Hamahashi et al.
Progress in Earth and Planetary Science(2024) 11:8
https://doi.org/10.1186/s40645-024-00607-0

RESEARCH ARTICLE

Progress in Earth andPlanetary Science

Open Access

Shallow structure and late quaternary slip
rate of the Osaka Bay fault, western Japan
Mari Hamahashi1,8* , Hironori Otsuka1, Yoshiaki Suzuki2, Jun Arimoto2, Tetsuo Matsuno1, Nobukazu Seama1,3,
Yuzuru Yamamoto1,3, Hiroko Sugioka1,3, Stephen A. Bowden1,4, Satoshi Shimizu1,6, Hikaru Iwamaru1,5,Mamoru Sano1,5, Keita Suzuki1,5, Katsuya Kaneko1,3, Kazuo Nakahigashi7 and Yoshiyuki Tatsumi1

Abstract
The Osaka Bay is situated at a seismically active region north of the Median Tectonic Line and east of Awaji Island
in western Japan, known as part of the Kinki Triangle and the Niigata–Kobe Tectonic Zone. Dense distributionof active faults and high geodetic strain rates characterize the region, posing a major seismic hazard potential
to the coastal and metropolitan areas of the Kansai region. To investigate the shallow structure and recent deformation history of active faults in the Osaka Bay, we acquired 15 high-resolution seismic profiles using a Mini-GI airgun
and a Boomer as active sources, together with multi-beam bathymetry data across the Osaka Bay Fault. Our seismic
sections image a ~ 0.1 to 3.7 km-wide asymmetric anticline forelimb above the Osaka Bay Fault at shallow depths,
coupled with a ~ 2.6 km-wide syncline to the west, and a broad, ~ 11 km-wide syncline in the footwall to the east. The
synclinal axial surface at shallow depths measured in this study ranges 75°–89°. We observe the vertical displacementof the Osaka Bay Fault increasing northwards along strike. The sediment thickness on the hanging wall, however,is variable, modified by non-tectonic processes such as by tidal currents, affecting the geometry of growth strata.
The most recent deformation by the Osaka Bay Fault reaches to near the seafloor by active folding, with large vertical offsets of 8–14 m over the last ~ 11 ka, and 5–11 m over the last ~ 5 ka. By combining with previously reported
borehole age data, the average uplift rate on the Osaka Bay Fault is estimated to be ~ 1.0 to 1.7 m/ka during the Latest Pleistocene to Holocene. The inferred slip of the Osaka Bay Fault during the Holocene is likely to account for > 5%of the regional geodetic strain accumulation within the Kinki Triangle. Further studies to evaluate the Holocene slip
rates of regional faults are necessary to assess the seismic hazards and the internal strain budgets within the Kinki
Triangle and the Niigata–Kobe Tectonic Zone.
Keywords Seismic reflection survey, Multi-beam bathymetry, Osaka Bay fault, Slip rate, Late quaternary sediments,
Growth strata, Kinki triangle, Niigata–Kobe tectonic zone, Boreholes, Sedimentary basins*Correspondence:
Mari Hamahashi
mhamahashi@yamaguchi-u.ac.jp
Full list of author information is available at the end of the article© The Author(s) 2024. Open Access This article is licensed under a Creative Commons Attribution 4.0 International License, which
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Graphical abstract(See figure on next page.)
Fig. 1 Tectonic setting of the Kinki Triangle, western Japan. (a) Distribution of active faults (red lines) and the Kinki Triangle (yellow) after National
Institute of Advanced Industrial Science and Technology (2012). Shaded relief and bathymetry (10-m and 500-m horizontal grid-spacing,respectively) are based on digital elevation map by the Geospatial Information Authority of Japan. White triangles: Quaternary volcanoes(Committee for Catalog of Quaternary Volcanoes in Japan 1999). Box shows location of study area closed-up in Fig. 2a. Dotted line shows locationof cross-section in Fig. 1c. (b) Interseismic secular velocity at onland GNSS and offshore GPS-A stations modified from Nishimura et al. (2018). Vectors
are velocities relative to the stable Amurian plate with 95% confidence limit ellipsoids (Nishimura et al. 2018). The shaded yellow region and pinkline are the Niigata–Kobe Tectonic Zone (Sagiya et al. 2000) and the Median Tectonic Line, respectively. Depth contour (km) of the upper boundaryof the Philippine Sea and Pacific Plates are from Nakajima and Hasegawa (2007). (c) Schematic cartoon of the E–W cross-section and inferredgeometry of faults at ~ 34°N based on previous studies modified after Sato et al. (1998, 2009). Shaded gray: sedimentary basin in this region. OsF:
Osaka Bay Fault. KaF: Kariya Fault. NoF: Nojima Fault. UeF: Uemachi Fault. IkF: Ikoma Fault. NbF: Nara Basin East Edge Fault. TgF: Taguchi Fault. IsF: Isshi
Fault. YiF: Yokkaichi Fault. The upper/lower crust boundary is inferred at the downdip limit of the seismogenic zone (cross-hatched area) (e.g., Tanaka
2004). A possible detachment is speculated. Box shows region of study area closed-up in Fig. 2a

