Shimada et al.
Progress in Earth and Planetary Science(2023) 10:49
https://doi.org/10.1186/s40645-023-00577-9

RESEARCH ARTICLE

Progress in Earth andPlanetary Science

Open AccessMarine inundation history during the last3000 years at Lake Kogare‑ike, a coastal lakeon the Pacific coast of central Japan
Yumi Shimada1,2* , Yuki Sawai1, Dan Matsumoto1, Koichiro Tanigawa1, Kazumi Ito1, Toru Tamura1,2,
Yuichi Namegaya1, Masanobu Shishikura1 and Shigehiro Fujino3

Abstract
Sediment cores collected at Lake Kogare-ike, a coastal lake on the Pacific coast of central Japan, record the marine
inundation history during the last 3000 years. The sediments consist mainly of organic mud, sand, gravel, inorganic
mud, and volcanic ash, and inundation events were recognized as 19 event deposits (E1–E19, from top to bottom)
interbedded with the organic mud. Visual observation by naked eyes and X-ray computed tomography (CT) images
identified 16 event deposits based on quantitative and qualitative changes in sand contents and changes in the textures and colors of the sediment samples (E1–E3, E5, E6, E8, E9, and E11–E19). The other three event deposits (E4, E7,
and E10) were identified only on the CT images as layers with higher radiodensity than the underlying and overlying
organic mud layers. The sedimentary features, the spatial bias of the event deposits toward seaward areas, the diatom
assemblages, and the frequency of inundation events suggest that 13 (E1–E10 and E12–E14) of the 19 event deposits
were formed by tsunamis or extraordinary storms. To constrain the depositional ages of the event deposits, Bayesianage–depth models were constructed based on radiocarbon dating of plant macrofossils and concentrated fossil pollen and the 137Cs profile. The depositional ages of the event deposits indicate that five or possibly six event deposits
can be correlated with historical tsunamis along the Nankai Trough: E2, either of E3 or E4, E5, E7, and E9 correspondto the 1707 CE Hoei, the 1605 CE Keicho, the 1498 CE Meio, the 1096 CE Eicho, and the 684 CE Hakuho tsunamis,
respectively. E1 was possibly formed by the 1944 CE Showa-Tonankai tsunami, the 1854 Ansei–Tokai tsunami,
the 1959 Isewan typhoon, or a combination of two or all three events.
Keywords Bayesian age–depth model, Extreme events, Japan, Lacustrine sediments, Nankai trough, Storm surges,
Tsunami*Correspondence:
Yumi Shimada
yumi.shimada@aist.go.jp
1
Geological Survey of Japan, National Institute of Advanced Industrial
Science and Technology (AIST), Higashi 1‑1‑1‑C7, Tsukuba, Ibaraki
305‑8567, Japan
2
Graduate School of Frontier Sciences, The University of Tokyo,
Kashiwanoha 5‑1‑5, Kashiwa, Chiba 277‑8561, Japan
3
Faculty of Life and Environmental Sciences, University of Tsukuba,
Tennodai 1‑1‑1, Tsukuba, Ibaraki 305‑8572, Japan

1 Introduction
Infrequent but extraordinary extreme waves, such as tsunamis and storms, impact coastal environments. Instrumental records (e.g., from tide gauges and oceanographic
observation buoys) are generally available only for within
the last century. Historical documents are also used tocomprehend the history of past extreme wave events,
but such events are difficult to reconstruct in detail on
millennial or centennial timescales because of the shortness and/or incompleteness of historical records. To
supplement these records and to extend the history of
extreme events, geological records have been employed© The Author(s) 2023. Open Access This article is licensed under a Creative Commons Attribution 4.0 International License, which
permits use, sharing, adaptation, distribution and reproduction in any medium or format, as long as you give appropriate credit to theoriginal author(s) and the source, provide a link to the Creative Commons licence, and indicate if changes were made. The images orother third party material in this article are included in the article’s Creative Commons licence, unless indicated otherwise in a credit line
to the material. If material is not included in the article’s Creative Commons licence and your intended use is not permitted by statutory
regulation or exceeds the permitted use, you will need to obtain permission directly from the copyright holder. To view a copy of thislicence, visit http://​creat​iveco​mmons.​org/​licen​ses/​by/4.​0/.

Shimada et al. Progress in Earth and Planetary Science(2023) 10:49

in tsunami geology (Hindson et al. 1996; Dawson et al.
1988, 2020; Atwater et al. 2005; Cisternas et al. 2005; Jankaew et al. 2008; Costa and Andrade 2020; Sawai 2020;
Ishimura et al. 2022) and paleotempestology (Liu andFearn 1993, 2000; Woodruff et al. 2009; Yao et al. 2020;
Bianchette et al. 2022) worldwide to better understandthe history of extreme waves and aid with prediction of
future such events. In both research fields, event deposits within lacustrine and marshy sediments have served
as a guide to reconstruct inundation areas (Nanayama
et al. 2003; Sawai et al. 2012; Pilarczyk et al. 2021) and therecurrence history of extreme waves (Liu and Fearn 1993;
Nanayama et al. 2003; Sawai et al. 2009a).
The Pacific coast of central and western Japan has
been repeatedly inundated by tsunamis and storms(Arakawa et al. 1961; Watanabe 1998; Usami et al. 2013;
Ishibashi 2014), but the inundation history over the past
thousands of years has not been fully reconstructed. Fortsunami history, the governmental long-term evaluationof subduction-zone earthquakes in this region is based
mainly on historical documents recording the nine greatearthquakes and their accompanying tsunamis that haveoccurred during the last 1300 years along the Nankai
Trough (The Headquarters for Earthquake Research Promotion 2013; Fig. 1d). However, records of earthquakes
before the seventeenth century are fewer and more fragmentary than those of earthquakes after the eighteenth
century (Koyama 1999). Koyama (1999) pointed out that
there is a lack of records for the medieval period or earlier. In addition, the 1300 years of history recorded in
written documents is too short to evaluate earthquakes
because unusually giant earthquakes recur at intervals of
hundreds of years (Satake and Atwater 2007). In evaluations of storm surges and waves, the shortness of the historical and instrumental records is the most problematic
issue. Written documents for the past 300 years show that
the Pacific coastal areas of Japan have been repeatedly
affected by storms (Arakawa et al. 1961), but there is little information on the exact values of inundation heights
and distances. Specific data from instrumental records areavailable only for the past 70 years. Recent instrumental
records have demonstrated that typhoons, as shown, forexample, by the Isewan Typhoon in 1959, can cause widespread inundation of the Pacific coast of central Japan,
but the frequency of such extraordinary typhoons is notknown. Datasets based on geological records are therefore required for long-term assessment of coastal hazards
resulting from marine inundation in this coastal area.
Geological evidence for tsunamis and storms has beenreported from more than 80 areas along the Nankai
Trough (e.g., Tsuji et al. 2002; Okamura and Matsuoka
2012; Kitamura et al. 2013); however, inter-regional correlation among these areas remains difficult because ofPage 2 of 28

sparse dates (Okamura and Matsuoka 2012) and discontinuities in the sedimentary record (Fujino et al. 2018;
Tanigawa et al. 2018; Shimada et al. 2019). In this study,
we obtained sediment cores containing a continuousrecord from 3000 years ago to the present day from Lake
Kogare-ike, a coastal lake facing the central part of the
Nankai Trough (Fig. 1). Application of X-ray computed
topographic imaging and construction of Bayesian age–depth models for the sediment samples enabled us torecognize faint geological records of extreme wave eventsin historical and prehistoric time. The inundation history deduced in this study can provide new insights into
inter-regional correlations with previously reported tsunami deposits along the Nankai Trough and contributeto the reconstruction of prehistoric and historic tsunami
history.

2 Setting
2.1 Tectonics

Subduction-zone earthquakes occur repeatedly in the
Nankai Trough, where the Philippine Sea plate is subducting beneath the Eurasian plate (Ishibashi 2004; Sangawa 2007; Usami et al. 2013; Fig. 1a). The coastal areas
along the Nankai Trough are conventionally divided
into six parts, Z and A–E (Fig. 1d) or, more simply, four
segments (the Hyuga-nada segment, Z; the Nankai segment, A and B; the Tonankai segment, C and D; and theTokai segment, E). Our study site is located in part C of
the Tonankai segment. Instrumental and written records
indicate that earthquakes have occurred with variablerupture zones and the associated tsunamis have inundated coastal areas in this region (Central Meteorological Observatory 1945; Ishibashi and Satake 1998; Fig. 1d).
The latest two earthquakes on the Tonankai segment(the 1944 CE Showa-Tonankai and 1854 CE Ansei–Tokai
earthquakes) caused coseismic coastal subsidence in the
study area in Minami-ise town (Central Meteorological Observatory 1945; Hatori 1978; Education Center ofNanto Town 2000) (Fig. 1b). The geological record indicates that earthquakes and tsunamis have occurred along
the Nankai Trough not only in the historical period but
also over the past few thousand years (Garrett et al. 2016;
Fujiwara et al. 2020a).
2.2 Historical records of extreme waves around the studysite
2.2.1 Tsunamis

Instrumental records, written documents, tsunami monuments, and oral tradition record the latest three historical tsunamis associated with earthquakes along the
Nankai trough: the 1944 CE Showa-Tonankai; the 1854
CE Ansei–Tokai; and the 1707 CE Hoei tsunamis (e.g.,
Editing Committee of the History of Nanto Town 1985;

Shimada et al. Progress in Earth and Planetary Science(2023) 10:49

Page 3 of 28

Fig. 1 a Index map of Japan. b Map of eastern Mie Prefecture showing the study site. c Location of the “Lake Kogare-ike” study site on the eastern
shore of Kamizaki Bay. d Ages and estimated rupture zones of historical tsunamis along the Nankai Trough (modified from The Headquarters
for Earthquake Research Promotion 2013). Red solid lines, red dotted lines, and gray dotted lines represent certain, probable, and possible rupturezones, respectively. The 1605 Keicho earthquake (dotted blue line) is regarded as a tsunami earthquake or an earthquake off an adjacent trench(Yamamoto and Hagiwara 1995; Seno 2002; Ishibashi and Harada 2013; Tsuji 2016). Generic Mapping Tools (Wessel et al. 2013) was partly used
to create Fig. 1d. e Sediment sampling locations and topographic setting of the study site. The photograph was captured by an unmanned aerial
vehicle. f Aerial photograph of the study site showing the survey transect from the shoreline (point A) to the inland end of Lake Kogare-ike (point
G). The photograph was taken by the Geospatial Authority of Japan on 24 September 2008 (photograph number: CKK-2008-1X-C13A-13). g
Topographic profile from the shoreline to the study site. Elevations are expressed with respect to mean sea level at Tokyo Bay (Tokyo Peil, T.P.)

