Inoue et al. Progress in Earth and Planetary Science (2017) 4:42
DOI 10.1186/s40645-017-0158-1

RESEARCH ARTICLE

Progress in Earth andPlanetary Science

Open Access

Paleo-tsunami history along the northern
Japan Trench: evidence from Noda Village,
northern Sanriku coast, JapanTaiga Inoue1*, Kazuhisa Goto2, Yuichi Nishimura3, Masashi Watanabe4, Yasutaka Iijima1,5 and Daisuke Sugawara2,6

Abstract
Throughout history, large tsunamis have frequently affected the Sanriku area of the Pacific coast of the Tohoku
region, Japan, which faces the Japan Trench. Although a few studies have examined paleo-tsunami deposits along theSanriku coast, additional studies of paleo-earthquakes and tsunamis are needed to improve our knowledge of thetiming, recurrence interval, and size of historical and pre-historic tsunamis. At Noda Village, in Iwate Prefecture on thenorthern Sanriku coast, we found at least four distinct gravelly sand layers based on correlation and chronological data.
Sedimentary features such as grain size and thickness suggest that extreme waves from the sea formed these layers.
Numerical modeling of storm waves further confirmed that even extremely large storm waves cannot account for the
distribution of the gravelly sand layers, suggesting that these deposits are highly likely to have formed by tsunami
waves. The numerical method of storm waves can be useful to identify sand layers as tsunami deposits if the deposits
are observed far inland or at high elevations. The depositional age of the youngest tsunami deposit is consistent with
the AD 869 Jogan earthquake tsunami, a possible predecessor of the AD 2011 Tohoku-oki tsunami. If this is the case,then the study site currently defines the possible northern extent of the AD 869 Jogan tsunami deposit, which is an
important step in improving the tsunami source model of the AD 869 Jogan tsunami. Our results suggest that fourlarge tsunamis struck the Noda site between 1100 and 2700 cal BP. The local tsunami sizes are comparable to the AD
2011 and AD 1896 Meiji Sanriku tsunamis, considering the landward extent of each tsunami deposit.
Keywords: Paleo-tsunami, Sanriku coast, Japan, Tsunami deposit identification, AD 869 Jogan tsunami, Storm wave,
Numerical modeling

Introduction
The Tohoku-oki earthquake (Mw = 9.0) and tsunami thatstruck on March 11, 2011, generated severe damage along
the Pacific coast of eastern Japan. The threat of large tsunamis is of great concern to Japan and other countries(e.g., Goto et al. 2014). Studies of tsunami deposits are important in estimating the size and recurrence interval of
historical and pre-historical tsunamis, and in assessing therisk of large low-frequency tsunamis (e.g., Goto et al.
2014). Many studies have described tsunami deposits
around the world (e.g., Dawson and Shi 2000; Pinegina
and Bourgeois 2001; Peterson et al. 2011).

* Correspondence: taiga.inoue.r7@dc.tohoku.ac.jp
1
Department of Earth Sciences, Graduate School of Science, Tohoku
University, Sendai 980-8578, Japan
Full list of author information is available at the end of the article

Studies of paleo-tsunami deposits along the Pacific coastof the Tohoku region (Fig. 1), which faces the Japan
Trench, are the result of global collaboration. The Sendai
and Ishinomaki plains, located in the southern part of theSanriku coast, are representative locations examined inthose studies and have yielded evidence for historical andpre-historical tsunamis (Minoura and Nakaya 1991;
Minoura et al. 2001; Sugawara et al. 2001, 2012, 2013;
Sawai et al. 2007, 2008, 2012, 2015; Shishikura et al. 2007;
Satake et al. 2008; Namegaya et al. 2010; Matsumoto et al.
2012). Shishikura et al. (2007), for example, identified fivepossible tsunami deposits that formed during the past
3000 years and estimated a tsunami recurrence interval of
500–1000 years. Sawai et al. (2015) proposed that thetsunami recurrence period in this region is ~ 600 years.
The Sendai and Ishinomaki plains also contain deposits of
the AD 869 Jogan tsunami (e.g., Abe et al. 1990; Minoura© The Author(s). 2017 Open Access This article is distributed under the terms of the Creative Commons Attribution 4.0
International License (http://creativecommons.org/licenses/by/4.0/), which permits unrestricted use, distribution, andreproduction in any medium, provided you give appropriate credit to the original author(s) and the source, provide a link to
the Creative Commons license, and indicate if changes were made.

Inoue et al. Progress in Earth and Planetary Science (2017) 4:42

Fig. 1 Location of the study area. Dark gray area denotes the Tohoku
region. Tsunami source models of the 869 Jogan (Sawai et al. 2012),
1611 Keicho (Hatori 2009), 1896 Meiji Sanriku (Tanioka and Satake 1996;
Tanioka and Seno 2001), 1933 Sanriku (Kanamori 1971), and 2011
Tohoku-oki earthquakes (Ozawa et al. 2011) are shown by colored lines
off the Sanriku coast. The broken lines indicate that the extent of the
rupture model remains unknown. Thickness distributions (cm) of the
Towada-a (To-a) tephra of AD 915 (Machida and Arai 2003) and theBaegdusan-Tomakomai (B-Tm) tephra of AD 946 (Machida and Arai
1992) are shown by broken orange and purple lines, respectively

