as an example. According to the spectrograms in the near-field, i.e., event (a) and (b), strong phase having 0.1 to 0.2 Hz is obviously observed during the tsunamigenic earthquake. Tsunamigenic earthquake out of the PGs, i.e., event (c), their characteristic phase appeared after the earthquake rather than during the earthquake. This is

10<sup>1</sup>

10<sup>2</sup>

10<sup>3</sup>

Pressure (Pa)

10<sup>4</sup>

10<sup>5</sup>

Cross section profiles of the spectrogram during the earthquakes are plotted in Fig. 7. Tsuamigenic events have peak from 0.1 to 0.2 Hz, which correspond to a natural frequency uniquely depending on the water depth. This is an acoustic resonant wave, i.e., a standing wave forming between the ocean bottom and the sea surface caused by the coseismic deformation (e.g., Nosov & Kolesov (2007)). The larger earthquake magnitude becomes, the

Thus the tsunamigenic earthquake has a peak of frequency between 0.1 Hz and 0.2 Hz in the case of the water depth about 2000 m. And its peak attenuates in duration of 20 s. The same peak of frequency between 0.1 Hz and 0.2 Hz is involved during the non-tsunamigenic earthquake, but its peak is lower than the high frequency peaks associated with seismic

Maximum water pressure Pmax in the case of abrupt bottom deformation resulting in

Because density and sound velocity are constant, i.e., 1.03 kg/m3 and 1500 m/s, resptctively, Eq. (1) provide the bottom velocity. For example, (a) the 2003 and (b) the 2008 earthquake cases, the bottom deformation velocity are given to be 0.13 m/s and 0.03 m/s, respectively. On the other hand, the empirical relation between earthquake magnitude *M* and rise-time of

*c v* (1)

0.01 0.1 1 10

Frequency (Hz)

100 10 1 0.1

Period (s)

 2003.9.26 (M8.0) 2008.9.11 (M7.1) 2004.11.29 (M7.1) 2003.9.29 (M6.5)

, sound

tsunami generation process is expressed as multiplication of density of water

*Pmax =* 

because this phase is reproduced in the tsunami source area, and then it propagates.

Fig. 7. Cross section profiles of spectrograms during the earthquakes.

0.01 0.1 1 10

Frequency (Hz)

100 10 1 0.1

Period (s)

larger water pressure amplitude responses in its narrow band.

velocity in water *v*, and the bottom deformation velocity *v*,

is proposed by Sato (1979),

waves.

101

10<sup>2</sup>

10<sup>3</sup>

Pressure (Pa)

10<sup>4</sup>

10<sup>5</sup>

the seismic faulting

**3.3 Implication of water pressure** 

Fig. 6. Pressure waveforms' spectrograms and its original waveforms during the earthquakes.


Spectrograms with the original pressure waveforms during the tsunamigenic earthquakes on (a) 26 September 2003 (M8.0), (b) 11 September 2008(M7.1), and (c) 29 November 2004 (M7.1) are plotted in Fig. 6, and the largest non-tsunamigenic earthquake on (d) 29 September 2003

PG1 PG2

0.01 0.1 1 10

Frequency (Hz)

Pressure (x105Pa)


0.01 0.1 1 10

Frequency (Hz)

Pressure (x105Pa)


0.01 0.1 1 10

Frequency (Hz)

Pressure (x105Pa)


0.01 0.1 1 10

Frequency (Hz)

Pressure (x105Pa)


Period (s)

Period (s)

Period (s)

Period (s)

0 50 100 150 200 250 300

Time (s)

0 2.0E3 4.0E3 6.0E3 8.0E3

0 50 100 150 200 250 300

Time (s)

0 5.6E3 1.1E4 1.7E4 2.0E4

0 50 100 150 200 250 300

Time (s)

0 3.0E3 6.0E3 9.0E3 1.2E4

0 50 100 150 200 250 300

Time (s)

0 50 100 150 200 250 300 Time (s)

0 50 100 150 200 250 300 Time (s)

0 50 100 150 200 250 300 Time (s)

0 50 100 150 200 250 300 Time (s)

0 2.5E4 5.0E4 7.5E4 1.0E5

Period (s)

Period (s)

Fig. 6. Pressure waveforms' spectrograms and its original waveforms during the

Period (s)

Period (s)

3. Band-pass filtering of each section above is applied to the entire 5min dataset.

amplitude, and spectrogram of PGs during the earthquake can be made.

2. We divide the frequency from 0.01 Hz to 10 Hz into 40 sections as formed by

4. Envelopes of the filtered waveforms for each section are layout to get absolute

Spectrograms with the original pressure waveforms during the tsunamigenic earthquakes on (a) 26 September 2003 (M8.0), (b) 11 September 2008(M7.1), and (c) 29 November 2004 (M7.1) are plotted in Fig. 6, and the largest non-tsunamigenic earthquake on (d) 29 September 2003

1. 5min dataset of PG including each earthquake is collected.

exponentially (i.e., linearly in logarithmic scale).

0 50 100 150 200 250 300

Time (s)

0 50 100 150 200 250 300 Time (s)

0 1.7E3 3.4E3 5.0E3 6.0E3

0 50 100 150 200 250 300

Time (s)

0 50 100 150 200 250 300 Time (s)

0 50 100 150 200 250 300

Time (s)

0 3.4E3 6.7E3 1.0E4 1.2E4

0 50 100 150 200 250 300 Time (s)

0 9.8E3 2.0E4 2.9E4 3.5E4

0 50 100 150 200 250 300

Time (s)

0 50 100 150 200 250 300 Time (s)

0 3.9E4 7.8E4 1.2E5 1.4E5

(a) 26 September 2003 (M8.0)

(b) 11 September 2008 (M7.1)

(c) 29 November 2004 (M7.1)

(d) 29 September 2003 (M6.5)

earthquakes.

0.01 0.1 1 10

0.01 0.1 1 10

Frequency (Hz)

Pressure (x105Pa)


Frequency (Hz)

Pressure (x105Pa)


0.01 0.1 1 10

Frequency (Hz)

Pressure (x105Pa)


0.01 0.1 1 10

Frequency (Hz)

Pressure (x105Pa)


Fig. 7. Cross section profiles of spectrograms during the earthquakes.

(M6.5) is also displayed as an example. According to the spectrograms in the near-field, i.e., event (a) and (b), strong phase having 0.1 to 0.2 Hz is obviously observed during the tsunamigenic earthquake. Tsunamigenic earthquake out of the PGs, i.e., event (c), their characteristic phase appeared after the earthquake rather than during the earthquake. This is because this phase is reproduced in the tsunami source area, and then it propagates.

Cross section profiles of the spectrogram during the earthquakes are plotted in Fig. 7. Tsuamigenic events have peak from 0.1 to 0.2 Hz, which correspond to a natural frequency uniquely depending on the water depth. This is an acoustic resonant wave, i.e., a standing wave forming between the ocean bottom and the sea surface caused by the coseismic deformation (e.g., Nosov & Kolesov (2007)). The larger earthquake magnitude becomes, the larger water pressure amplitude responses in its narrow band.

Thus the tsunamigenic earthquake has a peak of frequency between 0.1 Hz and 0.2 Hz in the case of the water depth about 2000 m. And its peak attenuates in duration of 20 s. The same peak of frequency between 0.1 Hz and 0.2 Hz is involved during the non-tsunamigenic earthquake, but its peak is lower than the high frequency peaks associated with seismic waves.

#### **3.3 Implication of water pressure**

Maximum water pressure Pmax in the case of abrupt bottom deformation resulting in tsunami generation process is expressed as multiplication of density of water , sound velocity in water *v*, and the bottom deformation velocity *v*,

$$P\_{\text{max}} = \rho \circledast \upsilon \tag{1}$$

Because density and sound velocity are constant, i.e., 1.03 kg/m3 and 1500 m/s, resptctively, Eq. (1) provide the bottom velocity. For example, (a) the 2003 and (b) the 2008 earthquake cases, the bottom deformation velocity are given to be 0.13 m/s and 0.03 m/s, respectively. On the other hand, the empirical relation between earthquake magnitude *M* and rise-time of the seismic faulting is proposed by Sato (1979),

Advances for Tsunami Measurement Technologies and Its Applications 13

earthquake, respectively. Because more than 60 years have past since the last earthquake, Japanese government evaluates that the probability of the next presumed megathrust earthquake along the Nankai trough is estimated to be 60-70 % within the next 30 years. Japanese government has constructed an offshore observatory network, which consists of dense 20 bottom seismic sensors and bottom pressure sensors in total in order for monitoring seismic activity and its consequence, megathrust earthquake, and followed by tsunami. JAMSTEC is operating the offshore observatory network. The observatory layout is shown in Fig. 10. As of May 2011, 17 observatories have been deployed, and it started to acquire their data in real-time. If megathrust earthquake and accompanied giant tsunami would be predicted before their arrival nearby the coast and effective warning would be issued, it must contribute to mitigate earthquake and tsunami related disasters. We should establish measurement technology including data processing and accumulate technical know-how for future meagathrust earthquake and tsunami in advance; hence we carry out

Parameters Value

Location 33.277 °N, 136.394 °E Depth 10 km Strike 226 ° Dip 10 ° Rake 90 ° Length 130 km Width 60 km Dislocation 2 m Rise time 5 s Rupture velocity 3 km/s Table 3. Fault parameters used in the tsunami computation from the Tonankai earthquake.

Fig. 9. Deformation patterns from the seismic faults' dislocation

tsunami computation of the last 1944 Tonankai earthquake.

$$r = 10^{1.5M \cdot 1.4} / 80$$

Eq. (2) provides that the rise-times for (a) the 2003 and (b) the 2008 are 5.0 s and 1.7 s, respectively. Assuming the duration time of bottom deformation coincides with the risetime of the seismic faulting, deformation is given by its velocity integrated by the rise-time. Thus derived deformations at the location of PG1 are estimated to be (a) 0.65 m and (b) 0.06 m, respectively. These values almost coincide with (a) the static deformation from the fault plate model by Geospatial Information Authority of Japan (GSI) (2003) and (b) the point source equivalent to the seismic moment (Fig. 8). This means that the displacement of the location of the pressure sensor deployment can be roughly estimated in terms of the water pressure amplitude.

Fig. 8. Deformation patterns from the seismic faults' dislocation

An early tsunami detection approach based on a physical phenomenon uniquely observed in the source during the tsunamigenic earthquake was presented in this section. Tsunami initial waveform is mostly depended on the static deformation of the ocean bottom. Hence the amplitude of the water pressure associated with the acoustic resonant wave may be a potential indicator of the tsunami generation.

#### **4. Tsunami prediction along the Nankai trough**

The first offshore observatory in Japan has been deployed in the Suruga trough targetting the presumed Tokai earthquake, central Japan, and followed by seven cabled observatories. The newest system is being operated in the presumed Tonankai earthquake sourece area by JMA and JAMSTEC off Kii peninsula (Fig. 9).

#### **4.1 Tsunami monitoring system in the Nankai trough**

The Nankai trough is one of the palte subduction zones in Japan, where the last megathrust earthquakes took place in 1944 and 1946, namely the Tonankai eartqhauke and the Nankai

Eq. (2) provides that the rise-times for (a) the 2003 and (b) the 2008 are 5.0 s and 1.7 s, respectively. Assuming the duration time of bottom deformation coincides with the risetime of the seismic faulting, deformation is given by its velocity integrated by the rise-time. Thus derived deformations at the location of PG1 are estimated to be (a) 0.65 m and (b) 0.06 m, respectively. These values almost coincide with (a) the static deformation from the fault plate model by Geospatial Information Authority of Japan (GSI) (2003) and (b) the point source equivalent to the seismic moment (Fig. 8). This means that the displacement of the location of the pressure sensor deployment can be roughly estimated in terms of the water

 *=101.5M-1.4/80* (2)

(a) 26 September 2003 (M8.0) (b) 11 September 2008 (M7.1)

Fig. 8. Deformation patterns from the seismic faults' dislocation

potential indicator of the tsunami generation.

JMA and JAMSTEC off Kii peninsula (Fig. 9).

**4. Tsunami prediction along the Nankai trough** 

**4.1 Tsunami monitoring system in the Nankai trough** 

An early tsunami detection approach based on a physical phenomenon uniquely observed in the source during the tsunamigenic earthquake was presented in this section. Tsunami initial waveform is mostly depended on the static deformation of the ocean bottom. Hence the amplitude of the water pressure associated with the acoustic resonant wave may be a

The first offshore observatory in Japan has been deployed in the Suruga trough targetting the presumed Tokai earthquake, central Japan, and followed by seven cabled observatories. The newest system is being operated in the presumed Tonankai earthquake sourece area by

The Nankai trough is one of the palte subduction zones in Japan, where the last megathrust earthquakes took place in 1944 and 1946, namely the Tonankai eartqhauke and the Nankai

pressure amplitude.

Fig. 9. Deformation patterns from the seismic faults' dislocation

earthquake, respectively. Because more than 60 years have past since the last earthquake, Japanese government evaluates that the probability of the next presumed megathrust earthquake along the Nankai trough is estimated to be 60-70 % within the next 30 years. Japanese government has constructed an offshore observatory network, which consists of dense 20 bottom seismic sensors and bottom pressure sensors in total in order for monitoring seismic activity and its consequence, megathrust earthquake, and followed by tsunami. JAMSTEC is operating the offshore observatory network. The observatory layout is shown in Fig. 10. As of May 2011, 17 observatories have been deployed, and it started to acquire their data in real-time. If megathrust earthquake and accompanied giant tsunami would be predicted before their arrival nearby the coast and effective warning would be issued, it must contribute to mitigate earthquake and tsunami related disasters. We should establish measurement technology including data processing and accumulate technical know-how for future meagathrust earthquake and tsunami in advance; hence we carry out tsunami computation of the last 1944 Tonankai earthquake.


Table 3. Fault parameters used in the tsunami computation from the Tonankai earthquake.

Advances for Tsunami Measurement Technologies and Its Applications 15

T = 20 s T = 20 s

T = 40 s T = 40 s

Snapshots of the tsunami generation are compared in Fig. 11. Although tsunami is generated after the fault rupture halt at 20 s in the both models, tsunami height in the source area is different. In the source area, dynamic effect is significantly appeared. This is because the acoustic wave by the seismic wave is superposed. At 40 s, the water wave propagating to SE direction is computed, which is Rayleigh wave. As for amplitude and source area of

Pressure (x105Pa)

Fig. 12. Computed pressure waveforms during the tsunami generation at each observatory.

Time (s)

0 50 100 150 200 250 300

Deformation + Seismic wave

Time histories of water pressure at the observatories are shown in Fig. 12. In the case that only the bottom deformation is input, acoustic resonant wave is reproduced during the

Fig. 11. Snapshots of wave height during the tsunami generation.

the tsunami are not so different each other.

Deformation

Pressure (x104Pa)

Numberings represent that of the observatories in Fig. 10.

Time (s)

0 50 100 150 200 250 300

Deformation Deformation + Seismic wave

Fig. 10. Observaroty layout (open circles) and coseismic deformation caused by the seismic fault.

#### **4.2 Tsunami computation from the presumed Tonankai earthquake**

The latest study implies that the splay fault might contribute to the tsunami generataion process in addition to the main fault during the 1944 Tonankai earthquake (Park et al., 2002), however a simplified fault model is assumed and the pressure waveform is computed at the 20 observatories in the present study. Fault parameters are based on what has been estimated by Kanamori (1972) and they are listed in Table 3. Geometric relation between the fault plane and the observatories is shown in Fig. 10. It is assumed that the fault rupture starts from the bottom at the fault plane and propagate toward the top along the width direction, meaning uni-lateral faulting. Dynamic tsunami computation developed by Ohmachi et al. (2001) is applied to the present scenario. Dynamic tsunami computation can demonstrate fluid dynamic response due to the seismic fault rupture considering both the static deformation and the seismic wave in the tsunami computation. Dynamic tsunami computation can reproduce the bottom pressure because realistic 3D fluid domain is modeled.

Two different tsunami generation models are computed in the present study. One model is that the static deformation is given as a ramp-time function into the bottom of the fluid domain. The duration time, i.e., elapsed time to generate tsunami initial shape is assumed to be equal to the source time of the seismic faulting. Because the fault width and rupture velocity are assumed to be 60 km and 3 km/s, respectively, the time duration is solved to be 20 s divided by two parameters. In this model, the dynamic contribution of the ocean bottom is considered, but the seismic wave associated with the fault rupturing is not considered. Another model is that the seismic wave due to the fault rupturing is also considered, in which the ocean bottom is not displaced simultaneously in the tsunami source area. This model can demonstrate more realistic tsunami generation process than the former one.

Fig. 10. Observaroty layout (open circles) and coseismic deformation caused by the seismic

The latest study implies that the splay fault might contribute to the tsunami generataion process in addition to the main fault during the 1944 Tonankai earthquake (Park et al., 2002), however a simplified fault model is assumed and the pressure waveform is computed at the 20 observatories in the present study. Fault parameters are based on what has been estimated by Kanamori (1972) and they are listed in Table 3. Geometric relation between the fault plane and the observatories is shown in Fig. 10. It is assumed that the fault rupture starts from the bottom at the fault plane and propagate toward the top along the width direction, meaning uni-lateral faulting. Dynamic tsunami computation developed by Ohmachi et al. (2001) is applied to the present scenario. Dynamic tsunami computation can demonstrate fluid dynamic response due to the seismic fault rupture considering both the static deformation and the seismic wave in the tsunami computation. Dynamic tsunami computation can reproduce the bottom pressure because realistic 3D fluid domain is

Two different tsunami generation models are computed in the present study. One model is that the static deformation is given as a ramp-time function into the bottom of the fluid domain. The duration time, i.e., elapsed time to generate tsunami initial shape is assumed to be equal to the source time of the seismic faulting. Because the fault width and rupture velocity are assumed to be 60 km and 3 km/s, respectively, the time duration is solved to be 20 s divided by two parameters. In this model, the dynamic contribution of the ocean bottom is considered, but the seismic wave associated with the fault rupturing is not considered. Another model is that the seismic wave due to the fault rupturing is also considered, in which the ocean bottom is not displaced simultaneously in the tsunami source area. This model can demonstrate more realistic tsunami generation process than the

**4.2 Tsunami computation from the presumed Tonankai earthquake** 

fault.

modeled.

former one.

Fig. 11. Snapshots of wave height during the tsunami generation.

Snapshots of the tsunami generation are compared in Fig. 11. Although tsunami is generated after the fault rupture halt at 20 s in the both models, tsunami height in the source area is different. In the source area, dynamic effect is significantly appeared. This is because the acoustic wave by the seismic wave is superposed. At 40 s, the water wave propagating to SE direction is computed, which is Rayleigh wave. As for amplitude and source area of the tsunami are not so different each other.

Fig. 12. Computed pressure waveforms during the tsunami generation at each observatory. Numberings represent that of the observatories in Fig. 10.

Time histories of water pressure at the observatories are shown in Fig. 12. In the case that only the bottom deformation is input, acoustic resonant wave is reproduced during the

Advances for Tsunami Measurement Technologies and Its Applications 17

Geospatial Information Authority of Japan. (2003). Press release on 26 September 2003.

Gonzalez, F. I, Bernard, E. N., Meinig, C., Eble, M., Mofjeld, H. O. & Stalin S. (2005) The

Hino, R., Tnioka, Y., Kanazawa, T., Sakai, S, Nishino, M, & Suyehiro, K. (2001). Micro-

Hirata, K., Aoyagi, M., Mikada, H., Kawaguchi, K., Kaiho, Y., Iwase, R., Morita, S., Fujisawa,

Kanamori, H. (1972). Tectonic implication of the 1944 Tonankai and the 1946 Nankaido

Kato, T., Terada, Y., Kinosita, M., Kakimoto, H., Issiki, H., Matsuishi, M., Yokoyama, A., &

Kato, T., Terada, Y, Ito, K., Hattori, R., Abe, T., Miyake, T., Koshimura, S., & Nagai, T. (2005).

Münch, U., Rudoloff, A. & Lauterjung, J. (2011). The GITEWS Project – results, summary

Matsumoto, H., & Mikada, H. (2005). Fault geometry of the 2004 off the Kii peninsula

Matsumoto, H., Tanioka, Y., Nishimura, Y., Tsuji, Y., Namegaya, Y., Nakasu, T., & Iwasaki,

Merrifield, M. A., Firing, Y. L., Aarup, T., Agricole, W., Brundrit, G., Chang-Seng, D., Farre,

Namaegaya, Y., Tanioka, Y., Abe, K., Satake, K., Hirata, K., Okada, M. & Gusman, A., R.

recorded by a new GPS buoy, *Earth Planets Space*, 57, 297-301

and outlook, *Nat. Hazards Earth Syst. Sci.*, 11, 765-769

observation in northeastern Japan, *Geophys. Res. Lett.*, 28, 3533-3536

the southern Kurile subduction zone, *IEEE J. Ocean. Eng.*, 27, 170-181 Hirata K., Takahashi, H., Geist, E. L., Satake, K., Tanioka, Y., Sugioka, H., & Mikada, H.

tsunami from a local interplate earthquake detected by cabled offshore tsunami

I., Sugioka, H., Mitsuzawa, K., Suyehiro, K., Kinoshita, H., & Fujiwara, N. (2002). Real-time geophysical measurements on the deep seafloor using submarine cable in

(2003). Source depth dependence of micro-tsunamis recorded with ocean-bottom pressure gauges: the January 28, 2000 Mw 6.8 earthquake off Nemuro Peninsula,

Tanno, T. (2000). Real-time observation of tsunami by RTK-GPS, *Earth Planets Space*,

Tsunami due to the 2004 September 5th off the Kii peninsula earthquake, Japan,

earthquake inferred from offshore pressure waveforms, *Earth Planets Space*, 57, 161-

S. (2009). Review of tide gauge records in the Indian Ocean, *J. Earthquake Tsunami*,

R., Kilonsky, B., Knight, W, Kong, L., Magori, C., Manurung, P., McCreery C., Mitchell, W., Pillay, S., Schindele, F., Shillington, F., Testut, L., Wijeratne, E. M. S., Galdwell, P., Jardin, J., Nakahara S., Porter F. Y., & Tuetsky, N. (2005). Tide gauge observations of the Indian Ocean tsunami, December 26, 2004, *Geophys. Res. Lett.*,

(2009). In situ Measurements of Tide Gauge Response and Corrections of Tsunami Waveforms from the Niigataken Chuetsu-oki Earthquake in 2007, *Pure Appl.* 

Filloux, J. H. (1982). Tsunami recorded on the open ocean floor, *Geophs. Res. Lett.*, 9, 25-28 Geospatial Information Authority of Japan. (online). History of Tide Gauges, available from

http://www.gsi.go.jp/WNEW/PRESS-RELEASE/2003-0926-2.html

NTHMP tsunameter network, *Nat. Hazards*, 35, 25-39

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3, 1-15

tsunami generation. The amplitude corresponds to the distribution of the static deformation due to the seismic faulting. On the other hand, in the case that seismic wave is input to the fluid domain either, water pressure fluctuation associated with the Rayleigh wave is reproduced in addition to the water resonant wave. It should be noted that the maximum amplitude (~ a few of 105 Pa) is fairly equal to that of the experienced in the 2003 earthquake discussed in the previous section. The amplitude is obviously large at the offshore observatories such as 9, 10, 11, and 12 sites in the Nankai trough. These observatories are located in deeper area than others, hence the water pressures tend to be amplified by the long period Rayleigh wave. The acoustic resonant wave is an unique phenomenon during the tsunami generation process. Precise measurement of the acoustic resonant would provide tsunami generation prediction in advance.