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Fig. 2 Geologic map of study area. (a) Distribution of active faults (red lines) around the Osaka Bay (e.g., Kato et al. 2008). Black lines: seismicreflection survey lines in this study closed-up in (b). Dotted lines: deep seismic profiles by previous studies. Orange dots: location of boreholes
referred in this study (Nakaseko 1987; Masuda et al. 2000; Masuda & Miyahara 2000; Nanayama et al. 2001; Kitada et al. 2001). OsF: Osaka Bay Fault.
The Awaji Island Fault zone consist of the KaF: Kariya Fault, NoF: Nojima Fault, KsF: Kusumoto Fault, HgF: Higashiura Fault, KmF: Kamakuchi Fault, NdF:
Nodao Fault, and SnF: Senzan Fault. WaF: Wadamisaki Fault. MaF: Maya Fault. RIF: Rokko Island Fault. YoF: Yokooyama Fault. TaF: Takatsukayama Fault.
OtF: Otsuki Fault. R1-R10: Rokko Fault zone (R1: Nobata Fault. R2: Itami Fault. R3: Koyaike graben. R4: Nishinomiya flexure. R5: Koyo Fault. R6: Ashiya
Fault. R7: Gosukebashi Fault. R8: Okamoto Fault. R9: Suwayama Fault. R10: Motomachi and Itayado flexures. R11: Suma Fault.). A1-A5: Arima-Takatsuki
Fault zone (A1: Makami Faults. A2: Ai Fault. A3: Bojima Fault. A4: Kiyoshi-kojin Fault. A5: Rokko Fault). UeF: Uemachi Fault. IkF: Ikoma Fault. Blue stars
show representative earthquakes. Focal mechanisms are based on Imanishi et al. (2020). Geologic map is modified from Geological Survey of Japan,
AIST (2022). (b) Close-up of seismic reflection survey lines in this study. Green and black lines were acquired using an Airgun and a Boomer as active
source, respectively

1 Introduction
Further than 100 km inboard from the subduction interface in central-western Japan, the crustal seismogenic
zone is 5–20 km depth, based on the cut-off depths ofseismicity and on heat flow gradients (Tanaka 2004;
Nakajima and Hasegawa 2007). Active faulting is manifested under the contemporary compressional regime
formed by the oblique subduction of the Philippine
Sea Plate under the Japan arc at a current velocity of

40–65 mm/yr (Seno et al. 1993; Miyazaki and Heki 2001;
Zang et al. 2002). The “Kinki Triangle” (Huzita 1962) isa seismically active region north of the Median Tectonic
Line and east of Awaji Island, characterized by a dense
distribution of active faults and high geodetic strain rates,
also known as part of the “Niigata–Kobe Tectonic Zone”(Sagiya et al. 2000; Nishimura et al. 2018), posing a major
seismic hazard potential to the coastal and metropolitan
areas of the Kansai region (Fig. 1a, b). Moderate to large,