and Nakaya 1991; Sawai et al. 2012), which was a possible
predecessor of the 2011 Tohoku-oki tsunami (e.g., Goto etal. 2011; Namegaya and Satake 2014).
In contrast to the well-studied Sendai and Ishinomaki
regions, few geological studies have described tsunami
deposits along the Sanriku coast. Consequently, it is necessary to elucidate the long-term tsunami history along
the Sanriku coast from geological evidence. However,this region has the disadvantage of low sedimentary potential and a lack of lowlands necessary to form tsunami
deposits, due to the development of a ria coast. It has
also been intensely modified by agricultural development. Observation records and other historical documents indicate that large tsunamis have frequently
occurred in this region (e.g., Utsu 2004; Fig. 1). In fact,
within the past 400 years, the AD 1611, 1896, 1933, and2011 tsunamis have severely affected the Sanriku coast,with maximum recorded inundation heights of 25.0,
38.2, 29.2, and 33.0 m, respectively (Japan Tsunami
Trace database: http://tsunami-db.irides.tohoku.ac.jp).
Along the Sanriku coast, Goto et al. (2015) studied thenorth area of Taro (Fig. 1) and identified possible tsunami deposits from the last 600 years. Ishimura andMiyauchi (2015) and Ishimura (2017) reported 11

Page 2 of 15

tsunami deposits, including a possible AD 869 Jogan
tsunami deposit, from the past 4000 years at Koyadori(Fig. 1). Ishimura (2017) estimated that the average recurrence interval of these deposits is 350–390 years. Farther north, on the Shimokita Peninsula (Fig. 1),
Tanigawa et al. (2014) reconstructed the geological record of the past ~ 6000 years. Minoura et al. (1994, 2013)
and Tanigawa et al. (2014) reported evidence of tsunami
deposits that were possibly formed by the AD 1611
Keicho tsunami. Between Miyako and the Shimokita
Peninsula, only two studies, Hirakawa (2012) and Takada
et al. (2016), have described possible tsunami deposits. It
is necessary to collect more geological evidence along
the northern Sanriku coast to elucidate the size and recurrence interval of paleo-tsunamis that are generated
along the Japan Trench.
It is difficult to distinguish tsunami deposits fromstorm and flood deposits (e.g., Morton et al. 2007; Goffet al. 2012). Therefore, a multi-proxy approach is recommended (Goff et al. 2012). Although diatom analysis
might be useful to infer the origins of deposits (HenphillHaley 1996), it cannot distinguish tsunami deposits fromstorm deposits. Numerical methods can be used to distinguish deposits formed by tsunami and other extreme
wave events. For instance, the inundation distance ofstorm waves is known to be generally shorter than that oftsunami waves because of the difference in wavelength(e.g., Morton et al. 2007; Watanabe et al. 2017). Therefore,
in a lowland sedimentary environment, the distance thatevent deposits can be transported will differ for storms
and tsunamis. If the event layers were deposited far beyond the numerically estimated maximum limit of inundation by storm waves, then it can be inferred that thedeposits were formed by tsunami waves.
In this study, we conducted two field surveys at Noda
Village, Iwate Prefecture, on the northern Sanriku coast(Fig. 1). We coupled a sedimentological approach, to accurately identify the origin of event layers at the study
site, with applied numerical modeling of storm waves rather than tsunami modeling, which allowed us to test
whether the distribution of event deposits can be explained by large storm waves.

Study area
Noda Village is located on the northern Sanriku coast(Fig. 1), where the geology is dominated by the Upper
Cretaceous Kuji Group (Terui and Nagahama 1995).
The alluvial lowland developed along rivers and marine
terraces, corresponding to marine isotope stage (MIS)
1:14 ka across the Sanriku coast area (Koike andMachida 2001). In coastal areas, cliffs and shore reefs
extend seaward. The beach consists of sand and gravel,
which is a rare sedimentary environment on theInoue et al. Progress in Earth and Planetary Science (2017) 4:42

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northern Sanriku coast, where a ria coast is developed.
Paddy fields and urban areas are distributed in the lowland area.
The study site, in the Maita area, is a narrow valley
among forested hills (Fig. 2a). Present-day beach sediments consist of well-rounded gravel and sand (Fig. 2b).
Wave breakers are set about 50 m from the shoreline.
The Sanriku railway runs north–south at ~ 300 m inlandfrom the shoreline. These artificial objects were built
well before the 2011 Tohoku-oki tsunami. The Maita
River is ~ 5 m wide, and the banks are covered by revetments. The study site is used as a rice paddy field andcontains agricultural terraces (Fig. 3).
The historical tsunamis that have struck the Sanriku
coast are summarized by Ishimura and Miyauchi (2015).
Among historical tsunamis, it is known that the AD
1611, 1896, 1933, and 2011 tsunamis were particularly
large. In our study area specifically, the 1896 Meiji
Sanriku tsunami reached up to ~ 880 m inland and 16 m
in elevation, according to local people. No written records of large tsunamis are available before the 1896
event. The 2011 Tohoku Earthquake Tsunami Joint Survey (TTJS) Group reported that tsunami height was
22.84 m near the shoreline at this site and reached up to~ 720 m inland and 12 m in elevation (http://www.coastal.jp/ttjt/) (Fig. 3b). The inundation distances and runup heights of the 1896 and 2011 tsunamis cannot becompared directly because the wave breakers were notconstructed until after the 1896 tsunami.

A

141°49’ 20 E

141°49’ 40 E

N

The gently sloping topography and low-energy wave
environment of the study area promotes the preservation of tsunami deposits. In July 2011, we confirmed thatthe 2011 tsunami deposits were formed in the area, as
shown in Fig. 2c. The 2011 tsunami deposit consists ofgravel to very fine sand layers of 5–10 cm thickness. The
deposits show normal grading and lamination and contain shell fragments. The 2011 tsunami deposit was removed during the recovery process, several months after
this field survey.