#### **5. Conclusion**

This chapter reviews some tsunami measurements being in operation. Traditional tide gauge deployed at the coast is unable to perform early tsunami detection because of its deployed location. Recent offshore tsunami measurement technologies such as bottom pressure sensor and kinematic GPS buoy enabled to detect far-field tsunamis before its arrival at the coast. As an on-going study, HF radar to detect tsunami current approaching coast at long ranges is being developed and theoretically examined in the Atlantic Ocean (Dzvonkovskaya and Gurgel, 2009). More recently, electromagnetic (EM) sensors eventually could detect tsunami signals associated with its water mass passage from the 2006 and 2007 Kuil Is. earthquakes (Toh et al., 2011). Thus new tsunami measurement technologies and relevant sensors have been developed and applied for early tsunami detection in order for improving conventional tsunami warning system using bottom pressure sensors.

Then the present chapter introduces the actual observation of the bottom pressure sensors deployed in the tsunami source area. The acoustic resonant wave that is significantly produced in the tsunami generation process may contribute to the early tsunami detection scheme. Tsunami from the presumed Tonankai earthquake being though to take place within a next few decades is computed, which predict pressure waveforms at the offshore observatory in the source area. Acoustic resonant wave is computed in the tsunami source area, suggesting that its large amplitude implies large deformation. This gives an opportunity to issue an automated tsunami alert during the real-time monitoring by the bottom pressure sensors in the tsunami source area.

#### **6. Acknowledgment**

This study was partly supported by Grant-in-Aid for Young Scientist (B) 22710175 of the Ministry of Education, Culture, Sports, Science and Technology (MEXT), Japan. Some figures were prepared by Generic Mapping Tools (Wessel and Smith, 1995).

#### **7. References**

Dzvonkovskaya, A., & Gurgel, K. W. (2009). Future contribution of HF radar WERA to tsunami early warning systems, *European J. Navigation*, 7 (2), 17-23

Eble, M. C., & Gonzalez, F. I. (1991). Deep-Ocean bottom pressure measurements in the northeast Pacific, *J. Atom. Ocean. Tech.*, 8, 221-233

tsunami generation. The amplitude corresponds to the distribution of the static deformation due to the seismic faulting. On the other hand, in the case that seismic wave is input to the fluid domain either, water pressure fluctuation associated with the Rayleigh wave is reproduced in addition to the water resonant wave. It should be noted that the maximum amplitude (~ a few of 105 Pa) is fairly equal to that of the experienced in the 2003 earthquake discussed in the previous section. The amplitude is obviously large at the offshore observatories such as 9, 10, 11, and 12 sites in the Nankai trough. These observatories are located in deeper area than others, hence the water pressures tend to be amplified by the long period Rayleigh wave. The acoustic resonant wave is an unique phenomenon during the tsunami generation process. Precise measurement of the acoustic resonant would

This chapter reviews some tsunami measurements being in operation. Traditional tide gauge deployed at the coast is unable to perform early tsunami detection because of its deployed location. Recent offshore tsunami measurement technologies such as bottom pressure sensor and kinematic GPS buoy enabled to detect far-field tsunamis before its arrival at the coast. As an on-going study, HF radar to detect tsunami current approaching coast at long ranges is being developed and theoretically examined in the Atlantic Ocean (Dzvonkovskaya and Gurgel, 2009). More recently, electromagnetic (EM) sensors eventually could detect tsunami signals associated with its water mass passage from the 2006 and 2007 Kuil Is. earthquakes (Toh et al., 2011). Thus new tsunami measurement technologies and relevant sensors have been developed and applied for early tsunami detection in order for

Then the present chapter introduces the actual observation of the bottom pressure sensors deployed in the tsunami source area. The acoustic resonant wave that is significantly produced in the tsunami generation process may contribute to the early tsunami detection scheme. Tsunami from the presumed Tonankai earthquake being though to take place within a next few decades is computed, which predict pressure waveforms at the offshore observatory in the source area. Acoustic resonant wave is computed in the tsunami source area, suggesting that its large amplitude implies large deformation. This gives an opportunity to issue an automated tsunami alert during the real-time monitoring by the

This study was partly supported by Grant-in-Aid for Young Scientist (B) 22710175 of the Ministry of Education, Culture, Sports, Science and Technology (MEXT), Japan. Some

Dzvonkovskaya, A., & Gurgel, K. W. (2009). Future contribution of HF radar WERA to

Eble, M. C., & Gonzalez, F. I. (1991). Deep-Ocean bottom pressure measurements in the

tsunami early warning systems, *European J. Navigation*, 7 (2), 17-23

figures were prepared by Generic Mapping Tools (Wessel and Smith, 1995).

northeast Pacific, *J. Atom. Ocean. Tech.*, 8, 221-233

improving conventional tsunami warning system using bottom pressure sensors.

provide tsunami generation prediction in advance.

bottom pressure sensors in the tsunami source area.

**6. Acknowledgment** 

**7. References** 

**5. Conclusion** 

Filloux, J. H. (1982). Tsunami recorded on the open ocean floor, *Geophs. Res. Lett.*, 9, 25-28


http://www.gsi.go.jp/WNEW/PRESS-RELEASE/2003-0926-2.html


**0**

**2**

*France*

Giovanni Occhipinti

*Institut de Physique du Globe de Paris*

**Tsunami Detection by Ionospheric Sounding:**

After the Great Sumatra Earthquake and the consequent Indian Ocean Tsunami scientific community puts their attention to alternative methods in ocean monitoring to improve the

Improvement of classic techniques, as the seismic source estimation (e.g., Ammon et al., 2006) and densification of number of buoys over the oceans (Gonzalez et al., 2005), was supported by a new effort in remote sensing: nominally the space altimetry observation of the tsunami in the open sea (Okal et al., 1999; Smith et al., 2005) and the tsunami detection by ionospheric monitoring (e.g., Occhipinti et al., 2006). Today the recent tsunamis declare, one times more,

The indirect tsunami observation by ionospheric sounding is based on the idea anticipated in the past by Hines (1972) and Peltier & Hines (1976) that tsunamis produce internal gravity waves (IGWs) in the overlooking atmosphere. During the upward propagation the IGWs are strongly amplified by the effect of the decrease of the density. The interaction of IGWs with the plasma at the ionospheric height produces strongly variation in the plasma velocity and

This chapter i) resumes the moderne debate based on the Sumatra event (2004) about the tsunami detection by ionospheric sounding to demonstrate the hypothesis anticipated by Peltier & Hines (1976), and identifies the technics that proved and validated it, nominally altimeters and GPS. ii) Supports, with the recent theoretical works, the coupling between the ocean, the neutral atmosphere and the ionospheric plasma during the tsunami propagation and explores, based on the numerical modeling, the remote sensing possibility with additional techniques as the over the horizon radar (OTH-R). iii) Presents the ionospheric observations of the recents tsunamis to prove the systematic detection capability; nominally we review the following tsunamigenic earthquakes: 12 September, 2007, in Sumatra; the 14 November, 2007, in Chile; the 29 September, 2009, in Samoa; and the recent Tohoku-Oki (Japan) earthquake on 11 Mars 2011. We anticipate here that this last event also allow to prove that the signature of

tsunami in the ionosphere can be also detected by optical camera *via* the airglow.

It finally concludes discussing the role of ionospheric sounding and remote sensing in the

**1. Introduction**

response of the tsunami warning systems.

the importance to go forward in this direction.

plasma density observable by ionospheric sounding (Figure 1).

modern evolution of tsunami detection and warning systems.

**New Tools for Oceanic Monitoring**


### **Tsunami Detection by Ionospheric Sounding: New Tools for Oceanic Monitoring**

Giovanni Occhipinti *Institut de Physique du Globe de Paris France*

#### **1. Introduction**

18 Tsunami – A Growing Disaster

Nosov, M. A., & Kolesov, S. V. (2007). Elastic ocsillations of water column in the 2003

Ohmachi, T., Tsukiyama, H., & Matsumoto, H. (2001). Simulation of tsunami induced by

Okada, M. (1985). Response of some tide-wells in Japan to tsunamis, Proc. Int. Tsunami

Park, J. O., Tsuru, T., Kodaira, S., Cummins, P. R., & Kaneda, Y. (2002). Splay fault branching

Rudloff, A, Lauterjung, J. Münch, U. & Tinti, S. (2009). The GITEWS Project (German-

Satake, K., Baba, T., Hirata, K., Iwasaki, S., Kato, T., Koshimura, S., Takenaka, J., & Terada,

Toh, H., Satake, K., Hmano, Y., Fujii, Y. & Goto, T. (2011). Tsunami signals from the 2006

Watanabe, T., Matsumoto, H., Sugioka, H., Mikada, H., Suyehiro, K., & Otsuka, R. (2004)

Wessel, P. & Smith W. H. F. (1995). New version of the Generic Mapping Tools released,

along the Nankai subduction zone, *Science*, 297, 1157-1160

*Nat. Hazards Earth Syst. Sci.*, 7, 243-249

magnitude, *J. Phys. Earth*, 27, 353-372

island, *EOS Trans. AGU*, 85, 14

*EOS Trans. AGU*, 76, 329.

*Geophs. Res.*, 116, B02104, doi:10.1029/2010JB007873

1898-1909

1882

Symp., 208-213

Tokachi-oki tsunami sourece : in-situ measument and 3-D numerical modeling,

dynamic displacement of seabed due to seismic faulting, *Bull. Seism. Soc. Am.*, 91,

Indonesian Tsunami Early Warning System), *Nat. Hazards Earth Syst. Sci.*, 9, 1381-

Y. (2005). Tsunami source of the 2004 off the Kii Peninsura earthquakes inferred from offshore tsunami oand coastal tide gages, *Earth Planets Space*, 57, 173-178 Sato, R. (1979). Theoretical baseis on relationship between focal parameters and earthquak

and 2007 Kuril earthquakes detected at a seafloor geomagnetic observatory, *J.* 

Offshore monitoring system records recent earthquake off Japan's northernmost

After the Great Sumatra Earthquake and the consequent Indian Ocean Tsunami scientific community puts their attention to alternative methods in ocean monitoring to improve the response of the tsunami warning systems.

Improvement of classic techniques, as the seismic source estimation (e.g., Ammon et al., 2006) and densification of number of buoys over the oceans (Gonzalez et al., 2005), was supported by a new effort in remote sensing: nominally the space altimetry observation of the tsunami in the open sea (Okal et al., 1999; Smith et al., 2005) and the tsunami detection by ionospheric monitoring (e.g., Occhipinti et al., 2006). Today the recent tsunamis declare, one times more, the importance to go forward in this direction.

The indirect tsunami observation by ionospheric sounding is based on the idea anticipated in the past by Hines (1972) and Peltier & Hines (1976) that tsunamis produce internal gravity waves (IGWs) in the overlooking atmosphere. During the upward propagation the IGWs are strongly amplified by the effect of the decrease of the density. The interaction of IGWs with the plasma at the ionospheric height produces strongly variation in the plasma velocity and plasma density observable by ionospheric sounding (Figure 1).

This chapter i) resumes the moderne debate based on the Sumatra event (2004) about the tsunami detection by ionospheric sounding to demonstrate the hypothesis anticipated by Peltier & Hines (1976), and identifies the technics that proved and validated it, nominally altimeters and GPS. ii) Supports, with the recent theoretical works, the coupling between the ocean, the neutral atmosphere and the ionospheric plasma during the tsunami propagation and explores, based on the numerical modeling, the remote sensing possibility with additional techniques as the over the horizon radar (OTH-R). iii) Presents the ionospheric observations of the recents tsunamis to prove the systematic detection capability; nominally we review the following tsunamigenic earthquakes: 12 September, 2007, in Sumatra; the 14 November, 2007, in Chile; the 29 September, 2009, in Samoa; and the recent Tohoku-Oki (Japan) earthquake on 11 Mars 2011. We anticipate here that this last event also allow to prove that the signature of tsunami in the ionosphere can be also detected by optical camera *via* the airglow.

It finally concludes discussing the role of ionospheric sounding and remote sensing in the modern evolution of tsunami detection and warning systems.

Fig. 2. Left-small-panels: TEC variations plotted at the ionospheric piercing points. A wave-like disturbance is propagating towards the coast of Japan. The perturbation presents characteristics of a tsunami IGW, and arrives approximately at the same time as the tsunami. Right-small-panels: Waves observed on the TEC maps throughout June 24th, 2001. The thickness of the arrows indicate the approximate amplitude of the wave (lower than 0.75 TECU, between 0.75 and 1.5 TECU, and between 1.5 and 2.25 TECU). The direction is the azimuth, and the lenght is proportional to the speed. Finally, the color indicate the time of observation (reddish colors are the local day time, blue is nighttime). The ellipse shows the

Tsunami Detection by Ionospheric Sounding: New Tools for Oceanic Monitoring 21

possible tsunami signal showed in the left-panels. Figure after Artru et al. (2005).

al., 2006a;b; Lognonné et al., 2006; Occhipinti et al., 2006; 2008b).

atmospheric gravity waves generated at the epicenter.

<sup>1</sup> The TEC is expressed in TEC units (TECU); 1 TECU = 1016*e*−/*m*2.

The giant tsunami following the Sumatra-Andaman event (Mw=9.3, 0:58:50 UT, 26 December, 2004 (Lay et al., 2005)), an order of magnitude larger than the Peruvian tsunami, provided worldwide remote sensing observations in the ionosphere, giving the opportunity to explore ionospheric tsunami detection with a vast data set (Fig. 3). In addition to seismic waves detected by global seismic networks (Park et al., 2005); co-seismic displacement measured by GPS (Vigny et al., 2005); oceanic sea surface variations measured by altimetry (Smith et al., 2005); detection of magnetic anomaly (Balasis & Mandea, 2007; Iyemori et al., 2005) and acoustic-gravity waves (Le Pichon et al., 2005); a series of ionospheric disturbances, observed with different techniques, have been reported in the literature (DasGupta et al., 2006; Liu et

Two ionospheric anomalies in the plasma velocities were detected North of the epicenter by a Doppler sounding network in Taiwan (Liu et al., 2006a). The first was triggered by the vertical displacement induced by Rayleigh waves. The second, arriving one hour later with a longer period, is interpreted by Liu *et al.* (2006a) as the response of ionospheric plasma to the

A similarly long period perturbation, with an amplitude of 4 TECU1 peak-to-peak, was observed by GPS stations located on the coast of India (DasGupta et al., 2006). These perturbations could be the ionospheric signature of IGWs coupled at sea level with the tsunami or the atmospheric gravity waves generated at the epicenter. Comparable TEC observations were done for five GPS stations (twelve station-satellite couples) scattered in the Indian Ocean (Liu et al., 2006b). The 30 sec differential amplitudes are equal to or smaller than

Fig. 1. Schematic view of the coupling mechanism and the ionospheric sounding by GPS. The vertical displacement of the ground floor (1) produced by an earthquake is directly transfered at the sea surface (2) following the incompressible hypothesis. The sea surface displacement initiate an internal gravity wave (IGW) propagating into the ocean (tsunami) as well as into the overlooking atmosphere. During the upward propagation the atmospheric IGW interact with the ionospheric plasma (3) creating perturbation in the plasma density and consequently in the local refraction index. The electromagnetic waves emitted by GPS satellites (4) to the ground stations (5) are perturbed by the plasma density variations and are able to image the signature of the IGW in the ionosphere.

#### **2. The modern debate**

The encouraging work of Artru et al. (2005) on the detection of the peruvian tsunamigenic quake on 23 June, 2001 (M=8.4 at 20:33 UT) in the total electron content (TEC) measured by the japanese dense GPS network GEONET opens the modern debate about the feasibility of tsunami detection by ionospheric sounding.

In essence, Artru et al. (2005) shows ionospheric traveling waves reaching the Japanese coast 22 hours after the peruvian tsunamigenic quake, with an azimuth and arrival time consistent with tsunami propagation (Fig. 2). Moreover, a period between 22 and 33 min, consistent with the tsunami, was identified in the observed TEC signals. The tsunami generated internal gravity waves (IGWs) were, however, superimposed by other signals associated with traveling ionospheric disturbances (TIDs) (Balthazor & Moffett, 1997). The ionospheric noise is large in the gravity domain (Garcia et al., 2005), consequently the identification of the tsunami signature in the TEC could be doubtful, and the debate still open.

2 Will-be-set-by-IN-TECH

Fig. 1. Schematic view of the coupling mechanism and the ionospheric sounding by GPS. The

The encouraging work of Artru et al. (2005) on the detection of the peruvian tsunamigenic quake on 23 June, 2001 (M=8.4 at 20:33 UT) in the total electron content (TEC) measured by the japanese dense GPS network GEONET opens the modern debate about the feasibility of

In essence, Artru et al. (2005) shows ionospheric traveling waves reaching the Japanese coast 22 hours after the peruvian tsunamigenic quake, with an azimuth and arrival time consistent with tsunami propagation (Fig. 2). Moreover, a period between 22 and 33 min, consistent with the tsunami, was identified in the observed TEC signals. The tsunami generated internal gravity waves (IGWs) were, however, superimposed by other signals associated with traveling ionospheric disturbances (TIDs) (Balthazor & Moffett, 1997). The ionospheric noise is large in the gravity domain (Garcia et al., 2005), consequently the identification of the

tsunami signature in the TEC could be doubtful, and the debate still open.

vertical displacement of the ground floor (1) produced by an earthquake is directly transfered at the sea surface (2) following the incompressible hypothesis. The sea surface displacement initiate an internal gravity wave (IGW) propagating into the ocean (tsunami) as well as into the overlooking atmosphere. During the upward propagation the atmospheric IGW interact with the ionospheric plasma (3) creating perturbation in the plasma density and consequently in the local refraction index. The electromagnetic waves emitted by GPS satellites (4) to the ground stations (5) are perturbed by the plasma density variations and are

able to image the signature of the IGW in the ionosphere.

tsunami detection by ionospheric sounding.

**2. The modern debate**

Fig. 2. Left-small-panels: TEC variations plotted at the ionospheric piercing points. A wave-like disturbance is propagating towards the coast of Japan. The perturbation presents characteristics of a tsunami IGW, and arrives approximately at the same time as the tsunami. Right-small-panels: Waves observed on the TEC maps throughout June 24th, 2001. The thickness of the arrows indicate the approximate amplitude of the wave (lower than 0.75 TECU, between 0.75 and 1.5 TECU, and between 1.5 and 2.25 TECU). The direction is the azimuth, and the lenght is proportional to the speed. Finally, the color indicate the time of observation (reddish colors are the local day time, blue is nighttime). The ellipse shows the possible tsunami signal showed in the left-panels. Figure after Artru et al. (2005).

The giant tsunami following the Sumatra-Andaman event (Mw=9.3, 0:58:50 UT, 26 December, 2004 (Lay et al., 2005)), an order of magnitude larger than the Peruvian tsunami, provided worldwide remote sensing observations in the ionosphere, giving the opportunity to explore ionospheric tsunami detection with a vast data set (Fig. 3). In addition to seismic waves detected by global seismic networks (Park et al., 2005); co-seismic displacement measured by GPS (Vigny et al., 2005); oceanic sea surface variations measured by altimetry (Smith et al., 2005); detection of magnetic anomaly (Balasis & Mandea, 2007; Iyemori et al., 2005) and acoustic-gravity waves (Le Pichon et al., 2005); a series of ionospheric disturbances, observed with different techniques, have been reported in the literature (DasGupta et al., 2006; Liu et al., 2006a;b; Lognonné et al., 2006; Occhipinti et al., 2006; 2008b).

Two ionospheric anomalies in the plasma velocities were detected North of the epicenter by a Doppler sounding network in Taiwan (Liu et al., 2006a). The first was triggered by the vertical displacement induced by Rayleigh waves. The second, arriving one hour later with a longer period, is interpreted by Liu *et al.* (2006a) as the response of ionospheric plasma to the atmospheric gravity waves generated at the epicenter.

A similarly long period perturbation, with an amplitude of 4 TECU1 peak-to-peak, was observed by GPS stations located on the coast of India (DasGupta et al., 2006). These perturbations could be the ionospheric signature of IGWs coupled at sea level with the tsunami or the atmospheric gravity waves generated at the epicenter. Comparable TEC observations were done for five GPS stations (twelve station-satellite couples) scattered in the Indian Ocean (Liu et al., 2006b). The 30 sec differential amplitudes are equal to or smaller than

<sup>1</sup> The TEC is expressed in TEC units (TECU); 1 TECU = 1016*e*−/*m*2.

