Geological evidence of recurrent great Kanto earthquakes at the Miura Peninsula, Japan

JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 116,, doi:10.1029/2011jb008639, 2011 Geological evidence of recurrent great Kanto earthquakes at the Miura Peninsula, Japan K. Shimazaki, 1 H. Y. Kim, 1 T. Chiba, 2 and K. Satake 1 Received 29 June 2011; revised 11 October 2011; accepted 13 October 2011; published 30 December 2011. [1] The Tokyo metropolitan area s well documented earthquake history is dominated by the 1703 and 1923 great Kanto earthquakes produced by slip on the boundary between the subducting Philippine Sea plate and the overlying plate. Both earthquakes caused 1.5 m of uplift at the Miura Peninsula directly above the inferred fault rupture, and both were followed by tsunamis with heights of 5 m. We examined cores 2 m long from 8 tidal flat sites at the head of a small bay on the peninsula. The cores penetrated two to four layers of shelly gravel, as much as 0.5 m thick, with abundant shell fragments and mud clasts. The presence of gravel indicates strong tractive currents. Muddy bay deposits that bound the gravel layers show vertical changes in grain size and diatom assemblages consistent with abrupt shoaling at the times of the currents. The changes may further suggest gradual deepening of the bay during the intervals between the strong currents. We infer, based on 137 Cs, 14 C, and 210 Pb dating, that the top two shelly gravel layers represent tsunamis associated with the 1703 and 1923 great Kanto earthquakes, and that the third layer was deposited by a tsunami during an earlier earthquake. The age range of this layer, AD 1060 1400, includes the time of an earthquake that occurred in 1293 according to a historical document. If so, the recurrence interval before the 1703 earthquake was almost twice as long as the interval between the 1703 and 1923 earthquakes. Citation: Shimazaki, K., H. Y. Kim, T. Chiba, and K. Satake (2011), Geological evidence of recurrent great Kanto earthquakes at the Miura Peninsula, Japan, J. Geophys. Res., 116,, doi:10.1029/2011jb008639. 1. Introduction [2] Two great Kanto earthquakes, one in 1923 and the other in 1703, are among the most devastating in Japanese history and affected the Tokyo (known as Edo until 1868) metropolitan area (Figure 1). These were interplate earthquakes associated with the subduction of the Philippine Sea plate beneath the Japanese islands. Other types of earthquakes have also caused damage to the Tokyo area, and the rates of recurrence of such large inland earthquakes is known to have increase in the decades prior to the great Kanto earthquakes [Okada, 2001; Central Disaster Management Council, Shuto chokkazisin taisaku ni tsuite (translated as Measures against Tokyo metropolitan earthquakes ) [in Japanese], 2004, available at http://www.bousai.go.jp/jishin/chubou/taisaku_ syuto/pdf/gaiyou/gaiyou.pdf]. There are also late Holocene faults such as the Kozu Matsuda Fault and the Miura Fault Group above the source area of the great Kanto earthquakes (Figure 1) [Research Group for Active Faults of Japan, 1991]. To assess the seismic risk to the Tokyo area, it is critical to determine when the next great Kanto earthquake is likely to occur and when to expect an increase in the rate of large earthquakes. 1 Earthquake Research Institute, University of Tokyo, Tokyo, Japan. 2 Graduate School of Frontier Science, University of Tokyo, Chiba, Japan. Copyright 2011 by the American Geophysical Union. 0148 0227/11/2011JB008639 [3] The most recent great Kanto earthquake (M7.9), which occurred on September 1, 1923, caused more than 105,000 casualties, mostly as a result of shaking induced fires in the Tokyo area [Moroi and Takemura, 2004]. Leveling surveys [Land Survey Department, 1926; Miyabe, 1931] showed that uplift was largest at the southwestern tips of the Miura and Boso Peninsulas ( 1.5 m), and gradually decreased toward the northeast (Figure 2a), indicating that the source of the earthquake was a shallowly dipping thrust fault along the Sagami Trough (Figure 1) [Ando, 1971; Kanamori, 1971]. The earthquake also produced a large tsunami with a height of 5 m or more along the coast at Sagami bay, which led to more than 300 casualties [Hatori et al., 1973] (Figure 2b). [4] The earlier great Kanto earthquake (M 8.1), which occurred on December 31, 1703, caused more than 10,000 casualties [Usami, 2003]. This earthquake was also accompanied by the vertical uplift of the coast. The maximum uplift at the Miura peninsula was 1.5 m, similar to that during the 1923 earthquake, while the maximum uplift at the Boso Peninsula was much larger and estimated to be 4.0 to 6.0 m [Matsuda et al., 1978; Shishikura, 2003] (Figure 2c). The tsunami heights in 1703 along the coast at Sagami Bay were also similar to those following the 1923 earthquake, but were larger along the Pacific coast of the Boso Peninsula (Figures 2b and 2d) [Hatori et al., 1973; Ono and Tsuji,2008]. These facts suggest that an additional fault off the Boso Peninsula was involved in the 1703 earthquake [Matsuda et al., 1978]. 1of16

Figure 1. Map of Kanto area, showing the plate configuration and the location of Koajiro Bay. The stars show the instrumental epicenter of the 1923 Kanto earthquake (M7.9) [Kanamori and Miyamura, 1970] and the center of the source area of the 1703 Kanto earthquake (M 8.0) estimated from damage [Usami, 2003]. The rectangle indicates the 1923 fault area of Ando [1971]. PHS and PAC indicate the Philippine Sea Plate and the Pacific Plate, respectively. KMF and MFG are two of the major active faults, the Kozu Matsuda Fault and the Miura Fault Group, respectively [Research Group for Active Faults of Japan, 1991]. [5] The average recurrence interval of great Kanto earthquakes is estimated to be 200 400 years on the basis of seismological, geodetic, geological and geomorphological data (Earthquake Research Committee, Long term evaluation of the seismicity along the Sagami Trough [in Japanese], Headquarters of Earthquake Research Promotion, Ministry of Education, Culture, Sports, Science and Technology, Japan, 2004, http://www.jishin.go.jp/main/chousa/04aug_sagami/ index.htm). In contrast, the recurrence interval inferred from the uplifted marine terraces at