Methods/experimental
Field survey

In November 2014 and March 2015, we conducted geological surveys and examined sediments in the coastal
lowland of the study area. We set a survey transect perpendicular to the shoreline, 400–700 m from the sea. To retrieve cores, we used a Handy Geoslicer sampling tool of
0.6 and 1 m long (Takada et al. 2002) and a peat sampler.
Coring sites were selected to evaluate the continuity ofevent deposits (NM1 to NM20 and NM-A to NM-L;
Fig. 3a). Samples for analyses were taken from the cores
using rectangular plastic cases (5 × 20 cm). In the field andlaboratory, we conducted lithological observations, madesketches, and took photographs of the cores. As for somecoring sites, also used a peat sampler that can obtain samples from greater depths than the ones by Handy Geoslicer,
we made a composite columnar section from both cores.

B

40°6’00 N

Fig.2C

bre
ve
Wa

ay
Sanriku railw

M
ai

ta

riv

er

ake

r

Fig.2B

0

.3B

Fig

15 (cm)C

Fig.3A40°5’30 N

0

100

200
(m)

Fig. 2 a Aerial photograph of the survey area at Noda Village (CTO20112-C23–22, provided by the Geospatial Information Authority of Japan). b
Sediments deposited on the nearby beach. c Typical 2011 Tohoku-oki tsunami deposits (photograph taken July 29, 2011)

Inoue et al. Progress in Earth and Planetary Science (2017) 4:42

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A

B

Fig. 3 a Locations of survey points NM1 to NM20 and NM-A to NM-L. b Cross-sectional topography along the transect shown in Fig. 2a. Maximum
run-up heights of the 2011 Tohoku-oki and 1896 Meiji Sanriku tsunamis are shown. Tsunami height at the coast of the 2011 Tohoku-oki tsunami isbased on the 2011 Tohoku Earthquake Tsunami Joint Survey (TTJS) Group

Laboratory analyses

We analyzed the grain size of the gravelly sand deposits
using a CAMSIZER (Verder Scientific) at Tsukuba University, Japan. This instrument provides particle size datafrom 30 μm to 30 mm (Yamada and Fujino 2013). Samples for the analysis were collected from the gravelly
sand layers at 1 cm intervals vertically, although somesamples were collected as bulk samples because they
were thin (0.5–1.5 cm in thickness).
We first dried samples and added 10% hydrogen peroxide to remove organic material. Then, ultrasoniccleaning was performed after drying. We removed the
mud component using a 63-μm sieve, as the target sediment is composed mostly of coarse sand to gravel. We
chose to investigate the sand and gravel grain sizes only.
After removing the mud component, we dried the samples and measured the weight. The grain size analysis
results show the volume frequency of the grain size
distribution.

We used radiocarbon dating to correlate event deposits among cores and to estimate the depositional agesof event layers. We performed age-dating on 20 bulk
peat sediments obtained below and above the gravelly
sand layers. Radiocarbon dating was conducted using accelerator mass spectroscopy (AMS) at Beta Analytic Inc.
in Miami, USA. Radiocarbon ages were calibrated to calendar years using the approach of Vogel et al. (1993)
and the calibration curve IntCal13 (Reimer et al. 2013).
We recognized and analyzed tephra layers in somecores, which are useful in determining the absolute age.
The tephra layers were sampled from cores at 600–
650 m from the shoreline. To ascertain the origin oftephra quantitatively, the major element compositions
were determined by energy dispersive X-ray spectrometer (EDS, JSM-5310; JEOL) at Hokkaido University,
Japan, following the method of Nakamura (2016). We
used TiO2 and K2O contents to identify the tephra,
employing the database of Quaternary marker tephra

Inoue et al. Progress in Earth and Planetary Science (2017) 4:42

provided by Aoki and Machida (2006) and data of Nakamura (2016).
Numerical method to identify tsunami deposits

We performed a numerical simulation of wave propagation and run-up using a combination of Delft-3D (Deltares2011) and SWAN (Booij et al. 1999) software. The numerical method followed Bricker et al. (2014) and Watanabe
et al. (2017). When using Delft-3D and SWAN, the hydrodynamic model repeatedly passes the water level andcurrent fields to the wave model, which then calculates
the wave field considering currents and storm surges(Bricker et al. 2014). Then, the wave model passes the radiation stress field back to the hydrodynamic model,
which calculates the breaking-wave-induced setup andnearshore currents (Bricker et al. 2014).
To ensure coverage of the survey area, the numerical
domain was 3-km long and 6-km wide. Topographic
data on a 5-m grid, provided by the Geospatial Information Authority of Japan, was used for the land area, andbathymetric data on a 4-m grid, from the Central Public
Disaster Prevention Council of Japan, was used for themarine area. We removed artificial embankments suchas wave breakers, the Sanriku Railway, and the artificial
topography created by reclamation and banking fromthe whole numerical domain as they are contemporary
features. The computational grid size was 10 m. Timesteps of 0.01 s were used to ensure computational stability. The propagation and inundation extent of storm
waves were simulated over a period of 4 h.
According to the measurement records of NOWPHAS(1970–2013) (Kawaguchi et al. 2015), the previous maximum significant wave observed at Kuji, located justnorth of the study area (Fig. 1), had a 9.53-m wave
height, with a 12.1 s period (recorded in 2010). Our
study site is located in the northern part of Japan, where
the impacts of typhoons are generally lower than in thesouthern subtropical region (Kawaguchi et al. 2015).
Other regions in southern Japan have been affected bymuch larger waves: the maximum recorded significant
wave heights and periods are 13.5 m and 15.8 s at Murotsu (33°16′18′′N, 134°08′50′′E) in 2004, 12.49 m and16.4 s at Kochi (33°2857′′N, 133°35′13′′E) in 2004,
14.44 m and 16.1 s at Omaezaki-oki (34°24′12′′N, 138°
16′30′′E) in 2009, and 15.14 m and 14.4 s at Owase-oki(33°54′08′′N, 136°15′34′′E) in 2004, respectively.
The highest sea-level rise caused by low atmosphericpressure during a typhoon since 1950 was recorded by
the Japan Meteorological Agency at 3.89 m at Nagoya
port (35°05′00′′N, 136°53′00′′E) in 1959. Using thestorm records along the entire coast of Japan, we set a
parameter, assuming a large size and an extremely large
size of storm waves, and the sea-level rise for boundary
conditions (Table 1). We also adopted several