0.4 TECU (which generates amplitudes comparable to the DasGupta et al. (2006) observations for periods of ≈165 min, *i.e.* 30 points) and the arrival times coherent with the tsunami propagation. The observed satellites were located approximately at the station zenith. Comparison between oceanic sea-level measured by tide-gauge at Coco Island and the TEC measured by the co-located GPS shows similarity in the waveform suggesting that the ionosphere is sensitive to the tsunami propagation as well as the ocean (Occhipinti et al., 2008b). We highlight that the tsunami reaches Coco Island 3 hours after the tsunami generation, this is the first oceanic observation of the Sumatra tsunami (Titov et al., 2005). Close to these observations, the Topex/Poseidon and Jason-1 satellites acquired the key observations of the Sumatra tsunami with altimetry profiles. The measured sea level displacement is well explained by tsunami propagation models with realistic bathymetry, and provides useful constraints on source mechanism inversions (e.g. Song *et al.*, 2005). In addition, the inferred TEC data, required to remove the ionospheric effects from the altimetric measurements (Imel, 1994), showed strong anomalies in the integrated electron

Tsunami Detection by Ionospheric Sounding: New Tools for Oceanic Monitoring 23

In essence, altimetric data from Topex/Poseidon and Jason-1 shows at the same time the tsunami signature on the sea surface and the supposed tsunami signature in the ionosphere (Fig. 4). By a three-dimentional numerical modeling Occhipinti et al. (2006) compute the atmospheric IGWs generated by the Sumatra tsunami and their interaction with the ionospheric plasma. The quantitative approach reproduces the TEC observed by Topex/Poseidon and Jason-1 in the Indian Ocean the 26 December 2004. Consequently, Occhipinti et al. (2006) closed the debate about the nature and the existence of the tsunami signature in the ionosphere. The results obtained by Occhipinti et al. (2006) was recently

The TEC observation close to the epicenter using the local GPS network SEAMARGES, shows an early signal appearing at around 20 min after the tsunami generation and observable during 1 hour (Fig. 3). This signal could be contain both, an acoustic-gravity wave perturbation directly link to the vertical displacement at the source, and the tsunami signature in the ionosphere (Occhipinti et al., 2011b). The systematic observation of this early TEC perturbation could be used for tsunami warning system purpose. Anyway, we highlight that today the acoustic-gravity wave signature in the TEC observed close to the epicenter has not

Tsunamis are long period oceanic gravity waves (Satake, 2002): their frequency is generally much smaller than the atmosheric Brünt-Vaïsalla frequency and, in the limit of linear analysis, they generate internal gravity waves in the overlying atmosphere (Hines, 1972; Lognonné et al., 1998; Occhipinti et al., 2006; 2008a). In other words, the coupling mechanism does not transfer a significant propagating energy in the acoustic domain. As a consequence of this theoretical hypothesis and the slow propagation velocity of IGW, a Bussinesq approximation, equivalent to incompressible fluid (Spiegel & Veronis, 1960), can be used in

Following those hypothesis Occhipinti et al. (2006; 2008a) developed a vertical

continuity equation and the incompressible hypothesis) explicitly described by the following

*dz* = *A* · *V* (based on the Navier Stokes equations, the

the ocean-atmosphere coupling mechanism and tsunami-IGW propagation.

density (Occhipinti et al., 2006).

reproduced by Mai & Kiang (2009).

been reproduced by modeling.

pseudo-spectral propagator *dV*

**3. Theoretical works**

Fig. 3. Main: Maximum sea-level perturbation model produced by the propagation os the Sumatra tsunami (26 December, 2004). Top: TEC perturbation appearing within 15 min after the tsunami generation. The ionospheric piercing points (IPPs) obtained by satellites PRN01, 03, 13, 19, 20 and 23 coupled with the SEAMARGES network (red point in the main figure) are showed here during 40 min and highlight a clear early perturbation moving from the epicenter (the red star) to the North of Sumatra. Fugure after Occhipinti et al. (2011b). Middle: TEC perturbation observed by DasGupta et al. (2006) with the 3 satellite-station couples showed in the main figure by colored diamonds (station location) and lines (satellites) in the main figure. DasGupta et al. (2006) explain this signal as the IGW generated at the source by the vertical displacement but not link to the tsunami. Bottom-right: Average horizontal speeds of TIDs (red line) and tsunami (black line). The used GPS stations are indicated by triangle in the main figure. Figure after Liu et al. (2006b). Bottom-left: Tsunami signal measured at Coco Island by the tide gauge (red) and by the co-located GPS receiver (blue). The tide gauge measures the sea level displacement (tsunami + tide) and the GPS measures the TEC perturbation in the ionosphere. Both waveforms are similar in showing the sensitivity of ionosphere to the tsunami structure. Figure after Occhipinti et al. (2008b)

0.4 TECU (which generates amplitudes comparable to the DasGupta et al. (2006) observations for periods of ≈165 min, *i.e.* 30 points) and the arrival times coherent with the tsunami propagation. The observed satellites were located approximately at the station zenith.

Comparison between oceanic sea-level measured by tide-gauge at Coco Island and the TEC measured by the co-located GPS shows similarity in the waveform suggesting that the ionosphere is sensitive to the tsunami propagation as well as the ocean (Occhipinti et al., 2008b). We highlight that the tsunami reaches Coco Island 3 hours after the tsunami generation, this is the first oceanic observation of the Sumatra tsunami (Titov et al., 2005).

Close to these observations, the Topex/Poseidon and Jason-1 satellites acquired the key observations of the Sumatra tsunami with altimetry profiles. The measured sea level displacement is well explained by tsunami propagation models with realistic bathymetry, and provides useful constraints on source mechanism inversions (e.g. Song *et al.*, 2005). In addition, the inferred TEC data, required to remove the ionospheric effects from the altimetric measurements (Imel, 1994), showed strong anomalies in the integrated electron density (Occhipinti et al., 2006).

In essence, altimetric data from Topex/Poseidon and Jason-1 shows at the same time the tsunami signature on the sea surface and the supposed tsunami signature in the ionosphere (Fig. 4). By a three-dimentional numerical modeling Occhipinti et al. (2006) compute the atmospheric IGWs generated by the Sumatra tsunami and their interaction with the ionospheric plasma. The quantitative approach reproduces the TEC observed by Topex/Poseidon and Jason-1 in the Indian Ocean the 26 December 2004. Consequently, Occhipinti et al. (2006) closed the debate about the nature and the existence of the tsunami signature in the ionosphere. The results obtained by Occhipinti et al. (2006) was recently reproduced by Mai & Kiang (2009).

The TEC observation close to the epicenter using the local GPS network SEAMARGES, shows an early signal appearing at around 20 min after the tsunami generation and observable during 1 hour (Fig. 3). This signal could be contain both, an acoustic-gravity wave perturbation directly link to the vertical displacement at the source, and the tsunami signature in the ionosphere (Occhipinti et al., 2011b). The systematic observation of this early TEC perturbation could be used for tsunami warning system purpose. Anyway, we highlight that today the acoustic-gravity wave signature in the TEC observed close to the epicenter has not been reproduced by modeling.

#### **3. Theoretical works**

4 Will-be-set-by-IN-TECH

Fig. 3. Main: Maximum sea-level perturbation model produced by the propagation os the Sumatra tsunami (26 December, 2004). Top: TEC perturbation appearing within 15 min after the tsunami generation. The ionospheric piercing points (IPPs) obtained by satellites PRN01, 03, 13, 19, 20 and 23 coupled with the SEAMARGES network (red point in the main figure) are showed here during 40 min and highlight a clear early perturbation moving from the epicenter (the red star) to the North of Sumatra. Fugure after Occhipinti et al. (2011b). Middle: TEC perturbation observed by DasGupta et al. (2006) with the 3 satellite-station couples showed in the main figure by colored diamonds (station location) and lines

(satellites) in the main figure. DasGupta et al. (2006) explain this signal as the IGW generated at the source by the vertical displacement but not link to the tsunami. Bottom-right: Average horizontal speeds of TIDs (red line) and tsunami (black line). The used GPS stations are indicated by triangle in the main figure. Figure after Liu et al. (2006b). Bottom-left: Tsunami signal measured at Coco Island by the tide gauge (red) and by the co-located GPS receiver (blue). The tide gauge measures the sea level displacement (tsunami + tide) and the GPS measures the TEC perturbation in the ionosphere. Both waveforms are similar in showing the sensitivity of ionosphere to the tsunami structure. Figure after Occhipinti et al. (2008b)

Tsunamis are long period oceanic gravity waves (Satake, 2002): their frequency is generally much smaller than the atmosheric Brünt-Vaïsalla frequency and, in the limit of linear analysis, they generate internal gravity waves in the overlying atmosphere (Hines, 1972; Lognonné et al., 1998; Occhipinti et al., 2006; 2008a). In other words, the coupling mechanism does not transfer a significant propagating energy in the acoustic domain. As a consequence of this theoretical hypothesis and the slow propagation velocity of IGW, a Bussinesq approximation, equivalent to incompressible fluid (Spiegel & Veronis, 1960), can be used in the ocean-atmosphere coupling mechanism and tsunami-IGW propagation.

Following those hypothesis Occhipinti et al. (2006; 2008a) developed a vertical pseudo-spectral propagator *dV dz* = *A* · *V* (based on the Navier Stokes equations, the continuity equation and the incompressible hypothesis) explicitly described by the following

vector *V* and matrix *A*:

Where *u*˜∗

et al. (2008a).

(Fig. 5):

computed from *u*˜*<sup>z</sup>* and *P*˜ as follow:

*V* = *u*˜∗ *z P*∗ ;

*A* = − 1 Ω *kx dux*<sup>0</sup> *dz* + *ky*

*i* Ω + *<sup>g</sup>* Ω *d* ln *ρ*<sup>0</sup> *dz*

*<sup>u</sup>*˜*<sup>x</sup>* <sup>=</sup> <sup>1</sup> <sup>Ω</sup>√*ρ*<sup>0</sup>

*<sup>u</sup>*˜*<sup>y</sup>* <sup>=</sup> <sup>1</sup> <sup>Ω</sup>√*ρ*<sup>0</sup>

*<sup>ρ</sup>*˜ <sup>=</sup> <sup>−</sup>*<sup>i</sup>* <sup>Ω</sup>√*ρ*<sup>0</sup>

take the form (1) and consequently the dispersion equation the form (2).

*k*2 *h N*<sup>2</sup> *<sup>ω</sup>*<sup>2</sup> <sup>−</sup> <sup>1</sup>

*<sup>ω</sup>*<sup>2</sup> <sup>=</sup> *<sup>k</sup>*<sup>2</sup>

*k*2 *<sup>z</sup>* + *k*<sup>2</sup> *<sup>h</sup>* + *<sup>N</sup>*<sup>2</sup> 2*g*

*kz* =

 − *i dux*<sup>0</sup> *dz <sup>u</sup>*˜ ∗ *<sup>z</sup>* + *kxP*˜<sup>∗</sup>

 − *i duy*<sup>0</sup> *dz <sup>u</sup>*˜ ∗ *<sup>z</sup>* + *kyP*˜<sup>∗</sup>

*dρ*<sup>0</sup> *dz <sup>u</sup>*˜ ∗ *z*

Following Occhipinti et al. (2008a; 2011b), in the case of linearized theory for a realistic atmosphere with horizontal stratification and no-background wind, the vertical *k*-number *kz*

*hN*<sup>2</sup>

Consequently it is possible to evaluate the vertical and horizontal group velocity *v<sup>z</sup>*

 − *N*<sup>2</sup> 2*g*

*duy*<sup>0</sup> *dz* 1 2 *d* ln *ρ*<sup>0</sup> *dz* <sup>−</sup> *<sup>i</sup>*(*k*<sup>2</sup>

Tsunami Detection by Ionospheric Sounding: New Tools for Oceanic Monitoring 25

*omega-k* domain: in essence the propagating plane waves with horizontal wave-numbers *kx*, *ky* and angular frequency *ω*; *g* is the gravity, *ρ*<sup>0</sup> is the unperturbed atmospheric density, *ux*<sup>0</sup> and *uy*<sup>0</sup> are the meridional and zonal background winds, and Ω = *ω* − *ux*0*kx* − *uy*0*ky* is the intrinsic frequency relative to the flow induced by the winds (Nappo, 2002). The effect of the wind on the IGW propagation is fully explored by Sun et al. (2007): in essence the IGW propagating against the wind is amplified, and slow-down compared to the IGW going in the same direction of the wind. This result is corroborated here by figure XX following Occhipinti

The methodology entirely described by Occhipinti et al. (2008a) can be simply resumed as follow: the vertical velocity field, induced by the sea motion during the tsunami propagation, is decomposed in planar waves by a three-dimensional Fourier-transform in the Cartesian frame (*x*, *y*) and time, where *x*ˆ and *y*ˆ are eastward and northward. In essence, Occhipinti et al. (2008a) imposes the continuity of vertical displacement between the ocean and the atmosphere. Injected in the propagator described above, it produce the tsunami-related IGW. The other two components of the perturbed velocity (*u*˜*x*, *u*˜*y*), and density perturbation *ρ*˜ are

<sup>−</sup><sup>1</sup>

*<sup>z</sup>* <sup>=</sup> <sup>√</sup>*ρ*0*u*˜*<sup>z</sup>* and *<sup>P</sup>*˜<sup>∗</sup> <sup>=</sup> *<sup>P</sup>*˜ <sup>√</sup>*ρ*<sup>0</sup> are normalized vertical velocity *<sup>u</sup>*˜*<sup>z</sup>* and pressure *<sup>P</sup>*˜ in the

2 *d* ln *ρ*<sup>0</sup> *dz*

*x*+*k*<sup>2</sup> *y* ) Ω

<sup>2</sup>

<sup>2</sup> (2)

(1)

*<sup>g</sup>* and *v<sup>h</sup> g*

Fig. 4. Top: Altimetric and TEC signatures of the Sumatra tsunami. The modelled and observed TEC are shown for (left) Jason-1 and (right) Topex/Poseidon: synthetic TEC (top-panels) without production-recombination-diffusion effects (blue), with production-recombination (red), and production-recombination-diffusion (green). The Topex/Poseidon synthetic TEC has been shifted up by 2 TEC units. (bottom-panels) The altimetric measurements of the ocean surface (black) are plotted for the Jason-1 and Topex/Poseidon satellites, respectively. The synthetic ocean displacements, used as the source of IGWs in the neutral atmosphere, are shown in red. For each plot from the latitude and corresponding Universal Time are shown. Bottom: Tsunami signature (right) in the TEC at 3:18 UT and (left) the unperturbed TEC. The TEC images have been computed by vertical integration of the perturbed and unperturbed electron density fields. The TEC perturbation induced by tsunami-coupled IGW is superimposed on a broad local-time (sunrise) TEC structure. The broken lines represent the Topex/Poseidon (left) and Jason-1 (right) trajectories. The blue contours represent the magnetic field inclination. Figures after Occhipinti et al. (2006).

vector *V* and matrix *A*:

6 Will-be-set-by-IN-TECH

Fig. 4. Top: Altimetric and TEC signatures of the Sumatra tsunami. The modelled and observed TEC are shown for (left) Jason-1 and (right) Topex/Poseidon: synthetic TEC

production-recombination (red), and production-recombination-diffusion (green). The Topex/Poseidon synthetic TEC has been shifted up by 2 TEC units. (bottom-panels) The altimetric measurements of the ocean surface (black) are plotted for the Jason-1 and Topex/Poseidon satellites, respectively. The synthetic ocean displacements, used as the source of IGWs in the neutral atmosphere, are shown in red. For each plot from the latitude and corresponding Universal Time are shown. Bottom: Tsunami signature (right) in the TEC at 3:18 UT and (left) the unperturbed TEC. The TEC images have been computed by vertical integration of the perturbed and unperturbed electron density fields. The TEC perturbation induced by tsunami-coupled IGW is superimposed on a broad local-time (sunrise) TEC structure. The broken lines represent the Topex/Poseidon (left) and Jason-1 (right) trajectories. The blue contours represent the magnetic field inclination. Figures after

(top-panels) without production-recombination-diffusion effects (blue), with

Occhipinti et al. (2006).

$$\begin{aligned} V &= \begin{pmatrix} \tilde{\mathfrak{u}}\_z^\* \\ P^\* \end{pmatrix}; \\ A &= \begin{pmatrix} -\frac{1}{\Omega} \left( k\_x \frac{d\mu\_{00}}{dz} + k\_y \frac{d\mu\_{00}}{dz} \right) & \frac{1}{2} \frac{d\ln\rho\_0}{dz} - \frac{i(k\_x^2 + k\_y^2)}{\Omega} \\ i \left( \Omega + \frac{g}{\Omega} \frac{d\ln\rho\_0}{dz} \right) & -\frac{1}{2} \frac{d\ln\rho\_0}{dz} \end{pmatrix}. \end{aligned}$$

Where *u*˜∗ *<sup>z</sup>* <sup>=</sup> <sup>√</sup>*ρ*0*u*˜*<sup>z</sup>* and *<sup>P</sup>*˜<sup>∗</sup> <sup>=</sup> *<sup>P</sup>*˜ <sup>√</sup>*ρ*<sup>0</sup> are normalized vertical velocity *<sup>u</sup>*˜*<sup>z</sup>* and pressure *<sup>P</sup>*˜ in the *omega-k* domain: in essence the propagating plane waves with horizontal wave-numbers *kx*, *ky* and angular frequency *ω*; *g* is the gravity, *ρ*<sup>0</sup> is the unperturbed atmospheric density, *ux*<sup>0</sup> and *uy*<sup>0</sup> are the meridional and zonal background winds, and Ω = *ω* − *ux*0*kx* − *uy*0*ky* is the intrinsic frequency relative to the flow induced by the winds (Nappo, 2002). The effect of the wind on the IGW propagation is fully explored by Sun et al. (2007): in essence the IGW propagating against the wind is amplified, and slow-down compared to the IGW going in the same direction of the wind. This result is corroborated here by figure XX following Occhipinti et al. (2008a).

The methodology entirely described by Occhipinti et al. (2008a) can be simply resumed as follow: the vertical velocity field, induced by the sea motion during the tsunami propagation, is decomposed in planar waves by a three-dimensional Fourier-transform in the Cartesian frame (*x*, *y*) and time, where *x*ˆ and *y*ˆ are eastward and northward. In essence, Occhipinti et al. (2008a) imposes the continuity of vertical displacement between the ocean and the atmosphere. Injected in the propagator described above, it produce the tsunami-related IGW. The other two components of the perturbed velocity (*u*˜*x*, *u*˜*y*), and density perturbation *ρ*˜ are computed from *u*˜*<sup>z</sup>* and *P*˜ as follow:

$$\begin{aligned} \tilde{u}\_{\chi} &= \frac{1}{\Omega \sqrt{\rho\_0}} \left( -i \frac{d\mu\_{x0}}{dz} \tilde{u}\_z^\* + k\_{\chi} \tilde{P}^\* \right) \\ \tilde{u}\_{\mathcal{Y}} &= \frac{1}{\Omega \sqrt{\rho\_0}} \left( -i \frac{d\mu\_{y0}}{dz} \tilde{u}\_z^\* + k\_{\mathcal{Y}} \tilde{P}^\* \right) \\ \tilde{\rho} &= \frac{-i}{\Omega \sqrt{\rho\_0}} \frac{d\rho\_0}{dz} \tilde{u}\_z^\* \end{aligned}$$

Following Occhipinti et al. (2008a; 2011b), in the case of linearized theory for a realistic atmosphere with horizontal stratification and no-background wind, the vertical *k*-number *kz* take the form (1) and consequently the dispersion equation the form (2).

$$k\_z = \sqrt{k\_h^2 \left(\frac{N^2}{\omega^2} - 1\right) - \left(\frac{N^2}{2g}\right)^2} \tag{1}$$

$$
\omega^2 = \frac{k\_h^2 N^2}{k\_z^2 + k\_h^2 + \left(\frac{N^2}{2g}\right)^2} \tag{2}
$$

Consequently it is possible to evaluate the vertical and horizontal group velocity *v<sup>z</sup> <sup>g</sup>* and *v<sup>h</sup> g* (Fig. 5):

Fig. 6. Vertical cross section of the modeled tsunami-related electron density perturbation and ray-paths computed using a 10 MHz OTH radar signal at 19*o*, 30*<sup>o</sup>* and 35*<sup>o</sup>* elevation (dashed gray lines). Dash-dotted purple line indicates a possible geometry of a nearby GPS station for a satellite at 25*<sup>o</sup>* elevation. Arrows indicate the tsunami and IGW energy directions of propagation. Note that the vertical scale has been exaggerated. Figure after

Tsunami Detection by Ionospheric Sounding: New Tools for Oceanic Monitoring 27

*ni vi* 

*dt* <sup>=</sup> −∇*pi* <sup>+</sup> *<sup>ρ</sup><sup>i</sup><sup>g</sup>* <sup>+</sup> *niqi*

3 ∑ *i*=1

The method developped by Occhipinti et al. (2006; 2008a) is also used to estimate the role of the geomagnetic field in the tsunami signature at the E-region and F-region (Occhipinti et al., 2008a). Nominally the authors show that the amplification of the electron density perturbation in the ionospheric plasma at the F-region is strongly dependent by geomagnetic inclination as well as by the direction of propagation of the tsunami. This effect is explained by the Lorenz force term in the momentum equation explaining the neutral plasma coupling (eq. 4). Consequently, the detection of tsunamigenic perturbation in the F-region-plasma is easily observed at equatorial and mid-latitude then the hight latitude. The heterogenic amplification drove by the magnetic field is not observable in the E-region, consequently detection at low altitude by Doppler sounding and over-the-horizon (OTH) radar are not affected by geographical location. The theoretical possibility of detection by OTH radar is explored by Coïsson et al. (2011) for a simple tsunami-related IGW (Fig. 6) propagating in an dynamic ionosphere. Coïsson et al. (2011) demonstrate that, in absence of noise, the 3-dimentional pattern of the emission/reception beam of the OTH radar don't hide the tsunami signature

*ne* =

<sup>=</sup> <sup>±</sup> *<sup>β</sup>ni* <sup>−</sup> *<sup>α</sup>n*<sup>2</sup>

<sup>−</sup> *<sup>ρ</sup>iμin*

 *<sup>E</sup>* <sup>+</sup>*vi* <sup>×</sup> *B* 

*vi* −*vn*

*<sup>i</sup>* (3)

*ni* (5)

(4)

Coïsson et al. (2011).