Boso Peninsula is >400 years [e.g., Matsuda et al., 1978; Nakata et al., 1980; Kumaki, 1985, 1999; Shishikura et al., 2001; Shishikura, 2003], while that inferred from tsunami deposits in the same region is 150 300 years for the period between 8200 and 6900 years BP [Fujiwara et al., 1999, 2000; Komatsubara and Fujiwara, 2007]. Based on leveling and triangulation data obtained between 1972 and 1990, and GPS data obtained between 1996 and 2000, previous studies [Yoshioka et al., 1994; Sagiya, 2004] have estimated the recurrence interval as 245 years and 200 to 300 years, respectively. [6] The occurrence time of the most recent pre 1703 great Kanto earthquake is not well known due to the lack of unequivocal written records. An abundance of documents exits for Edo only after 1603, when it became the capital of the Tokugawa shogunate. However, an earlier shogunate was headquartered in nearby Kamakura, on the Miura Peninsula, and written records from this capital city are available for much of the period between 1180 and 1455. Two earthquakes described in the Kamakura records, one in 1293 and the other in 1433, have been suggested as potential predecessors Figure 2. Vertical crustal deformation and tsunami height for the 1923 earthquake (M7.9) and the 1703 earthquake (M 8.0). The data for 1923 is based on leveling data [Miyabe, 1931]. The vertical coseismic movements for the 1703 earthquake are based on geomorphic shoreline studies [Matsuda et al., 1978]. Tsunami heights are taken from Hatori et al. [1973] except for that labeled T, which is taken from Ono and Tsuji [2008]. 2of16

Figure 3. (a) Koajiro Bay and (b) geoslice sites. Location of Koajiro is shown in Figures 1 and 2. Geological columns at each site are shown in Figure 5. The location of the tide gauge station (Aburatsubo) is shown in Figure 3a. The contour interval is 10 m in topographic map (Figure 3a). to the 1703 great Kanto earthquake [Ishibashi, 1991, 1994]. Since the effects of these earthquakes on the Kanto plain are not well described in the written records, it has been difficult to rule out fault sources other than the plate boundary. Shishikura et al. [2001] inferred that a pre 1703 great Kanto earthquake occurred around 1050 (AD 899 1277) based on the 14 C age from a humic soil in a marsh just landward of a beach ridge, on the coseismically uplifted coastline of the western Boso Peninsula. [7] In the present study, we sought to supplement historical evidence with geological evidence for the most recent pre 1703 great Kanto earthquake at Koajiro Bay, on the southern Miura Peninsula (Figures 1 and 3). We found three sandy gravel layers with abundant shell fragments, each bounded by subtidal deposits (Figures 4 and 5). We infer the shelly gravel layers to be tsunami deposits produced by the past three great Kanto earthquakes based on stratigraphy, chronology, and associated land level changes. Based on radiocarbon dating, the date of the pre 1703 great Kanto earthquake is estimated to be between AD 1060 and 1400. By combining our stratigraphic evidence with the information from historical documents, we suggest that the pre 1703 great earthquake occurred in 1293. If so, our findings indicate that the recurrence interval of the great Kanto earthquakes varies from about 200 to 400 years. 2. Study Area [8] Koajiro Bay, at the southwestern tip of the Miura Peninsula, faces Sagami Bay (Figures 1 and 2). During the 1923 earthquake, the sea first receded and then a tsunami with a height of 1.2 to 1.8 m inundated coastal areas and caused five deaths and two injuries in Koajiro [Kanagawa Fisheries Experimental Station, 1924; Tanakadate, 1926]. [9] Koajiro Bay is 2,000 m long and 500 m wide becoming narrower toward inland and is flanked by a rocky coast (Figure 3a). The bay head contains a modern tidal flat, 300 m long and 100 m wide and the beach is gravelly with shell fragments. [10] A stream flows into the tidal flat from the adjacent hills (Figure 3b). The main rock exposed along the coast and in the hilly area is Neogene marine siltstone and mudstone, comprising tuffaceous silt and tephra [Kodama and Oka, 1980]. In the hilly area, the Neogene rock is incised by the stream and exposed on the streambed. Sand and gravel are partially distributed on the streambed (Figure 3b). In the bay area, the tidal flat is mostly silty sand with some gravel. The tidal range is about 2 m between the highest and the lowest tides based on measurements at a nearby tidal station, Aburatsubo, 1 km to the southwest (Figure 3b). [11] Koajiro Bay was uplifted during the earthquakes of 1703 and 1923, and slowly subsided during the previous and current interseismic periods. Tanakadate [1926] estimated that the coseismic uplift was about 1.2 m during the 1923 earthquake, and reported an eyewitness account that the bay had subsided about 0.3 m during the 30 year period prior to 1923. At Aburatsubo, the recorded coseismic uplift was 1.4 m during the 1923 great Kanto earthquake, and subsidence in the almost 90 year period since 1923 is about 0.4 m (Figure 4). The amount of coseismic uplift is reflected in the topography around the tidal flat in two ways: uplifted wave cut benches and bedrock notches (Figure 3b). Two steps shown by both notches and benches are observed at 0.8 to 1.2 m and at 1.5 to 2.1 m above mean sea level. The amounts of uplift inferred from the elevations of these emergent coasts are similar to the recorded uplift at the Figure 4. Annual mean sea level at Aburatsubo tide gauge station. Station location in Figure 3a. The interseismic subsidence rate for the period of 1954 2009 is estimated by Coastal Movements Data Center (see http://cais.gsi.go.jp/ cmdc/centerindex.html) based on the method of Kato and Tsumura [1979]. 3of16

Figure 5. Geological columns at geoslice sites with calibrated 14 C calendar date and vertical variations of excess 210 Pb and 137 Cs. Site locations are shown in Figure 3b. Sandy gravel layers T1, T2, and T3 are correlated on the basis of dating results. The broken lines show the inferred correlation from stratigraphy. All the dating samples for 14 C are terrestrial plants. Sample numbers and the dating results with a confidence range of 2s are shown. Modern limits (1900s) are omitted for simplicity. The details are presented in Table 1. The bars attached to the 210 Pb and 137 Cs data indicate a confidence limit of 1s. The solid curve for 210 Pb is calculated theoretically, as described in the text. 4of16