Watanabe 1998; Education Center of Nanto Town 2000;
Editing Committee of the History of Nansei Town 2004;
Namegaya and Tsuji 2005). Instrumental records showthat the wave height of the 1944 CE Showa-Tonankai
tsunami was 5.4 m on the eastern shore of Kamizaki Bay(Editing Committee of the History of Nanto Town 1985;

Shimada et al. Progress in Earth and Planetary Science(2023) 10:49

Fig. 1c). Namegaya and Tsuji (2005) proposed, basedon a local historical document, that the western shoreof Kamizaki Bay was inundated by the 1854 CE Ansei–
Tokai tsunami with a run-up height of 5.2 m Tokyo Peil
(T.P.). They also estimated the run-up height of the 1707
CE Hoei tsunami on the eastern shore of Kamizaki Bay
to have been 8.6 m, from an oral tradition that the tsunami reached Higashi-Shinta, approximately 1 km fromthe shoreline (Fig. 1c).
2.2.2 Storms

According to historical records, at least 55 storms
attacked the area around Ise Bay (locally called “Isewan”)
between 1600 and 1865 CE (Arakawa et al. 1961; Fig. 1b,d; Additional file 1: Table S1). Instrumental records ofwind directions, storm surges, and damage are available
for storms since the mid-twentieth century. As many as
29 storms affected Mie Prefecture (Fig. 1a) between 1945
and 2017 CE (Japan Meteorological Agency 2017). The
largest storm was the 1959 CE Isewan Typhoon, during
which the highest storm surge in the observation history
in Japan was recorded (Committees for Technical Investigation on Lessons Learned from Past Disasters 2008).
The storm surge resulted in hazardous coastal inundation along Ise Bay and its surroundings. The inundation
reached more than 15 km inland in the coastal areas
north of Ise Bay (Fig. 1d) in association with a sea-level
rise to 4.27 m T.P. (Japan Meteorological Agency 1961).
Around the study site, sea level rose to 2.8 m T.P. andinundation extended more than 10 km inland at Gokasho
Bay (Fig. 1b). According to one observation report, sealevel rose to 2.3 m T.P. on the western shore of Kamizaki
Bay (Editing Committee of the History of Nanto Town
1985; Fig. 1c).
2.3 Geomorphology

Coastal lakes and ponds are useful archives for reconstructing the long-term and continuous history of inundation by extreme waves (e.g., Sugawara et al. 2008;
Baranes et al. 2016). These water bodies are usually stable depositional environments, in which coarser-grained
event deposits sporadically emplaced by extreme waves
are interbedded with muddy lacustrine sediments.
Coastal wetlands such as swales and marshes are also
commonly studied to identify event deposits (e.g., Atwater 1987; Cisternas et al. 2005; Pilarczyk et al. 2021;
Okada et al. 2022), but coastal lakes/ponds have theadvantage of a high preservation potential due to theirstable, calm environment, and unfilled accommodation
space (Szczuciński 2012; Kempf et al. 2017). This setting
usually prevents event deposits from being completely
eroded and yields reliable materials for dating.

Page 4 of 28

Our study site was a coastal lake called Kogare-ike,which is isolated from the Pacific Ocean by a sand barrier(Fig. 1e, f ). The height of the barrier is around 5 m T.P.,
though the maximum height is 4.8 m T.P. on the surveyed
transect (Fig. 1f ). Kamizaki Bay is a part of a complex riatype coast, and the lake is surrounded by hills. The hills
are composed of Upper Cretaceous pelitic mixed rockthat includes blocks of basalt, chert, reddish siliceous
mudstone, and sandstone (Nishioka et al. 2010). A narrow drainage channel connects the lake to the sea, but
inflow and outflow are controlled by a sluice at present(Fig. 1e). The lake is more than 1 m deep in most parts
and becomes deeper toward the center (Fig. 1f ). The
maximum depth is 2.3 m, and the elevation of the lakebottom at that point is approximately − 1.5 m T.P. The
mean sea level at the study site is approximately equal toT.P.; the mean tide level at Owase, the nearest tide station to the study site, has been + 0.07 m T.P. for the past
10 years (from 2012 to 2021 CE) (Japan Meteorological
Agency 2022a, 2022b). Agricultural activity around ourstudy site can be recognized in the land use map for 1892,
and aerial photographs show that farming was abandoned around 2003 (Additional file 2: Figure S1). Thisdate is consistent with information from a landownerthat the area of fringe marshes was used for agriculture
before it was abandoned 10–15 years ago. There is nolarge freshwater input to the lake.
2.4 Previous studies on tsunami deposits in the study area

The inundation history at Lake Kogare-ike has not been
explored previously, except for a preliminary study by
Komatsubara et al. (2007), who collected three short(150-cm-long) cores from the lake bottom. They identified a sandy event deposit interbedded with muddy
sediments by visual observation. A radiocarbon age of
980 ± 20 BP (recalibrated by this study to 958–791 cal BP)
was obtained approximately 20 cm below the sand layer.
The geological evidence for earthquakes and tsunamis along the Nankai Trough has been summarized in asystematic review by Garrett et al. (2016). Here we summarize the geological record in areas adjacent to ourstudy area, including reports published after Garrett
et al. (2016). In the area east of our study site, tsunami
deposits have been reported by Fujino et al. (2018) andOkahashi et al. (2005). Fujino et al. (2018) documented
10 sandy tsunami deposits dating to between 4500 and
500 cal BP in the Shijima lowland (Fig. 1b). The tsunami
deposits were identified within muddy sediments by visual inspection and on X-ray radiographs. Three of thesetsunami deposits were correlated with historical tsunamis (the 684 Hakuho tsunami, 1096 Eicho tsunami, and
1498 Meio tsunami). Okahashi et al. (2005) identified 12
tsunami deposits in a marsh at Osatsu (Fig. 1b) based on

Shimada et al. Progress in Earth and Planetary Science(2023) 10:49

sedimentary structures and micropaleontological evidence. They obtained radiocarbon ages from wood fragments and cones, but did not construct a chronology
or estimate the depositional ages of each event deposit.
To the west of our study site, Komatsubara et al. (2007)
reported their preliminary results of stratigraphic observation of two short cores. They found 5–11 event deposits within peaty deposits at the Shihara lowland (Fig. 1d),
but did not carry out detailed analysis of the samples.

3 Methods
3.1 Collection of sediment samples

We collected lacustrine sediments from seven locations at Lake Kogare-ike (K-1 to K-7; Fig. 1e, f ) in 2018and 2019. We assembled a raft with a hole in the centeras a platform and pushed a 50-cm-long Russian sampler
down to 4.75 m depth below the lake bottom. To recovercontinuous sedimentary records, we collected alternate
overlapping samples from three or four different parallel
holes. The holes were less than 1 m apart and overlapsof 25 cm were applied. In addition, push-core samples
were obtained at K-1, K-4, and K-5 to recover soft surface sediments. The elevations of the ground surface
and the present sea level were obtained with a network
Real Time Kinematic-Global Navigation Satellite System(RTK-GNSS) survey system from Leica Geosystems Inc.
(Norcross, Georgia, USA). The standard deviation (1SD) of the leveling in our study was 0.005–0.017 m horizontally and 0.010–0.043 m vertically. The elevations of
the lake bottom were obtained by subtracting the water
depth measured with a leveling rod from the elevation of
the lake surface measured with RTK-GNSS.
3.2 Observation and grain size analysis of sediment
samples

We observed the collected sediment samples in both thefield and the laboratory. In the laboratory, X-ray computed tomography (CT) scans of the cores were obtainedwith a Hitachi Supria Grande Premium CT scanner(Hitachi, Ltd., Tokyo, Japan) at the Geological Surveyof Japan (Tsukuba, Ibaraki, Japan). The CT images werereferred to during visual observation and used to guidesubsampling for radiocarbon dating, gamma-ray spectrometry, grain size analysis, and diatom assemblage
analysis.
Grain size analyses of the samples were conductedusing a Retsch Camsizer (Verder Scientific, Haan, Germany), an image analyzer capturing a wide range of grain
size from silt to pebbles. All samples were sieved in waterto remove the mud component and then treated withhydrogen peroxide solution to remove the organic components. During the procedure, large pieces of organic
matter (roots, leaves, and wood fragments) were removed

Page 5 of 28

by hand. The sand content of each sample was calculated
from the resulting weight reduction. The measurement
range of grain size analysis on the residual sand fractions
was set from − 5.25 to + 6.25 phi at intervals of 0.25 phi.
Grain size distribution properties were calculated using
the logarithmic graphical method of Folk and Ward
(1957).
3.3 Analysis of volcanic glass shards

The elemental concentrations of volcanic glass shards
in event deposit E11 at locations K-3, K-4, K-6, and K-7
were measured with an energy dispersive X-ray microanalyzer (EDX) by Furusawa Geological Survey Inc.,
Okazaki, Japan. Identification of glass sources was carried out by referring to previously published regional
studies (Sugihara 1984; Machida and Arai 2003; Sugiuchi
and Fukuoka 2005; Tsukui et al. 2006; Kobayashi et al.
2020; Murata et al. 2021).
3.4 Diatom analysis

Fossil diatom assemblages in 127 samples from the corefrom K-4 were studied to reconstruct the paleoenvironment of the study site. The samples were prepared by thebleaching method (Nagumo and Kobayasi 1990; Nagumo
1995). Fossil diatoms were identified under an optical
microscope. More than 200 valves were counted per sample. Identification and ecological information followed
standard (e.g., Patrick and Reimer 1966, 1975; Simonsen
1987; Krammer and Lange-Bertalot 1986, 1988, 1991a, b)
and regional (Kosugi 1988; Kobayasi et al. 2006) sources.
Fossil diatoms were divided into seven groups on thebasis of their autoecology: freshwater plankton, freshwater–brackish plankton, freshwater tychoplankton, freshwater benthos, freshwater–brackish benthos, brackishbenthos, and unknown.
Stratigraphically constrained cluster analysis (Grimm
1987) was performed to statistically identify the zonationof fossil diatom assemblages. The cluster analysis was
carried out using the Bray–Curtis distance among theassemblages and constraining the stratigraphic order ofthe samples. The cluster analysis was executed using the
package “rioja” in R version 4.2.2 (Juggins 2022).
3.5 Radiocarbon dating, gamma spectrometry,
and Bayesian age–depth modelingRadiocarbon dating of plant macrofossils and fossil pollen grains was conducted to estimate depositional ages.
Macrofossils were selected under a binocular microscope. Fossil pollen grains were extracted from sediment
samples at the Research Centre for Palaeoclimatology,
Ritsumeikan University (Kusatsu, Shiga, Japan) following the procedure of Yamada et al. (2021). Radiocarbon
dating of plant macrofossils and pollen grains was carriedShimada et al. Progress in Earth and Planetary Science(2023) 10:49

out by Beta Analytic (Miami, Florida, USA) and the University of Tokyo (Tokyo, Japan), respectively.
Cesium-137 activities of 29 samples from shallower
than 140 cm depth in the sediment core collected at K-4
were measured by high-resolution gamma spectrometry(GCW2022; Mirion Technologies [Canberra], Inc, Meriden, Connecticut, USA). Cesium-137 is one of the mostcommon radionuclides used to establish chronologies
for sediments during the last 80 years. Cesium-137 is ananthropogenic radionuclide that was produced by atmospheric nuclear weapons testing from 1950 to ∼1975. In
the northern hemisphere, undisturbed sedimentary
records generally show the first occurrence (onset) andthe peak of 137Cs as time markers for 1954 and 1963–
1964, respectively (Ritchie and McHenry 1990; Appleby
2001), though 137Cs onset might have occurred earlier inwestern Japan than in other regions in the northern hemisphere because of the nuclear bomb attacks on Hiroshima and Nagasaki in World War II (see discussion).