(Fig. 7).

the ionosperic plasma (eq. 5).

*∂ni <sup>∂</sup><sup>t</sup>* <sup>+</sup> ∇ ·

*ρi dvi*

Fig. 5. Vertical (left) and horizontal (right) group velocity of the internal gravity wave coupled at the sea surface with tsunamis (generated at different oceanic deep *h*, see gray-scale) and a characteristic period *T* of 10 min. Tsunamis move at the speed defined by the relation *vtsuna* = *hg*, where *<sup>g</sup>* is the gravity. Consequently, the horizontal *<sup>k</sup>*-vector *kh* that the tsunami transfer to the atmospheric internal gravity wave also depend by *h* following the relation *kh* = <sup>2</sup>*<sup>π</sup> T* <sup>√</sup>*hg* . Note that tsunamis generated/moving in the deeper oceanic zone produce faster IGW.

$$v\_{\mathcal{S}}^h = \frac{\delta \omega}{\delta k\_h} = \frac{k\_h N^2 (D - k\_h^2)}{\omega D^2} \qquad \qquad \qquad v\_{\mathcal{S}}^z = \frac{\delta \omega}{\delta k\_z} = \frac{k\_z k\_h^2 N^2}{\omega D^2}$$

where *D* = *k*<sup>2</sup> *<sup>z</sup>* + *k*<sup>2</sup> *<sup>h</sup>* + *<sup>N</sup>*<sup>2</sup> 2*g* 2 is the denominator of the dispersion equation (2). The horizontal group velocity don't play a role in the vertical propagation delay but it is useful to estimate the epicentral distance where the internal gravity waves start to interact with the ionosphere as well as the delay between the tsunami propagating at the sea surface and the internal gravity wave propagating in the atmosphere at the ionospheric altitude: *e.g.*, for a period of 10 min, the vertical propagation to reach 300 km is in the order of 1 hour, the horizontal epicentral distance 600 km and the delay between the tsunami and ionospheric IGW wavefronts is in order of 10 min.

The following interaction of IGW with the ionospheric plasma induces perturbation in the plasma density and plasma velocity. In essence, the variation in the neutral velocity *vn* produced by IGW propagation in the atmosphere produces by dynamic and electromagnetic effect the ions movement with a perturbed speed*vi* (eq. 4) that induce ion density variation *ni* (eq. 3). The principal effect is produced by collisions between the neutral molecules and ions, secondly the ions drag the electrons by charge attraction to satisfy the neutral proprieties of

Fig. 6. Vertical cross section of the modeled tsunami-related electron density perturbation and ray-paths computed using a 10 MHz OTH radar signal at 19*o*, 30*<sup>o</sup>* and 35*<sup>o</sup>* elevation (dashed gray lines). Dash-dotted purple line indicates a possible geometry of a nearby GPS station for a satellite at 25*<sup>o</sup>* elevation. Arrows indicate the tsunami and IGW energy directions of propagation. Note that the vertical scale has been exaggerated. Figure after Coïsson et al. (2011).

the ionosperic plasma (eq. 5).

8 Will-be-set-by-IN-TECH

**50**

<sup>√</sup>*hg* . Note that tsunamis generated/moving in the deeper

Fig. 5. Vertical (left) and horizontal (right) group velocity of the internal gravity wave coupled at the sea surface with tsunamis (generated at different oceanic deep *h*, see

<sup>=</sup> *khN*<sup>2</sup>(*<sup>D</sup>* <sup>−</sup> *<sup>k</sup>*<sup>2</sup>

gray-scale) and a characteristic period *T* of 10 min. Tsunamis move at the speed defined by the relation *vtsuna* = *hg*, where *<sup>g</sup>* is the gravity. Consequently, the horizontal *<sup>k</sup>*-vector *kh* that the tsunami transfer to the atmospheric internal gravity wave also depend by *h*

*h*)

The horizontal group velocity don't play a role in the vertical propagation delay but it is useful to estimate the epicentral distance where the internal gravity waves start to interact with the ionosphere as well as the delay between the tsunami propagating at the sea surface and the internal gravity wave propagating in the atmosphere at the ionospheric altitude: *e.g.*, for a period of 10 min, the vertical propagation to reach 300 km is in the order of 1 hour, the horizontal epicentral distance 600 km and the delay between the tsunami and ionospheric

The following interaction of IGW with the ionospheric plasma induces perturbation in the plasma density and plasma velocity. In essence, the variation in the neutral velocity *vn* produced by IGW propagation in the atmosphere produces by dynamic and electromagnetic effect the ions movement with a perturbed speed*vi* (eq. 4) that induce ion density variation *ni* (eq. 3). The principal effect is produced by collisions between the neutral molecules and ions, secondly the ions drag the electrons by charge attraction to satisfy the neutral proprieties of

*<sup>ω</sup>D*<sup>2</sup> *<sup>v</sup><sup>z</sup>*

is the denominator of the dispersion equation (2).

**100**

**150**

**200**

**250**

**300**

**<sup>50</sup> <sup>100</sup> <sup>150</sup> <sup>200</sup> <sup>250</sup> <sup>0</sup>**

**Horizontal Velocity (m/s)**

*<sup>g</sup>* <sup>=</sup> *δω δkz* <sup>=</sup> *kzk*<sup>2</sup> *hN*<sup>2</sup> *ωD*<sup>2</sup>

**1000**

**2000**

**3000**

**4000**

**5000**

**6000**

**7000**

**<sup>20</sup> <sup>40</sup> <sup>60</sup> <sup>80</sup> <sup>100</sup> <sup>120</sup> <sup>0</sup>**

**Vertical Velocity (m/s)**

*T*

*vh <sup>g</sup>* <sup>=</sup> *δω δkh*

**50**

following the relation *kh* = <sup>2</sup>*<sup>π</sup>*

*<sup>z</sup>* + *k*<sup>2</sup> *<sup>h</sup>* + *<sup>N</sup>*<sup>2</sup> 2*g* 2

IGW wavefronts is in order of 10 min.

where *D* = *k*<sup>2</sup>

oceanic zone produce faster IGW.

**100**

**150**

**Altitude (deg)** 

**200**

**250**

**300**

$$\frac{\partial n\_i}{\partial t} + \nabla \cdot \left( n\_i \vec{v}\_i \right) = \pm \beta n\_i - \alpha n\_i^2 \tag{3}$$
 
$$\vec{d}\vec{v}\_i$$

$$
\rho\_{\vec{i}} \frac{d\vec{v}\_{\vec{i}}}{dt} = -\nabla p\_{\vec{i}} + \rho\_{\vec{i}} \vec{g} + \eta\_{\vec{i}} \eta\_{\vec{i}} (\vec{E} + \vec{v}\_{\vec{i}} \times \vec{B})
$$

$$
\tag{4}
$$

$$-\rho\_i \mu\_{in} \left(\vec{v}\_i - \vec{v}\_n\right) \tag{4}$$

$$m\_{\mathcal{E}} = \sum^3 n\_i \tag{5}$$

The method developed by Occhipinit et al. (2006; 2008) is also used to estimate the role of the geomagnetic field in the tsumani signature at the E-region and F-region (Cobulli et al., 2008a). Normally the authors show that the amplification of the electron density perturbation in the incompressible plasma at the F-region is strongly dependent by geomagnetic inclination as well as by the direction of propagation of the tsumami. This effect is explained by the Lorenz force term in the momentum equation explained by the neutral plasma coupling (e.g. 4). Consequently, the detection of tsunamgingic perturbation in the F-region-plasma is easily observed at equatorial and mid-latitude then the high latitude. The heterogeneous amplification force by the magnetic field is not observable in the E-region, consequently detection at low altitude by Doppler sounding and over-the-horizon (OTH) radar are not affected by geomparplical location. The theoretical possibility of detection by OTH radar is explored by Co-tissue et al. (2011) for a simple tsumani-related UGW (Fig. 6) propagating in an dynamic ionsoperhe. Coison et al. (2011) for a simple tsumani-related UGW (Fig. 6) propagating in an dynamic mixture of the emission/recception beam of the OTH radar don't hide the tsumani signature (Fig. 7).

*i*=1

Fig. 9. Top: Instantaneous vTEC plotted about 1 hour after the theoretical tsunami arrival at sea-level, from left to right: after Kuril, Samoa and Chile earthquakes respectively. Bottom: Travel-time diagrams of the vTEC time-series at the time of tsunami arrival off Hawaii, from left to right: for Kuril, Samoa and Chile events, respectively. Time is related to the tsunami travel time to highlight coherence with the tsunami model. Figure after Rolland et al. (2010).

Tsunami Detection by Ionospheric Sounding: New Tools for Oceanic Monitoring 29

2011, 05:47:32 UT), detailed in the next section, allows to validate this hypothesis comparing

After the ionospheric detection of the great Sumatra earthquake (26 December 2004) and the consequent tsunami, several works focalized on the minor tsunamis in order to generalize and validate the tsunami detection by ionospheric sounding (Galvan et al., 2011; Makela et

Rolland et al. (2010) clearly showed the ionospheric detection in the far field for three tsunami events with a moderate magnitude compared to Sumatra: Kuril Islands 2006 (15 November, 2006, Mw 8.3), Samoa 2009 (29 September, 2009, Mw 8.1) and Chile 2010 (27 February 2010, Mw 8.8). Using the Hawaiian GPS network (50 stations) this work highlight the tsunami signature in the TEC. The ionospheric observations, supported by oceanic measurements by DART buoys, showed a signal coherent in arrival time and frequency-signature with the tsunami propagating in the ocean (Fig. 9). Galvan et al. (2011) showed similar coherent observations for Samoa 2009 and Chile 2010 not only in Hawaii but also in Japan, using the

Both those studies confirm, one time more, that detection of tsunami by ionospheric sounding

Close to the epicenter, the coupling mechanism is more complex as the vertical displacement induced by the seismic rupture induces, in the same time, the generation of a propagating acoustic-gravity pulse in the atmosphere (Aframovich et al., 2010; Heki & Ping, 2005; Rolland et al., 2011a) as well as the tsunami formation if the epicenter is in oceanic regions (Occhipinti et al., 2011b). The following tsunami propagation induces, as described above, the formation of gravity wave propagating in the atmosphere and perturbing the ionosphere. Taking into account the theoretical vertical and horizontal speed of tsunami-related IGW (5), Occhipinti et al. (2011b) highlights that the signature of the tsunami with, e.g., a main period of 10 min, in the ionosphere is visible only after 1 hour and at the epicentral distance of 600 km, the

measurement and modeling (Makela et al., 2011; Occhipinti et al., 2011a).

al., 2011; Occhipinti et al., 2011a; Rolland et al., 2010; 2011).

is systematic and possible during the propagation in the open ocean.

**4. Recent tsunamis observations**

dense japanese GPS network GEONET.

Fig. 7. Synthetic OTH radar record from 01:00 to 0605:00 UT at 270*<sup>o</sup>* azimuth 30*<sup>o</sup>* elevation during tsunami related IGW propagation showed in Fig. 6. Left: unperturbed ionosphere. Right: ionosphere with IGW perturbation. White points indicate the maximum signal strength at each UT. Figure after Coïsson et al. (2011).

Fig. 8. Left: Relative electron density perturbation induced by tsunami related IGW. Right: Mean electron density (m-3), mean OI 6300 ˛A VER (photons/s/m3), and mean O 1356 ˛A VER (photons/s/m3). Figures after Hickey et al. (2009; 2010).

The effect of dissipation, nominally viscosity and thermal conduction have been taken into account in the tsunami atmosphere/ionosphere modeling (Hickey et al., 2009) showing that their effect become non-neglectable above 200 km of altitude (Fig. 8). Consequently, the main theoretical and numerical objective in near future is combine the attenuation effects with a full 3-dimentional modeling.

Theoretical works appeared recently, explore the possible detection by airglow monitoring (Hickey et al., 2010). The recent dramatic event of Tohoku Earthquake (Mw=9.3, 11 March,

Fig. 9. Top: Instantaneous vTEC plotted about 1 hour after the theoretical tsunami arrival at sea-level, from left to right: after Kuril, Samoa and Chile earthquakes respectively. Bottom: Travel-time diagrams of the vTEC time-series at the time of tsunami arrival off Hawaii, from left to right: for Kuril, Samoa and Chile events, respectively. Time is related to the tsunami travel time to highlight coherence with the tsunami model. Figure after Rolland et al. (2010).

2011, 05:47:32 UT), detailed in the next section, allows to validate this hypothesis comparing measurement and modeling (Makela et al., 2011; Occhipinti et al., 2011a).

#### **4. Recent tsunamis observations**

10 Will-be-set-by-IN-TECH

Fig. 7. Synthetic OTH radar record from 01:00 to 0605:00 UT at 270*<sup>o</sup>* azimuth 30*<sup>o</sup>* elevation during tsunami related IGW propagation showed in Fig. 6. Left: unperturbed ionosphere. Right: ionosphere with IGW perturbation. White points indicate the maximum signal

Fig. 8. Left: Relative electron density perturbation induced by tsunami related IGW. Right: Mean electron density (m-3), mean OI 6300 ˛A VER (photons/s/m3), and mean O 1356 ˛A

The effect of dissipation, nominally viscosity and thermal conduction have been taken into account in the tsunami atmosphere/ionosphere modeling (Hickey et al., 2009) showing that their effect become non-neglectable above 200 km of altitude (Fig. 8). Consequently, the main theoretical and numerical objective in near future is combine the attenuation effects with a full

Theoretical works appeared recently, explore the possible detection by airglow monitoring (Hickey et al., 2010). The recent dramatic event of Tohoku Earthquake (Mw=9.3, 11 March,

VER (photons/s/m3). Figures after Hickey et al. (2009; 2010).

3-dimentional modeling.

strength at each UT. Figure after Coïsson et al. (2011).

After the ionospheric detection of the great Sumatra earthquake (26 December 2004) and the consequent tsunami, several works focalized on the minor tsunamis in order to generalize and validate the tsunami detection by ionospheric sounding (Galvan et al., 2011; Makela et al., 2011; Occhipinti et al., 2011a; Rolland et al., 2010; 2011).

Rolland et al. (2010) clearly showed the ionospheric detection in the far field for three tsunami events with a moderate magnitude compared to Sumatra: Kuril Islands 2006 (15 November, 2006, Mw 8.3), Samoa 2009 (29 September, 2009, Mw 8.1) and Chile 2010 (27 February 2010, Mw 8.8). Using the Hawaiian GPS network (50 stations) this work highlight the tsunami signature in the TEC. The ionospheric observations, supported by oceanic measurements by DART buoys, showed a signal coherent in arrival time and frequency-signature with the tsunami propagating in the ocean (Fig. 9). Galvan et al. (2011) showed similar coherent observations for Samoa 2009 and Chile 2010 not only in Hawaii but also in Japan, using the dense japanese GPS network GEONET.

Both those studies confirm, one time more, that detection of tsunami by ionospheric sounding is systematic and possible during the propagation in the open ocean.

Close to the epicenter, the coupling mechanism is more complex as the vertical displacement induced by the seismic rupture induces, in the same time, the generation of a propagating acoustic-gravity pulse in the atmosphere (Aframovich et al., 2010; Heki & Ping, 2005; Rolland et al., 2011a) as well as the tsunami formation if the epicenter is in oceanic regions (Occhipinti et al., 2011b). The following tsunami propagation induces, as described above, the formation of gravity wave propagating in the atmosphere and perturbing the ionosphere. Taking into account the theoretical vertical and horizontal speed of tsunami-related IGW (5), Occhipinti et al. (2011b) highlights that the signature of the tsunami with, e.g., a main period of 10 min, in the ionosphere is visible only after 1 hour and at the epicentral distance of 600 km, the

**5. Conclusion and perspectives**

sensing in the future tsunami warning systems.

by the PNTS/INSU. This is IPGP contribution 3245.

*Geophys. Res. Lett., 33*, L14314, 2006.

431, doi:10.1016/j.tecto.2006.05.038.

airglow.

IGW.

detection technique.

**6. Acknowledgments**

**7. References**

The analysis of tsunamigenic ionospheric perturbation observed after major events provides valuable information for understanding the physical processes and explore new techniques for tsunami warning systems. Along this chapter it has shown that early detection of tsunamigenic IGWs is possible using a bunch of remote sensing techniques as the TEC measurement by radar-altimetry and GPS, as well as the observation of the atmospheric

Tsunami Detection by Ionospheric Sounding: New Tools for Oceanic Monitoring 31

If the TEC measurement of tsunami seems today an established technique for tsunami observation, it present a large number of limits, primarily the observation geometry, that have

The preliminary result about ariglow observation highlights that remote sensing of tsunamis via the atmospheric/ionospheric monitoring by ground-based or on-board camera could be a mature technique for oceanic monitoring and have to find a place in the future of tsunami

By numerical modeling recents works prove the ability of additional techniques as the OTH-Radar who measure the perturbation at the ionospheric E-region (around 150 km) reducing, if used close to the epicenter, the response time of detection of the tsunami related

Anyway resuming, some of this techniques are able to highlight the presence of a tsunami several hours before that the wave hits the coast and could play a revolutionary role of remote

The works presented here are supported by the French Space Agency CNES and by the Unite States Office of Naval Research (ONR) Global under contract IONONAMI-N07-25, as well as

Ammon, C. J., A. A. Velasco, T. Lay, Rapid estimation of first-order rupture characteristics

Afraimovich, E. L., D. Feng, V. V. Kiryushkin, E. I. Astafyeva, S. Jin, and V. A. Sankov (2010),

Balasis G. and Mandea M., 2007. Can electromagnetic disturbances related to the recent

Balthazor, R. L., and R. J. Moffett, 1997. A study of atmospheric gravity waves and travelling ionospheric disturbances at equatorial latitudes, *Ann. Geophysicae*, 15, 1048-1056.

Bilitza, D., 2001. International Reference Ionosphere 2000, *Radio Science*, 36, 2, 261-275. P. Coïsson, G. Occhipinti, P. Lognonné, L. M. Rolland, Tsunami signature in the ionosphere:

the innovative role of OTH radar, *Radio Sci.*, Under Revision, 2011.

gravity waves induced by tsunamis. *J. Geophys. Res.*, 160, 840.

for large earthquakes using surface waves: 2004 Sumatra-Andaman earthquake,

TEC response to the 2008 Wenchuan earthquake in comparison with other strong earthquakes, *Int. J. Remote Sens.*, 31, 3601–3613, doi:10.1080/01431161003727747. Artru, J., V. Ducic, H. Kanamori, P. Lognonné, M. Murakami, 2005. Ionospheric detection of

great earthquakes be detected by satellite magnetometers?, *special issue Mechanical and Electromagnetic Phenomena Accompanying Preseismic Deformation: from Laboratory to Geophysical Scale*, ed. by K. Eftaxias, T. Chelidze and V. Sgrigna, Tectonophysics,

to be taken into account for an eventual application in the oceanic monitoring.

Fig. 10. a) IGW imaged by the 630.0-nm ground-based airglow camera located on the Haleakala Volcano on Maui, Hawaii, at 13:20 UT. b) normalized vertical velocity of the modeled IGW ((*kg*/*m*3) 1 <sup>2</sup> *m*/*s*). The Y structure as well as the longer wavelength anticipating the Y (X) are present in both, airglow data and AGW synthetics. Those structures are observed between 12:12 to 13:32. The white dotted line in a and b shows the tsunami wavefront line at 13:20 UT. c) graphical estimation of the shift induced by the wind between model and data: the grid and the white lines are estimated on b, then shifted on a to fit with the position of Y and X. The estimated shift of 2◦ is coherent with previous observations Occhipinti et al. (2006). Figure after Occhipinti et al. (2011a).

theoretical estimation is supported by the observation of several tsunamigenic earthquakes: the 26 December, 2004, and 12 September, 2007, in Sumatra; the 14 November, 2007, in Chile; the 29 September, 2009, in Samoa; and the recent Tohoku-Oki (Japan) earthquake on 11 Mars 2011. Additionally, this last work highlight how the sensitivity of the TEC measurement is affected by the inclination angle of the station-satellite line-of-site: as a consequence of the integrated nature of TEC, the low inclination measurements have stronger sensitivity to tsunami signature in the ionosphere.

During the Tohoku-Oki event, close to the TEC measurements by GPS (Rolland et al., 2011) or altimeters (Jason-1), the tsunami related IGW propagating over the Pacific Ocean has also been detected for the first time by the airglow wide-angle camera system located at the top of the Haleakala Volcano on Maui, Hawaii Makela et al. (2011).

In essence, the camera is observing the airglow layer at approximately 250 km in altitude caused by the dissociative recombination of O2+ [Link and Cogger, 1988], which emits photons at 630.0 nm has predicted by Hickey et al. (2010).

Numerical modeling of IGW reproduces the main features observed in the airglow images (Fig. 10) showing interesting likenesses between the model and data, and explaining the nature of the airglow observation (Occhipinti et al., 2011a).

The tsunamic nature of the airglow observation is, first, clearly explained by the presence of a Y shape appearing in both synthetics and data; second, by the presence of a wave with a longer wavelength (indicated by X in Fig. 10) that is arriving before the tsunami front-wave. This observation is theoretically explained by Occhipinti et al. (2011a) as the combined effect of the low bathymetry around Hawaii and the period-dependence of the horizontal IGW speed propagation. In essence, the tsunami related IGW with longer-period go faster than shorter-period, consequently, when the tsunami slow-down by the effect of the low bathymetry close to the Hawaiian arcipelagos, the longer-period IGW goes over the tsunami wavefront.