Aburatsubo tide gauge station in 1923, indicating that they were formed during the 1923 and 1703 earthquakes, respectively. 3. Methods [12] We obtained tabular cores from deposits beneath the tidal flat, to a depth of 2.5 m, using the geoslicer sampler [Nakata and Shimazaki, 1997; Takada et al., 2002] that was 3.0 m long, 0.1 m wide and 0.03 m thick. Sampling cores were collected at eight sites: sites A to E along the longitudinal axis of the tidal flat and sites B1 to B3 across the mouth of the flat (Figure 3b). At each site, coring produced minimal compaction of surface sediment and we did not observe other signs of significant disturbance of the sediment slices; the slicer obtained almost complete cores to the desired depth. 3.1. Criteria for Identifying Tsunami Deposits [13] We sought indicators of tsunami deposits that we could correlate with historical evidence of past great Kanto earthquakes. The structure of sediment deposited during short lived events such as storms and tsunamis may provide insights into the flow characteristics at the time of the deposition [e.g., Lowe, 1982; Hiscott, 1994; Einsele et al., 1996; Hand, 1997; Sohn, 1997; O. Fujiwara et al., 2003; Fujiwara and Kamataki, 2007; Bourgeois, 2009]. We characterized the sedimentary structure of our sediment slices with depositional fabric, grain size, grading, sorting, thickness, layering, mud clast, sedimentary contact, and fossils. [14] The discrimination between tsunami deposits and others, such as storm deposits, is usually problematic [e.g., Witter et al., 2001; O. Fujiwara et al., 2003; Fujiwara and Kamataki, 2007; Goff et al., 2004; Morton et al., 2007; Switzer and Jones, 2008]. One characteristic that may help distinguish storm from tsunami deposits is that if the sedimentary environment changes due to sudden uplift, the stratigraphy may show an abrupt change. If such an abrupt stratigraphic contact coincides with tsunami deposits, the deposits are most likely to be tsunami deposits associated with a great subduction zone earthquake [e.g., Atwater, 1987; Nelson et al., 1996b, 2008]. 3.2. Grain Size Analysis [15] At the head of the bay, the main source of bay sediment is streamflow (Figure 3). Because sediment grain size is sensitive to changes in flow conditions, grain size is an indicator of inflow from rivers, tides and the waves in the bay [e.g., Wright, 1977]. The strength and velocity of flow is usually reduced by the depth and distance from the river. [16] In the present study, measurements were carried out at 2.5 5 cm intervals on sediment slices from sites A, B, C, D and E (Figure 3). Four grain size categories were considered: silt, fine sand, medium sand and coarse sand. At site B2, measurements were taken at 2 cm intervals, and the grains were categorized into 10 different sizes at phi scale intervals of 0.5. 3.3. Diatom Analysis [17] Diatom assemblages in and around the Miura Peninsula are good indicators of modern and ancient wet environments, and provide proxy information on salinity, temperature, current velocity, depth, sea grass and water plants, sediment substrate and geomorphic setting [e.g., Kosugi, 1987, 1988; Yanagisawa, 1996]. For example, diatom assemblages at Koajiro exhibit an increase in the ratio of benthic to planktonic diatoms with the increasing water depth. Changes in diatom assemblages in tidal sediment sequences at other subduction zones have been used to reconstruct the relative sea level changes associated with crustal movements during the earthquake cycle [e.g., Nelson et al., 1996a, 1996b, 2008; Shennan et al., 1996, 2006; Atwater and Hemphill Haley, 1997; Sawai, 2001; Sawai et al., 2004]. [18] We sampled the diatoms at 2 cm intervals in the geoslice samples from site B2 (Figure 8). At least 300 diatom valves were analyzed under an oil immersion microscope for each sample. Details of species with references are given in Appendix A with the results of analysis. 3.4. Dating [19] To help correlate the strata and estimate the ages of deposits, we obtained AMS 14 C ages on samples. Since it is difficult to obtain accurate dates after about AD 1650 with 14 C dating, 210 Pb and 137 Cs were measured to help estimate ages for modern samples. [20] For AMS 14 C dating, we selected terrestrial plant fragments, mainly comprising wood and charcoal, from the bay deposits and dated 42 samples (Figure 5, Table 1). The age of dated materials were calculated in 2s calendar year range (cal. AD or cal. BC) from a dendrochronologically calibrated database [Reimer et al., 2004] using the program OxCal v4.1 [Bronk Ramsey, 2009] (Figure 5, Table 1). In the text of this paper, AD or BC is added to the year for 14 C ages, but not for dates determined from historical documents. Radiation levels from 137 Cs and 210 Pb were also measured for dating purposes. Eight samples were taken at intervals of 0.1 m beneath the sea bottom surface at site B3 (Figure 5). The half lives of 137 Cs and 210 Pb are 30.3 and 22.2 years, respectively. The 137 Cs was produced by nuclear experiments after the 1950s and a peak is observed in the 1960s. Dating using 210 Pb is usually carried out based on the half lives for fine grained deposits produced during the past 100 years. 4. Stratigraphy and Age [21] The bay deposits in the geoslice samples are classified into two types: the upper deposits composed of silty sand, and the lower deposits composed of sandy silt (Figures 5 and 6). As will be described later, the upper deposits are tidal flat deposits and the lower deposits will be called as subtidal deposits. We also identified two to four sandy gravel layers with abundant shell fragments in the subtidal deposits (Figure 6). Organic matter comprised 15 to 22% of the shelly gravel layers. Shelly gravel layers are usually deposited during large, sudden events such as storms or tsunamis. Stratigraphy and ages helped us correlate upper three shelly gravel layers among the geoslice sites. We refer to the layers as the T1, T2 and T3 layers in order of increasing depth (Figure 5). 4.1. Stratigraphy [22] The tidal flat deposits have thicknesses of 0.3 0.5 m at all the sites studied, except for site B3 (Figure 5). Site B3 is located in shallow water (depth of 0.4 0.5 m at low tide) and 5of16