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A Bayesian age–depth model of the sediments at K-4
was constructed based on the radiocarbon ages. We
applied Bchron (Haslett and Parnell 2008; Parnell et al.
2008), a package written for the free open-source language R (R Core Team 2021). Age–depth models wereconstructed by using IntCal20 (Reimer et al. 2020) as
the radiocarbon calibration curve. Event deposits weretreated as instantaneous, and their thicknesses were
removed from the original core log of K-4 to construct anevent-free stratigraphic column (Kempf et al. 2017, 2020;
Sabatier et al. 2017). For this reason, the radiocarbon age
measured within event deposit E1 (no. 20 in Table 1) was
not used in the age–depth model. Because of the stochastic nature of the Bchron model, which uses random sampling, each run produces slightly different depositional
ages; therefore, the depositional ages of event deposits
E1–E19 were determined by averaging the ages obtained
from the three trials.

Table 1 Radiocarbon ages from the study site
No. Core location Elevation (m T.P.) Position

Conventional Calibrated ageage (14C BP)
2σ range (calBP)

Calibrated age
Material
2σ range (cal CE/
BCE)

Lab no.

1K-1− 2.51 to − 2.45

Between E3
and E4304 ± 34

461–295

1489–1655 CE

Concentrated
pollenTKA-23301

2

K-1− 2.98 to − 2.94

Below E5

777 ± 33

730–668

1220–1282 CE

Concentrated
pollenTKA-23296

3

K-1− 3.73 to − 3.69

Above E7590 ± 30

647–540

1303–1410 CE

Leaves, Cladium chinense,and seeds of unidentified taxa

Beta-581028

4K-1− 4.05 to − 4.01

Between E7
and E8870 ± 30

901–6901049–1260 CE

A leaf of unidenti- Beta-581029
fied taxon

5

K-1− 4.45 to − 4.43

Below E91130 ± 30

1174–958

776–992 CE

Najas spp.,
Potamogeton
spp., Trapa spp.,
and unidentified
taxa

Beta-531173

6K-1− 4.85 to − 4.81

Below 101430 ± 30

1367–1294

583–656 CE

Leaves and seedsof unidentified
taxa

Beta-581030

7K-1− 5.06 to − 5.04

Above E131720 ± 30

1700–1539

250–411 CE

Fruits of unidenti- Beta-531172
fied taxa

8

K-1− 5.39 to − 5.35

Below E14

1841 ± 36

1829–1627

121–323 CE

Concentrated
pollenTKA-23290

9K-2− 2.08 to − 2.06

Above E2

150 ± 30

282–0

1668–present

BranchBeta-531171

10

K-2

− 2.53 to − 2.50

Below E3273 ± 50475–0

1475–present

Concentrated
pollenTKA-23294

11

K-2− 3.27 to − 3.24

Between E5
and E6624 ± 33

655–551

1295–1399 CE

Concentrated
pollenTKA-23302

12

K-2− 3.94 to − 3.90E7470 ± 30

541–492

1409–1458 CE

Leaves and seedsof unidentified
taxa

Beta-581031

13

K-2

− 4.18 to − 4.14

Below E71100 ± 30

1062–933

888–1017 CE

Leaves and seedsof unidentified
taxa

Beta-581032

Shimada et al. Progress in Earth and Planetary Science(2023) 10:49

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Table 1 (continued)
No. Core location Elevation (m T.P.) Position

Conventional Calibrated ageage (14C BP)
2σ range (calBP)

Calibrated age
Material
2σ range (cal CE/
BCE)

Lab no.

14

K-2− 4.68 to − 4.65

Above E91290 ± 30

1290–1154

660–796 CE

Leaf

Beta-583061

15

K-2

− 4.81 to − 4.77

Between E9
and E101330 ± 30

1300–1177

650–773 CE

Leaves and seedsof unidentified
taxa

Beta-581033

16

K-2

− 5.28 to − 5.23

Above 12560 ± 30

640–524

1310–1426 CE

Leaves and seedsof unidentified
taxa

Beta-581034

17

K-3

− 3.41 to − 3.36

Above E5780 ± 30

729–672

1221–1278 CE

Najas spp., Seeds Beta-664442
of unidentified
taxa, a fruit of unidentified taxon,and leaves

18

K-3

− 3.55 to − 3.51

Below E6900 ± 30

908–732

1042–1218 CE

Seeds of unidenti- Beta-665871
fied taxa, leaves

19

K-3

− 4.59 to − 4.55

Between E10
and E111490 ± 30

1404–1309

546–641 CE

Leaves, Cladium chinense,and seeds of unidentified taxa

Beta-583062

20

K-4

− 2.09 to − 2.05

Within E1

287 ± 28

446–158

1504–1792 CE

Concentrated
pollenTKA-23297

21

K-4− 2.72 to − 2.68

Above E2

209 ± 37

311–0

1639–present

Concentrated
pollenTKA-23298

22

K-4− 3.09 to − 3.05

Below E3

303 ± 32

458–295

1492–1655 CE

Concentrated
pollenTKA-21913

23

K-4− 3.33 to − 3.34

Below E5260 ± 30

435–0

1515–present

Trapa spp.

Beta-545282

24

K-4

− 3.46 to − 3.42

Between E5
and E6653 ± 38

670–555

1280–1395 CE

Concentrated
pollenTKA-24604

25

K-4− 3.59 to − 3.55

Below E6901 ± 45915–7291035–1221 CE

Concentrated
pollenTKA-21914

26

K-4− 3.9 to − 3.88

Below E7820 ± 30

782–678

1168–1272 CE

Leaves and seedsof unidentified
taxa

Beta-528350

27

K-4

− 4.05 to − 4.00

Above E8660 ± 30

670–557

1280–1393 CE

LeavesBeta-545283

28

K-4

− 4.06 to − 4.02

Above E81371 ± 27

1344–1179

606–771 CE

Concentrated
pollenTKA-24607

29

K-4− 4.28 to − 4.23

Between E8
and E91292 ± 39

1294–1126

656–824 CE

Concentrated
pollenTKA-21911

30

K-4− 4.36 to − 4.35

Below E91300 ± 30

1289–1177

661–773 CE

LeavesBeta-528349

31

K-4

− 4.54 to − 4.51

Below E101490 ± 30

1404–1309

546–641 CE

LeavesBeta-528351

32

K-4

− 4.75 to − 4.69

Between E11
and E121540 ± 30

1516–1353

434–597 CE

Leaves and seedsof unidentified
taxa

Beta-545279

33

K-4

− 4.84 to − 4.82

Above E121480 ± 30

1398–1308

552–642 CE

Leaves, branch,
and fruits of unidentified taxa

Beta-528352

34

K-4

− 5.03 to − 5.02

Below E141650 ± 30

1687–1414

263–536 CE

BranchBeta-528353

35

K-4

− 5.14 to − 5.03

Below E141500 ± 30

1466–1308

484–642 CE

LeavesBeta-528354

36

K-4

− 5.23 to − 5.17

Between E14
and E151980 ± 30

1991–1831

41 BCE–119 CE

Fruits or seeds
Beta-545280
of an unidentified
taxon

37

K-4

− 5.33 to − 5.28

Above E15

2040 ± 27

2098–1891

148 BCE–59 CE

Concentrated
pollenTKA-24606

38

K-4− 5.60 to − 5.57

Between E15
and E162420 ± 30

2695–2352

Najas spp., Fruitsof unidentified
taxa, and leaves

Beta-545281

745–402 BCE

Shimada et al. Progress in Earth and Planetary Science(2023) 10:49

Page 8 of 28

Table 1 (continued)
No. Core location Elevation (m T.P.) Position

Conventional Calibrated ageage (14C BP)
2σ range (calBP)

Calibrated age
Material
2σ range (cal CE/
BCE)

Lab no.

39

K-4− 5.70 to − 5.66

Below E162598 ± 40

2836–2516

886–566 BCE

Concentrated
pollenTKA-23299

40

K-5− 2.41 to − 2.37

Below the 3rd
shallowest event
deposit319 ± 30

461–306

1489–1644 CE

Concentrated
pollenTKA-23300

41

K-5− 3.07 to − 3.03

Below the 3rd
deepest event
deposit

411 ± 35

522–326

1428–1624 CE

Concentrated
pollenTKA-23289

42

K-5− 4.22 to − 4.18

Below E8970 ± 30

928–798

1022–1157 CE

Leaves of Dicotyledoneae sp.

Beta-581027

43

K-5

− 4.55 to − 4.50

Below the deepest event deposit

1567 ± 37

1523–1383

427–567 CE

Concentrated
pollenTKA-23292

44

K-6− 3.57 to − 3.52

Below E5620 ± 30

651–551

1299–1399 CE

Seeds of unidentified taxa,
and leaves

Beta-664440

45

K-6

− 5.16 to − 5.12

Above E121310 ± 30

1293–1177

657–773 CE

LeavesBeta-664441

46

K-7

− 3.96 to − 3.92

Above E8720 ± 30

719–568

1231–1382 CE

Seeds of unidentified taxa, a fruitof unidentified
taxon, and leaves

Beta-664438

47

K-7

− 5.43 to − 5.39

Below E14

1820 ± 30

1821–1624

129–326 CE

Najas spp., Seeds
of unidentified
taxa

Beta-664439

Conventional ages were converted to cal BP with Bchron (Haslett and Parnell 2008; Parnell et al. 2008) using the calibration curve of IntCal20 (Reimer et al. 2020)

4 Results
4.1 Stratigraphy and ages of sediment samples
from Kogare‑ike
4.1.1 Detailed description of the stratigraphy at the centerof the lake

The sediments from Lake Kogare-ike were composed oforganic mud, sand, gravel, inorganic mud, and volcanicash. Sand, sandy mud (with/without gravel), inorganic
mud, and white volcanic ash layers were observable with
the naked eye both in the field and in the laboratory
as distinctive layers interbedded with organic mud. In
addition to these, there were 10 high-radiodensity layers that could be recognized only in X-ray CT images.
The high-radiodensity layers possibly possessed different mineral compositions (Boespflug et al. 1995) orlarger grain size (Godbout et al. 2019) than the organic
mud layers. We regard these sand, sandy mud, inorganic
mud clasts, volcanic ash, and high-radiodensity layers as
event deposits (Figs. 2, 3 and 4; Additional file 3: Figure S2, Additional file 4: Figure S3, and Additional file 5:
Figure S4; Additional file 6: Table S2) and have named
them E1–E19 from top to bottom. Details of each event
deposit at K-4 are described in the following paragraphs.
Details of the contents of mud, organic material, andsand and the results of grain size analysis in each layer
are provided in Fig. 4 and Additional file 7: Table S3.