#### **5. Conclusion and perspectives**

12 Will-be-set-by-IN-TECH

Fig. 10. a) IGW imaged by the 630.0-nm ground-based airglow camera located on the Haleakala Volcano on Maui, Hawaii, at 13:20 UT. b) normalized vertical velocity of the

the Y (X) are present in both, airglow data and AGW synthetics. Those structures are observed between 12:12 to 13:32. The white dotted line in a and b shows the tsunami wavefront line at 13:20 UT. c) graphical estimation of the shift induced by the wind between model and data: the grid and the white lines are estimated on b, then shifted on a to fit with the position of Y and X. The estimated shift of 2◦ is coherent with previous observations

theoretical estimation is supported by the observation of several tsunamigenic earthquakes: the 26 December, 2004, and 12 September, 2007, in Sumatra; the 14 November, 2007, in Chile; the 29 September, 2009, in Samoa; and the recent Tohoku-Oki (Japan) earthquake on 11 Mars 2011. Additionally, this last work highlight how the sensitivity of the TEC measurement is affected by the inclination angle of the station-satellite line-of-site: as a consequence of the integrated nature of TEC, the low inclination measurements have stronger sensitivity to

During the Tohoku-Oki event, close to the TEC measurements by GPS (Rolland et al., 2011) or altimeters (Jason-1), the tsunami related IGW propagating over the Pacific Ocean has also been detected for the first time by the airglow wide-angle camera system located at the top of

In essence, the camera is observing the airglow layer at approximately 250 km in altitude caused by the dissociative recombination of O2+ [Link and Cogger, 1988], which emits

Numerical modeling of IGW reproduces the main features observed in the airglow images (Fig. 10) showing interesting likenesses between the model and data, and explaining the

The tsunamic nature of the airglow observation is, first, clearly explained by the presence of a Y shape appearing in both synthetics and data; second, by the presence of a wave with a longer wavelength (indicated by X in Fig. 10) that is arriving before the tsunami front-wave. This observation is theoretically explained by Occhipinti et al. (2011a) as the combined effect of the low bathymetry around Hawaii and the period-dependence of the horizontal IGW speed propagation. In essence, the tsunami related IGW with longer-period go faster than shorter-period, consequently, when the tsunami slow-down by the effect of the low bathymetry close to the Hawaiian arcipelagos, the longer-period IGW goes over the

<sup>2</sup> *m*/*s*). The Y structure as well as the longer wavelength anticipating

modeled IGW ((*kg*/*m*3)

tsunami signature in the ionosphere.

tsunami wavefront.

1

Occhipinti et al. (2006). Figure after Occhipinti et al. (2011a).

the Haleakala Volcano on Maui, Hawaii Makela et al. (2011).

photons at 630.0 nm has predicted by Hickey et al. (2010).

nature of the airglow observation (Occhipinti et al., 2011a).

The analysis of tsunamigenic ionospheric perturbation observed after major events provides valuable information for understanding the physical processes and explore new techniques for tsunami warning systems. Along this chapter it has shown that early detection of tsunamigenic IGWs is possible using a bunch of remote sensing techniques as the TEC measurement by radar-altimetry and GPS, as well as the observation of the atmospheric airglow.

If the TEC measurement of tsunami seems today an established technique for tsunami observation, it present a large number of limits, primarily the observation geometry, that have to be taken into account for an eventual application in the oceanic monitoring.

The preliminary result about ariglow observation highlights that remote sensing of tsunamis via the atmospheric/ionospheric monitoring by ground-based or on-board camera could be a mature technique for oceanic monitoring and have to find a place in the future of tsunami detection technique.

By numerical modeling recents works prove the ability of additional techniques as the OTH-Radar who measure the perturbation at the ionospheric E-region (around 150 km) reducing, if used close to the epicenter, the response time of detection of the tsunami related IGW.

Anyway resuming, some of this techniques are able to highlight the presence of a tsunami several hours before that the wave hits the coast and could play a revolutionary role of remote sensing in the future tsunami warning systems.

#### **6. Acknowledgments**

The works presented here are supported by the French Space Agency CNES and by the Unite States Office of Naval Research (ONR) Global under contract IONONAMI-N07-25, as well as by the PNTS/INSU. This is IPGP contribution 3245.

#### **7. References**


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**3** 

*USA* 

**Proximal Records of Paleotsunami Runup in** 

**Great Earthquakes in the Central Cascadia** 

**Suduction Zone, Oregon, USA** 

Curt D. Peterson and Kenneth M. Cruikshank

*Department of Geology, Portland State University, Portland, OR,* 

**Barrage Creek Floodplains from Late-Holocene** 

A 100 km section of the central Oregon coast (Fig. 1) was surveyed for proximal creek floodplains, located at less than 500 m distance from the ocean shoreline, that could host sand sheet records of paleotsunami inundation (Peterson & Cruikshank, 2007). The study region is located near the center of the Cascadia margin, an active subduction zone that spans about 1000 km distance in the central west coast of North America (Atwater et al., 1995; Darienzo et al., 1994). Previous work in two distal floodplain localities within the study region, Neskowin and Beaver Creek, showed multiple paleotsunami inundations of one to several kilometers distance landward during the last ~ 2,500 years (see section 2.2) (Peterson et al., 2010a). It was not known how those distal records might relate to proximal or shoreline runup heights of the same paleotsunami events in the region. Such proximal or near-shoreline runup heights are needed to 1) demonstrate flooding hazards in coastal areas that have not suffered catastrophic flooding in historic time (Dengler, 2006), and 2) independently test flooding predictions based on assumed fault displacements and

In this paper we document anomalous sand sheet layers in 8 small creek floodplains that exceed 6 m elevation thresholds for tsunami inundation. Target sand sheets are examined for evidence of marine shell fragments, tracers of marine surge origins, and the landward limits of sand sheet extent. The time span of continuous deposition in one representative floodplain locality is dated by radiocarbon. An adjacent floodplain locality is examined for landward trends of sand sheet composition, sand sheet thickness, microscopic tracers of marine deposits, and sand sheet radiocarbon age. Maximum sand sheet extent in the proximal floodplain locality is compared to maximum sand sheet extent in a distal floodplain locality (Peterson et al., 2010a) to yield a landward runup height gradient for the most recent event of large magnitude runup. The runup attenuation gradient is tested against previously reported runup elevations in another central Cascadia study area, Cannon Beach, Oregon (Fig. 1) (Peterson et al., 2008). The methods reported here should be applicable in similar settings to the documentation of paleotsunami runup heights in other

numerical tsunami runup models (González et al., 2009).

susceptible coastlines around the world.

**1. Introduction** 


### **Proximal Records of Paleotsunami Runup in Barrage Creek Floodplains from Late-Holocene Great Earthquakes in the Central Cascadia Suduction Zone, Oregon, USA**

Curt D. Peterson and Kenneth M. Cruikshank *Department of Geology, Portland State University, Portland, OR, USA* 

#### **1. Introduction**

16 Will-be-set-by-IN-TECH

34 Tsunami – A Growing Disaster

Titov, V., A. B. Rabinovich, H. O. Mofjeld, R. E. Thomson, F. I. González, The global reach of the 26 December 2004 Sumatra tsunami, *Science, 309*, 2045-2048, 2005. Vigny, C., W. J. F. Simons, S. Abu, R. Bamphenyu, C. Satirapod, N. Choosakul, C. Subarya,

436, 201-206.

A. Socquet, K. Omar, H. Z. Abidin, B. A. C. Ambrosius, 2005. Insight into the 2004 Sumatra-Andaman earthquake from GPS measurement in southeast Asia, *Nature*,

> A 100 km section of the central Oregon coast (Fig. 1) was surveyed for proximal creek floodplains, located at less than 500 m distance from the ocean shoreline, that could host sand sheet records of paleotsunami inundation (Peterson & Cruikshank, 2007). The study region is located near the center of the Cascadia margin, an active subduction zone that spans about 1000 km distance in the central west coast of North America (Atwater et al., 1995; Darienzo et al., 1994). Previous work in two distal floodplain localities within the study region, Neskowin and Beaver Creek, showed multiple paleotsunami inundations of one to several kilometers distance landward during the last ~ 2,500 years (see section 2.2) (Peterson et al., 2010a). It was not known how those distal records might relate to proximal or shoreline runup heights of the same paleotsunami events in the region. Such proximal or near-shoreline runup heights are needed to 1) demonstrate flooding hazards in coastal areas that have not suffered catastrophic flooding in historic time (Dengler, 2006), and 2) independently test flooding predictions based on assumed fault displacements and numerical tsunami runup models (González et al., 2009).

> In this paper we document anomalous sand sheet layers in 8 small creek floodplains that exceed 6 m elevation thresholds for tsunami inundation. Target sand sheets are examined for evidence of marine shell fragments, tracers of marine surge origins, and the landward limits of sand sheet extent. The time span of continuous deposition in one representative floodplain locality is dated by radiocarbon. An adjacent floodplain locality is examined for landward trends of sand sheet composition, sand sheet thickness, microscopic tracers of marine deposits, and sand sheet radiocarbon age. Maximum sand sheet extent in the proximal floodplain locality is compared to maximum sand sheet extent in a distal floodplain locality (Peterson et al., 2010a) to yield a landward runup height gradient for the most recent event of large magnitude runup. The runup attenuation gradient is tested against previously reported runup elevations in another central Cascadia study area, Cannon Beach, Oregon (Fig. 1) (Peterson et al., 2008). The methods reported here should be applicable in similar settings to the documentation of paleotsunami runup heights in other susceptible coastlines around the world.

Proximal Records of Paleotsunami Runup in Barrage Creek Floodplains from

in the study area for the rupture events L and J at ~ 2.8 and 3.2 ka, respectively.

Paleotsunami records have been reported from two alluvial floodplain settings in the central Oregon coast, Neskowin and Beaver Creek (Fig. 1)(Table 2) (Peterson et al., 2010a). The two localities differ in landward floodplain gradient and protective barrier ridges. A lowgradient floodplain (0 to 3 m elevation) in the Beaver Creek Valley records landward thinning sand sheets (20 to 1 cm thickness) over inundation distances of 1 to 4 km from the beach. In this paper all reported elevations are in the NAVD88 datum, which is about 1

The two longest paleotsunami runups in Beaver Creek correspond to dated paleotsunami deposits at 1520-1700 BP and 2960-3220 BP (Table 2). The age of the tsunami remobilized organic debris should predate the Cascadia rupture event, so we assign the two events to tsunami #3 and tsunami #5 or #6. The oldest three rupture events (H, J, and L) have short recurrence intervals (Table 1), so the radiocarbon ages of their associated tsunami deposits

The Neskowin back-barrier wetlands are fronted by a barrier dune ridge (5–8 m elevation) and are backed by an uplifted terrace, which is dissected by small creeks. Sand sheets from four nearfield paleotsunami were traced across the back-barrier wetland (3 m elevation) and into a high-gradient creek floodplain, Hawk Creek, at elevations of 3 to 8 m at distances of ~ 0.5 to 1.0 km from the ocean shoreline (Table 2). The longest and highest tsunami runup at Neskowin, dated at 1114-1300 BP, is correlated to tsunami #3, which is coincident with the

The central Oregon coast was selected for this study based on a straight coastline, with low marine terraces that are dissected my numerous small creek valleys. Gouge coring (2.5 cm

**2.2 Distal paleotsunami runup records** 

meter below mean sea level (MSL).

youngest of the two longest runup events in Beaver Creek.

might overlap.

**3. Methods** 

Late-Holocene Great Earthquakes in the Central Cascadia Suduction Zone, Oregon, USA 37

**Event Atwater Tsunami # Age**  Subsidence / Tsunami Y 1 0.3 (AD1700) ? / Tsunami - 2a ~ 0.9 Subsidence / Tsunami W 2b ~ 1.1 Subsidence / Tsunami U 3 ~ 1.3 Subsidence / Tsunami S 4 ~ 1.7 Subsidence / Tsunami H 5 ~ 2.6 Subsidence / Tsunami L - ~ 2.8 Subsidence / ? J - ~ 3.2 Table 1. Central Cascadia Margin rupture and tsunami events. Subsidence related rupture events and ages are from Atwater et al., (2004). Tsunami correlations (Tsunami #) in the central Cascadia margin are from Peterson et al., (2010a). The source of the 2nd tsunami sand layer in the central Cascadia tsunami is not known (Peterson et al., 2008), but it might reflect a partial rupture in the northern part of the Cascadia margin (Clague et al., 2000; Schlichting & Peterson, 2006). Due to difficulties of discriminating the 2nd and 3rd tsunami events in some floodplain settings we assign them numbers of #2a and #2b, where they can be discriminated, and #2 where only one layer can be discriminated. Tsunami inundation has yet to be verified

Fig. 1. The paleotsunami study region (~ 100 km in length) is located in the central Cascadia subduction zone (map at right). The Cascadia megathrust daylights along the trench (dashed line). The focused study area (~ 35 km in length) is shown in central Oregon coast (boxed area in inset map at left). Coastal communities (solid circles) and small estuaries (Bays) are named. Two localities investigated for paleotsunami inundation in distal floodplain settings are shown at Neskowin and Beaver Creeks (solid squares). Map coordinates for the study region (inset map at left) are in UTM x 1,000 m. Basemap is from Google (2011).

#### **2. Study region**

#### **2.1 Paleoseismic evidence**

A total of six or seven great earthquakes that ruptured though the central Cascadia margin during the last ~ 3,000 years are well recorded in Washington and northernmost Oregon where coseismic subsidence of least 1.0 m predominates (Table 1). Investigations of the associated nearfield paleotsunami inundations moved from the coastal marshes to overland inundation recorded in lakes, beach plains, back-barrier wetlands, and barrage lakes (Hutchinson et al., 1997; Kelsey et al., 2005; Schlichting & Peterson, 2006). The geologic records of paleotsunami sand sheet deposition that are located landward of foredunes, beach ridges, and other unstable coastal barriers provide minimum estimates of runup height (Peterson et al., 2006). To overcome these limitations some recent investigations have been redirected towards mapping paleotsunami sand sheets in small alluvial floodplains that are located on relative stable marine terraces (see Section 2.2 below).


Table 1. Central Cascadia Margin rupture and tsunami events. Subsidence related rupture events and ages are from Atwater et al., (2004). Tsunami correlations (Tsunami #) in the central Cascadia margin are from Peterson et al., (2010a). The source of the 2nd tsunami sand layer in the central Cascadia tsunami is not known (Peterson et al., 2008), but it might reflect a partial rupture in the northern part of the Cascadia margin (Clague et al., 2000; Schlichting & Peterson, 2006). Due to difficulties of discriminating the 2nd and 3rd tsunami events in some floodplain settings we assign them numbers of #2a and #2b, where they can be discriminated, and #2 where only one layer can be discriminated. Tsunami inundation has yet to be verified in the study area for the rupture events L and J at ~ 2.8 and 3.2 ka, respectively.

#### **2.2 Distal paleotsunami runup records**

Paleotsunami records have been reported from two alluvial floodplain settings in the central Oregon coast, Neskowin and Beaver Creek (Fig. 1)(Table 2) (Peterson et al., 2010a). The two localities differ in landward floodplain gradient and protective barrier ridges. A lowgradient floodplain (0 to 3 m elevation) in the Beaver Creek Valley records landward thinning sand sheets (20 to 1 cm thickness) over inundation distances of 1 to 4 km from the beach. In this paper all reported elevations are in the NAVD88 datum, which is about 1 meter below mean sea level (MSL).

The two longest paleotsunami runups in Beaver Creek correspond to dated paleotsunami deposits at 1520-1700 BP and 2960-3220 BP (Table 2). The age of the tsunami remobilized organic debris should predate the Cascadia rupture event, so we assign the two events to tsunami #3 and tsunami #5 or #6. The oldest three rupture events (H, J, and L) have short recurrence intervals (Table 1), so the radiocarbon ages of their associated tsunami deposits might overlap.

The Neskowin back-barrier wetlands are fronted by a barrier dune ridge (5–8 m elevation) and are backed by an uplifted terrace, which is dissected by small creeks. Sand sheets from four nearfield paleotsunami were traced across the back-barrier wetland (3 m elevation) and into a high-gradient creek floodplain, Hawk Creek, at elevations of 3 to 8 m at distances of ~ 0.5 to 1.0 km from the ocean shoreline (Table 2). The longest and highest tsunami runup at Neskowin, dated at 1114-1300 BP, is correlated to tsunami #3, which is coincident with the youngest of the two longest runup events in Beaver Creek.

#### **3. Methods**

36 Tsunami – A Growing Disaster

Fig. 1. The paleotsunami study region (~ 100 km in length) is located in the central Cascadia

A total of six or seven great earthquakes that ruptured though the central Cascadia margin during the last ~ 3,000 years are well recorded in Washington and northernmost Oregon where coseismic subsidence of least 1.0 m predominates (Table 1). Investigations of the associated nearfield paleotsunami inundations moved from the coastal marshes to overland inundation recorded in lakes, beach plains, back-barrier wetlands, and barrage lakes (Hutchinson et al., 1997; Kelsey et al., 2005; Schlichting & Peterson, 2006). The geologic records of paleotsunami sand sheet deposition that are located landward of foredunes, beach ridges, and other unstable coastal barriers provide minimum estimates of runup height (Peterson et al., 2006). To overcome these limitations some recent investigations have been redirected towards mapping paleotsunami sand sheets in small alluvial floodplains that are located on relative stable marine terraces (see Section 2.2

subduction zone (map at right). The Cascadia megathrust daylights along the trench (dashed line). The focused study area (~ 35 km in length) is shown in central Oregon coast (boxed area in inset map at left). Coastal communities (solid circles) and small estuaries (Bays) are named. Two localities investigated for paleotsunami inundation in distal floodplain settings are shown at Neskowin and Beaver Creeks (solid squares). Map coordinates for the study region (inset map at left) are in UTM x 1,000 m. Basemap is from

Google (2011).

below).

**2. Study region** 

**2.1 Paleoseismic evidence** 

The central Oregon coast was selected for this study based on a straight coastline, with low marine terraces that are dissected my numerous small creek valleys. Gouge coring (2.5 cm

Proximal Records of Paleotsunami Runup in Barrage Creek Floodplains from

provided by Schilchting (2000).

molar HCL for dissolution testing.

analytical error level by Beta Analytic Inc.

**3.1 High-gradient proximal floodplains** 

landward thinning layers of beach sand.

Folk (1980).

Late-Holocene Great Earthquakes in the Central Cascadia Suduction Zone, Oregon, USA 39

et al., 1997; Kelsey et al., 2005). However, unexpected traces of marine diatoms in some control intervals, or non-tsunami deposits, from the central Cascadia beach plains and floodplains have suggested marine diatom transport by ocean wind/spray to distances of 2 km inland from the beach (Peterson et al., 2010a; Schlichting & Peterson, 2006). To further test these findings some target paleotsunami deposits and control intervals in the creek floodplains were examined for diatom taxa abundances following methods

Carbonate shell fragments provide an alternative marine source tracer for the target paleotsunami sand layers in proximal localities where the presence of small gravel size fractions (> 2 mm diameter) include marine shell fragments. Shell fragments in the granule and small pebble size ranges can usually be identified in the field with hand lens and 7.0

Another technique for indentifying carbonate shell fragments was used for the finer sand size fractions in the more landward sites of the proximal localities. Small carbonate fragments (0.2-0.5 mm diameter) in the sand size fraction are present at trace abundances (1,000 total grain counts per sample). They are identified in grain mounts by moderately high relief and high dispersion in polarized light under petrographic microscopy at 250x. Sediment grain sizes are analyzed in the target paleotsunami deposits to establish any vertical trends and landward trends in grain size distributions. Upward grain-size fining trends are established by high-resolution digital photography at ~ 50x magnification. Landward fining trends are based on petrographic microscopy (250x) of grain counts (50 total grains per slide) of samples taken from the middle of target tsunami sand layers. Standard statistical methods for characterizing the sample grain size populations follow

Samples of organic detritus, including leaves and twigs, were collected from tsunami sand sheet layers for radiocarbon dating by AMS method. Small samples (< 0.5 g) were air dried, weighed and submitted to Beta Analytic Inc. for dating. Sample dates are provided in isotope adjusted radiocarbon age and in calibrated radiocarbon years BP at the 2–

In this study we target high-gradient floodplains in the central Oregon coast for sand sheet records of proximal paleotsunami runup (Fig. 2**;** Table 3). The small creek floodplains rise from 5 m to as much as 15 m in elevation within short distances (0.5-1.0 km) from the beach. The incised creek valleys were downcut into Pleistocene dune sheets and underlying marine terraces during the last sea level lowstand. The perched creek floodplains formed after late-Holocene beach sand ramped against the sea cliffs and creek mouths in the study area (Hart & Peterson, 2007). The onset of surplus beach sand supply began between 4.0 and 4.5 ka based on dated buried stumps on the beach platform at Grant Creek and Deer Creek, and from a sea cliff dune ramp at Lost Creek (Fig. 2; Table 3). The excess supply of beach sand slowed or stopped by about 3.0 ka in the study area, based on a radiocarbon date from the top of the dune ramp that blocked Quail Creek. Seasonal flooding and shallow groundwater surfaces in the creek valleys permitted accumulations of peaty mud in the floodplains, which serve as hosting deposits for


Table 2. Overland paleotsunami runup records in distal alluvial flood plain settings, central Oregon coast. Rupture ages for the central Cascadia margin are from Atwater et al., (2004). Neskowin and North Beaver Creek data are from Peterson et al., (2010a) and Schlichting (2000). Distances from the ocean shoreline (km) are based on observed tsunami sand deposition, so they might underestimate maximum flooding distance. Elevations of terminal deposits (m NAVD88) shown here are not adjusted for late Holocene rise of sea level height. A net relative sea-level rise of 1m/1000 yr can be applied to the older paleotsunami events to correct for increased runup heights at the time of inundation. Radiocarbon ages (calibrated at ±2 sig. yr BP) are based on transported tsunami debris, so they should predate the corresponding rupture event age. Beaver Creek and Neskowin Creek floodplain localities are shown in Fig. 1.

diameter x 2.0 m core lengths) and/or ram coring (7.5 cm diameter x 2.0 m core lengths) were used to test the creek floodplain deposits for anomalous landward-thinning sand sheets. The surveyed core sites were investigated with 2 to 5 core retrievals prior to 1) core photography (50 mm macro-lens on a 10 megapixel DSLR), 2) subsampling for sand size, carbonate shell fragments, diatoms, and radiocarbon dating, and 3) logging at the 1 cm length scale. Core site positions are established by 12 channel WAAS-enabled GPS (e.p.e. 2.5-5.0 m). Locality elevations are taken from LiDAR (U. S. Geological Survey, 2011). Selected core sites are surveyed into registered benchmarks by EDM-Total Station for precise elevation control (elevation error ± 5 cm). Photos of target tsunami deposits, site position data, elevation surveying data, and initial field logs are archived in the Oregon Tsunami Database (Cruikshank & Peterson, 2011)

In the previous studies of paleotsunami runup in distal alluvial floodplains (see Section 2.2) heavy mineral tracers were used discriminate between beach sand and river sand sources to the target paleotsunami sand sheets (Peterson et al., 2010a). That approach could not be used in the proximal alluvial floodplains, due the presence of uplifted Pleistocene beach and dune deposits in the small alluvial drainages, which could contribute beach sand minerals to the creek sand bedload. The presence of marine diatoms has previously been used to establish paleotsunami inundation in some upland or inland freshwater settings of the Cascadia margin (Hemphill-Haley, 1996; Hutchinson et al., 1997; Kelsey et al., 2005). However, unexpected traces of marine diatoms in some control intervals, or non-tsunami deposits, from the central Cascadia beach plains and floodplains have suggested marine diatom transport by ocean wind/spray to distances of 2 km inland from the beach (Peterson et al., 2010a; Schlichting & Peterson, 2006). To further test these findings some target paleotsunami deposits and control intervals in the creek floodplains were examined for diatom taxa abundances following methods provided by Schilchting (2000).