Table 1. Results of Radiocarbon Analysis in Koajiro Bay a Site Depth (m) Material Radiocarbon Age (ybp) Calibrated Age Calendar Year Range (2s) Sample Samples Used, Cases a or b Lab Number IAAA T1/T2 A 0.91 charcoal 155 ± 32 1660 1960 1 60017 A 1.05 wood 75 ± 33 1680 1930 2 52759 B 0.80 wood 154 ± 33 1660 1960 3 52761 C 0.95 wood 203 ± 33 1640 1960 11 52765 B1 0.74 charcoal 110 ± 28 1680 1940 23 61981 B1 0.78 wood 53 ± 26 1690 1960 24 61982 T2/T3 B 1.15 wood 132 ± 31 1670 1950 4 a, b 52762 B 1.28 wood 129 ± 32 1670 1950 5 a, b 52763 B 1.44 charcoal 564 ± 33 1300 1430 6 a, b 60019 B 1.59 charcoal 41 ± 35 1690 1960 7 a, b 60020 B 1.61 charcoal 999 ± 32 980 1160 8 a 60021 B 1.64 charcoal 66 ± 35 1680 1940 9 a, b 60022 B 1.67 wood 917 ± 37 1020 1210 10 a 52764 C 1.22 wood 241 ± 31 1520 1960 12 a, b 52766 D 0.74 chacoal 546 ± 37 1300 1440 17 a, b 52771 B1 0.88 chacoal 149 ± 27 1660 1960 25 a, b 61983 B1 0.92 wood 344 ± 27 1460 1640 26 a, b 61984 B1 0.99 wood 153 ± 28 1660 1960 27 a, b 61994 B1 1.02 bark 66 ± 28 1690 1930 28 a, b 61985 B1 1.04 wood 108 ± 26 1680 1940 29 61986 B1 1.16 wood 135 ± 31 1660 1950 30 a, b 61987 B1 1.20 chacoal 182 ± 27 1650 1960 31 a, b 61988 B1 1.26 chacoal 129 ± 26 1670 1950 32 a, b 61989 B1 1.44 chacoal 870 ± 27 1040 1260 33 a 61990 B1 1.46 wood 342 ± 37 1460 1650 34 a, b 61991 B1 1.54 chacoal 926 ± 28 1020 1170 35 a 61992 B1 1.58 chacoal 905 ± 29 1030 1210 36 a 61993 B2 1.12 leaf 140 ± 27 1670 1940 37 a, b 80189 B2 1.16 wood 307 ± 28 1490 1650 38 a, b 80190 B2 1.22 wood 116 ± 28 1680 1940 39 a, b 80191 B2 1.24 wood 144 ± 27 1670 1950 40 a, b 80192 <T3 C 1.81 wood 3111 ± 36 1490BC 1260BC 13 a, b 52767 C 1.83 wood 1709 ± 32 250 410 14 a, b 52768 C 1.85 wood 5260 ± 38 4230BC 3970BC 15 a, b 52769 C 1.89 wood 4168 ± 41 2890BC 2620BC 16 a, b 52770 D 1.47 chacoal 1948 ± 37 50BC 130 18 a, b 52772 D 1.70 wood 1665 ± 36 250 530 19 a, b 52773 B2 2.02 wood 920 ± 30 1030 1190 41 a, b 80193 B2 2.07 wood 845 ± 28 1060 1260 42 a, b 80194 <Older Event D 2.23 chacoal 3178 ± 37 1530BC 1380BC 20 52774 Subtidal Deposits E 0.94 chacoal 847 ± 33 1050 1270 21 60028 E 1.08 chacoal 2479 ± 40 770BC 410BC 22 60029 a Sampling location and depth are shown in Figure 5. Lab number: Institute of Acceleratory Analysis Ltd., Shirakawa, Fukushima, Japan. Radiocarbon ageis measured by the accelerator mass spectrometry method. It is corrected by d13c and calculated by using the Libby half life of 5568 years. Calendar year range is determined from dendrochronlogically calibrated probable age ranges using the program IntCal04 [Reimer et al., 2004]. Calibration was carried out with confidence limits of 2s and rounded to the nearest decade. For estimating probable year range for T3, two cases, a and b are considered. In case a, all the samples are used. In case a, samples above T3, but showing ages similar to or older than sample #42, are removed because they might be reworked samples. the top unit is composed of fine sand and silt. At other sites, the tidal flat deposits are composed of silty sand with mostly medium to coarse grains, but also include some fine sand and silt (Figure 6). The uniform grain size distribution in these units results in massive sedimentary structure. [23] The subtidal deposits are of similar lithology but have variable sedimentary structures among the sites. Subtidal deposits are finer grained than tidal flat deposits and have weak parallel laminae at several levels. The contact between the tidal flat and subtidal deposits is sharp and distinct (Figure 5). [24] Except for sites A and E, the three shelly gravel layers, T1, T2 and T3, are observed at three depth horizons of 0.5 to 0.7 m, i.e., 0.9 to 1.3 below the mean sea level (MSL), 0.8 to 1.2 m (1.4 to 1.8 m below MSL) and 1.1 to 1.9 m (1.7 to 2.4 m below MSL), respectively (Figures 5, 6, and 8). At site D, an additional shelly gravel layer is observed at depths of 1.7 to 2.1 m (2.1 to 2.5 m below MSL). At site B1, an 6of16

Figure 6. Grain size distribution at sites A to E. Site locations are shown in Figure 3b. Sandy gravel deposits, T1, T2, and T3 layers are considered to be event deposits. additional lens like sandy gravel sheet with a thickness of less than 0.04 m is observed at a depth of 1.1 m. The subtidal deposits are divided into subunits by these shelly gravel layers. [25] In each subunit of the subtidal deposits, 10 to 60% of the deposits are medium to coarse sands. The subtidal deposits are thought to have originated mainly from streamflow from the hilly area (Figures 3b, 6, and 8). The sedimentary structure exhibits an increase in the fraction of coarse grains with increasing depth in each subunit, except the uppermost subunit at site B2 where such a tendency cannot be observed, probably due to the effects of sea waves (Figures 6 and 8). The grain size in the subtidal deposits is larger above than below each shelly gravel layer (Figures 6 and 8), except for T2 at sites A and B, and T3 at site C. [26] At site A, located at the mouth of the tidal flat, fine materials are dominant with a partial ratio of 70 to 90% for silt and fine sand (Figures 3 and 6). Above the T2 layer, the amount of coarse sand and medium sand is extremely small in the depth range of 0.5 to 1.0 m ( 1.1 to 1.6 m below MSL). The subtidal deposits at this site are therefore clearly matrix supported and well sorted (Figure 6), and exhibit gradual reverse grading, unlike at other sites. Beneath the T2 layer, at depths of 1.2 to 1.8 m, (1.6 to 1.8 m below MSL), coarse sand comprises up to 30% of subtidal deposits and is more abundant than above T2. The grain size above T2 is generally finer than that below T2. The sedimentary structure shows gradual normal grading. [27] At site E near the outflow mouth of a stream, no shelly gravel layer nor fluvial gravel are observed in subtidal deposits composed of silty sand (Figures 5 and 6). The sed- imentary structure is relatively poorly sorted and grading is indistinct compared with that at site A (Figure 6). 