E1, the uppermost event deposit, was gray in color andranged from 46.9 to 77.1 cm core depth. This event deposit
was visible as a stack of three subunits both to the naked
eye and in the CT images (Fig. 4). We could not recognize
an erosional lower contact with the underlying organic
mud in the CT images. Parallel laminae were distinct
within the subunits. The sand content of E1 was between
0.9 and 3.4 wt.% (Fig. 4 and Additional file 7: Table S3).
E2 appeared as a grayish faint inorganic mud layer tothe naked eye, and was visible as a high-radiodensity
layer from 140.0–141.8 cm depth in the CT images(Fig. 4). The sand content was low (< 1 wt.%) above,
below, and within event deposit E2 (Fig. 4 and Additional
file 7: Table S3).
E3 was a sandy layer at 153.6–156.2 cm core depth. The
lower part of E3 consisted of medium sand (mean grain
size = 1.55 φ). The upper part of the event deposit contained multiple layers of gray inorganic mud and brownorganic mud. Changes in the sand contents show grading
in grain size of this event deposit. The upper part was difficult to recognize with the naked eye but was clear on
the CT images (Fig. 4).
E4 and E5 were recognized as high-radiodensity layers only on the CT images at 172.9–173.8 cm and 183.9–
184.5 cm core depth, respectively. Event deposit E4
showed very faint lateral continuity within the core. The

Shimada et al. Progress in Earth and Planetary Science(2023) 10:49

Page 9 of 28

Fig. 2 Lithostratigraphy from K-1 to K-5 (locations in Fig. 1e). Radiocarbon ages are listed in Table 1. The correlations of event deposits between coresites are shown as solid lines, with tentative correlations indicated by dotted lines. Stacks of X-ray CT images are shown on Additional file 3: Figure
S2 and Additional file 4: Figure S3. A photograph of the core from K-4 is provided in Additional file 3: Figure S2

sand contents of the two event deposits were less than 1
wt.% above, below, and within the layers (Fig. 4 and Additional file 7: Table S3).

E6 occurred between 207.5 and 219.8 cm core depth.
As shown in the CT images, this event deposit was notclearly divided into subunits, in contrast to the younger

Shimada et al. Progress in Earth and Planetary Science(2023) 10:49

Page 10 of 28▸

Fig. 3 Lithostratigraphy of a transect connecting K-6, K-4, and K-7(locations in Fig. 1e). Radiocarbon ages are listed in Table 1. The
correlations of event deposits between core sites are shown
as solid lines, with tentative correlations indicated by dotted lines.
Stacks of X-ray CT images are shown on Additional file 3: Figure S2
and Additional file 4: Figure S3. A photograph of the core from K-4
is provided in Additional file 3: Figure S2

events. It was difficult determine from visual observation
and CT images whether the non-layered structure of E6
was created by obvious bioturbation or deformation of
the sand layer after rapid deposition (e.g., load casting).
The sand content of E6 ranged from 10.5 to 31.2 wt.%. E6
contained a small proportion of granules, but the mean
grain size of the sieved residual sand fractions was fine–coarse sand (2.49–0.25 φ, average = 1.17 φ). The sandcontent and the results of grain size analyses show grading within the event layer.
A faint event deposit (E7) at 231.0–231.3 cm core depth
was recognized only on CT images. This deposit was visible as a rather patchy horizontal distribution within thecore, rather than a stratified high-radiodensity layer. The
sand content was low (less than 1 wt.%) above, below, andwithin event deposit E7.
E8 was a light grayish inorganic mud layer between
269.1 and 273.6 cm core depth. A few mud clasts wererecognized both in the field and on CT images at the baseof the event deposit. The CT images showed an upwardchange in radiodensity from high to low, which may represent upward fining of grain size; however, no grading
was recognized in the sand content or grain size analysis.
E9 and E10 were visible only on the CT images at
283.4–284.3 cm and 303.4–304.2 cm core depth, respectively. The two layers exhibited high radiodensity andwere distinctively stratified on the images, but there wasno significant change in the sand content (< 1 wt.%) or
the grain size of the residual sand fractions.
E11 appeared at 313.3–317.9 cm core depth. The bottom part of the event deposit was a very distinctive whitelayer in the field. The middle and upper parts of the layer
were hard to recognize with the naked eye but visible on
the CT images. This event layer consisted of pumice andglass shards. EDX analyses indicated that the shards hadvarious origins (see below), suggesting erosion and redeposition from older strata.
A faint event deposit, E12, was visible only on CT
images at 337.2–338.4 cm core depth. This event deposit
was apparent as a patchy distribution within the core,
and the sand content of the sediment was less than 1
wt.% above, below, and within the event deposit.

Shimada et al. Progress in Earth and Planetary Science(2023) 10:49

Page 11 of 28

Fig. 4 Photographs, X-ray CT images, and sketches of the event deposits identified at K-4 with the results of grain size analysis (sand content
and mean grain size of residual sand fractions). The mean grain size, shown as black circles, represents that of samples with very small sand contents(10% or less)Shimada et al. Progress in Earth and Planetary Science(2023) 10:49

Page 12 of 28

Fig. 4 continued

E13 was distributed at 346.5–350.9 cm core depth. To
the naked eye, the upper part of the event layer was visible as a light grayish inorganic mud layer. The upper part
had higher density than the lower part in the colored CT

image. There was no significant difference in sand content between the upper and lower parts.
E14 was identified only on the CT images, and occurred
between 352.6 and 354.8 cm core depth. The CT images

Shimada et al. Progress in Earth and Planetary Science(2023) 10:49

showed higher radiodensity at the bottom of the layer
than in the middle and upper parts, but no difference
was recognizable with the naked eye. The sand content
was persistently low (less than 1 wt.%) above, below, andwithin the event deposit.
E15 occurred between 388.5 and 406.0 cm core
depth. The CT images showed three sandy and muddy
subunits within the event deposit. The middle subunit
appeared not to be stratified, but we could not determine whether the lack of strata was caused by bioturbation or load casting. The sand contents and grain size
analyses of the subunits indicated that the lowest subunit contained much more sand and coarser fractions
than the higher subunits. The mean grain size of theuppermost, middle, and lowest subunits was fine sand(mean grain size is 2.68–2.12 φ), very fine sand (meangrain size = 3.57–3.45 φ), and medium sand (meangrain size = 1.37 φ), respectively. Debris filled the space
between the upper and middle subunits. The lower contact with the underlying deposit was sharp (less than
5 mm).
E16 was visible to the naked eye as a distinct grayish
sandy layer at 416.2–417.4 cm core depth. The sand content and mean grain size of the event were 48.5 wt.% andcoarse sand (mean grain size = 0.85 φ), respectively. The
upper and lower contacts with the overlying and underlying deposits were sharp.
E17 was a grayish sandy mud layer at 436.2–446.7 cm
core depth. The CT images showed clear parallel laminae in the upper and basal parts of the event deposit.
The sand content in the basal part of the event deposit
was low sand content (< 1 wt.%), whereas the upper threesamples within the event deposit showed relatively highsand content (11.6 to 39.7 wt.%). The mean grain size ofthe residual sand fractions of the upper three samples
was fine–coarse sand (2.72– 0.53 φ, 1.52 φ on average).
The sand contents and mean grain size indicated upwardfining of grain size. The well-preserved clear laminae may
suggest that this event layer had not been vertically disturbed, for example, by benthic fauna, after deposition.
E18 was recognized as a grayish sandy mud layer at449.1–459.1 cm core depth. The CT images enabled us torecognize clear laminae in the upper and lower parts of
the layer. The sample in the upper part had low sand content (7.5 wt.%). The sand contents in other four samples
ranged from 12.6 to 48.4 wt.%. The mean grain size ofthe residual sand fractions in these four samples was veryfine to medium sand (1.79 to 3.22 φ, average 2.57 φ). The
CT images and changes in the sand content and mean
grain size indicated inverse and normal grading within
the event layer.

Page 13 of 28

The lowest event deposit, E19, was a gray distinctive layer that consisted of sandy mud, muddy sand, andgranules. The CT images showed laminae, visible as
high-radiodensity layers. The sand content was 1.2–64.4
wt.%. Grain size analyses showed that the residual sandfractions were in the medium sand range (mean grain
size of the four samples with a sand content of more than
10% = 1.08 phi). The top of E19 was observed at 461.3 cmcore depth, but the bottom was not clearly confirmed in
our fieldwork.
4.1.2 Major‑element composition of volcanic glass shards

The major-element compositions of individual glass
shards from the volcanic ash layer (event deposit E11)
showed various sources and ages of eruptions (Additional
file 8: Fig. S5). Major elements, especially MgO, CaO,­Na2O, and ­K2O, are useful to distinguish the sources of
volcanic ash layers around the study site (Sugihara 1984;
Machida and Arai 2003; Kobayashi et al. 2020; Murata
et al. 2021). Scatter plots of MgO–SiO2, CaO–SiO2,­Na2O–SiO2, and ­K2O–SiO2 for volcanic ash corresponding to E11 in four cores (K-3, K-4, K-6, and K-7) recognized four clusters of element compositions. Three of theclusters correspond to widespread tephras; specifically,
the Aira-Tn tephra (AT; 29–26 cal kBP), the Kikai-Akahoya tephra (K-Ah; 7.3 cal kBP), and the Amagi-Kawagodaira tephra (Kg; 3.1 cal kBP) (Sugihara 1984; Machida
and Arai 2003). The other cluster is similar to the element composition of Izu Island eruptions (low CaO andhigh ­Na2O and ­K2O), especially the Kozushima-Tenjosan
tephra (Iz-Kt; AD 838) (Sugihara 1984; Sugiuchi andFukuoka 2005; Tsukui et al. 2006; Kobayashi et al. 2020;
Murata et al. 2021), though this age exhibits a discrepancy with the radiocarbon dating.
4.1.3 Dating by radioactive isotopes

The obtained radiocarbon ages ranged from 2836–2516
to 282–0 cal BP (Figs. 2 and 3; Table 1). Forty-seven radiocarbon ages were obtained, 30 of which were from plant
macrofossils and the remaining 17 from concentrated
pollen fossils.
Cesium-137 was detectable in samples above 93 cm
in depth at K-4, but its concentration in deeper samples was below the detection limit (Fig. 5 and Additional file 9: Table S4). The concentration of 137Cs ranged
from 0.8 ± 0.2 to 1.9 ± 0.2 Bq/kg in samples between
93 cm and 82.5 cm core depth (below E1), but exhibited a pronounced increase just above E1. The highest
137
Cs concentration was recognized at 42.5 cm depth
below the lake bottom (28.5 ± 0.6 Bq/kg). Thereafter, theShimada et al. Progress in Earth and Planetary Science(2023) 10:49

Fig. 5 The 137Cs profile in the top of the core from K-4. Detailsof the analytical results are listed in Additional file 9: Table S4

concentration decreased consistently upward, except fora slight increase from 30.5 to 27.5 cm core depth.
4.1.4 Intra‑site correlation of event deposits among cores
taken from Kogare‑ike

The stratigraphic features and positions of event deposits, chemical composition of glass shards, and radiocarbon ages permitted us to correlate event deposits E1–E16
at K-4 with event deposits in the cores taken from theother six locations at Kogare-ike (K-1, K-2, K-3, K-5, K-6,
and K-7; Figs. 2 and 3). No event deposits correlativewith E17–E19 were present in any cores except for K-4,
because of short recovery of the sediment cores takenfrom K-1, K-2, K-3, K-5, K-6, and K-7. The correlations of
the other events (E1–E16) are described below; the characteristics of each event are summarized in Additional
file 6: Table S2.
Event deposit E1 found at K-4 correlated among K-1–
K-6 and possibly K-7. The correlation is made on thebasis of the depth of the lower contact of the event at
K-1–K-7 (about 50–100 cm from the lake bottom) andthe stratigraphic features at K-1 and K-3–K-6 (multiple
subunits of high radiodensity and parallel laminae).