Page 5 of 15

Table 1 Boundary conditions for the numerical simulation
Case

Sea level (m)Wave height (m)Period (s)

1A

0

10

10

1B

0

15

15

1C

0

20

20

2A

5

10

10

2B

5

15

15

2C

5

20

20

Sea level, significant wave height, and period are set as the main conditions
on the boundary of the numerical domain

unrealistically large assumptions for the storm wave conditions and sea-level rise at the study site (e.g., cases 1Cand 2C in Table 1) to evaluate whether the lateral extent
of the event layers can be explained by such exaggerated
circumstances.

Results
Lithology

The stratigraphy in the study region consists mainly of
black peat and black peaty silt deposited in a marsh environment (Fig. 4). The 2011 tsunami deposits werealready removed at that time of this study, and the surface sediment was artificially disturbed. At each coring
site, one or several gravel sand layers were interbedded
within peat and peaty silt. At 20–30 cm below the surface, single-layer or patchy tephra were found. The
gravelly sand consists of well-rounded gravel andcoarse to medium sand (Fig. 5a, b) that locally shows
normal grading. At site NM-K, we identified four gravelly sand layers that are named layers I to IV in descending order (Fig. 6).
Differences were found in the lithostratigraphy between the seaward and landward sides of the transect.
Here, we describe the details of the lithology of NM2(seaward side: 467 m inland; Fig. 5a) and NM-K (landward side: 655 m inland; Fig. 6) as representative cores.
The characteristics of all the gravelly sand layers, such as
thickness, mean grain size, and sedimentary structures,are presented in Additional file 1, Table A1.
NM2 (Fig. 5a): In this 50-cm-long core, agriculturally
disturbed sediment and black peat lie at respective
depths of 0–22 and 22–50 cm from the surface. At
depths of 28–30 and 38–46 cm, layers of gravelly sandare interbedded with the peat. Normal grading was observed in the lower layer of the gravelly sand.
NM-K (Fig. 6): In this 58-cm-long core, agriculturally
disturbed sediment and black peat occur at respective
depths of 0–6 and 6–58 cm from the surface. At depthsof 17–18, 33–34, 43–45, and 53–55 cm, gravelly sandlayers are interbedded with peat or peaty silt. A thin (<1 cm) tephra layer lies 1 cm above the uppermost gravelly sand layer. The peat layer at 22–29 cm depth isInoue et al. Progress in Earth and Planetary Science (2017) 4:42

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Fig. 4 Correlation of event layers (I, II, III, and IV) among the cores, based on lithology and chronological data. Red arrowheads indicate thelocations of samples for radiocarbon dating and sample numbers (Table 2)

brown in color, and that at 34–47 cm depth is black,and the peat gradually becomes peaty silt with increasing
depth below the gravelly sand layer. At 47–53 cm depth,
the peat is mainly black in color, except for the lowest
part, which is brown. The peat at 55–58 cm depth isbrown. We used these characteristics to identify key
beds for stratigraphic correlation.
The number and thickness of gravelly sand layers vary
among the cores. At some sites, no gravelly sand layers
were found (see Additional file 1, Table A1). The stratigraphic correlation of the gravelly sand layers between
cores was based on the following criteria: (1) layer I isdeposited immediately below patchy tephra or a layer oftephra, (2) brown peat occurs between layers II and III,(3) layer III is thicker than the other layer I, II, and IV,
and interbedded within black peaty silt with sharpboundaries, and (4) layer IV is finer in grain size thanlayer III and brown peat occurs below layer IV.
Radiocarbon ages, as described below, are used to improve the accuracy of the correlation, as described in the“Discussion” section.

− 3.41 ϕ (in core NM-B). The lower event layer in coreNM2 shows macroscopically normal grading. However,grain size analysis further revealed that variations in themode of the diameter in the lower event layer represent
repeated normal and inverse grading. In other event
layers, the vertical variation in the mode of the diameter
indicates various trends such as normal and inverse
grading, and massive structure (Fig. 7). The NM-L coreis the landward limit at which we identified gravelly
sands along the transect. The maximum grain size at
this site was − 2.83 ϕ.
Radiocarbon dating

Radiocarbon ages were used to confirm the continuity ofsand layers and to constrain depositional ages. The ages
range from 1200 to 2700 cal BP (Table 2; Fig. 4). Detailed age measurements were undertaken for the gravelly sand layers in NM-K, yielding ages of (in descendingorder): ~ 1290, 2150–1930, 2310–2070, and 2750–
2385 cal BP (Fig. 6).
Tephra analysis