Carbonate shell fragments provide an alternative marine source tracer for the target paleotsunami sand layers in proximal localities where the presence of small gravel size fractions (> 2 mm diameter) include marine shell fragments. Shell fragments in the granule and small pebble size ranges can usually be identified in the field with hand lens and 7.0 molar HCL for dissolution testing.

Another technique for indentifying carbonate shell fragments was used for the finer sand size fractions in the more landward sites of the proximal localities. Small carbonate fragments (0.2-0.5 mm diameter) in the sand size fraction are present at trace abundances (1,000 total grain counts per sample). They are identified in grain mounts by moderately high relief and high dispersion in polarized light under petrographic microscopy at 250x.

Sediment grain sizes are analyzed in the target paleotsunami deposits to establish any vertical trends and landward trends in grain size distributions. Upward grain-size fining trends are established by high-resolution digital photography at ~ 50x magnification. Landward fining trends are based on petrographic microscopy (250x) of grain counts (50 total grains per slide) of samples taken from the middle of target tsunami sand layers. Standard statistical methods for characterizing the sample grain size populations follow Folk (1980).

Samples of organic detritus, including leaves and twigs, were collected from tsunami sand sheet layers for radiocarbon dating by AMS method. Small samples (< 0.5 g) were air dried, weighed and submitted to Beta Analytic Inc. for dating. Sample dates are provided in isotope adjusted radiocarbon age and in calibrated radiocarbon years BP at the 2– analytical error level by Beta Analytic Inc.

#### **3.1 High-gradient proximal floodplains**

38 Tsunami – A Growing Disaster

Beaver Ck 0.3 1 20 2.5 0.0 320-520

1.3 3 16 4.1 1.5 1520-1700

 2.6-2.8 5-6 6 4.0 0.7 2960-3220 Neskowin 0.3 1 20 0.6 3.0 - 0.8 2a 5 0.6 3.0 - 1.1 2b 8 0.8 6.5 940-1140 1.3 3 40 1.0 8.3 1140-1300 Table 2. Overland paleotsunami runup records in distal alluvial flood plain settings, central Oregon coast. Rupture ages for the central Cascadia margin are from Atwater et al., (2004). Neskowin and North Beaver Creek data are from Peterson et al., (2010a) and Schlichting (2000). Distances from the ocean shoreline (km) are based on observed tsunami sand

deposition, so they might underestimate maximum flooding distance. Elevations of terminal deposits (m NAVD88) shown here are not adjusted for late Holocene rise of sea level height. A net relative sea-level rise of 1m/1000 yr can be applied to the older paleotsunami events

(calibrated at ±2 sig. yr BP) are based on transported tsunami debris, so they should predate

diameter x 2.0 m core lengths) and/or ram coring (7.5 cm diameter x 2.0 m core lengths) were used to test the creek floodplain deposits for anomalous landward-thinning sand sheets. The surveyed core sites were investigated with 2 to 5 core retrievals prior to 1) core photography (50 mm macro-lens on a 10 megapixel DSLR), 2) subsampling for sand size, carbonate shell fragments, diatoms, and radiocarbon dating, and 3) logging at the 1 cm length scale. Core site positions are established by 12 channel WAAS-enabled GPS (e.p.e. 2.5-5.0 m). Locality elevations are taken from LiDAR (U. S. Geological Survey, 2011). Selected core sites are surveyed into registered benchmarks by EDM-Total Station for precise elevation control (elevation error ± 5 cm). Photos of target tsunami deposits, site position data, elevation surveying data, and initial field logs are archived in the Oregon

In the previous studies of paleotsunami runup in distal alluvial floodplains (see Section 2.2) heavy mineral tracers were used discriminate between beach sand and river sand sources to the target paleotsunami sand sheets (Peterson et al., 2010a). That approach could not be used in the proximal alluvial floodplains, due the presence of uplifted Pleistocene beach and dune deposits in the small alluvial drainages, which could contribute beach sand minerals to the creek sand bedload. The presence of marine diatoms has previously been used to establish paleotsunami inundation in some upland or inland freshwater settings of the Cascadia margin (Hemphill-Haley, 1996; Hutchinson

to correct for increased runup heights at the time of inundation. Radiocarbon ages

the corresponding rupture event age. Beaver Creek and Neskowin Creek floodplain

**Pinchout distance (km)** 

**Terminal elevation (m, NAVD88)** 

**Calibrated Radiocarbon (yr BP ± 2 )** 

**Max. TSL thickness (cm)** 

 0.9 2a 8 0.5-1.0 0.0 1.1 2b 12 1.5 0.0

1.7 4 5 1.5-2.0 0.0

**Distal Record Locality**  **Rupture Age (ka)** 

localities are shown in Fig. 1.

Tsunami Database (Cruikshank & Peterson, 2011)

**TSL event (#)** 

> In this study we target high-gradient floodplains in the central Oregon coast for sand sheet records of proximal paleotsunami runup (Fig. 2**;** Table 3). The small creek floodplains rise from 5 m to as much as 15 m in elevation within short distances (0.5-1.0 km) from the beach. The incised creek valleys were downcut into Pleistocene dune sheets and underlying marine terraces during the last sea level lowstand. The perched creek floodplains formed after late-Holocene beach sand ramped against the sea cliffs and creek mouths in the study area (Hart & Peterson, 2007). The onset of surplus beach sand supply began between 4.0 and 4.5 ka based on dated buried stumps on the beach platform at Grant Creek and Deer Creek, and from a sea cliff dune ramp at Lost Creek (Fig. 2; Table 3). The excess supply of beach sand slowed or stopped by about 3.0 ka in the study area, based on a radiocarbon date from the top of the dune ramp that blocked Quail Creek. Seasonal flooding and shallow groundwater surfaces in the creek valleys permitted accumulations of peaty mud in the floodplains, which serve as hosting deposits for landward thinning layers of beach sand.

Proximal Records of Paleotsunami Runup in Barrage Creek Floodplains from

**UTM Easting (m)** 

**Locality/ Core Sites** 

**UTM Northing (m)** 

Cruikshank and Peterson (2011).

study area (see Section 4.1).

last several thousand years.

Late-Holocene Great Earthquakes in the Central Cascadia Suduction Zone, Oregon, USA 41

**Threshold /core elev. (m)** 

**TSL number**  **TSL max. thickness (cm)** 

**Shoreline distance (m)** 

BIGC2/3 4945660 416690 400 6.5/6.5 2 7 HEND3/4 4938210 415680 420 6.0/5.5 3 12 GRAN5/6 4936930 415250 120 6.0/8.0 3 10 MOOR3/4 4936080 415140 80 7.5/8.0 3 36 THEI1A/2 4935090 415110 100 6.0/5.5 3 12 PUMP5/6 4931050 414870 340 6.0/6.5 3 12 BEAV\_D1 4930160 414990 400 4.5/2 4 13 TILL1/1b 4916010 413630 250 6.0/5.5 2 5 VING1/1b 4910390 412680 170 6.0/7.0 3 5 Table 3. Representative proximal runup core sites in the central Oregon coast. Core site elevation is based on core site surface elevations (m, NAVD88) taken from LiDAR (U. S. Geological Survey, 2011). Distance is the straight-line distance (m) of surge flow from beach backshore or sea cliff to the representative core sites. Threshold elevation (m, NAVD88) is based on the mean across-valley elevation at or near the shoreline. TSL number is the number of observed tsunami sand layers of least 1 cm sand thickness in each runup locality. TSL maximum thickness is the thickest target tsunami sand layer (cm) observed in the runup locality. Core sites in creek floodplain localities are shown in Fig. 2. Data from

the Grant Creek silt alluvium, located just above basal creek gravel, is consistent with an expected onset of creek damming by dune ramp barrages that occurred at about 4 ka in the

The tsunami sand layers (TSL) decrease in number and thickness with increasing landward distance and elevation in each runup locality (Fig. 2). The tsunami sand layers fine upward, terminating with thin laminae of very-fine sand or silt at the tops of the sandy intervals (Fig. 3). Sandy organic debris layers (TDL), which commonly cap tsunami sand sheets in distal alluvial runup localities of the Cascadia margin (Carver et al., 1998; Peterson et al., 2008), are uncommon in the proximal runup localities observed in the central Oregon study area. These data demonstrate that as many as four paleotsunami inundations substantially overtopped the barrage dune barriers (6.0 to 7.5 m elevation) at the mouths of corresponding creeks during the

Grant Creek beach Beach platform stump 3750±60 3988-4228 B118658 Lost Creek beach Beach dune ramp 3420±60 3600-3720 B148094 Deer Creek beach Beach platform stump 3920±60 4255-4421 B81340 Quail Creek mouth Top of barrage dune 2930±40 3018-3152 B172772 Table 4. Radiocarbon dates for dune ramps and creek barrage dunes in the study area. Data for Grant Creek beach platform stump (UTM 4936010n 415050e), Lost Creek beach dune ramp (UTM 4933080n 414770e), Deer Creek beach platform stump (UTM 4928880n 414110e, and Quail Creek top of barrage dune (UTM 4926060n 413880e) are from (Hart & Peterson, 2007). The dated sites of buried beach platform stumps and sea cliff dune ramps are shown in Fig. 2**.** 

**yr ± error** 

**Cal RC 2- (yr BP)** 

**Lab** 

**Site Setting adjC14** 

### **4. Results**

Fig. 2. Study area includes 9 creek floodplain localities (maps), and buried beach platform stumps or dune ramp forest soils at Grant Creek, Deer Creek, Lost Creek, and Quail Creek. Beach platform and dune ramp radiocarbon dates are shown in Table 4. Representative core sites are summarized in Table 3. Core logs and core site position data are provided in Cruikshank and Peterson (2011).

#### **4.1 Paleotsunami records in proximal high-gradient alluvial wetlands**

Eight proximal runup localities with inundation thresholds of 6.0 to 7.5 m elevation (NAVD88) in the study area are found to contain multiple anomalous sand sheets in creek floodplain deposits (Table 3). Alluvial mud hosts the anomalous sand sheets, which fine upward and thin landward in the small floodplain settings (Fig. 3). Additional creek floodplains exist within the 100 km long study region, but the tight grouping of eight localities within a 35 km study area permits comparisons of different catchment responses to similar conditions of tsunami surge forcing. The southern grouping of alluvial creek wetlands is divided on either side of the low gradient Beaver Creek valley that contains the longest overland inundation records reported to date for the central Cascadia margin (Table 2).

Six out of the eight creek localities record three or four prominent sand sheets, suggesting a similarity in age span of the alluvial hosting deposits. Radiocarbon dating of the alluvial section in Grant Creek, containing three sand sheets and one sandy debris layer, yields an age range from post-modern carbon (107pmc) to 3480-3690 BP (Table 4). The age span for

Fig. 2. Study area includes 9 creek floodplain localities (maps), and buried beach platform stumps or dune ramp forest soils at Grant Creek, Deer Creek, Lost Creek, and Quail Creek. Beach platform and dune ramp radiocarbon dates are shown in Table 4. Representative core sites are summarized in Table 3. Core logs and core site position data are provided in

Eight proximal runup localities with inundation thresholds of 6.0 to 7.5 m elevation (NAVD88) in the study area are found to contain multiple anomalous sand sheets in creek floodplain deposits (Table 3). Alluvial mud hosts the anomalous sand sheets, which fine upward and thin landward in the small floodplain settings (Fig. 3). Additional creek floodplains exist within the 100 km long study region, but the tight grouping of eight localities within a 35 km study area permits comparisons of different catchment responses to similar conditions of tsunami surge forcing. The southern grouping of alluvial creek wetlands is divided on either side of the low gradient Beaver Creek valley that contains the longest

overland inundation records reported to date for the central Cascadia margin (Table 2).

Six out of the eight creek localities record three or four prominent sand sheets, suggesting a similarity in age span of the alluvial hosting deposits. Radiocarbon dating of the alluvial section in Grant Creek, containing three sand sheets and one sandy debris layer, yields an age range from post-modern carbon (107pmc) to 3480-3690 BP (Table 4). The age span for

**4.1 Paleotsunami records in proximal high-gradient alluvial wetlands** 

**4. Results** 

Cruikshank and Peterson (2011).


Table 3. Representative proximal runup core sites in the central Oregon coast. Core site elevation is based on core site surface elevations (m, NAVD88) taken from LiDAR (U. S. Geological Survey, 2011). Distance is the straight-line distance (m) of surge flow from beach backshore or sea cliff to the representative core sites. Threshold elevation (m, NAVD88) is based on the mean across-valley elevation at or near the shoreline. TSL number is the number of observed tsunami sand layers of least 1 cm sand thickness in each runup locality. TSL maximum thickness is the thickest target tsunami sand layer (cm) observed in the runup locality. Core sites in creek floodplain localities are shown in Fig. 2. Data from Cruikshank and Peterson (2011).

the Grant Creek silt alluvium, located just above basal creek gravel, is consistent with an expected onset of creek damming by dune ramp barrages that occurred at about 4 ka in the study area (see Section 4.1).

The tsunami sand layers (TSL) decrease in number and thickness with increasing landward distance and elevation in each runup locality (Fig. 2). The tsunami sand layers fine upward, terminating with thin laminae of very-fine sand or silt at the tops of the sandy intervals (Fig. 3). Sandy organic debris layers (TDL), which commonly cap tsunami sand sheets in distal alluvial runup localities of the Cascadia margin (Carver et al., 1998; Peterson et al., 2008), are uncommon in the proximal runup localities observed in the central Oregon study area. These data demonstrate that as many as four paleotsunami inundations substantially overtopped the barrage dune barriers (6.0 to 7.5 m elevation) at the mouths of corresponding creeks during the last several thousand years.


Table 4. Radiocarbon dates for dune ramps and creek barrage dunes in the study area. Data for Grant Creek beach platform stump (UTM 4936010n 415050e), Lost Creek beach dune ramp (UTM 4933080n 414770e), Deer Creek beach platform stump (UTM 4928880n 414110e, and Quail Creek top of barrage dune (UTM 4926060n 413880e) are from (Hart & Peterson, 2007). The dated sites of buried beach platform stumps and sea cliff dune ramps are shown in Fig. 2**.** 

Proximal Records of Paleotsunami Runup in Barrage Creek Floodplains from

Late-Holocene Great Earthquakes in the Central Cascadia Suduction Zone, Oregon, USA 43

Several recent paleotsunami with limited inundations in distal runup localities, Neskowin and Beaver Creek, include the last three central Cascadia tsunami events TSL #1, #2a and #2b, (Table 3). Paleotsunami events with expected low runup heights, including events #1 and #2a, are recorded in the low elevation wetlands at the mouth of Beaver Creek in core site BEAV\_D1 (Fig. 4). These lower runup events did broadly overtop minimum threshold elevations of the beach backshore at 4.5 m elevation that fronts the mouth of Beaver Creek (Fig. 2). Based on these minimum threshold elevations of recorded paleotsunami inundation we estimate minimum runup heights of at least 5.0 m for nearfield tsunami in the central Cascadia margin.

Fig. 4. Representative core logs from nine creek runup localities in the central Oregon coast study area, including Big Creek (BIGC), Henderson Creek (HEND), Grant Creek (GRAN), Moore Creek (MOOR), Theil Creek (THEI), Pumphouse Creek (PUMP), Beaver Creek (BEAV), Tillicum Creek (TILL), and Vingie Creek (VING). Target tsunami layers are verified

Three runup localities in the study area, including Grant, Moore, Theil, and Vingie Creeks, contain anomalous sandy silt layers that can be correlated between successive target

in two adjacent core sites from each elevated floodplain, except Moore Creek, where additional core logs are shown in later sections of the paper. The position data for the localities and corresponding core sites are shown in Fig. 2 and Table 3. Core site elevations (m) are relative to the NAVD88 datum. One proximal locality (BEAV\_D1) is a low elevation back-barrier wetland, whereas the other eight creek floodplain localities all exceed 6 m

elevations at their shoreline barrage-dune thresholds.

Fig. 3. Contrasting paleotsunami sand layers shown between black arrows for (A) Grant Creek (40-47 cm depth at core site GRAN1) and (B) Tillicum Creek (151-156 cm depth at core site TILL1). The sand sheet deposit from Grant Creek (5-6 cm thick) shows 1) an undisturbed tsunami sand layer (TSL), 2) clear contrast with oxidized hosting alluvium, 3) a sharp bottom contact, 4) fining-upward grain-size trend, and 5) a dark organic-rich debris layer (TDL). The sand sheet deposit from Tillicum Creek (~ 3-4 cm thick) is much disturbed by root bioturbation and the quartz-rich beach sand is obscured by the dark reducing hosting mud. However it does have a sharp lower contact, and it fines upward in grain-size, though no capping debris layer is apparent. Red tape rule is scaled at 1 mm intervals. See Fig. 2 for core site locations.

Fig. 3. Contrasting paleotsunami sand layers shown between black arrows for (A) Grant Creek (40-47 cm depth at core site GRAN1) and (B) Tillicum Creek (151-156 cm depth at core

undisturbed tsunami sand layer (TSL), 2) clear contrast with oxidized hosting alluvium, 3) a sharp bottom contact, 4) fining-upward grain-size trend, and 5) a dark organic-rich debris layer (TDL). The sand sheet deposit from Tillicum Creek (~ 3-4 cm thick) is much disturbed by root bioturbation and the quartz-rich beach sand is obscured by the dark reducing hosting mud. However it does have a sharp lower contact, and it fines upward in grain-size, though no capping debris layer is apparent. Red tape rule is scaled at 1 mm intervals. See

site TILL1). The sand sheet deposit from Grant Creek (5-6 cm thick) shows 1) an

Fig. 2 for core site locations.

Several recent paleotsunami with limited inundations in distal runup localities, Neskowin and Beaver Creek, include the last three central Cascadia tsunami events TSL #1, #2a and #2b, (Table 3). Paleotsunami events with expected low runup heights, including events #1 and #2a, are recorded in the low elevation wetlands at the mouth of Beaver Creek in core site BEAV\_D1 (Fig. 4). These lower runup events did broadly overtop minimum threshold elevations of the beach backshore at 4.5 m elevation that fronts the mouth of Beaver Creek (Fig. 2). Based on these minimum threshold elevations of recorded paleotsunami inundation we estimate minimum runup heights of at least 5.0 m for nearfield tsunami in the central Cascadia margin.

Fig. 4. Representative core logs from nine creek runup localities in the central Oregon coast study area, including Big Creek (BIGC), Henderson Creek (HEND), Grant Creek (GRAN), Moore Creek (MOOR), Theil Creek (THEI), Pumphouse Creek (PUMP), Beaver Creek (BEAV), Tillicum Creek (TILL), and Vingie Creek (VING). Target tsunami layers are verified in two adjacent core sites from each elevated floodplain, except Moore Creek, where additional core logs are shown in later sections of the paper. The position data for the localities and corresponding core sites are shown in Fig. 2 and Table 3. Core site elevations (m) are relative to the NAVD88 datum. One proximal locality (BEAV\_D1) is a low elevation back-barrier wetland, whereas the other eight creek floodplain localities all exceed 6 m elevations at their shoreline barrage-dune thresholds.

Three runup localities in the study area, including Grant, Moore, Theil, and Vingie Creeks, contain anomalous sandy silt layers that can be correlated between successive target

Proximal Records of Paleotsunami Runup in Barrage Creek Floodplains from

Late-Holocene Great Earthquakes in the Central Cascadia Suduction Zone, Oregon, USA 45

Fig. 6. Photos of representative tsunami sand layers (TSL) from the Moore Creek locality. Core sections are listed by core site number and depth (cm). Tsunami sand layers are

tsunami debris layer. See Fig. 7 for core logs.

bracketed by white arrows. The tsunami sand layer in MOOR01 continues 154–163 cm depth in the photo to 187 cm depth below the photo, totaling 33 cm in thickness. The thin sandy layer in MOOR12 at 91-92 cm depth is transitional between a tsunami sand layer and sandy

tsunami sand layers in adjacent core sites (Cruikshank & Peterson, 2011). Such isolated tsunami debris layers (TDL) are shown in core sites GRAN5/6 and MOOR3 (Fig. 4). These layers might represent terminal inundations by smaller tsunami events. Other smaller target layers (< 1 cm thickness) might exist in the seaward sites of the study area localities, but are not further addressed in this study.


Table 5. Radiocarbon dates for paleotsunami in proximal runup sites, central Oregon coast. Radiocarbon dates are taken from base of tsunami sand layers, as shown by core depth (cm). See Fig. 2 for core sites and Fig. 4 and Fig. 7 for core logs.

Fig. 5. Moore Creek core site map including 12 representative core sites ranging from 8 to 13 m in elevation (NAVD88). Contour base map is based on LIDAR DEM (U. S. Geological Survey, 2011). 1m contour interval, with bold contours at 10, 20, and 30 m elevation. See Fig. 2 for location of Moore Creek locality in the study area. Core photos and core logs are shown in Fig. 6 and Fig. 7, respectively.