Page 14 of 28

Event deposits E2 and E3 at K-4 were found at K-1–K-4
and K-6. The two event deposits were present in organic
mud younger than 461–295 cal BP (no. 1 in Table 1) at
K-1 and within organic mud dated between 282 and0 cal BP (no. 9 in Table 1) and 475–0 cal BP (no. 10 inTable 1) at K-2 and 311–0 cal BP (no. 21 in Table 1) and
458–295 cal BP (no. 22 in Table 1) at K-4. Two event layers at each of K-3 and K-6 were correlated with E2 and E3
on the basis of their depth and stratigraphic features: E2
is a very faint layer and E3 has multiple laminae with highradiodensity.
Event deposits E4 and E5 were correlated among K-1–
K-4 and K-6. The two events are located within organic
mud dated between 461 and 295 cal BP (no. 1 in Table 1)
and 730–668 cal BP (no. 2 in Table 1) at K-1, between
475 and 0 cal BP (no. 10 in Table 1) and 655–551 cal BP(no. 11 in Table 1) at K-2, and between 458–295 cal BP(no. 23 in Table 1) and 670–555 cal BP (no. 25 in Table 1)
at K-4. The correlations are also based on a radiocarbon age between E4 and E5 at K-3 (729–672 cal BP, no.
17 in Table 1) and a limiting maximum age of E5 at K-6(651–551 cal BP, no. 44 in Table 1). These deposits’ stratigraphic features (i.e., a very thin layer, or visible only on
CT images) are consistent among the cores.
Event deposit E6 of K-4 can probably be correlated with
event deposits at K-1–K-3, and possibly K-6 and K-7. This
correlation is based on stratigraphic features, depths ofevent layers, and three radiocarbon ages. Stratigraphic
features visible as multiple layers of high radiodensity
in the CT images are very similar between event layers
at K-2 and K-3. However, high-radiodensity layers correlative to deposit E6 are very thin at K-1. We correlated
event deposit E6 of K-4 with those of K-1–K-3 on thebasis of the depths of the event deposits and four radiocarbon ages (655–551 cal BP, no. 11 in Table 1; 908–
732 cal BP, no. 18 in Table 1; 670–555 cal BP, no. 24 inTable 1; and 915–729 cal BP, no. 25 in Table 1).
A faint event deposit, E7, is found at K-1–K-4, andpossibly K-6 and K-7. The depositional age of E7 at
K-4 is constrained by three limiting maximum ages
(782–678 cal BP, no. 26 in Table 1; 670–557 cal BP, no.
27 in Table 1; and 1344–1179 cal BP, no. 28 in Table 1)
obtained from samples below the event deposit E7 and alimiting minimum age (915–729 cal BP; no. 26 in Table 1)
from a sample above E7. At K-1 and K-2, two high-radiodensity layers correspond to the constrained age of E7
at K-4. Either of these two layers may be correlated with
E7 of K-4; alternatively, the two high-radiodensity layers may represent a single event such as a single tsunami
deposit having multiple subunits due to multiple waves.
Event layers at depths of ~ 284 cm at K-6 and ~ 229 cm at
K-7 are possibly correlated either E6 or E7 based on theirShimada et al. Progress in Earth and Planetary Science(2023) 10:49

depths on CT images and stratigraphic position above
the light grayish inorganic mud layer (E8).
Event deposit E8 was correlated among K-1–K-4, K-6,
and K-7 by means of sedimentary features (light grayish inorganic mud and mud clasts). Radiocarbon ages
from samples above E8 at K-1 (901–690 cal BP, no. 4 inTable 1) and below E8 at K-2 (1290–1154 cal BP, no. 14
in Table 1) is consistent with this visual correlation. An
event deposit at 362.1–367.0 cm depth at K-5 may also
correlate with E8, but an age from a sample below theevent (928–798 cal BP, no. 42 in Table 1) is younger thanevent deposit E8 at K-4.
Event layers corresponding to event deposit E9 at K-4
were found at K-1, K-2, K-3, K-6, and K-7. The depositional age of E9 at K-4 is constrained by a limiting maximum age (1289–1177 cal BP, no. 30 in Table 1) and alimiting minimum age (1294–1126 cal BP, no. 29 inTable 1). This age is consistent with a radiocarbon age
from a sample below the correlated event deposit of K-1(1174–958 cal BP, no. 5 in Table 1) and two ages fromsamples above and below the correlated event deposit of
K-2 (1290–1154 cal BP, no. 14 in Table 1; 1300–1177 calBP, no. 15 in Table 1). The event deposits at K-3, K-6, andK-7 are common in their relative position below the lightgrayish inorganic mud layer (E8).
Event deposit E10 at K-4 is found at all locations but
K-5. The depositional age of E10 at K-4 is constrainedby two radiocarbon ages (1289–1177 cal BP, no. 30 in
Table 1; 1404–1309 cal BP, no. 31 in Table 1). This age isconsistent with radiocarbon ages below an event depositat K-1 (1367–1294 cal BP, no. 6 in Table 1), above an
event deposit at K-2 (1300–1177 cal BP, no. 15 in Table 1),
and below an event deposit at K-3 (1404–1309 cal BP, no.
19 in Table 1). The event deposits observed at K-6 andK-7 were correlated with event deposit E10 at K-4 on thebasis of their relative position to the white volcanic ash
layer (E11).
Event deposit E11 of K-4, the volcanic ash layer withvarious sources, correlated with event layers found in
K-2, K-3, K-6, and K-7. Because the event deposits werevisible as white ash layers, preliminary correlation amongcores was possible in the field. The event deposits consisting of volcanic ash at K-2, K-3, K-4, K-6, and K-7 werecomposed of glass shards. As shown by major chemical component analysis, the event deposits at K-3, K-4,
K-6, and K-7 had in common that they contained glass
erupted by different ages and different sources. The similar compositions of the volcanic glass shards within these
event deposits and the radiocarbon ages from below
and above the event deposits (1404–1309 cal BP, no. 19
in Table 1; 1404–1309 cal BP, no. 31 in Table 1; 1516–
1353 cal BP, no. 32 in Table 1; and 1293–1177 cal BP,

Page 15 of 28

no. 45 in Table 1) enable correlation of the event layers
among K-3, K-4, K-6, and K-7.
Event deposits E12, E13, and E14 of K-4 were foundat K-1, K-2, K-6, and K-7. The three event deposits arepresent in organic mud dated between 1398–1308 cal
BP (no. 33 in Table 1) and 1687–1414 cal BP (no. 34
in Table 1) at K-4, between 1367–1294 cal BP (no. 6 inTable 1) and 1829–1627 cal BP (no. 8 in Table 1) at K-1,
older than 1293–1177 cal BP (no. 45 in Table 1) at K-6,
and younger than 1821–1624 cal BP (no. 47 in Table 1)
at K-7. These three event layers at K-2 were correlated bytheir depth in the cores and being stacks of three event
layers.
E15 and E16 at K-4 are possibly correlated with two
event deposits found at K-6. This correlation is based
only on their stratigraphic positions (depth below thelake bottom).
4.2 Fossil diatom assemblages

We identified a total of 845 diatom taxa from 127 samples of core K-4 (Fig. 6 and Additional file 10: Table S5).
The fossil diatom assemblages from Lake Kogare-ike
were divided into freshwater plankton, freshwater–brackish plankton, freshwater tychoplankton, freshwater
benthos, freshwater–brackish benthos, and brackish benthos based on their autecology (Fig. 6). Stratigraphicallyconstrained cluster analysis was applied to identify 10
biostratigraphic zones (Zones 1–10 from top to bottom;
Fig. 6). Detailed descriptions of each diatom zone areprovided below.
The topmost zone 1 (0–29 cm depth) is characterized
by a mixture of freshwater plankton (Aulacoseira spp. andDiscostella stelligera) and freshwater–brackish plankton(Cyclotella atomus, Cyclotella atomus var. gracilis, andCyclotella meneghiniana). Freshwater–brackish plankton(C. atomus, C. atomus var. gracilis, and C. meneghiniana) are dominant in organic mud samples of Zone 2(29–79 cm depth). Assemblages within E1 belonging toZone 2 contain the same dominant species of freshwater–brackish plankton as the assemblages in the organic
mud of this zone. The brackish benthic diatom Navicula
gregaria was also identified within E1. In Zones 3 (79–134.5 cm depth) and Zone 4 (134.5–163.5 cm depth),
freshwater diatoms are dominant, except in the uppermost part of Zone 3 and within E3 of Zone 4. Zone 3 ischaracterized by freshwater tychoplankton (Staurosira
binodis, Staurosira construens, and Staurosira venter)and freshwater benthos (Gogorevia exilis). Exceptionally, freshwater–brackish plankton (C. atomus, C. atomus
var. gracilis, and C. meneghiniana) exhibits high relative abundance in the uppermost part of Zone 3. Zone
4 is dominated by the freshwater–brackish benthic species Tabellaria fenestrata, but assemblages within E3 areShimada et al. Progress in Earth and Planetary Science(2023) 10:49

Page 16 of 28

Fig. 6 Fossil diatom assemblages in the core from K-4. The term “relative abundance” denotes the percentage of the total diatom valves counted.
The rightmost column shows the biostratigraphic zones based on stratigraphically constrained cluster analysis (Grimm 1987). Additional file 10:
Table S5 shows the list of all diatom taxa and the count data

Shimada et al. Progress in Earth and Planetary Science(2023) 10:49

composed mainly of freshwater (D. stelligera) and freshwater–brackish (C. meneghiniana) plankton.
The assemblages in Zones 5–9 consist of freshwater
species. In most assemblages classified in Zone 5 (163.5–
250.5 cm depth), including the ones within E6, tychoplankton (S. binodis, S. construens, and S. venter) display
high relative abundance. Zone 6 (250.5–266.5 cm depth)
is dominated by freshwater plankton (Aulacoseira spp.,
D. stelligera, and Lindavia praetermissa). Tabellaria
fenestrata and G. exilis are also recognized in this zone.
Zone 7 (266.5–311.3 cm depth) is characterized by a mixture of the freshwater planktic species D. stelligera andthe freshwater benthic species T. fenestrata. The assemblages within E8 are similar in composition to those in
the organic mud samples of Zone 7. Assemblages within
E9 and E10 were similar to those in the organic mud layerof zone 7 except for a slightly smaller relative abundanceof D. stelligera. It should be noted that subsamples for
E9, E10, and underlying and overlying layers were hard
to be separated completely due to faint distributions of
the event layers. Zone 8 (311.3–402.4 cm depth) andZone 9 (402.4–428.3 cm depth) are dominated by freshwater benthos such as Encyonema spp., Eunotia spp., andT. fenestrata. The assemblages within E14 consist of thesame taxa as those in Zone 8. In Zone 9, plankton exhibita very low relative abundance. The division between
Zone 8 and Zone 9 is identified within E15. Althoughboth assemblages within E15 are dominated by freshwater benthos, the assemblages in Zone 8 contain a higher
relative abundance of T. fenestrata and Eunotia spp. thanthose in Zone 9. This result is consistent with the character of the assemblages in the organic mud of each zone.
The assemblages within E16 display a similar composition to those in the organic mud samples of Zone 9.
The assemblages in Zone 10 (428.3–475 cm depth) consist of a mixture of freshwater (Lindavia praetermissa)
and freshwater–brackish (C. meneghiniana) plankton
and freshwater benthos such as Achnanthidium minutissimum. The brackish benthic species Navicula peregrinais also present. The assemblages within E17–E19 have thesame composition.
Diatom assemblages within the other 7 event deposits (E2, E4, E5, E7, and E11–E13) were not prepared for
microscopic observation because it was not possible tosample those deposits separately from the inorganic mud
above and below.