Grain size and thickness of gravelly sand layers

A typical example of the vertical variation in grain size isshown for the NM2 core (Fig. 5). The gravelly sands are
generally poorly sorted (Fig. 5), and some show a bimodal grain-size distribution. The coarsest grain size isRadiocarbon dating results from the NM-K core showthat the tephra layer was deposited after 1290 cal BP.
Around the study site, the AD 915 Towada-a (To-a) and
AD 946 Baedusan-Tomakomai (B-Tm) tephra are candidates during this period (Machida and Arai 2003; Aoki

A

0

Depth [cm]Inoue et al. Progress in Earth and Planetary Science (2017) 4:42

20

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B

II

1

40
III

-3.0

Photograph ofcore sample

Stratigraphy

-1.5

0

1.5

3.0

Grain size [φ]CT image60
Disturbed soil

fine

Volume %mode diameter

grading

Gravelly sand
Peat

fine

coarseRadiocarbon dating

0

5

10

15

20

25

30

35

40

45

50

volume [%]Fig. 5 a Photographs, stratigraphy, CT images, and grain size of core NM2. For the grain size diagram, the horizontal axis is grain size (Φ) and thevertical axis is volume frequency (%). The gray scale indicates the mode of grain diameter. The volume frequency (%) of each grain size is indicated by
the color scale. b Photograph of particles of layer III in core NM2 (depth 38–46 cm). Red arrowheads indicate the location of a sample for radiocarbon
dating and the sample number (Table 2)

and Machida 2006; Nakamura 2016). The major element
composition (TiO2–KiO2) of the tephra layers at the
study site is similar to that of reference samples fromthe To-a and B-Tm tephras investigated at Tagajo (38°18′24′′N, 140°59′18′′E) and Tomakomai (42°37′53′′N,141°36′34′′E) (Aoki and Machida 2006), respectively(Fig. 8). Therefore, we suggest that the tephras at the
study site are the To-a and B-Tm tephras.
Although a single-tephra layer was observed, and two
different kinds of tephra glass were identified in the layer,probably because the tephra layers were formed during avery short time interval. Moreover, the tephra is locallyintermixed, showing patch-like features, suggesting thatminor reworking occurred after sedimentation.

near the shoreline were both recorded in case 2C. The
inundation limit of storm waves is up to 450 m from theshoreline (7.33 m in elevation) if we assume an extreme
case (i.e., case 2C; Fig. 9d). The inundation limit covers asmall portion of the seaward section of the transect, inwhich the seaward end was set 410 m inland of thecoastline and 7 m in elevation. The modeled inundation
by storm waves is much shorter and lower than the distance and elevation of the landward end of the site withgravelly sand layers (705 m inland and 11 m in elevation). It is important to note that this inundation can be
explained mainly by the assumed sea-level rise (5 m) rather than the storm wave.

Numerical simulation of storm waves

Discussion

Cases 1C and 2C in Fig. 9 are assumed to the most extremely large storm waves and sea-level rise due to low
atmospheric pressure among the study cases of the simulated waves. The results show that the maximum significant wave height (5.46 m) and flow speed (1.79 m/s)

Correlation of event deposits

The gravelly sand layers have clear upper and lower
boundaries. Some layers show macroscopic normal grading. These features suggested that the gravelly sands
were allochthonous and deposited by high-energy

Inoue et al. Progress in Earth and Planetary Science (2017) 4:42

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Fig. 6 Photograph, stratigraphy, and the radiocarbon dating results of core NM-K. From top to bottom, the deposits in cores NM-K are labeled as
event layers I, II, III, and IV. Red arrowheads indicate the locations of samples for radiocarbon dating

events. Hereinafter, we designate these gravelly sandlayers as event layers (Figs. 4 and 6).
Each event layer was correlated between cores basedon lithology and radiocarbon ages (see Table A1 for a
description of the event layers and criteria for stratigraphic correlation). In the seaward half of the transect(from NM-3 to NM-9), the upper and lower boundariesof the event layers are clear. The event layers along
NM3 to NM-C display a remarkable difference in grain
size (Table A1), which made it possible to correlate each
event layer based on observations. In contrast, correlation of event layers in the middle part of the transect(from NM-E to NM12) was not easy because somelayers were unclear or not visible in some cores. In the
landward half of the transect (from NM-F to NM-L), thethickness of the event layers decreased landward. Consequently, correlation using only sedimentological features
was not straightforward. However, radiocarbon dating
results, tephra layers, brown peats, and the black peaty
silt were useful as key beds to correlate event layers. Figure 4 shows the correlation of the event layers between
different cores.
Based on the correlation, we clarified that event layer III
is the most distinct layer in terms of lateral continuity andthickness (Fig. 4). Conversely, event layer I was only
observed in the landward area, 650 m inland. Event layers
II and IV were deposited widely but lacked continuity.

Sources and identification of tsunami deposits

Given that event layer III is the most distinct and continuous event layer, we discuss this layer first. In eachcoring site, some gravelly sand layers show overall normal grading (Figs. 5a and 7), indicating that event layer
III was deposited under the influence of flow. Moreover,a coarser fraction (particles larger than − 3 ϕ) was notobserved in the landward part of the transect (> 588 mfrom the shoreline), as shown by the maximum grain
sizes and their thicknesses of the layers. Grain size becomes finer inland, and layers become thinner (Fig. 7),
which indicates the transportation of material from thesea-side such as beach and/or ocean (Dawson 1994).
Sediment supply from river flooding or the surrounding cliffs can be ruled out for the following reasons.
Firstly, the event layer particles are well rounded (Fig. 5b)and similar to particles on the beach (Fig. 2b), which
supports the possibility of sediment transportation fromthe sea. In contrast, sediments from terrestrial cliffs areusually composed of brecciated particles (e.g., Nishina et
al. 2013). Thus, the sediments of nearby cliffs are unlikely to have been the source of the event layers. Secondly, although flooding from the Maita River might
have occurred in the past, such flooding was probably
insufficient to explain the distribution of event layers because of the small size of the river and the small volume
of material carried by the river. Deposition from river