4.2. Identification of Traction Current Deposits [28] The three shelly gravel layers, T1, T2 and T3, in the subtidal deposits probably were deposited in the sudden and strong current (Figure 5). The lower contact with the subtidal deposits is mainly erosive at most of the sites. The lower contact of T1 at site B1 has a sharp contact with the subtidal deposits, although no erosion surface is found (Figure 7a). The upper contact is also distinguishable from the subtidal and tidal flat deposits with relatively sharp contacts being present (see Figures 7b, 7c, and 7d). The thickness of the shelly gravel layers varies from place to place, ranging from 0.1 to 0.5 m (Figure 5). [29] The shelly gravel layers are composed of coarse sand, granules and pebbles with abundant shell fragments and calcareous sessile organisms (Figure 7). Most bivalve shells have broken into small pieces. The gravels originate from Pliocene sandy mud that is found in the surrounding coastal rocks, and mud clasts are also present (Figure 7b). In addition, a mud clast in the T2 layer at site B is found to be adhered to a fossilized specimen of Pomatoleios kraussii, which had been living in the tidal zone of the rocky coast in and around the Miura Peninsula [Kayanne and Yoshikawa, 1986]. [30] The depositional fabric shows that the various deposits such as pebbles, granules, sand and shell fragments are mutually mixed within each shelly gravel layer (Figure 7). Shell fragments and gravels have apparently been rolled over before being mixed (e.g., Figures 7a and 7c). Spiral shells are found mixed with sandy gravel in the T1 layer, but original living position (Figure 7a). In addition, mud clasts, in which 7 of 16

layer, and lap directly onto the erosional surface (Figures 7b and 7c). The T1 layer at site B1 and the T3 layer at site B show that sand and gravel lap directly onto the bottom subtidal deposits with a sharp or erosional contact (Figures 7a and 7d). 4.3. Correlation and Age [32] Results of 137 Cs, 210 Pb, and 14 C dating support our correlation of the three gravelly deposits, T1, T2 and T3. The beds were deposited in the early to middle AD 1900s for T1, after AD 1650 and before the AD 1900s for T2, and after AD 1060 and before the AD 1900s for T3, if we assume that only the youngest 14 C age is the most accurate age in each subunit of the subtidal deposits (Figure 5). [33] The results of the 137 Cs and 210 Pb analyses at site B3 indicate that T1 was deposited in the mid 1900s. 137 Cs is detected only in the subtidal bay deposits directly above T1 and never beneath, therefore T1 must have been deposited before 1950s. The amount of excess 210 Pb is found to decrease with increasing depth, from 2.1 dpm/g near the sea bottom to 0.9 dpm/g directly above the T1 layer. A 210 Pb profile was calculated based on a depositional age of 1924 just above T1, and then compared with the measured 210 Pb profile (Figure 5). There is reasonably good agreement between the measured and calculated values of 210 Pb profiles. The 137 Cs and 210 Pb thus suggest an age of about 1923 for the T1 layer. [34] The deposition age of the T2 layer is estimated as later than AD 1650 from the calibrated 14 C results (Figure 5 and Table 1). From the region between the T2 and T3 layers, a total of 25 plant fragments are dated. Their ages range from AD 980 to later than AD 1650. Among them, 15 samples are dated as later than AD 1650 (Table 1). [35] Below the T3 layer, eight 14 C dates are obtained at sites B2, C, and D (Figure 5). The deposition age of the T3 layer is constrained to after AD 1060 1260 (Table 1), based on the results for sample #42 at B2. Another sample located slightly above #41 shows a similar age of AD 1030 1190. The other six samples show much older ages and are almost certainly reworked from older deposits. [36] At site D, we consider the third layer to be T3 and the lowest layer to be an older tractive current event layer, for which the deposition age is not well constrained. The two ages above the oldest layer in slice D are probably reworked and so do not help in estimating its age. Figure 7. Sedimentary structures of T1, T2, and T3 layers shown by close up photographs and sketches. Unconformities are drawn with thicker lines than other contacts. many granules with diameters of about 2 3 cm are mixed in mud, are observed in the upper part of the T2 layer at site A (Figure 7b) and within the T3 layer at sites B2 and D. [31] Inverse normal grading is often observed in the shelly gravel layers. The T2 layer at all sites exhibits inverse normal grading, i.e., relatively coarse materials such as pebbles are deposited near the center of the layer, and fine materials such as sand and granules are present in the lower part of the 4.4. Diatom Analysis [37] The assemblage and total number of fossil diatom valves are analyzed at site B2 (Figure 8). The concentration of diatoms is very low in the T1, T2 and T3 layers. Most valves in these layers are fragmented. In contrast, the tidal flat and subtidal deposits above T3 have a large number of intact valves. [38] Freshwater diatoms such as Epithemia adnata (Kützing) Brébissonand, and brackish water diatoms such as Navicula cryptotenella Lange Bertalot, make up 10% 20% of the diatom assemblages in the tidal deposits. However, marine species are dominant, especially Cocconeis scutellum Ehrenberg, Tryblionella lanceola Grunow ex Cleve. Very few fresh water species are found in the subtidal deposits while only a few brackish water species are observed between T1 and T2, and T2 and T3 (Figure 8). 8of16