5 Discussion
5.1 Environmental history at the study site during the late
Holocene

The fossil diatom assemblages reveal the paleoenvironmental changes recorded in Lake Kogare-ike sediments during the last 3000 years. A freshwater–brackish

Page 17 of 28

lagoonal environment is indicated by the freshwater andfreshwater–brackish plankton in Zone 10 (Fig. 6). The
decrease in these plankton and the increase in freshwater
benthic taxa recognized at the boundary between Zones
10 and 9 show that the environment changed to a freshwater swamp/marsh that was isolated from the sea. The
swampy environment lasted until 1300–1400 cal BP, theboundary between Zones 8 and 7. Plankton reappeared
during deposition of the Zone 7 sediments, and the site
consequently became a standing freshwater environment, as indicated by the dominance of freshwater plankton and tychoplankton in Zones 6–3. This environment
continued until slightly before the deposition of E1.
The upward change in diatom assemblages from samples just below E1 to the top may represent eutrophication at this site. One of the most striking changes indiatom assemblages is the appearance of freshwater–brackish plankton (C. atomus, C. atomus var. gracilis, andC. meneghiniana). This appearance may simply reflect
the increasing depth of the lake, but similar patterns havebeen reported in hypereutrophic conditions (Wolin andDuthie 1999; Wolin and Stone 2010). Nutrient increases
caused by agricultural activity encourage phytoplankton growth, resulting in light limitation at the bottom of
the lake. The light limitation disturbs growth of epipelic
diatoms in the surface sediment and also macrophyte
growth; these plants act as a substrate for epiphytic diatoms at the bottom of the lake. Consequently, the productivity of benthic (epiphytic and epipelic) diatoms at
a eutrophic site sometimes approaches zero as plankton
productivity and light limitation increase (Wolin andDuthie 1999; Wolin and Stone 2010). Agricultural activity around our study site can be recognized in the land
use map for 1892, and aerial photographs from different times show that the paddy fields and farmland wereabandoned around 2003 (Additional file 2: Fig. S1). The
recent change in land use is consistent with the hypothesis of eutrophication due to human activity.
5.2 Age–depth model and depositional ages

A Bayesian age–depth model was constructed by using
Bchron (Haslett and Parnell 2008; Parnell et al. 2008)
based on the stratigraphy and radiocarbon ages and thepeak of 137Cs concentration from the core at K-4 (Fig. 7).
The highest concentration of 137Cs at 39.5 cm depth
below the lake bottom corresponded to the global fallout
peak in 1963 CE (Ritchie and McHenry 1990; Appleby
2001). The script is shown in Additional file 11: Text S1.
The age–depth model provides the depositional ages of
event deposits E1–E16 with 95% credible intervals (Fig. 7;
Table 2; Additional file 12: Table S6). The deepest three
event deposits E17–E19 were not applied to the model
because we were unable to constrain their maximum

Shimada et al. Progress in Earth and Planetary Science(2023) 10:49

Page 18 of 28

Fig. 7 Age–depth model reconstructed by using Bchron (Haslett and Parnell 2008; Parnell et al. 2008) for Bayesian chronological analysis. Eventdeposits were treated as instantaneous deposits, and their thicknesses were removed from the original core log of K-4. The scale on the y-axis showsthe event-free depth. We constructed the model three times, and the averages of the three trials were used as the depositional ages. The resultsof each trial are provided in Additional file 12: Table S6

Shimada et al. Progress in Earth and Planetary Science(2023) 10:49

Page 19 of 28

Table 2 Depositional ages of the event deposits (E1–E16) estimated by the Bayesian age-depth modelCore location

Name of event
deposits

Corrected depth below thelake bottom (cm)Calibrated depositional age (95%credible interval, cal BP)

Calibrated depositional age(95% credible interval, cal CE/
BCE)

K-4

E1

46.9

100-present

1850–1960 CE

K-4

E2

109.8

330–164

1620–1786 CE

K-4

E3

121.7

394–277

1556–1673 CE

K-4

E4

138.4

430–338

1520–1612 CE

K-4

E5

148.5

535–4121415–1538 CE

K-4

E6

171.5

818–675

1132–1275 CE
1079–1246 CE

K-4

E7

182.7

871–704

K-4

E8

220.5

1243–1070

707–880 CE

K-4

E9

230.3

1273–1154

677–796 CE

K-4

E10

249.5

1354–1279

596–671 CE

K-4

E11

258.6

1396–1320

554–630 CE

K-4

E12

277.9

1503–1367

447–583 CE

K-4

E13

286

1558–1399

392–551 CE

K-4

E14

287.7

1569–1404

381–546 CE

K-4

E15

321.4

2267–1976

317–26 BCE

K-4

E16

331.6

2615–2385

665–435 BCE

The model was consructed using Bchron (Haslett and Parnell 2008; Parnell et al. 2008). Event deposits were treated as instantaneous deposits, and their thickness
were removed from the original core log of K-4. “Corrected depth below the lake bottom” represents the event-free depth

ages. We note that our model has uncertainties, in thatit does not take erosion and redeposition prior to formation of event deposits into consideration.
The onset of 137Cs also has the potential to constrain
depositional ages, but was not applied to the age–depth
model in this study for the following reasons. First, theonset of 137Cs is regarded generally as corresponding
to 1954 CE globally, when atmospheric nuclear weapons testing started (Pennington et al. 1973); however, insouthwestern Japan, 137Cs may have begun to accumulate
in 1945 CE when the two atomic bombs were dropped
on Hiroshima and Nagasaki (Saito-Kokubu et al. 2008).
Second, the depth of the onset may be uncertain because
cesium can migrate downward (e.g., Schimmack et al.
1997; Matisoff et al. 2011; Fujii et al. 2014), for example,
via pore water (Takada et al. 2017).
5.3 Origin of event deposits
5.3.1 Tsunamis and storms

Event deposits formed by extreme waves are generally
identified as relatively coarse sediment layers (sand andgravel) within fine deposits (mud and peat) in coastal
lowlands (e.g., Atwater and Moore 1992; Donnelly andWoodruff 2007; Jankaew et al. 2008; Nelson et al. 2008;
Dura et al. 2015). Tsunami geologists have extensively
studied event deposits to differentiate tsunami fromother event deposits, such as those formed by storms
and flash flooding (e.g., Costa and Andrade 2020). Coarse
deposits can be recognized as tsunami deposits by meansof multiple sedimentological, paleontological, and geochemical techniques; for example, the sedimentaryfeatures of event deposits have been compared with features observed in recent tsunami deposits (e.g., the 2004
Indian Ocean tsunami: Moore et al. 2006 and Matsumoto
et al. 2010; the 2011 Tohoku-oki tsunami: Goto et al.
2011 and Naruse et al. 2012). The presence of marineorganisms and chemical indicators within event deposits
has also been used to constrain the origin of those deposits (e.g., shells: Moore et al. 2006 and Donato et al. 2008;
foraminifera: Nanayama and Shigeno 2006 and Pilarczyk
et al. 2012, 2021; diatoms: Dawson 2007 and Sawai et al.
2009b, 2012; salt components: Szczuciński et al. 2005
and Chagué-Goff et al. 2012; biomarkers: Shinozaki et al.
2015 and Bellanova et al. 2020). We applied some of thesetypes of evidence to infer the origin of the event deposits,as discussed below, although not all previously reported
criteria are applicable to our data.
5.3.1.1 Sedimentological features of the event deposits Some sedimentological features are common inpreviously reported modern tsunami and storm deposits. Normal and inverse grading of grain size has beenreported in both tsunami (e.g., Paris et al. 2007; Choowong
et al. 2008a, b; Richmond et al. 2012; Matsumoto et al.
2023) and storm deposits (e.g., Wang and Horwitz 2007;
Williams 2009). Rip-up clasts within event deposits arealso common (e.g., Kortekaas and Dawson 2007; Goto
et al. 2011; Putra et al. 2013), suggesting erosion of surface

Shimada et al. Progress in Earth and Planetary Science(2023) 10:49

sediments and its redeposition by high-velocity currents.
In this study, changes in sand contents and the results ofgrain size analysis show that six event deposits at K-4 (E3,
E6, E15, and E17–E19) exhibit either normal or inverse
grading. Event deposit E8 includes rip-up clasts formed bylight grayish inorganic mud at its base. Because the composition of the rip-up clasts (gray inorganic mud) is different from that of the contemporary lake bottom (brownorganic mud), they may have been transported via muddy
environments such as a channel and the bottom of the sea.
5.3.1.2 Spatial bias of the event deposits As has been
demonstrated for recent tsunami and storm deposits oncoastal lowlands, the spatial distributions of marine-originated event deposits are limited to seaward locations (Liuand Fearn 2000; Goff et al. 2004; Nakamura et al. 2012;
Phantuwongraj and Choowong 2012; Pilarczyk et al.
2016) and the vicinity of inlets and barriers in coastal lakes(Matsumoto et al. 2010). This limitation is simply due tothe decreased sediment transport capacity of tsunami orstorm waves as they move away from the contemporary
shoreline. At our study site, 13 of the 19 event deposits(E1–E10, E12–E14) occurred near the barrier (cores K-1,
K-2, and K-3) and around the center of the lake (cores K-4,
K-6, and K-7) (Figs. 2 and 3). The most inland core (K-5)
included fewer event deposits than the other cores andcontained event E1 and possibly E8 in samples younger
than 1523–1383 cal BP (Fig. 2 and Table 1). The spatial
distributions of the other events (E2–E7 and E9–E14)
were not estimated, mainly because of the lack of cores
between K-5 and the center of the lake, but they certainly
did not have extensive distributions and were limited toaround the center of the lake, which might suggest thatthe sediment was transported from the seaward direction.
5.3.1.3 Diatom
assemblages Diatom
assemblages
within event deposits can be applied to infer the originof the deposits because diatom species segregate according to their autecology and preferred substrate (HemphillHaley 1996; Dura and Hemphill-Haley 2020). The diatom
assemblages within E1 contained higher relative abundances of brackish benthic species (Navicula gregaria) andfreshwater–brackish planktonic taxa (Cyclotella atomus,
Cyclotella atomus var. gracilis, and C. meneghiniana) thanthe muddy organic sediments below and above the event
deposit. E3 also contained freshwater–brackish plankton (C. meneghiniana) within the event deposit, whereas
samples below/above E3 were dominated by freshwaterplankton (e.g., Aulacoseira spp., Discostella stelligera,and Lindavia praetermissa) and freshwater benthos (e.g.,
Tabellaria fenestrata). The diatom assemblages within
E17–E19 consisted of freshwater, freshwater–brackish,and brackish species. These assemblages suggest that the