Inoue et al. Progress in Earth and Planetary Science (2017) 4:42

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A
16
Event layer I
Event layer III

Thickness [cm]14

Event layer II
Event layer IV

12
10
8
6
4
2
0

400

450

500

550

600

650

700

Distance from shoreline [m]BMode diameter

NM2

Core name
Sample interval [1 cm]

Depth

upwardVolume % [Maximum 50% (*62.5%)]
fine

Depth [cm]

NM-D
NM7

NM8

NM-F

NM15
NM14*

NM-H

NM-I

NM-J
NM-L

0

-3

467

3 -3

0

3 -3

470

0

3 -3

489

0

509

3 -3

0

3 -3

555

0

584

0

3 -3

588

3 -3

0

3 -3

0

606

603

3 -3

0

3 -3

620

0

3

705

[m]

3

Grain size [φ]2Maximum grain size

Mode diameter

1
0
−1
−2
−3

400

450

500

550

600

650

700

Distance from shoreline [m]Fig. 7 a Thicknesses of event layers I, II, III, and IV with distance from shoreline. b Vertical (upper panel) and horizontal (lower panel) variations in
grain size (ϕ) of event layer III with distance from the shoreline. Each error bar shows the range of the mode of the diameter. Gray hatching in
the bottom panel indicates that the coarser fractions are not observed in the landward part of the transect

flooding is also contradictory to the landward fining andthinning of layer III. Therefore, we infer that it is highly
likely that event layer III was formed by an extreme
wave event.
Regarding the other event layers (I, II, and IV), event
layers II and IV are consistently thin (1–2 cm) throughout the transect (Fig. 7). The interpretation of event
layer I is uncertain because little evidence of the layer
was found in the survey area. Hence, although evidence

is limited for event layers I, II, and IV, we interpret thatthese layers were also formed by extreme wave events,
considering the similarities in lithology and distribution
to event layer III.
We further examined the type of wave that could haveformed these event layers; i.e., tsunami or storm waves.
In case 2C (Table 1), the size of the input waves was set
considerably larger than the measured largest waves inJapan. The numerical simulation results for case 2C

Inoue et al. Progress in Earth and Planetary Science (2017) 4:42

Page 10 of 15

Table 2 Radiocarbon dating results
SamplenumberCore

Depth (cm)Material

Conventional
radiocarbon age (yrBP)

Calibrated age(2σ) cal BP

δ13C (‰)

1

NM2

30–32

Organic sediment1910 ± 30

1920–1815− 27.2

2

NM8

25–27

Organic sediment2100 ± 30

2150–1995− 27.0

3

NM-D

35–37

Organic sediment

1740 ± 30

1715–1565− 27.6

4

NM12

30–32

Organic sediment2030 ± 30

2055–1900− 24.4

5

NM-F

17–19

Organic sediment1910 ± 30

1920–1815− 26.9

6

NM13

48–50

Organic sediment2610 ± 30

2765–2730− 24.1

7

NM-H

21–23

Organic sediment

1940 ± 30

1945–1825− 27.6

8

NM15

48–50

Organic sediment2170 ± 30

2305–2070− 27.6

9

NM-I

41–43

Organic sediment2000 ± 30

2000–1885− 27.3

10

NM16

42–44

Organic sediment2070 ± 30

2120–1950− 26.8

11

NM17

50–52

Organic sediment2000 ± 30

2000–1885− 26.8

12

NM-K

18–20

Organic sediment1300 ± 30

1290–1180− 27.4

13

NM-K

30–32

Organic sediment2050 ± 30

2115–1930− 26.6

14

NM-K

34–36

Organic sediment2100 ± 30

2150–1995− 26.5

15

NM-K

40–42

Organic sediment2170 ± 30

2305–2070− 26.2

16

NM-K

44–45

Organic sediment2190 ± 30

2310–2120− 24.3

17

NM-K

51–53

Organic sediment2480 ± 30

2720–2385− 25.2

18

NM-K

54–55

Organic sediment2570 ± 30

2750–2710− 26.6

19

NM18

25–27

Organic sediment2150 ± 30

2300–2055− 26.9

20

NM19

41–43

Organic sediment2160 ± 30

2305–2065− 24.7All data were obtained by accelerator mass spectroscopy (AMS). Sample numbers correspond to those in Figs. 4 and 5[Wt %]
6%

This study (Noda)

4%

To-a

K2O

Aoki and Machida (2006)
B-Tm
Aoki and Machida (2006)
Nakamura (2016)

2%

0%
0%

0.4 %

[Wt %]

0.8 %

TiO2

Fig. 8 K2O–TiO2 diagram for glass shards. The red, blue, and green
symbols indicate marker tephras from Aoki and Machida (2006) andNakamura (2016). Error bars show the range of values

show that storm waves reached only 450 m inland of
the shoreline. The storm waves failed to inundate mostof the transect (distance from the shoreline, 410–
710 m). Therefore, it is likely that such storm conditions did not affect the study area. As for the strengthof storms, Mishima et al. (2009) proposed that seawatertemperature is the most important factor. According toKoizumi and Sakamoto (2010), the sea surface temperatures around Japan have been almost constant for the
past 3000 years. Multi-decadal climate fluctuations in
the NW Pacific Ocean also affect seawater temperature.
However, Mishima et al. (2009) showed that seawatertemperature, which has a cycle of 30–40 years, has had
less of an influence on multi-decadal climate fluctuations, based on records from the last several hundredyears. Therefore, it is likely that there has been no clear
change in storm strength over the past 3000 years.
Goto et al. (2013) also suggested, based on a study ofcoastal boulders deposited on fringing reefs by storm
waves in the Ryukyu Islands, southern Japan, that storm
wave impacts during at least the past 2300 years haveremained stable. Therefore, it is reasonable to infer thatthe extremely large storm waves assumed in this study
are not likely to have occurred over the past several
thousand years.