Figure 8. Grain size distribution and diatom assemblages of geoslice B2. For grain size; S, F, M, and C are silt, fine sand, medium sand and coarse sand, respectively. Coevent and interevent changes in sedimentary environment can be inferred from the vertical changes in the grain size distribution and diatom assemblages. [39] In the subtidal deposits, a large number of marine diatoms are identified. Marine benthic diatoms, such as Amphora immarginata Nagumo clearly decrease in the upward direction in each subunit between T1 and T2 and between T2 and T3 (see Appendix A). Marine planktonic diatoms, such as Cheatoceros resting spore clearly increase in the upward direction in each subunit between T1 and T2 and between T2 and T3 (see Appendix A). 5. Identification of Great Kanto Earthquakes [40] The shelly gravel layers, T1, T2 and T3, indicate the strong currents, whereas the sedimentary characteristics of tidal flat deposits and subtidal deposits indicate quiet flows. Also, the grain size and diatom distributions suggest abrupt changes before and after the deposition of T1, T2 and T3 layers. We infer that such abrupt changes of sedimentary environment were caused by coseismic crustal movements. In addition, interseismic subsidence is also inferred from the diatom and grain size distributions. On the basis of these results, it is suggested that the shelly gravel layers, T1, T2 and T3, are produced by the tsunamis that followed the past three great Kanto earthquakes. 5.1. Possible Causes of Traction Current [41] The sedimentary characteristics of the T1, T2 and T3 layers indicate that a strong, rapid traction current flowed on the bay bed. In Figures 7b, 7c, and 7d, the bottom surface is seen to be erosive. Erosional surfaces are usually caused by bottom shearing, which occurs in a strong flow. Some of the shelly gravel layers have inverse normal grading (Figure 7b and 7c). Such a grading structure is probably the result of a velocity gradient in the flow, indicating that the traction carpet flowed [Lowe, 1982; Hiscott, 1994; Sohn, 1997]. No 9of16

grading structure is found in the T3 layer at site B where pebbles are deposited directly on the erosive bottom bed, probably due to their heavy weight (Figure 7d). In the T1, T2 and T3 layers, there is a mixture of gravel and shell fragments, showing a type of traction current associated with turbulent flow. [42] Mud clasts are found at the upper part of the T2 layer at site A (Figure 7). These are produced by erosion of the surface bed in flow events such as tsunamis and redeposition during or after the flow event [e.g., Morton et al., 2007]. Since the mud clasts are seen around the boundary of the inversenormal grading, in the upper part of the shelly gravel layer, they may have been produced during a period of backward flow and deposited during a stagnant flow period. [43] The shelly gravel layers have not been directly transported by normal streamflow but rather by large scale events such as tsunamis and storms. No continuous gravel layer is identified in the bay deposits at site E (Figures 5 and 6), and the gravel is not always present at the modern tidal flat (Figure 3b). The only possible sources for the shelly gravel layers are the sea or the land. In the former case, they may have been raked and transported from the beach to the head of the bay by strong tsunami waves or storms smashing the surrounding rock, because breccia are included in the shelly gravel deposits (Figure 3b). In the latter case, shelly gravels that lay on the streambed may have been washed to the subtidal bay bed by the seaward currents (Figure 3b). It is difficult to determine which is the case simply from the structure of the geoslice samples. [44] The shelly gravel layers are not likely to have been the result of storms or abnormal climatic changes for three reasons. First, a recent large typhoon did not leave any such deposits. The 2009 MELOR typhoon (atmospheric pressure: 910 hpa, wind velocity: 55 m/s) was the largest and most violent during the past 50 years. The Aburatsubo tide gauge station (Figure 3a) recorded a high tide of 3.36 m. The waves attacked man made structures in and around Koajiro Bay, and a wooden warehouse was destroyed near the excavation site. No shelly gravel layer, however, was found after the storm. Second, the frequency of severe storms clearly exceeds the number of shelly gravel layers identified in this study. Historical records of abnormal weather events such as storms, thunderstorm, and droughts in Kanto, are to be found in various documents produced mainly after 1664 [e.g., Yoshimura, 1993; Yoshino, 2007]. Third, it is likely that the depositional duration is too short, and the amount of deposition too large, for the shelly gravel layer to be piled up by short period waves such as those associated with typhoons and storms [O. Fujiwara et al., 2003]. 5.2. Identification of Abrupt Environmental Change [45] As described above, the results of the present study indicate that the shelly gravel layers, T1, T2 and T3, are by the strong traction current. Associated with these layers accumulated, the rapid depositional change of environment is recorded in the sediments of the inner bay. We call three rapid changes of depositional environment corresponding to T1, T2 and T3 layers as events 1, 2, and 3, respectively. 5.2.1. Event 1 [46] The sedimentary environment was abruptly shifted from a subtidal to a tidal flat environment after the deposition of T1 layer at sites B2 and D. While at sites B and C, the T1 layer is covered by the about 10 cm thickness of the thin subtidal deposits before the deposition of the tidal flat deposits. Although such local influence is observed in change of the depositional environment, the grain size becomes coarser above the T1 layer at all sites (Figures 6 and 8), probably as a result of stronger currents carrying coarser sediment from the stream at the head of the bay. [47] The greater influence of the stream on the environment in the bay is also reflected in the diatom assemblages at site B2 (Figure 8). Within the tidal flat deposits above the T1 layer, the ratios of the freshwater and brackish species abruptly increase by 10 20% compared to those in the underlying subtidal deposits. Very few freshwater and brackish species are present within the subtidal deposits. In contrast, the ratios of marine planktonic species abruptly decrease from 50% in the subtidal deposits beneath the T1 layer to 5% in the tidal flat deposits above T1. 5.2.2. Event 2 [48] At most sites, an abrupt increase in the amount of coarser material is observed after the deposition of the T2 layer. At sites B2, C and D, the ratios of the coarser materials (coarse and medium sands) to the finer ones (fine sand and silt) increase abruptly above the T2 layer by 10% compared to those below (Figures 6 and 8). We infer that flow from the stream became stronger and more frequent after the deposition of the T2 layer, and consequently the amount of coarser sands increased. At site A, however, the vertical change in grain size shows the opposite pattern. The amount of finer materials dramatically increases above the T2 layer, and the sorting is good (Figure 6). Judging from the position of this site, near the mouth of the tidal flat, the subtidal deposits may have become finer due to wave sorting after event 2 (Figure 3b). [49] The ratio of marine benthic diatoms to planktonic diatoms abruptly increases above the T2 layer by 10% (Figure 8). Based on the modern analogs of the diatom species in Koajiro, such an increase in the ratio of benthic to planktonic species indicates a decrease in water depth. 