Page 20 of 28

sediments forming these three event deposits were transported by water from marine or brackish environments.
5.3.1.4 Simultaneity of event deposits and historical tsunamis and storms A common method to identify tsunami and storm deposits is to compare the constrained
ages of the event deposits with those of known earthquakes and storms from documented history. When thedepositional age of an event deposit corresponds to theage of a historical event, it has been accepted that theevent deposit is a tsunami deposit associated with theearthquake or a storm deposit, such as in the following
studies: the 1960 CE Chilean tsunami (Kempf et al. 2015);
the 1964 CE Alaska tsunami (Shennan et al. 2014); the1700 CE Cascadia tsunami (Atwater et al. 2005 as “TheOrphan Tsunami of 1700”); and 13th-century typhoons in
Kyushu (Woodruff et al. 2009).
The depositional ages of seven event deposits (E1–E5,
E7, and E9) cover the range of historical records of tsunamis along the Nankai Trough. The depositional age of E1(100 cal yr BP to the present; 1850–1960 CE) covers the1944 Showa-Tonankai and the 1854 Ansei–Tokai tsunamis. The depositional age of E2 (330–164 cal BP; 1620–
1786 CE) is consistent with the 1707 CE Hoei tsunami.
The depositional ages of E3 (394–277 cal BP; 1556–1673
CE) and E4 (430–338 cal BP; 1520–1612 CE) cover the1605 CE Keicho tsunami. The depositional ages of E5(535–412 cal BP; 1415–1538 CE), E7 (871–704 cal BP;
1079–1246 CE), and E9 (1273–1154 cal BP; 677–796 CE)
are consistent with the 1498 CE Meio, the 1096 CE Eicho(Kaho), and the 684 CE Hakuho earthquake, respectively.
The written record describes at least 55 storms that
attacked areas around Ise Bay (Fig. 1d) between 1600 and
1865 CE (Arakawa et al. 1961; Additional file 1: Table S1).
In addition, recent instrumental records show that 29
storms struck coastal areas of Mie Prefecture between
1945 and 2017 CE (Japan Meteorological Agency 2017).
As for the earthquake history, the depositional ages of E1,
E2, and E3 overlap the range of historical records of thesenumerous storms. Especially, the age of E1 (1850–1960
CE) covers the 1959 CE Isewan typhoon, which caused
inundation and sea-level rise around Ise Bay including
Minami-ise Town (Japan Meteorological Agency 1961;
Editing Committee of the History of Nanto Town 1985).
5.3.1.5 Difference in frequency between storms and event
deposit formation Storms occur at a much greater frequency than the event deposits around the study site. As
mentioned above, the written record describes at least 55
storms between 1600 and 1865 CE (Arakawa et al. 1961)
and 29 storms between 1945 and 2017 CE (Japan Meteorological Agency 2017). In contrast, the lacustrine sediments in this study record just 19 event deposits in theShimada et al. Progress in Earth and Planetary Science(2023) 10:49

last 3000 years. The frequency with which event deposits
formed at the study site is thus much lower than the frequency of recorded storms. We therefore consider thatthe storms that frequently occur in the study area do notleave event deposits in the area every time, and that tsunamis and extraordinary storms are candidates for thecauses of the event deposits. We could not estimate theprecise intensity, inundation area, or damage of storms
in historical records because of the fragmentary available information, but we are confident that the 1959 CE
Isewan Typhoon is the most likely to be an extraordinarystorm that had the potential to produce an event depositon the lake floor at the study site.
Considering the discussion above, we attribute 13 event
deposits (E1–E10 and E12–14) to marine inundation
caused by tsunamis or extraordinary storms. Five of the
13 (E2, either E3 or E4, E5, E7, and E9), the ages of which
are consistent with historical records of tsunamis, could
be correlated with tsunamis associated with earthquakes
along the Nankai Trough. The depositional age of E1(100 cal yr BP to the present; 1850–1960 CE) covers the1854 Ansei–Tokai tsunami, the 1944 CE Showa-Tonankai
tsunami, and the 1959 CE Isewan typhoon. Our chronological data cannot distinguish these three events. We
therefore conclude that E1 might be attributable to any of
the three events, or to a combination of two or all three
events. Furthermore, we are currently unable to definitively correlate event E8 with historical records. The 95%
credible interval of the age of event E8 (1243–1070 cal yr
BP; 707–880 CE) does not cover the 887 CE Ninna earthquake but is quite close. Considering analytical and statistical uncertainties of radiocarbon dating (Scott et al.
2007; Reimer et al. 2013), the possibility that E8 corresponds to the 887 CE event cannot be completely ruledout. For example, additional radiocarbon ages and subsequent reconstruction of the Bayesian age–depth modelin future work may show an age overlap of E8 and the
Ninna earthquake.
Either of E3 or E4 (one of which is not a tsunami
deposit associated with the 1605 CE Keicho earthquake)and E6 are likely to have been deposited by extraordinarystorm surges or unknown tsunamis. These event deposits
are not correlated with any historical earthquakes, eventhough their depositional ages belong to the historicalperiod. For E12–E14, it is unclear whether their origins
are tsunamis or extraordinary storms because criteria(iii) (diatom assemblages) and (iv) (simultaneity of event
deposits and historical tsunamis) cannot be applied.
5.3.2 Other causes

The volcanic ash layer (E11) probably resulted from
erosion and reposition by a debris flow, a flash flood,or another temporary inflow. The distribution of E11

Page 21 of 28

is limited to the central and landward parts of the lake
(Figs. 2 and 3), implying that it was not formed by marine
inundation. EDX analysis shows that E11 contains volcanic glass shards with different sources and eruption ages: glass shards from the Aira-Tn eruption (AT;
29–26 ka cal yr BP), the Kikai-Akahoya eruption (K-Ah;
7.3 ka cal BP), and the Amagi-Kawagodaira eruption (Kg;
3.1 ka cal yr BP), and shards similar in composition to
those from eruptions in Niijima and Kozushima such as
Iz-Kt (AD 838). This mixture of material indicates that
E11 is not a single fallout volcanic ash, but resulted from
redeposition. Therefore, we assume that older sediments
surrounding Lake Kogare-ike (Fig. 1e) were eroded byintense rainfall and transported to our study site.
The origins of the five oldest event deposits (E15–E19)
are also ambiguous in this study. These deposits were
found only at K-4 and could not be laterally correlated
because of the limited recovery of other cores. Because
the diatom assemblages within E17–E19 contained freshwater–brackish plankton and slightly brackish benthos,these three may have been deposited by a marine-originated event, such as a tsunami or a storm. As indicatedby the presence of freshwater–brackish and brackish
diatom species in samples below and above E17–E19,
the study site was more influenced by marine water during deposition of those units than at the present day,and might represent the process of change to a lagoonalenvironment.
5.4 Correlation with written documents and existinggeological records
5.4.1 Historical earthquakes and tsunamis

The Bayesian age–depth model allows us to correlate
the event deposits with historical tsunamis over the last
300 years that inundated our study site. Lake Kogareike records the 1707 Hoei tsunami (E2; 1620–1786 CE;
Table 2) and either the 1944 Showa tsunami or the 1854
CE Ansei–Tokai, or a combination of both tsunamis (E1;
1850–1960 CE; Table 2). As mentioned above, we are notable to distinguish between 1944 and 1854 for the originof the event layer. However, if E1 had been generated byonly one of those events, our findings would indicate thatthe other tsunami did not leave a tsunami deposit on thelake bottom at our study site. Or, the geological record
might have been erased as a result of bioturbation or
anthropogenic disturbance, but we could not recognize
clear evidence of either of these possibilities.
Either event deposit E3 (1556–1673 CE) or E4 (1520–
1612 CE) may be new geological evidence for the 1605
CE Keicho earthquake. According to historical documents, a tsunami in 1605 affected Ise city, about 30 km
northeast of our study site (Hatori 1975; Iida 1981). Thistsunami damaged the entire coast along the Nankai

Shimada et al. Progress in Earth and Planetary Science(2023) 10:49

Trough (Hatori 1975; Iida 1981; Murakami et al. 1996),
but its parent earthquake has remained uncertain. Somestudies have concluded that the event was a tsunami
earthquake (Yamamoto and Hagiwara 1995; Seno 2002),
whereas others have suggested an earthquake off thesouthern part of the Kanto region (Tsuji 2016; Fig. 1d)or an earthquake in the Izu-Bonin trench (Ishibashi andHarada 2013; Fig. 1a). Further discussion of the rupturein 1605 requires detailed geological as well as historicalresearch.
For historical tsunamis older than 1600 CE, the presence and absence of tsunami deposits at our study site
are consistent with previously reported geologicalrecords in the Tonankai and Tokai regions, apart fromthe 887 CE Ninna tsunami. Event deposits E5 (1415–
1538 CE; Table 2), E7 (1079–1246 CE), and E9 (677–796
CE) are consistent with the 1498 CE Meio tsunami, the1096 CE Eicho tsunami, and the 684 Hakuho tsunami,
respectively, and are correlated with tsunami deposits
in the coastal lowlands at Shijima (Fujino et al. 2018;
Fig. 1b) and Iwata (Fujiwara et al. 2020b; Fig. 1d). The
absence of a 1361 tsunami deposit is also consistent withprevious reports for the Tonankai and Tokai regions
by Fujino et al. (2018) and Fujiwara et al. (2020b),
respectively.
The predicted event age of E8 (707–880 CE) does notcover but is very close to the 887 Ninna earthquake. In
previous work near our study site, an absence of tsunami
deposits consistent with this earthquake was reported
by Fujino et al. (2018) in the Shijima lowland (Fig. 1b),whereas liquefaction features and tsunami deposits probably correlated with the Ninna earthquake have beenreported in the Tokai region (Sangawa 2013; Fujiwara
et al. 2020a). Historical records of shaking and inundation associated with this earthquake have been mainly
reported in Osaka and Kyoto (Fig. 1d), and damage in the
Tokai and Tonankai segments is not definite (Ishibashi
1999, 2004).
5.4.2 Prehistoric earthquakes and tsunamis

Our prehistoric event deposits (E10 and E12–E16) areconsistent with the geological records of tsunamis documented in previous studies. The depositional age of E10(1354–1279 cal BP; Table 2) is consistent with the ages oftsunami deposits Sand C (1360–1080 cal yr BP; 590–870
CE) and Sand D (1400–1300 cal BP) in the Shijima lowland (Fujino et al. 2018; Fig. 1b). The age of Sand D also
overlaps the age of our E12 (1503–1367 cal BP; Table 2).
E12 (1503–1367 cal BP; Table 2), E13 (1558–1399 cal BP),and E14 (1569–1404 cal BP) can be correlated with theage of tsunami deposit OS-1 (1563–1454 cal BP, constrained event age recalibrated in this study with Oxcal