Inoue et al. Progress in Earth and Planetary Science (2017) 4:42

Page 11 of 15

a

b

c

d

Fig. 9 a Domain of the numerical experiment. The red line denotes the transect. Numerical results are shown for maximum seawater level ofb case 1C (20 m wave height, 20 s period, and water level 0 m) and c case 2C (20 m wave height, 20 s period, and water level 5 m). Black line
denotes the shoreline. d Cross-sectional profile of the maximum seawater level (red points) above the mean sea level in case 2C. A minor discrepancy
between the red points and altitude data occurs because of the difference between our own measured altitude data along the transect and the
topographic data used for the numerical simulation

The inundation distance of former large storm waves
can be related to the paleo-shoreline position and paleoelevation. The shoreline position of the study area haschanged as the Sanriku coast has been eroded bydisproportionate littoral drift, resulting in a reduction in
sediment supply from rivers, and by the settlement of artificial embankments (Tanaka 1980; Kinokuni and Udo2015). This coastal erosion has occurred in the past 20–

Inoue et al. Progress in Earth and Planetary Science (2017) 4:42

100 years, so it may not be applicable directly tomillennium-scale environmental change. However, oursurvey site is a backland located behind a pocket beachsurrounded by rocky cliffs (Fig. 2a). Therefore, theshoreline might have been stable or shifted landwards
in response to past erosion of the rocky cliffs. This shiftsuggests that the paleo-shoreline would have been located farther offshore than its present position if we
consider coastal erosion. If we assume that the shoreline position was farther offshore than the present position, then the calculated inundation limits of the storm
waves in this study should be shifted seaward.
Pleistocene marine terraces along the Sanriku coast,
especially along the northern Sanriku coast, have beeninterpreted as evidence of long-term crustal uplift over
several ~ 200,000 years (Koike and Machida 2001). The
crustal uplift may change our numerical results of inundation distance by a storm wave. The marine terrace(MIS 5e, 125 ka) stretches continuously from Hachinohe
to Kuji, which is located just north of Noda (Niwa andSugai 2016). The uplift rate at Kuji is estimated to be ~0.2 mm/year based on the elevation of a marine terrace(Koike and Machida 2001). On the southern Sanriku
coast (south of Miyako), where there is little information on the marine terraces, Niwa et al. (2014, 2015)
reported that the Rikuzentakata and Kesennuma locations have undergone subsidence during the Holocene,yielding subsidence rates of 0.6–1.3 and 0.9–1.8 mm/
year, respectively. Niwa and Sugai (2016) also pointed
out that the Sanriku coast tilts from north (uplift) tosouth (subsidence). It is unclear whether Noda Village
has been affected by this tilting. As Noda Village is located between Kuji and the southern Sanriku coast, it isreasonable to assume that no great crustal uplift has
occurred near the study area during the past 2700 years,
although a future study of crustal movement in thisarea is required.
Based on this evidence, it is unlikely that the event
layers were formed by storm waves. The 2011 Tohokuoki and 1896 Meiji Sanriku tsunamis inundated far beyond the possible maximum limit of storm waves andfully covered the distribution area of the event layers(Fig. 3b). Therefore, while the event layers do not show
the laminations or shell fragments that are apparent in
the 2011 tsunami deposits, we infer that the event layers
were probably formed by tsunami waves.
Paleo-tsunami history and size

In core NM-K, the To-a (AD 915) and B-Tm (AD 946)
tephras lie immediately above layer I. The absence oftsunami deposits above the tephra layers might be theresult of recent agricultural disturbance. In fact, historical documents describe damage resulting from the AD
1611 Keicho tsunami near the study area (Imamura