Though flow height and duration are both important in surge magnitude, for this study we focus on estimating runup height in the proximal floodplain settings. To establish maximum-recorded runup elevations the tsunami sand sheets are traced to their most landward extent in several creek localities. The terminal sand sheet deposits that show the best preservation and continuity occur in the Moore Creek floodplain locality (Fig. 2).

tsunami sand layers in adjacent core sites (Cruikshank & Peterson, 2011). Such isolated tsunami debris layers (TDL) are shown in core sites GRAN5/6 and MOOR3 (Fig. 4). These layers might represent terminal inundations by smaller tsunami events. Other smaller target layers (< 1 cm thickness) might exist in the seaward sites of the study area localities, but are

> **adjC14 yr ± error**

GRAN6\_54 54 107±0.5pMC modern B281145 GRAN6\_161 161 3360±40 3,480-3,690 B281146 MOOR9\_86 86 1,830±40 1,640-1,650 B230659 MOOR9\_120 120 2,720±40 2,750-2,880 B228599 MOOR9\_129 128.5 2940±40 2,960-3,230 B222512 MOO9-200 200 3,390±40 3,560-3,710 B228600 Table 5. Radiocarbon dates for paleotsunami in proximal runup sites, central Oregon coast. Radiocarbon dates are taken from base of tsunami sand layers, as shown by core depth (cm).

Fig. 5. Moore Creek core site map including 12 representative core sites ranging from 8 to 13 m in elevation (NAVD88). Contour base map is based on LIDAR DEM (U. S. Geological Survey, 2011). 1m contour interval, with bold contours at 10, 20, and 30 m elevation. See Fig. 2 for location of Moore Creek locality in the study area. Core photos and core logs are

Though flow height and duration are both important in surge magnitude, for this study we focus on estimating runup height in the proximal floodplain settings. To establish maximum-recorded runup elevations the tsunami sand sheets are traced to their most landward extent in several creek localities. The terminal sand sheet deposits that show the best preservation and continuity occur in the Moore Creek floodplain locality (Fig. 2).

**Cal RC 2- (yr BP)** 

**Lab** 

not further addressed in this study.

**Locality/ Site** 

**Depth (cm)** 

See Fig. 2 for core sites and Fig. 4 and Fig. 7 for core logs.

shown in Fig. 6 and Fig. 7, respectively.

Fig. 6. Photos of representative tsunami sand layers (TSL) from the Moore Creek locality. Core sections are listed by core site number and depth (cm). Tsunami sand layers are bracketed by white arrows. The tsunami sand layer in MOOR01 continues 154–163 cm depth in the photo to 187 cm depth below the photo, totaling 33 cm in thickness. The thin sandy layer in MOOR12 at 91-92 cm depth is transitional between a tsunami sand layer and sandy tsunami debris layer. See Fig. 7 for core logs.

Proximal Records of Paleotsunami Runup in Barrage Creek Floodplains from

the target paleotsunami sand sheets in the proximal floodplain settings.

Polyhalobous (marine)

Mesohalobous (brackish)

(see Section 4.4 below).

**4.3 Marine surge tracers** 

Late-Holocene Great Earthquakes in the Central Cascadia Suduction Zone, Oregon, USA 47

deposit layer thinning with increasing distance landward (Fig. 7). The combination of all three features in the three target sand layers is consistent with catastrophic marine surges but not with river flooding, paleoliquefaction, or hillslope debris flows. Further tests of marine surge origin are provided by marine source tracers in the target tsunami sand sheets

The use of marine diatoms to confirm paleotsunami inundation in the proximal flood plain settings that were investigated in this study proved to be problematic. Diatoms occur in the silt size range, and are potentially very useful in discriminating marine surge sources of target paleotsunami layers in the Cascadia margin (Hemphill-Haley, 1996; Kelsey et al., 2005; Schlichting & Peterson, 2006). In this study a composite sample was taken from the top of the target tsunami layer where the fine to very-fine sand in the tsunami sand layer grades into sandy silt or sandy detrital organics in the tsunami debris layer. Ocean derived diatoms including full marine and brackish taxa are present in the paleotsunami layers from the Moore Creek locality (Table 6). As expected, diagnostic freshwater diatoms are also present in the flood plain soils. Diagnostic ocean derived diatoms are also present in very-low abundance in control intervals are that are not associated with target paleotsunami sand layers or paleotsunami debris layers in the Moore Creek locality. The presence of marine and brackish diatoms in trace to minor abundances in non-tsunami control layers, argues for ocean diatom transport in wind driven ocean foam and/or spray in proximal settings (100- 400 m from the beach). Additional evidence is needed to confirm the marine surge origins of

**Diatom Assemblage/ Taxa Core Site Depth (cm) Target layer** 

*Hyalodiscus scoticus* MOOR9 55-60 control *Coscinodiscus radiatus* MOOR12 54-58 TSL/TDL *Cocconeis scutellum* MOOR12 54-58 TSL/TDL *Endycta sp.* MOOR12 100-101 control

*Navicula lanceola* MOOR9 86-88 TSL *Pinnularia viridis* MOOR9 119-122 TSL *Pinnularia viridis* MOOR12 128-130 TSL/TDL *Opephora parva* MOOR12 128-130 TSL/TDL

Table 6. Presence of diagnostic diatom taxa in proximal alluvial deposits. Control layers are non-tsunami layers. TSL are target paleotsunami sand layers. TDL are target paleotsunami debris layers. Core sites and sample intervals are shown in Fig. 5 and Fig. 7, respectively.

*Amphora ovalis* MOOR9, 12 *Eunotia pectinal* MOOR9, 12 *Gomphonema parvulum* MOOR9, 12

Oligohalobous (freshwater) MOOR9, 12

#### **4.2 Moore creek runup locality**

A total of 12 representative core sites are logged to document the landward thinning of sand sheets hosted in alluvial silts of the Moore Creek locality (Fig. 5). The sand sheets thin to only a few centimeters in thickness with increasing landward distance and elevation gain in the high-gradient floodplain setting. However, the quartz-rich beach sand layers are distinctive in the gray silt alluvium (Fig. 6).

Fig. 7. Core Logs from the Moore Creek runup locality showing tsunami sand layers (TSL) and tsunami debris layers (TDL) in representative core sites. See Fig. 5 for core site positions and Table 8 for core site elevations.

The three sand sheets that largely span the length of the Moore Creek runup locality are dated by AMS radiocarbon analyses of organic debris in the base of the sand sheets in core site MOOR9 (Fig. 7**;** Table 4). The radiocarbon samples yield dates of 1640–1650 BP, 2750–2880 BP, and 2960–3230 BP, corresponding to Cascadia tsunami events #3, #5, and #6–7, respectively (Table 1). A radiocarbon date of 3560–3710 BP dates the onset of alluvial silt accumulation in the Moore Creek floodplain at site MOOR9. Several anomalous sand laminae (< 1 cm thickness) are apparent above or between the prominent sand sheets in the seaward core sites MOOR1, MOOR2, MOOR3 and MOOR4. The origins of these minor lamiane are not known, but could reflect terminal deposits of smaller runup events. Due to their lack of continuity in the landward core sites they are not further addressed in this paper.

Three depositional features are common in the three target tsunami deposit layers from the Moore Creek core sites, including sharp lower contacts, sediment fining-upward trends, and deposit layer thinning with increasing distance landward (Fig. 7). The combination of all three features in the three target sand layers is consistent with catastrophic marine surges but not with river flooding, paleoliquefaction, or hillslope debris flows. Further tests of marine surge origin are provided by marine source tracers in the target tsunami sand sheets (see Section 4.4 below).

#### **4.3 Marine surge tracers**

46 Tsunami – A Growing Disaster

A total of 12 representative core sites are logged to document the landward thinning of sand sheets hosted in alluvial silts of the Moore Creek locality (Fig. 5). The sand sheets thin to only a few centimeters in thickness with increasing landward distance and elevation gain in the high-gradient floodplain setting. However, the quartz-rich beach sand layers are

Fig. 7. Core Logs from the Moore Creek runup locality showing tsunami sand layers (TSL) and tsunami debris layers (TDL) in representative core sites. See Fig. 5 for core site positions

The three sand sheets that largely span the length of the Moore Creek runup locality are dated by AMS radiocarbon analyses of organic debris in the base of the sand sheets in core site MOOR9 (Fig. 7**;** Table 4). The radiocarbon samples yield dates of 1640–1650 BP, 2750–2880 BP, and 2960–3230 BP, corresponding to Cascadia tsunami events #3, #5, and #6–7, respectively (Table 1). A radiocarbon date of 3560–3710 BP dates the onset of alluvial silt accumulation in the Moore Creek floodplain at site MOOR9. Several anomalous sand laminae (< 1 cm thickness) are apparent above or between the prominent sand sheets in the seaward core sites MOOR1, MOOR2, MOOR3 and MOOR4. The origins of these minor lamiane are not known, but could reflect terminal deposits of smaller runup events. Due to their lack of continuity in

Three depositional features are common in the three target tsunami deposit layers from the Moore Creek core sites, including sharp lower contacts, sediment fining-upward trends, and

the landward core sites they are not further addressed in this paper.

**4.2 Moore creek runup locality** 

distinctive in the gray silt alluvium (Fig. 6).

and Table 8 for core site elevations.

The use of marine diatoms to confirm paleotsunami inundation in the proximal flood plain settings that were investigated in this study proved to be problematic. Diatoms occur in the silt size range, and are potentially very useful in discriminating marine surge sources of target paleotsunami layers in the Cascadia margin (Hemphill-Haley, 1996; Kelsey et al., 2005; Schlichting & Peterson, 2006). In this study a composite sample was taken from the top of the target tsunami layer where the fine to very-fine sand in the tsunami sand layer grades into sandy silt or sandy detrital organics in the tsunami debris layer. Ocean derived diatoms including full marine and brackish taxa are present in the paleotsunami layers from the Moore Creek locality (Table 6). As expected, diagnostic freshwater diatoms are also present in the flood plain soils. Diagnostic ocean derived diatoms are also present in very-low abundance in control intervals are that are not associated with target paleotsunami sand layers or paleotsunami debris layers in the Moore Creek locality. The presence of marine and brackish diatoms in trace to minor abundances in non-tsunami control layers, argues for ocean diatom transport in wind driven ocean foam and/or spray in proximal settings (100- 400 m from the beach). Additional evidence is needed to confirm the marine surge origins of the target paleotsunami sand sheets in the proximal floodplain settings.


Table 6. Presence of diagnostic diatom taxa in proximal alluvial deposits. Control layers are non-tsunami layers. TSL are target paleotsunami sand layers. TDL are target paleotsunami debris layers. Core sites and sample intervals are shown in Fig. 5 and Fig. 7, respectively.

Proximal Records of Paleotsunami Runup in Barrage Creek Floodplains from

elevations reach 11–12 m NAVD88 in core sites MOOR12 and MOOR13.

**5.1 Landward trends of paleotsunami deposits in proximal settings** 

sand layers from the bioturbated floodplain deposits.

10 m elevation at the Moore Creek runup locality.

**4.5 Runup elevations** 

**5. Discussion** 

**5.2 Event runup heights** 

Late-Holocene Great Earthquakes in the Central Cascadia Suduction Zone, Oregon, USA 49

Paleotsunami deposit elevations are taken from core site surface elevations and corresponding subsurface depths of tsunami deposits (Table 8**)**. There is a close correspondence between core site surface elevations as surveyed into registered benchmarks and as estimated by GPS located sites on the LIDAR digital elevation model (U. S. Geological Survey, 2011). For the runup elevation measurements shown here we use the benchmark survey elevation control (± 5 cm elevation accuracy) in all but 3 core sites. Elevation based on LIDAR is used for core sites MOOR13-MOOR15. Terminal sand sheet

The measured deposit elevations in the Moore Creek locality are corrected for paleo-sea level runup height based on the age of the paleotsunami event (Table 1) and an assumed net rate of sea level rise (1.0 m per 1000 years) for the study area during late Holocene time (Darienzo et al., 1994). Paleotsunami runup heights that are corrected for net sea level rise reach 13–15 m for tsunami events #3, #5 and #6 in the Moore Creek locality. It is assumed that the terminal sand sheet proxies for runup height underestimate actual flooding elevations. However, the apparent pinchouts of the corresponding paleotsunami debris layers near core site MOOR15 suggest that maximum recorded paleotsunami runup height is closely approximated by the terminal sand sheet deposits in the Moore Creek locality.

Several parameters including tsunami deposit sand: silt ratio, sand layer thickness, and abundance of coarse grained shell fragments are found to substantially decrease with elevation gain (from 8 m to 12 m) over the short inundation distance of 450 m in the Moore Creek locality (Fig. 7). A small, but statistically significant, decrease of mean sand size is also documented in the paleotsunami sand layers in the Moore Creek cores sites (Table 7). We attribute these landward trends to decreases in tsunami flow velocity and turbulence with lansdward increases in floodplain elevation (Fig. 8). Terminal sand sheet layers are on the order of only a few centimeters in thickness, requiring extensive coring to recover tsunami

A total of three sand sheets that reach 11-12 m elevation are recorded during the last 3.2 ka in the Moore Creek locality (Fig. 8). The last two events (#5 and # 6) are attributed to large magnitude runups in other central Cascadia tidal basins, back-barrier wetlands, and beach plains (Peterson et al., 2010a; Peterson et al., 2010b; Schlichting, 2000). These two paleotsunami are correlated to long inundation events in the adjacent Beaver Creek floodplain (see Section 5.3 below). When adjusted for Paleo-sea level at the time of inundation (Table 8) the three paleotsunami sand sheets reach 13-14 m runup height at the landward side of the Moore Creek runup locality. Three of the 6–7 nearfield tsunami produced by ruptures of the central Cascadia megathrust during the last 3.2 ka yielded large runup elevations in the study area. These runups would have reached 15 m above the current O m NAVD88 datum. The remaining 3–4 paleotsunami exceeded 5 m runup height, as shown by tsunami deposits at the BEAV\_D1 core site in the Beaver Creek locality (Fig. 4). These smaller magnitude tsunami did not leave sand sheet deposits above 9 m elevation in Moore Creek locality (Fig. 7). We assume that their maximum runup heights did not exceed


Table 7. Sand grain size and shell fragment frequency Tsunami sand layer samples are from three core sites (MOOR1, MOOR9 and MOOR12 in the Moore Creek runup locality. Population characteristics including maximum sand size (Max), minimum sand size (Min), average (mean), 1 standard deviation (Std Dev), and mean normalized standard deviation (MeanNStdDev) (Folk, 1980) are based on 50 grain counts. Carbonate shell fragment frequency (Shell frag %) are based on 1,000 grain counts. Core sites and sample intervals are shown in Fig. 5 and Fig. 7, respectively.

Carbonate shell fragments were identified in the gravel size fractions of target paleotsunami sand sheets in 7 out of the 8 runup localities (Fig. 4). Shell fragments were found in 9 of the 12 sand sheet samples that contained minor abundances of granules (size range 2-5 mm diameter). The granule abundance abruptly decreases to trace levels with increasing distance from the beach (200-400 m) and with increasing elevation (8–12 m) in the Moore Creek locality. Petrographic microscopy of the carbonate grains in the sand fractions of the target paleotsunami sand sheets in the Moore Creek locality are shown in Table 7. Though not as distinctive as the larger shell fragments, the trace abundances of the carbonate sandsize grains in representative tsunami sand layers from core sites MOOR1, MOOR9, and MOOR12, confirm that the thin sand layers were derived from littoral or inner-shelf sediment sources. Other investigators have examined mollusan shells in paleotsunami deposits (Fujiwara et al., 2003) and have recently reported the presence of foraminfera tests in tsunami deposits from the Sumatra 2004 tsunami (Hawkes et al., 2007). The use of foraminfera was not tested in this study, but it could provide a complimentary technique to the use of carbonate shell fragments, as reported here.

#### **4.4 Grain size**

The sand grain size distributions from representative samples of tsunami sand layers in the Moore Creek locality demonstrate landward fining trends (Table 7). Population means of the sand grain samples decrease from 341–375 µm in the seaward site MOOR1 to 229-245 µm in the landward site MOOR12. The apparent differences between the sand population means from the seaward and landward core sites are statistically different at the 95% confidence intervals. Both the separation in distance (~ 350 m) and elevation (4 m) between the seaward and landward sites are thought to contribute to the small, but significant, landward fining of the mean sand size.

#### **4.5 Runup elevations**

48 Tsunami – A Growing Disaster

Beach 460 180 269 56 0.21 0.27 Creek 1030 150 470 234 0.50 0.00 1\_100 cm #3 610 190 341 79 0.23 0.14 1\_126 cm #5 740 230 375 107 0.29 0.00 1\_151 cm #6-7 550 120 350 94 0.27 0.21 9\_86 cm #3 510 190 331 87 0.26 0.36 9\_119 cm #5 480 160 278 65 0.23 0.00 9\_128 cm #6-7 400 100 275 68 0.25 0.10 12\_54 cm #3 430 150 245 68 0.28 0.10 12\_82 cm #5 420 100 229 69 0.30 0.15 12\_91 cm #6-7 490 90 231 69 0.30 0.00 Table 7. Sand grain size and shell fragment frequency Tsunami sand layer samples are from

**Mean Size (µm)**  **Std Dev +/- (µm)** 

**MeanN StdDev** **Shell frag (%)** 

**Min. size (µm)** 

three core sites (MOOR1, MOOR9 and MOOR12 in the Moore Creek runup locality. Population characteristics including maximum sand size (Max), minimum sand size (Min), average (mean), 1 standard deviation (Std Dev), and mean normalized standard deviation (MeanNStdDev) (Folk, 1980) are based on 50 grain counts. Carbonate shell fragment

frequency (Shell frag %) are based on 1,000 grain counts. Core sites and sample intervals are

Carbonate shell fragments were identified in the gravel size fractions of target paleotsunami sand sheets in 7 out of the 8 runup localities (Fig. 4). Shell fragments were found in 9 of the 12 sand sheet samples that contained minor abundances of granules (size range 2-5 mm diameter). The granule abundance abruptly decreases to trace levels with increasing distance from the beach (200-400 m) and with increasing elevation (8–12 m) in the Moore Creek locality. Petrographic microscopy of the carbonate grains in the sand fractions of the target paleotsunami sand sheets in the Moore Creek locality are shown in Table 7. Though not as distinctive as the larger shell fragments, the trace abundances of the carbonate sandsize grains in representative tsunami sand layers from core sites MOOR1, MOOR9, and MOOR12, confirm that the thin sand layers were derived from littoral or inner-shelf sediment sources. Other investigators have examined mollusan shells in paleotsunami deposits (Fujiwara et al., 2003) and have recently reported the presence of foraminfera tests in tsunami deposits from the Sumatra 2004 tsunami (Hawkes et al., 2007). The use of foraminfera was not tested in this study, but it could provide a complimentary technique to

The sand grain size distributions from representative samples of tsunami sand layers in the Moore Creek locality demonstrate landward fining trends (Table 7). Population means of the sand grain samples decrease from 341–375 µm in the seaward site MOOR1 to 229-245 µm in the landward site MOOR12. The apparent differences between the sand population means from the seaward and landward core sites are statistically different at the 95% confidence intervals. Both the separation in distance (~ 350 m) and elevation (4 m) between the seaward and landward sites are thought to contribute to the small, but significant,

**TSL number**

shown in Fig. 5 and Fig. 7, respectively.

the use of carbonate shell fragments, as reported here.

landward fining of the mean sand size.

**4.4 Grain size** 

**Max. size (µm)** 

Paleotsunami deposit elevations are taken from core site surface elevations and corresponding subsurface depths of tsunami deposits (Table 8**)**. There is a close correspondence between core site surface elevations as surveyed into registered benchmarks and as estimated by GPS located sites on the LIDAR digital elevation model (U. S. Geological Survey, 2011). For the runup elevation measurements shown here we use the benchmark survey elevation control (± 5 cm elevation accuracy) in all but 3 core sites. Elevation based on LIDAR is used for core sites MOOR13-MOOR15. Terminal sand sheet elevations reach 11–12 m NAVD88 in core sites MOOR12 and MOOR13.

The measured deposit elevations in the Moore Creek locality are corrected for paleo-sea level runup height based on the age of the paleotsunami event (Table 1) and an assumed net rate of sea level rise (1.0 m per 1000 years) for the study area during late Holocene time (Darienzo et al., 1994). Paleotsunami runup heights that are corrected for net sea level rise reach 13–15 m for tsunami events #3, #5 and #6 in the Moore Creek locality. It is assumed that the terminal sand sheet proxies for runup height underestimate actual flooding elevations. However, the apparent pinchouts of the corresponding paleotsunami debris layers near core site MOOR15 suggest that maximum recorded paleotsunami runup height is closely approximated by the terminal sand sheet deposits in the Moore Creek locality.

#### **5. Discussion**

#### **5.1 Landward trends of paleotsunami deposits in proximal settings**

Several parameters including tsunami deposit sand: silt ratio, sand layer thickness, and abundance of coarse grained shell fragments are found to substantially decrease with elevation gain (from 8 m to 12 m) over the short inundation distance of 450 m in the Moore Creek locality (Fig. 7). A small, but statistically significant, decrease of mean sand size is also documented in the paleotsunami sand layers in the Moore Creek cores sites (Table 7). We attribute these landward trends to decreases in tsunami flow velocity and turbulence with lansdward increases in floodplain elevation (Fig. 8). Terminal sand sheet layers are on the order of only a few centimeters in thickness, requiring extensive coring to recover tsunami sand layers from the bioturbated floodplain deposits.

#### **5.2 Event runup heights**

A total of three sand sheets that reach 11-12 m elevation are recorded during the last 3.2 ka in the Moore Creek locality (Fig. 8). The last two events (#5 and # 6) are attributed to large magnitude runups in other central Cascadia tidal basins, back-barrier wetlands, and beach plains (Peterson et al., 2010a; Peterson et al., 2010b; Schlichting, 2000). These two paleotsunami are correlated to long inundation events in the adjacent Beaver Creek floodplain (see Section 5.3 below). When adjusted for Paleo-sea level at the time of inundation (Table 8) the three paleotsunami sand sheets reach 13-14 m runup height at the landward side of the Moore Creek runup locality. Three of the 6–7 nearfield tsunami produced by ruptures of the central Cascadia megathrust during the last 3.2 ka yielded large runup elevations in the study area. These runups would have reached 15 m above the current O m NAVD88 datum. The remaining 3–4 paleotsunami exceeded 5 m runup height, as shown by tsunami deposits at the BEAV\_D1 core site in the Beaver Creek locality (Fig. 4). These smaller magnitude tsunami did not leave sand sheet deposits above 9 m elevation in Moore Creek locality (Fig. 7). We assume that their maximum runup heights did not exceed 10 m elevation at the Moore Creek runup locality.