Page 22 of 28

4.4 (Bronk Ramsey 2009) with the IntCal20 radiocarbon
age calibration curve (Reimer et al. 2020) in the Osatsu
lowland (Okahashi et al. 2005; Fig. 1b). The age of fourthmost recent sand layer in Tsuji et al. (2002) was estimatedby recalibration and modeling by Garrett et al. (2016) as
1520–1260 cal BP. This age overlaps the depositional agesof our E10 and E12–E14. The depositional ages of E15(2267–1976 cal BP) and E16 (2615–2385 cal BP) overlap
with those of a tsunami deposit (Sand F; 2660–2190 cal
BP) in the Shijima lowland (Fujino et al. 2018) and withanother tsunami deposit (2500–2000 cal BP) in coastallake Sugari-oike, 30 km southwest of our study site (Okamura and Matsuoka 2012).
E15 is consistent with both tsunami deposits and withcoastal uplift inferred from emergence of sessile assemblages at 2200 cal BP on the southern coast of the Kii
Peninsula (Shishikura et al. 2008). We do not discuss correlations between E17–E19 and previous studies because
their depositional ages are not constrained because of alack of limiting maximum ages.
It should be noted that our age–depth model still has
difficulty in statistically distinguishing the ages of adjacent events between E10 and E14. The 95% credibleintervals of the ages of the events exhibit great overlap(Fig. 7), probably because too many extreme wave events(tsunamis and extraordinary storms) occurred during
a short period of time relative to other intervals. Morestudies in adjacent areas and clearer correlations withthose areas may lead to a better chronological reconstruction of marine inundations. We also note that thepossibility that extraordinary storms are the origin of ourprehistoric event deposits remains. At this stage, becauseof the large uncertainty of constrained event ages in bothprevious studies and this work, and the existence of only
a few case studies, we could not establish definitive correlation of event deposits among sites or estimate whether
each event represented extensive inundation.

6 Conclusions
Geological evidence for tsunami inundation along the
Nankai Trough has been discussed during the last few
decades; however, only a few studies have establishedwell-constrained chronologies for deposits, and individual chronologies have covered only limited periods; forexample, they lack records for the historical age becauseof anthropogenic modification. We obtained continuous
sediment cores recording the period from 3000 years ago
to the present and reconstructed the history of extreme
waves by integrating X-ray CT observation and Bayesianage–depth modeling.
This study revealed the marine inundation history
during the last 3000 years at Lake Kogare-ike, a coastal

Shimada et al. Progress in Earth and Planetary Science(2023) 10:49

lake in the Pacific coast of central Japan. Observationby the naked eye and X-ray CT images showed that thelacustrine sediments record 19 event deposits (E1–E19)
interbedded with organic muddy sediments. Thirteen of
the 19 event deposits (E1–E10 and E12–E14) were confirmed to occur near the barrier and around the central
part of the lake in the study site. E1 and E8 were possibly recognized at all sampling locations. The diatom
assemblages within E1 consisted of brackish and freshwater–brackish species, and those within E3 contained
freshwater–brackish species. The frequency of inundation events in this study was much lower than the frequency of storms known from the historical records.
On the basis mainly of changes in grain size, presenceof rip-up clasts, distributions of event deposits, diatom
assemblages, and information from historical records,
we conclude that 13 of the 19 event deposits (E1–E10
and E12–E14) were formed by tsunamis or extraordinary storms. Bayesian age–depth models indicated thatthe five event deposits (E2, either of E3 or E4, E5, E7,
and E9) are likely to be correlated with tsunami inundations associated with historical earthquakes along
the Nankai Trough. E2 (1620–1786 CE), either of E3(1556–1673 CE) or E4 (1520–1612 CE), E5 (1415–1538
CE), E7 (1079–1246 CE), and E9 (677–796 CE) are correlated with the 1707 CE Hoei, the 1605 CE Keicho,
the 1498 CE Meio, the 1096 CE Eicho (Kaho), and the
684 CE Hakuho tsunamis, respectively. E1 (1850–1960
CE) may be a geological record of the 1944 CE ShowaTonankai tsunami, the 1854 Ansei–Tokai tsunami, the1959 Isewan typhoon, or a combination of two or all
three events. There are no event deposits consistent withthe 1361 CE Shohei tsunami at our study site. Despite
such chronological correlation, it still remains possibility that some event deposits were formed by extraordinary storms. To solve this problem, it may be necessary
to carry out inter-site (within a few tens of kilometers)
and inter-regional (several tens of kilometers apart)
correlations using a similar approach. Because storms
generally affect only the specific area over which they
pass, whereas a tsunami inundates the coastline along
a distance of more than a hundred kilometers, storm
and tsunami deposits may be distinguished by careful
identification of event deposits and chronology. In thisprocess, reexamination of event deposits and their chronology (including selection of appropriate materials anda Bayesian age–depth model) at other sites will contribute to significant progress in regional correlation andreconstructing rupture length for particular earthquake
events, leading to reconstruction of the full history ofcoastal hazards along the Nankai Trough.

Page 23 of 28

Abbreviations
BCE	Before common era
cal BP	Calibrated years before the present
CE	Common Era
CT	Computed tomography
T.P.	Tokyo Peil

Supplementary Information
The online version contains supplementary material available at https://​doi.​
org/​10.​1186/​s40645-​023-​00577-9.
Additional file 1: Table S1. A list of storms and flooding around Ise bay,reported in historical written documents.
Additional file 2: Figure S1. Aerial photographs and topographic maps
showing topographic and land-use changes around Lake Kogare-ike. (a)Aerial photograph taken by the United States Army on 23 September
1947 (photograph number: USA-M500-75) (b) Aerial photograph taken by
the Geospatial Information Authority of Japan on 16 May 1967 (photograph number: KK-676Y-C9-3). (c) Aerial photograph taken by the Geospatial Information Authority of Japan on 23 October 1980 (photographnumber: KK-802X-C9-4). (d) Aerial photograph taken by the Geospatial
Information Authority of Japan on 30 April 1990 (photograph number:
KK-902X-C14-16). (e) Aerial photograph taken by the Geospatial Information Authority of Japan on 24 September 2008 (photograph number: CKK2008-1X-C13A-13). (f ) Topographic map at 1:20000 scale surveyed in 1892
and published in 1894 by the Land Survey Department of the Empire ofJapan (map number: s988, map name: Ugura mura). (g) Topographic map
at 1:50000 scale surveyed in 1892, modified in 1929, resurveyed in 1944,
repaired in 1949, and published in 1952 by the Geographical Survey Institute (map number: 1997-7-4, map name: Nie ura). (h) Topographic map at
1:25000 scale surveyed in 1967 and published in 1969 by the Geospatial
Information Authority of Japan (map number: 97-7-3-1, map name: Nieura). (i) Topographic map at 1:25000 scale surveyed in 1967, modified
based on an aerial photograph taken in 1980, correctively surveyed in
1982, and published in 1984 by the Geospatial Information Authority ofJapan (map number: 97-7-3-3, map name: Nie ura). (j) Topographic map at
1:25000 scale surveyed in 1967, modified based on an aerial photographtaken in 1988, correctively surveyed in 1990, and published in 1991 by the
Geospatial Information Authority of Japan (map number: 97-7-3-4B, map
name: Nie ura).
Additional file 3: Figure S2. Photographs and X-ray CT images of thecore sediments at K-4. The cores were mainly collected by using a Russian
sampler. The uppermost sediments were obtained by push-core sampling.
Additional file 4: Figure S3. X-ray CT images of core sediments at K-1,
K-2, K-3, K-5, K-6, and K-7. The cores were mainly collected by using a Russian sampler. The cores marked “*push-core” were obtained by push-coresampling.
Additional file 5: Figure S4. Photographs, X-ray CT images, and sketchesof the event deposits identified at K-4 with the results of grain size analysis(sand content and grain size distribution of residual sand fractions).
Additional file 6: Table S2. Sedimentary features of event deposits(E1–E19).
Additional file 7: Table S3. Mud content, organic content, sand conent,
and results of grain size analysis for redidual sand fractions.
Additional file 8: Figure S5. Scatter plots of MgO–SiO2, CaO–SiO2,­Na2O–SiO2, and ­K2O–SiO2 of volcanic glass shards in event deposit
E11 at sampling locations (a) K-3, (b) K-4, (c) K-6, and (d) K-7. Chemicalcompositions of the following tephras found in previous studies are also
shown: AT and K-Ah (Machida and Arai 2003); Kg (Sugihara 1984); Iz-Kt
and Nj-My (Sugihara 1984; Sugiuchi and Fukuoka 2005; Tsukui et al. 2006;
Kobayashi et al. 2020; Murata et al. 2021); Kz-An, Kz-CbA and Kz-CbA’(Murata et al. 2021). Abbreviations: At, Aira-Tn tephra ; K-Ah, Kikai-Akahoya
tephra; Kg, Amagi-Kawagodaira tephra; Iz-Kt, Kozushima-Tenjosan tephra;
Nj-My, Niijima-Mukaiyama tephra; Kz-An, Kozushima-Ananoyama tephra;Shimada et al. Progress in Earth and Planetary Science(2023) 10:49

Kz-CbA′, Kozushima-Chichibu-yama-A′ tephra; Kz-CbA, KozushimaChichibu-yama-A tephra.
Additional file 9: Table S4. Vertical change of 137Cs concentrations at K-4.
Additional file 10: Table S5. Count data of fossil diatom analysis at K-4.
Additional file 11: Text S1. The R script for construction of an age–depth
model using the “Bchron” library (Haslett and Parnell 2008; Parnell et al.
2008).
Additional file 12: Table S6. The depositional ages of event deposits (E1–
E16) estimated in the three trials.
Acknowledgements
We acknowledge the landowners and the staff of the government officeat Minami-ise town and the Ise-Shima National Park Management Office,Ministry of the Environment Government of Japan, who gave us permission
to collect sediments at Lake Kogare-ike. We thank Andrew Parnell for his kindadvice on constructing age–depth models with Bchron. The manuscript was
improved by reviews from Philipp Kempf and two anonymous reviewers.
Author contributions
YSh and YSa proposed the topic and conceived and designed the study. YSh,
YSa, DM, KT, YN, and MS collected sediment cores in the field. YN carried out a
topographic survey. YSh picked out plant macrofossils for radiocarbon dating.
YSa and YSh identified diatoms, and YSh counted diatoms. KI, YSh, and TT
carried out the 137Cs analysis. YSh wrote the initial manuscript, and YSa andTT produced the second version. YSh, YSa, DM, KT, KI, TT, YN, MS, and SF readand approved the final manuscript. A part of data in this study were taken as
master’s research of YSh. SF supervised YSh during the master’s course.
Funding
This work was financially supported by the Geological Survey of Japan and
the Ministry of Education, Culture, Sports and Technology of Japan. The latteris part of the Research Project for Compound Disaster Mitigation on the
Great Earthquakes and Tsunamis around the Nankai Trough Region and the
Research Project for Disaster Prevention on the great Earthquakes along the
Nankai trough.
Availability of data and materials
All data integral to the stated conclusions are presented within the paper andAdditional files (Figures S1–S5, Tables S1–S6, and Text S1).

Declarations
Competing interests
The authors declare that they have no competing interest.
Received: 18 August 2022 Accepted: 26 July 2023

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