5.2.3. Event 3 [50] Changes in grain size are recognized after event 3 as follows. At sites B and D, the ratio of coarse sand to finer sands abruptly increases by 30% above the T3 layer. At site C, an abrupt decrease of coarse sand is observed directly above the T3 layer. At site B2, on the other hand, no change in the mean grain size is observed after the event and an increase in the amount of silt is associated with the deposition of T3 layer. Very few diatoms are found below T3, which may reflect a low abundance of diatoms and/or a poor preservation environment due to some unknown local conditions (Figure 8). 5.3. Cause of Abrupt Environmental Change [51] We infer that the abrupt changes in the sedimentary environment, shown by the changes in grain size and diatom concentrations at the head of Koajiro Bay, were not caused by storms or a regional (eustatic) sea level change, but instead by crustal movements during the earthquake cycle on the subduction zone fault. If the environmental changes were caused by storms, the sedimentary environment would recover to or remain in a state similar to that before the storms. If the environmental changes were caused by a eustatic sea level change, the vertical changes in the grain size and the diatom 10 of 16

Figure 9. Occurrence time of the last two great Kanto earthquakes and probable ranges for the pre 1703 earthquake. The small closed circles show the occurrence time of candidate great Kanto earthquakes based on historical documents. The broken line indicates a time range of incomplete data. Events T1 and T2 are correlated with the 1923 and 1703 earthquakes, respectively, based on the 137 Cs and 210 Pb analysis (Figure 5). Probable time ranges for event T3, the most recent pre 1703 great Kanto earthquake, were estimated from 14 C analysis in two ways, i.e., cases a and b, using the software package OxCal v4.10 [Bronk Ramsey, 2009] and are shown by a probability density distribution. Case a uses all the samples while case b excludes five possible reworked samples. Confidence intervals of 2s are also shown by numerals. In the lower part, open and closed rectangles represent 2s range of ages for samples above and below the T3 layer, respectively. Attached numerals are sample numbers (Table 1). Dating materials are wood (indicated by w) and charcoal (unmarked). assemblages would be gradual in response to the gradual sea level changes. [52] As stated above, the 1923 Kanto earthquake caused an uplift of the Miura Peninsula by 1.4 m (Figures 3 and 4). After event 1, the stratigraphy changed abruptly from subtidal to tidal flat environment, accompanied by an abrupt change in the diatom assemblage (Figures 6 and 8). This shows that the bay bottom rose after event 1, forming the tidal flat. Such an abrupt change in lithology indicates sudden uplift. [53] Similar sedimentary environmental changes, for events 2 and 3, also are recorded in the subtidal deposits (Figures 6 and 8). The abrupt increases in the amount of coarser grains associated with these events reflect a change in the distance from the mouth of the stream and in its flow velocity. If the distance between the outflow mouth of a stream and the geoslice site in the bay became shorter due to coseismic uplift associated with the great Kanto earthquake, the influence of the flow from the stream might have become stronger. It has also been observed at other areas that the grain size of deposits in an open channel decreases with increasing distance from the river mouth [e.g., Wright, 1977; Eisma, 1991; Takashimizu et al., 1999]. 5.4. Interseismic Subsidence [54] We also infer interseismic subsidence during the periods between deposition of layers T1, T2, and T3 from the gradual changes in grain size and diatom assemblages. First, the thickness of the tidal flat deposits, 0.3 to 0.5 m, is similar to the subsidence of 0.4 m after the 1923 Kanto earthquake (Figures 4 and 5). These observations may suggest that the deposit thickness represents the amount of vertical crustal movements, although in general the thickness may be the result of many other factors rather than subsidence. Second, the gradual decrease in the ratio of marine benthic to planktonic diatoms indicates an increase in water depth (Figure 8). This may be related to a decrease in the amount of sunlight reaching the sea bottom. Also, an increased amount of flow from the sea might have caused a gradual increase in planktonic species. [55] The grain size of the bay sediments between the event deposits generally becomes finer over time, except above the T2 layer at site A and between the T1 and T2 layers at site B2. This is consistent with a gradually decreasing influence from the stream during periods of interseismic subsidence. 6. Pre 1703 Great Kanto Earthquake [56] We concluded that the T1, T2 and T3 layers are tsunami deposits generated by the past three great Kanto earthquakes. We estimate dates for these earthquakes by comparing the tsunami deposit ages with historical documents (Figure 9, Table 1). 6.1. Dates of Three Great Kanto Earthquakes [57] The deposition ages of the T1 and T2 layers were estimated at the early to mid 1900s, and after AD 1650, respectively, from the 137 Cs, 210 Pb and 14 C analyses (Figure 5). These ages correspond well to the 1923 and 1703 great Kanto earthquakes, which coincided with tsunamis and uplifts of the Miura Peninsula (Figure 2). [58] The timing of the pre 1703 event T3 is estimated to be before 1703 and after AD 1060 1260. In order to narrow 11 of 16