Proximal Records of Paleotsunami Runup in Barrage Creek Floodplains from

Late-Holocene Great Earthquakes in the Central Cascadia Suduction Zone, Oregon, USA 51

Fig. 8. Topographically corrected cross-section of stratigraphically correlated paleotsunami deposits (dotted lines) for events #3, #5, and #6–7 in the Moore Creek floodplain. Tsunami sand layers (TSL) pinchout to tsunami debris layers (TDL) with increasing distance and elevation from the present beach–terrace edge. Vertical exaggeration is ~ 100x. See Fig. 5 for

Using the maximum-recorded extents of tsunami sand sheet deposition as proxies for paleotsunami runup in Moore Creek and Beaver Creek (Fig. 2) we establish a runup height attenuation gradient for paleotsunami event #3 in the study area. Differences between paleotsunami event #3 sand layer elevations and runup distances in the proximal core site MOOR13 (12.3 m elevation at 0.45 km distance) (Table 8) and the distal core site BEAV03 (1.2 m elevation at 3.7 km distance) (Peterson et al., 2010a) yield an attenuation gradient of 3.4 m km-1 (Fig. 9). Using the attenuation gradient we extrapolate a shoreline runup elevation of 13.8 m at 0 m distance. Adjusting for Paleo-sea level (-1.0 m) at 1.3 ka we predict a runup height of 14.8 m based on paleotsunami sand sheet deposition. This extrapolated

The landward gradient of runup attenuation is tested in another central Cascadia locality, Cannon Beach, Oregon (Fig. 1), which was previously surveyed for distal and proximal

sand deposition height might underestimate actual surge height at the shoreline.

**5.4 Test of runup height attenuation gradient at Cannon Beach** 

**5.3 Landward runup height attenuation in Beaver Creek** 

core site locations.


Table 8. Core site and paleotsunami deposit elevations in the Moore Creek locality. Core site data: core top elevation from total station survey to registered benchmark/LIDAR (m) NAVD88 and overland inundation flow distance (m) from the shoreline. Tsunami deposit: event number (#), depth in core (m) and computed elevation relative to NAVD88 datum. Paleo-sea level estimated for event ages #3 (~ 1.3 ka), #5 (~ 2.6 ka), #6 (~ 2.8 ka), and # 7(~ 3.2 ka) assuming 1 m/1000 year relative sea level rise. Tsunami runup height (m) equivalent to NAVD88 datum but adjusted for lower paleo-sea level at the time of inundation. See Fig. 5 for core site locations.

1 7.98/8.0 10 # 3 TSL 1.00 6.98 -1.0 8.0 # 5 TSL 1.26 6.72 -2.5 9.2 #6-7 TSL 1.54 6.44 -3.0 9.4 2 8.14/8.2 20 #3 TSL 0.83 7.31 -1.0 8.3 #5 TSL 1.00 7.14 -2.5 9.6 #6-7 TSL 1.38 6.76 -3.0 9.8 3 8.29/8.2 90 #3 TSL 0.55 7.74 -1.0 8.7 TDL ? 1.00 7.29 -2.0 9.3 #5 TSL 1.15 7.14 -2.5 9.6 #6-7 TSL 1.53 6.76 -3.0 9.8 4/5 8.40/8.4 130 #3 TSL 0.39 8.01 -1.0 9.0 TDL ? 0.68 7.72 -2.0 9.7 #5 TSL 0.91 7.49 -2.5 10.0 #6-7 TSL 1.22 7.18 -3.0 10.2 8 9.33/9.5 180 #3 TSL 0.79 8.54 -1.0 9.5 #5 TSL 1.66 7.76 -2.5 10.3 #6-7 TSL 1.74 7.59 -3.0 10.6 9 9.69/9.7 180 #3 TSL 0.86 8.83 -1.0 9.8 #5 TSL 1.19 8.50 -2.5 11.0 #6-7 TSL 1.28 8.41 -3.0 11.4 10 10.57/11.9 270 #3 TSL 0.50 10.07 -1.0 11.1 #5 TSL 0.78 9.79 -2.5 12.3 #6-7 TSL 1.02 9.55 -3.0 12.5 11 10.52/10.6 280 #3 TSL 0.69 9.83 -1.0 10.8 #5 TSL 0.98 9.54 -2.5 12.0 12 12.17/13.3 400 #3 TSL 0.54 11.63 -1.0 12.6 #5 TSL 0.82 11.35 -2.5 13.8 #6-7 TSL 0.91 11.26 -3.0 14.3 13 na/12.8 450 #3 TSL 0.51 12.3 -1.0 13.3 #5 TSL 0.79 12.0 -2.5 14.5

**Core depth (m)** 

**Deposit elev. (m) NAVD88**

**Paleosea level (m)** 

0.98 11.8 -3.0 14.8

**Paleorunup height (m)** 

**Tsunami deposit event #** 

**Core Site** 

**Core top elev. (m)**  **Overland distance (m)** 

#6-7

5 for core site locations.

TDL

14 na/13.0 450 #3 TDL 0.4 12.6 -1.0 13.6 #5 TDL 0.97 12.0 -2.5 14.5 15 na/13.2 540 #3 TDL ? 0.5 12.7 -1.0 13.7 #5 TDL ? 0.85 12.3 -2.5 14.8 Table 8. Core site and paleotsunami deposit elevations in the Moore Creek locality. Core site data: core top elevation from total station survey to registered benchmark/LIDAR (m) NAVD88 and overland inundation flow distance (m) from the shoreline. Tsunami deposit: event number (#), depth in core (m) and computed elevation relative to NAVD88 datum. Paleo-sea level estimated for event ages #3 (~ 1.3 ka), #5 (~ 2.6 ka), #6 (~ 2.8 ka), and # 7(~ 3.2 ka) assuming 1 m/1000 year relative sea level rise. Tsunami runup height (m) equivalent to NAVD88 datum but adjusted for lower paleo-sea level at the time of inundation. See Fig.

Fig. 8. Topographically corrected cross-section of stratigraphically correlated paleotsunami deposits (dotted lines) for events #3, #5, and #6–7 in the Moore Creek floodplain. Tsunami sand layers (TSL) pinchout to tsunami debris layers (TDL) with increasing distance and elevation from the present beach–terrace edge. Vertical exaggeration is ~ 100x. See Fig. 5 for core site locations.

#### **5.3 Landward runup height attenuation in Beaver Creek**

Using the maximum-recorded extents of tsunami sand sheet deposition as proxies for paleotsunami runup in Moore Creek and Beaver Creek (Fig. 2) we establish a runup height attenuation gradient for paleotsunami event #3 in the study area. Differences between paleotsunami event #3 sand layer elevations and runup distances in the proximal core site MOOR13 (12.3 m elevation at 0.45 km distance) (Table 8) and the distal core site BEAV03 (1.2 m elevation at 3.7 km distance) (Peterson et al., 2010a) yield an attenuation gradient of 3.4 m km-1 (Fig. 9). Using the attenuation gradient we extrapolate a shoreline runup elevation of 13.8 m at 0 m distance. Adjusting for Paleo-sea level (-1.0 m) at 1.3 ka we predict a runup height of 14.8 m based on paleotsunami sand sheet deposition. This extrapolated sand deposition height might underestimate actual surge height at the shoreline.

#### **5.4 Test of runup height attenuation gradient at Cannon Beach**

The landward gradient of runup attenuation is tested in another central Cascadia locality, Cannon Beach, Oregon (Fig. 1), which was previously surveyed for distal and proximal

Proximal Records of Paleotsunami Runup in Barrage Creek Floodplains from

Late-Holocene Great Earthquakes in the Central Cascadia Suduction Zone, Oregon, USA 53

Fig. 10. Estimation of shoreline runup elevation for paleotsunami event #3 (~1.3 ka) in Cannon Beach, Oregon. Historic photograph of Cannon Beach shows Ecola Creek valley topography prior to extensive development in the middle to late 1900s. The proximal runup estimates are based on 1) terminal sand sheet deposition in core site C72 (solid circle in map inset) at 6.5 m elevation and 2.1 km flow inundation distance, and 2) reverse extrapolation of attenuation gradient 3.4 m km-1 (dashed line) to the shoreline (plot diagram). Adjustment for paleo-sea level yields a modern runup height of 15 m at the shoreline in Cannon Beach. Preliminary searches for possible paleotsunami sand sheets yielded two sand layers at CANU116 (UTM 5082190n 42542e) and a most landward extent of an anomalous sandy debris layer at CANU119 (UTM 5082054n 425584e) as shown in Fig. 11. No sand layer or sandy debris layers were observed at core sites CANU122 (UTM 5082134n 425662e) or CANU123 (UTM 5082078n 425654e). See Fig. 1 for the location of Cannon Beach in the

11.3 m and a landward distance of 300 m. No sand layers were observed at CANU122 at 12.7 m elevation, or CANU1213 at 13.5 m, both at a distance of 440 m from the beach. The anomalous sandy layers at CANU116 and CANU119 were not radiocarbon dated at the time of the reconnaissance survey (2006) due, in part, to a lack of target runup elevations. Such target elevations are now provided by the projected runups from the attenuation gradient

central Cascadia margin.

shown in Fig. 10.

Fig. 9. Runup attenuation for the #3 paleotsunami event in Moore Creek and Beaver Creek in the study area. The position and elevation of terminal sand sheet deposition are from Table 8 in Moore Creek and from (Peterson et al. (2010a)) in Beaver Creek. The proximal runup site (MOOR13) represents terminal sand sheet deposition in a high-gradient creek floodplain perched on a seaward facing hillslope. The distal site (BEAV03) represents terminal sand sheet deposition in a low gradient floodplain developed in a broad alluvial valley.

runup records (Peterson & Cruikshank, 2007; Peterson et al., 2008). The terminal extent of paleotsunami sand sheet deposition from event #3 (~ 1.3 ka) was traced to site C72 (6.5 m elevation at 2.1 km landward distance). Using the attenuation gradient of 3.4 m km-1 over the runup distance of 2.1 km and the core depth corrected elevation of 6.5 m for the event #3 sand layer at site C72 (Peterson et al., 2008) a predicted shoreline runup elevation of 13.6 m is calculated for Cannon Beach (Fig. 10). The projected elevation of shoreline runup at Cannon Beach (13.6 m) is similar to that projected for the Moore Creek locality (13.8 m) at the shoreline for the paleotsunami event # 3 (Fig. 9).

Upland terrace sites in Cannon Beach lack ideal hosting deposits for recording paleotsunami deposits. However two localities in a very small gully floodplain adjacent to Coolidge Avenue in Cannon Beach (sites CANU116 and CANU119) were identified as potential runup sites based on anomalous sandy intervals (Fig. 10). CANU116 at 8.4 m elevation, and 220 m landward distance, contained two prominent quartz-rich sand layers at 14–16 and 49– 51 cm depth (Fig. 11). The most landward extent of an anomalous sandy debris layer, including rounded granules, was observed at 64-67 cm depth in CANU119 at an elevation of

Fig. 9. Runup attenuation for the #3 paleotsunami event in Moore Creek and Beaver Creek in the study area. The position and elevation of terminal sand sheet deposition are from Table 8 in Moore Creek and from (Peterson et al. (2010a)) in Beaver Creek. The proximal runup site (MOOR13) represents terminal sand sheet deposition in a high-gradient creek floodplain perched on a seaward facing hillslope. The distal site (BEAV03) represents terminal sand sheet deposition in a low gradient floodplain developed in a broad alluvial

runup records (Peterson & Cruikshank, 2007; Peterson et al., 2008). The terminal extent of paleotsunami sand sheet deposition from event #3 (~ 1.3 ka) was traced to site C72 (6.5 m elevation at 2.1 km landward distance). Using the attenuation gradient of 3.4 m km-1 over the runup distance of 2.1 km and the core depth corrected elevation of 6.5 m for the event #3 sand layer at site C72 (Peterson et al., 2008) a predicted shoreline runup elevation of 13.6 m is calculated for Cannon Beach (Fig. 10). The projected elevation of shoreline runup at Cannon Beach (13.6 m) is similar to that projected for the Moore Creek locality (13.8 m) at

Upland terrace sites in Cannon Beach lack ideal hosting deposits for recording paleotsunami deposits. However two localities in a very small gully floodplain adjacent to Coolidge Avenue in Cannon Beach (sites CANU116 and CANU119) were identified as potential runup sites based on anomalous sandy intervals (Fig. 10). CANU116 at 8.4 m elevation, and 220 m landward distance, contained two prominent quartz-rich sand layers at 14–16 and 49– 51 cm depth (Fig. 11). The most landward extent of an anomalous sandy debris layer, including rounded granules, was observed at 64-67 cm depth in CANU119 at an elevation of

the shoreline for the paleotsunami event # 3 (Fig. 9).

valley.

Fig. 10. Estimation of shoreline runup elevation for paleotsunami event #3 (~1.3 ka) in Cannon Beach, Oregon. Historic photograph of Cannon Beach shows Ecola Creek valley topography prior to extensive development in the middle to late 1900s. The proximal runup estimates are based on 1) terminal sand sheet deposition in core site C72 (solid circle in map inset) at 6.5 m elevation and 2.1 km flow inundation distance, and 2) reverse extrapolation of attenuation gradient 3.4 m km-1 (dashed line) to the shoreline (plot diagram). Adjustment for paleo-sea level yields a modern runup height of 15 m at the shoreline in Cannon Beach. Preliminary searches for possible paleotsunami sand sheets yielded two sand layers at CANU116 (UTM 5082190n 42542e) and a most landward extent of an anomalous sandy debris layer at CANU119 (UTM 5082054n 425584e) as shown in Fig. 11. No sand layer or sandy debris layers were observed at core sites CANU122 (UTM 5082134n 425662e) or CANU123 (UTM 5082078n 425654e). See Fig. 1 for the location of Cannon Beach in the central Cascadia margin.

11.3 m and a landward distance of 300 m. No sand layers were observed at CANU122 at 12.7 m elevation, or CANU1213 at 13.5 m, both at a distance of 440 m from the beach. The anomalous sandy layers at CANU116 and CANU119 were not radiocarbon dated at the time of the reconnaissance survey (2006) due, in part, to a lack of target runup elevations. Such target elevations are now provided by the projected runups from the attenuation gradient shown in Fig. 10.

Proximal Records of Paleotsunami Runup in Barrage Creek Floodplains from

**6. Conclusion** 

inundation in historic time.

**7. Acknowledgments** 

of scientific exploration.

Schneider.

Late-Holocene Great Earthquakes in the Central Cascadia Suduction Zone, Oregon, USA 55

Small creek floodplains (< 500 m from the shoreline and > 6-7 m elevation) provide stable hosting settings for recording prehistoric tsunami inundation events in the central Cascadia subduction zone. The floodplain silts extend back to 3–4 ka in time, permitting the potential geologic recording of anomalous sand sheets from 6–7 nearfield tsunami during that time interval. A total of 3-4 paleotsunami deposits exceed 8 m in elevation, and 3 paleotsunami sand sheets can be traced to 12 m elevation. Adjusting for Paleo-sea level at the time of inundation the estimated runup for the 3 paleotsunami at the shoreline is 15 m in height. Minimum runups for nearfield tsunami in the study area are greater than 5 m height, based on the geologic record of smaller scale tsunami in a creek setting with a low-threshold elevation (~ 4.5 m) for inundation. The measured runup heights for central Cascadia paleotsunami during the last 3.2 ka are 10+/- 5 m. An attenuation gradient of 3.4 m km-1 is estimated from terminal sand sheet deposition produced by the last large runup event (at ~ 1.3 ka) as recorded in both proximal and distal floodplain settings in the study area. The attenuation gradient permits an extrapolation of distal terminal runup records to yield estimated shoreline runup heights for this tsunami event in other similar floodplain localities in the central Cascadia margin. The prehistoric runup records can be used to test and/or calibrate numerically modeled tsunami hazard in the region. The methodology used here should have broad use in other susceptible coastlines that have not experienced catastrophic tsunami

Roger Hart assisted with the first paleotsunami investigations in the Henderson creek floodplains in the study area. We dedicate this paper to Roger Hart (1940-2011) for his spirit

Robert Schlichting preformed the diatom analysis for this study. Anna Pilette assisted with detailed core logging, sampling, and photography of the Grant and Moore Creek tsunami sand sheets. Preliminary surveying of study area floodplains for target paleotsunami deposit included assistance from Steve Ahlquist, Andrea Adair, Tim Blazina, Adam Cambell, Charles Cannon, Annie Donehey, Brandon Ezzell, Ben Freudenberg, Christopher French, Jonathan Huster, Audra Inglish, Jenifer Lind, Robert Linscott, Fiona Seifert, Andrew Shaddox, Liza Shaw, Scott Waibel, Rex Whistler. Georg Grathoff and Karen Carroll provided XRD analyses of granule mineralogy from the Theil Creek paleotsunami deposits. David Percy provided assistance with LiDAR topographic imaging and elevation control of core sites. Meredith Savage and Derrick Tokos provided assistance with Newport City topographic interpretations. Stewart Cowy provided assistance with Lincoln County topographic flood maps and with interpretations of historic development features. The Oregon Department of Geology and Mineral Industries provided funding for field travel during the early phase of the creek floodplain surveys. The Office of Research at Portland State University provided some support for radiocarbon dating. LiDAR data was provided through USGS National Elevation Dataset (NED) with technical assistance from Sheri

Fig. 11. Photos of target paleotsunami sand layers from CANU116 at 14-16 cm depth (photo A), CANU116 at 49-51 cm depth (photo B), and CANU119 at 64-67 cm depth (photo C). A detrital wood fragment at 13 cm depth overlies the upper sand layer in CANU116 (photo A). The lower sand layer in CANU116 shows a sharp bottom contact and fining-upward sand grain size, both features are characteristic of paleotsunami deposits. The sandy debris layer at 64-67 cm depth in CANU119 includes rounded granules, well above wind-blown grain size thresholds. All three layers contain rounded quartz grains (beach sand source mineralogy) but the target paleotsunami sand layers from CANU116 have yet to be radiocarbon dated.

#### **6. Conclusion**

54 Tsunami – A Growing Disaster

Fig. 11. Photos of target paleotsunami sand layers from CANU116 at 14-16 cm depth (photo A), CANU116 at 49-51 cm depth (photo B), and CANU119 at 64-67 cm depth (photo C). A detrital wood fragment at 13 cm depth overlies the upper sand layer in CANU116 (photo A). The lower sand layer in CANU116 shows a sharp bottom contact and fining-upward sand grain size, both features are characteristic of paleotsunami deposits. The sandy debris layer at 64-67 cm depth in CANU119 includes rounded granules, well above wind-blown grain

size thresholds. All three layers contain rounded quartz grains (beach sand source mineralogy) but the target paleotsunami sand layers from CANU116 have yet to be

radiocarbon dated.

Small creek floodplains (< 500 m from the shoreline and > 6-7 m elevation) provide stable hosting settings for recording prehistoric tsunami inundation events in the central Cascadia subduction zone. The floodplain silts extend back to 3–4 ka in time, permitting the potential geologic recording of anomalous sand sheets from 6–7 nearfield tsunami during that time interval. A total of 3-4 paleotsunami deposits exceed 8 m in elevation, and 3 paleotsunami sand sheets can be traced to 12 m elevation. Adjusting for Paleo-sea level at the time of inundation the estimated runup for the 3 paleotsunami at the shoreline is 15 m in height. Minimum runups for nearfield tsunami in the study area are greater than 5 m height, based on the geologic record of smaller scale tsunami in a creek setting with a low-threshold elevation (~ 4.5 m) for inundation. The measured runup heights for central Cascadia paleotsunami during the last 3.2 ka are 10+/- 5 m. An attenuation gradient of 3.4 m km-1 is estimated from terminal sand sheet deposition produced by the last large runup event (at ~ 1.3 ka) as recorded in both proximal and distal floodplain settings in the study area. The attenuation gradient permits an extrapolation of distal terminal runup records to yield estimated shoreline runup heights for this tsunami event in other similar floodplain localities in the central Cascadia margin. The prehistoric runup records can be used to test and/or calibrate numerically modeled tsunami hazard in the region. The methodology used here should have broad use in other susceptible coastlines that have not experienced catastrophic tsunami inundation in historic time.

#### **7. Acknowledgments**

Roger Hart assisted with the first paleotsunami investigations in the Henderson creek floodplains in the study area. We dedicate this paper to Roger Hart (1940-2011) for his spirit of scientific exploration.

Robert Schlichting preformed the diatom analysis for this study. Anna Pilette assisted with detailed core logging, sampling, and photography of the Grant and Moore Creek tsunami sand sheets. Preliminary surveying of study area floodplains for target paleotsunami deposit included assistance from Steve Ahlquist, Andrea Adair, Tim Blazina, Adam Cambell, Charles Cannon, Annie Donehey, Brandon Ezzell, Ben Freudenberg, Christopher French, Jonathan Huster, Audra Inglish, Jenifer Lind, Robert Linscott, Fiona Seifert, Andrew Shaddox, Liza Shaw, Scott Waibel, Rex Whistler. Georg Grathoff and Karen Carroll provided XRD analyses of granule mineralogy from the Theil Creek paleotsunami deposits. David Percy provided assistance with LiDAR topographic imaging and elevation control of core sites. Meredith Savage and Derrick Tokos provided assistance with Newport City topographic interpretations. Stewart Cowy provided assistance with Lincoln County topographic flood maps and with interpretations of historic development features. The Oregon Department of Geology and Mineral Industries provided funding for field travel during the early phase of the creek floodplain surveys. The Office of Research at Portland State University provided some support for radiocarbon dating. LiDAR data was provided through USGS National Elevation Dataset (NED) with technical assistance from Sheri Schneider.

Proximal Records of Paleotsunami Runup in Barrage Creek Floodplains from

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**Tsunami Effect on Infrastructures** 

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