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destructive earthquakes frequently occur in this region,
represented by the devastating 1995 M­ w6.9 Kobe (Southern Hyogo Prefecture) earthquake (death toll: 6434) (e.g.,
Kikuchi and Kanamori 1996; Irikura et al. 1996), the 1596
Keicho-Fushimi Earthquake (death toll: > 600) (e.g., Ikeda
et al. 2019), the 2018 ­Mw5.6 Northern Osaka prefectureearthquake (death toll: 6) (e.g., Hirata and Kimura 2018),
and the 2013 M­ w5.8 Awaji Island earthquake (e.g., Imanishi et al. 2020) (Fig. 2a).
Within the Kinki Triangle, based on patterns of geodetic strain accumulation (interseismic secular velocity)from Ise Basin to the Osaka Bay, an ESE–WNW to ENE–
WSW shortening of ~ 10 to 20 mm/yr at latitude of 34°N
can be inferred (Fig. 1b; Sagiya et al. 2000; Sagiya 2004;
Loveless and Meade 2010; Nishimura et al. 2018, 2020).
During interseismic periods, the seismogenic zone is partially locked, and strain is released during coseismic periods by transferring slip onto reverse and strike-slip faults
in this region. Previous seismic reflection surveys haverevealed a NE–SW trending fold-and-thrust belt consisting of ~ 30 to 40° west- and east-dipping reverse faults
beneath the Ise Basin, Suzuka Mountain, Nara Basin,
and the Osaka Basin (Ishiyama 2003; Sato et al. 2009;
Kitada and Mimura 2016; Nishiwaki et al. 2021) (Fig. 1c).
Whereas, to the west, ~ 70 to 80° east-dipping, high-angle
reverse- and strike-slip faults characterize the structures
around the Awaji Island and the Osaka Bay, including theNojima Fault that ruptured in the 1995 earthquake, andthe offshore Osaka Bay Fault (Iwasaki et al. 1994; Iwabuchi et al. 1995, 2000; Yokokura et al. 1998; Sato et al.
1998) (Fig. 1c).
The largest sedimentary basin in the region comprises
the Osaka Basin and Osaka Bay. It is bounded by theE–W trending Median Tectonic Line to the south, theE–W trending Rokko-Awaji Island and Arima-Takatsuki
Fault zones to the north, the NE–SW trending Nojima
and Kariya Faults to the west, and the NE–SW trending
Ikoma and Uemachi Faults to the east (Fig. 2a; Kusumoto
et al. 2013; Itoh et al. 2013ab; Itoh and Takemura 2019c).
The Late Pliocene and Quaternary basin fill here is up
to ~ 3000 m thick (e.g., Yokokura et al. 1998). The subsidence and formation of this sedimentary basin is largely
thought to have been contributed to by the movement/
reactivation of faults within and surrounding the basin,
which have been active since the Late Pliocene (Yokokura
et al. 1998; Kusumoto et al. 2001, 2013; Itoh and Takemura 2019a, b, c). Here, active faults are buried by sediments and are often blind, complicating the assessmentof their distribution and slip histories. Although blind,the surface expression of these faults is often observedin the form of folding (Ishiyama 2003; Sato et al. 2009;
Kitada and Mimura 2016). The 2018 ­Mw5.6 Northern
Osaka prefecture earthquake occurred on a previously

unknown blind fault, demonstrating the damaging effectsof blind fault earthquakes, and the potential for presenceof many more hidden faults and/or off-fault seismicity(Hallo et al. 2019; Kato and Ueda 2019) (Fig. 2a).
Because these faults are continuously buried in rapidly subsiding basins, the progressive fault displacement
and syn-tectonic growth strata are defined by well-constrained stratigraphic sequences. Acquisition of high-resolution seismic reflection data enable the documentationof deformed strata and imaging of fault and fold geometries, which in combination with borehole data, provides
opportunities to characterize fault slip histories in detail(e.g., Ishiyama et al. 2007; Hamahashi et al. 2022). The
sediments that fill the Osaka Basin and the Osaka Bay
consist of fluvial, lacustrine, and marine sequences thatcan be correlated with global eustatic sea level changes,
and their sediment ages are well-constrained by abundant tephra derived from the volcanic arc, in addition
to marine microfossils, pollen assemblages and magnetostratigraphy (e.g., Hyodo et al. 1985; Masuda 1992;
Biswas et al. 1999; Masuda and Miyahara 2000; Masuda
et al. 2000; Nanayama et al. 2001; Itoh and Takemura
2019a).
The motivation for our research is to investigate theshallow structure and recent slip history of active faults
in the Osaka Bay. The Osaka Bay Fault is a blind fault
that extends for at least ~ 40 km across the bay, exhibiting one of the largest seismic and tsunami risks to thesurrounding Kansai megacities (Nakamura 1999; Kawata
et al. 2005; The Headquarters for Earthquake Research
Promotion 2014) (Fig. 2a). In this paper, we present
results from high-resolution seismic profile data that
we acquired along 15 survey lines in the central-eastern
portion of the Osaka Bay, using a Mini GI air gun (15/15­in3) and a Boomer (100/200 J) as active sources, together
with multi-beam bathymetry data collected onboard theT/S Fukaemaru and R/V Onokoro (Fig. 2b). By combining with published borehole data around the bay withknowledge of deep structures imaged by previous seismic surveys, our newly acquired data documents the Late
Pleistocene to Holocene deformation history at the central segment of the Osaka Bay Fault. Our results also havesignificant implications for changes in depositional environment, as well as for the formation of the sedimentarybasin and interactions with late Quaternary climate andtectonic events in this region, situated at the southernedge of the Kinki Triangle.