Page 12 of 15

1934; Tsuji and Ueda 1995). Hirakawa (2012), Ishizawa
et al. (2015), and Takada et al. (2016) described a possible AD 1611 tsunami deposit near the present study
site. Moreover, the AD 1896 tsunami certainly inundated
beyond the inland limit of the 2011 tsunami, but no geological evidence remains because of anthropogenic disturbance. Therefore, we discuss the tsunami history forthe period before deposition of the To-a and B-Tm
tephras in the tenth century.
The radiocarbon dating and tephra results show thatthe ages of layers I–IV in core NM-K are 1290–1004,
2150–1930, 2310–2070, and 2750–2385 cal BP, respectively, (Fig. 6). The age of layer I can probably be correlated with the AD 869 Jogan tsunami event. No
historical documentation has been found from this areathat describes the AD 869 Jogan earthquake and tsunami. In recent studies, possible AD 869 Jogan tsunami
deposits at Koyadori have been regarded as the northern
limit of its distribution (Ishimura and Miyauchi 2015;
Fig. 1). The identification of possible AD 869 Jogan tsunami deposits at the Noda Village site suggests that thetsunami affected areas farther north of the Sanriku region.
If this is the case, then our study site currently defines oneof the northern extents of the AD 869 Jogan tsunami. Thisis an important step towards improving the tsunami
source model in future work. However, it should be mentioned that Takada et al. (2016) suggested another possible
AD 869 Jogan tsunami deposit at Harashinai, ~ 25 km
north of Noda, but the age of the deposit is not well constrained by tephra chronology.
Our results are also useful for refining our estimation of the tsunami recurrence interval along the Sanriku coast. At the study site, it is likely that fourlarge tsunamis inundated the area between approximately 1100 and 2700 cal BP. It is noteworthy thatfour tsunami deposits are identified within the coresat approximately even intervals, although it might benecessary to consider the influence of sedimentation
rates and erosion. The recurrence interval is estimated at 400–600 years.
Ishimura (2017) reported that five tsunamis (E4 (the 869
Jogan tsunami) to E8) occurred during the 1779 years between 1081 and 2860 cal BP at Koyadori (Fig. 1), compared with the four events at the present study site.
Although the correlation of tsunami deposits between thetwo sites is not easy because of the wide age range, it islikely that the same tsunamis affected both sites. One possible reason for the one-fewer event bed at our study siteis that it is located at a higher elevation than Koyadori.
Therefore, it is possible that smaller tsunamis did notleave deposits in the study area. The location of the paleotsunami source and the difference in the direction of bay
entrances at Noda and Koyadori may have also influenced
the number of event layers.

Inoue et al. Progress in Earth and Planetary Science (2017) 4:42

It is noteworthy that all four tsunami deposits of thepresent study occur close to the inundation limit of the2011 Tohoku–oki tsunami. The 2011 tsunami inundation was affected by artificial embankments. For thatreason, it is difficult to compare the extent of inundationof the 2011 tsunami and older tsunamis directly. However, it is reasonable to speculate that the local sizes of
paleo-tsunamis were equivalent to those of the 2011
Tohoku-oki or 1896 Meiji Sanriku tsunamis when we
consider the landward extent of each tsunami deposit.
Future numerical modeling is needed to elucidate thesources of the pre-historical tsunami events.

Conclusions
We identified four tsunami deposits in Noda Village,
Iwate Prefecture, along the northern Sanriku coast,
based on the distribution and sedimentary features ofdeposits and numerical modeling results. The results
show that even unusually large storm waves with highsea-level rise caused by low atmospheric pressure fail toexplain the distribution of event deposits. Numericalmethods can be a useful tool in identifying tsunami deposits if the maximum possible storm waves are assumed. The ages of the tsunami deposits suggest thatfour large tsunamis occurred in this area during the1600 years between 1100 and 2700 cal BP, the youngestof which might have been the AD 869 Jogan event. Considering the landward extents of the respective tsunami
deposits, the local inundation extent of the tsunamis
might have been similar to that of the 2011 or 1896 tsunamis, although the tsunami sources are uncertain.
Additional file
Additional file 1: Table A1. Characteristics of gravelly sand deposits.
(PDF 368 kb)
Abbreviations
B-Tm: Baedusan-Tomakomai; MIS: Marine isotope stage; To-a: Towada-a
Acknowledgements
This research was supported by the JSPS research funds (no. 26242033) andfunds from IRIDeS, Tohoku University. CT image analysis was conducted at
the Center for Advanced Marine Core Research in Kochi University (accept
no. 14B041). We especially thank a landowner for giving us permission to
conduct surveys and for providing useful information. We thank T. Ishizawa
for his assistance during the field survey and M. Yamada and S. Fujino for
their assistance during grain size measurements. We thank J. D. Bricker forhelp with the numerical simulation and K. Edwards for helpful comments on
editing the English version of the manuscript. J. Matsumoto, K. Ikehara, W.
Szczuciński, and an anonymous reviewer provided valuable comments that
improved the manuscript. We thank all of these people for their support.
Funding
Not applicable
Authors’ contributions
All authors conducted field surveys, collected samples, and participated in
discussions during the research. Additionally, TI mainly performed thePage 13 of 15

experiments and wrote the paper. KG proposed the study topic, andconceived and designed the study. YN conducted the survey related to the2011 tsunami deposits and tephra analysis. MW and TI performed thenumerical simulation related to storm waves using Delft 3D. All authors readand approved the final manuscript.
Competing interests
The authors declare that they have no competing interests.

Publisher’s NoteSpringer Nature remains neutral with regard to jurisdictional claims inpublished maps and institutional affiliations.
Author details
1
Department of Earth Sciences, Graduate School of Science, Tohoku
University, Sendai 980-8578, Japan. 2International Research Institute ofDisaster Science, Tohoku University, 468-1 Aza-Aoba, Aramaki, Aoba-ku,
Sendai 980-0845, Japan. 3Institute of Seismology and Volcanology, Faculty of
Science, Hokkaido University, Sapporo, Hokkaido 060-0810, Japan.
4
Department of Civil and Environmental Engineering, Graduate School ofEngineering, Tohoku University, 468-1-305 Aza-Aoba, Aramaki, Aoba-Ku,
Sendai 980-0845, Japan. 5Department of Earth Sciences, Graduate School of
Science, Tohoku University, Sendai 980-8578, Japan. Present address: INPEXCorporation, Akasaka Biz Tower 5-3-1, Akasaka, Minato-ku, Tokyo 107-6332,
Japan. 6Department of Earth Sciences, Graduate School of Science, Tohoku
University, Sendai 980-8578, Japan. Present address: Museum of Natural andEnvironmental History, 5762 Oya, Suruga-ku, Shizuoka, Shizuoka Prefecture
422-8017, Japan.
Received: 7 January 2017 Accepted: 11 December 2017

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