the age range, we estimate two kinds of probable ages, cases a and b (Figure 9). In case a, all the ages obtained above and below the T3 layer are combined using OxCal [Bronk Ramsey, 2009]. The resulting probable deposition age for T3 is AD 1060 1150 (median: AD 1100). In case b, the ages on five samples that are probably reworked (#8: AD 980 1160, #10: AD 1020 1210. #33: AD 1040 1260, #35: AD 1020 1170, and #36: AD 1030 1210) are removed, because these samples taken from above the T3 layer show ages similar to or older than that of a sample taken from below the T3 layer (#42: AD 1060 1260) (Figure 5, Table 1). In case b, the probable formation date of T3 is AD 1160 1400 (median: AD 1290). [59] Combining these two cases, we estimate the date of the most recent pre 1703 earthquake to be AD 1060 1400. The pre 1703 great Kanto earthquake may be correlated to an earthquake around AD 1050 inferred from beach ridges on the western coast of the Boso Peninsula [Shishikura et al., 2001]. 6.2. Historical Earthquakes [60] The date of case a for the pre 1703 great Kanto earthquake is 1060 to 1150, but no damaging earthquake at Kamakura (Figure 1) is listed in Usami s [2003] catalog of historical earthquakes. This is because few documents exist for this time period, especially in Kanto. [61] For case b interval, several historical earthquakes occurred within this time range. Seven damaging earthquakes at Kamakura are listed, in 1213, 1227, 1230, 1240, 1241, 1257, and 1293. However, few historical records exist in this area between 887 and 1180. The last history book of the ancient Japanese government, Nihon Sandai Jitsuroku, ends in 887. The first history book about the Kamakura government, Azuma Kagami, starts its description in 1180. Various documents remain concerning events at Kamakura up to 1455, when a major war broke out there. [62] The other known historical earthquakes were probably not large enough to have produced gravelly beds like those deposited in 1703 and 1923. The most likely candidate for the pre 1703 great Kanto earthquake is the 1293 earthquake. As was pointed out by Ishibashi [1991], the earthquake caused severe damage at Kamakura and was accompanied by intense aftershock activity. A death toll of 23,024 was reported in a historical document. The word tsunami did not appear in any documents, but Ishibashi [1991] suggested tsunami damage on the basis of a description of 140 dead bodies along a sandy beach at Kamakura. The tsunami following the 1923 Kanto earthquake killed 24 people at Kamakura [Tanakadate, 1926]. Severe damage to temples such as Kenchoji is reported for both the 1293 and 1923 earthquakes [e.g., M. Fujiwara et al., 2003]. [63] The earthquake in 1241 was reportedly accompanied by a tsunami [Usami, 2003] but Ishibashi [2009] inferred that the damage was due to severe winds and waves. The magnitude of the earthquake in 1257 was estimated to be M7.0 7.5 by Usami [2003] on the basis of descriptions of the damage. No shrines or temples at Kamakura remained undamaged and all walls and fences were destroyed. Landslides, earthquake sounds, collapse of houses, liquefaction, and ground ruptures were reported. This earthquake may be another candidate, but there were no indications of a tsunami in the report. The remaining earthquakes in 1213, 1227, 1230, and 1240 appear to have been smaller than the 1257 earthquake. [64] Ishibashi [1991, 1994] suggested that the earthquakes in 878, 1433, and 1293 may have been great Kanto earthquakes. These events are shown in Figure 9 together with the earthquake of 1257. 6.3. Recurrence Interval [65] The average recurrence interval of the great Kanto earthquakes based on the dates of the past three events, i.e., 1923, 1703 and 1060 1400 ranges 260 to 430 years. These values almost agree with the estimation of 200 400 years by the Earthquake Research Committee (2004). However, the short interval between the 1703 and 1923 earthquakes (220 years) compared to the estimated preceding one (303 to 643 years), shows that the recurrence interval of the great Kanto earthquakes is variable. [66] The variable recurrence interval of the great Kanto earthquakes may be related to the source area or the fault slip in the individual earthquakes. The 1703 and 1923 earthquakes produced tsunamis with different heights and different amounts of coastal uplift in the Boso peninsula [e.g., Hatori et al., 1973; Matsuda et al., 1978; Kumaki, 1999; Shishikura, 2003] (Figure 2). The date range of the pre 1703 Kanto earthquake (AD 1060 1400) overlaps with the estimated date range, AD 1100 1350, of the most recent earthquake (Earthquake Research Committee, Long term forecast of earthquakes on the Kannawa/Kozu Matsuda fault [in Japanese], Headquarters of Earthquake Research Promotion, Ministry of Education, Culture, Sports, Science and Technology, Japan, 2009, http://www.jishin.go.jp/main/chousa/ 09jun_kannawa/index.htm) on the Kozu Matsuda fault (Figure 1). Because this fault is located on the landward extension of the plate boundary, the pre 1703 earthquake might have been extended on this fault. If the fault slip is also large, it would lead to a longer interval before the next recurrence [Shimazaki and Nakata, 1980]. However, we do not have any information on the fault slip or size of the pre 1703 earthquake. [67] Variability in size and recurrence interval of great earthquakes has been found for many subduction zones in the world [e.g., Satake and Atwater, 2007]. The recent gigantic (M 9) earthquakes demonstrate the variability. The 2004 Sumatra Andaman earthquake occurred in the region where smaller (M 8) have been historically known. The 2011 Tohoku earthquake (M 9.0) was also much larger than the characteristic earthquakes (M < 8) that repeated in the Miyagi oki region. Recurrence of great earthquakes along the Nankai trough [e.g., Satake and Atwater, 2007; Komatsubara and Fujiwara, 2007], along the Cascadia subduction zone [e.g., Nelson et al., 2006], along the Chilean trench [e.g., Cisternas et al., 2005] and along the Aleutian Islands [e.g., Shennan et al., 2009] also show some variability recorded in historical documents and coastal geology. 7. Conclusions [68] We identified three tsunami deposits at Koajiro Bay on the southwestern tip of the Miura Peninsula (Figure 1), the oldest of which was probably deposited following a great Kanto sudduction zone earthquake. We used sedimentary structure, grain size, diatom distribution and 14 C, 137 Cs, and 12 of 16