2 Background and geologic setting
2.1 The 1995 Kobe (southern Hyogo prefecture)earthquake

The ­Mw6.9 1995 Kobe (Southern Hyogo prefecture)
Earthquake (death toll: 6434) was the most destructive

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earthquake in Japan since the 1923 Kanto Earthquake
and the 1944/1946 Nankai-Tonankai Earthquakes, second to the 2011 Tohoku-oki Earthquake (Kikuchi and
Kanamori 1996; Baba and Cummins 2005; Kobayashi
and Koketsu 2005; Uchida and Bürgmann 2021). The
mainshock occurred at a depth of ~ 16 km beneath the
Akashi Strait north of Awaji Island in the Osaka Bay,with a right lateral strike-slip focal mechanism and a
WNW-ESE trending maximum compressional axis
(Japan Meteorological Agency/JMA 1995; Kikuchi and
Kanamori 1996; Katao et al. 1997) (Fig. 2a). The earthquake rupture propagated bilaterally towards the landin Kobe and Awaji Island along a ~ 60-km-long northeasterly trending zone consisting of multiple fault tracesof the Rokko-Awaji Island Fault zone; from the Nojima
and Kariya Faults to the Suma, Otsuki, Suwayama, andGosukebashi Faults, and to the Arima-Takatsuki Fault
zone (Hirata et al. 1996; Ide et al. 1996) (Fig. 2a). Dense
aftershock distributions occurred along and/or nearthese faults, and offshore northern Awaji Island in Osaka
Bay (Hirata et al. 1996; Katao et al. 1997). The propagation of the rupture produced long-period pulses in theobserved fault-normal ground motions, and a “damage
belt” where more than 30% of wooden-frame houses collapsed (seismic intensity VII defined by JMA) occurredalong the topographic boundaries between the cities andhills in the Kobe-Hanshin area (Furumura and Koketsu
1998). The location of the damage belt deviated from theearthquake distributions and was closer to the coast, andwas mainly explained by strong ground motions caused
by rupture directivity from basin edge effects (Furumura and Koketsu 1998; Hallier et al. 2008; Koketsu andMiyake 2008). Field surveys have found surface ruptures
occurring on a ~ 9 to 15-km-long segment of the preexisting right-lateral Nojima Fault on the Awaji Island witha maximum horizontal displacement of ~ 1.9 to 2.1 m
and vertical offsets of up to ~ 1.0 to 1.2 m, whereas surface ruptures were not reported in the Kobe area (Awata
et al. 1995; Ikeda et al. 1995; Nakata and Yomogida 1995;
Nakata et al. 1996; Sato et al. 1998).
In response to the disastrous 1995 earthquake, theJapanese government, national and private institutes
and universities conducted seismic reflection surveys toassess the structures of active faults and seismic hazards
in this region (Iwabuchi et al. 1995, 2000; Yokota et al.
1997; Yokokura et al. 1998; Sato et al. 1998; Nanayama
et al. 2000; Grothe et al. 2014; Itoh and Takemura 2019b).
Seismic reflection surveys around the Osaka Bay havecharacterized the distribution of high-angle (~ 80°), NE–
SW trending strike-slip/reverse faults responsible for the
1995 earthquake; the east-dipping Nojima Fault along thewest coast of Awaji Island, and the west-dipping Higashiura and Kariya Faults along the east coast (Yokota et al.
1997; Sato et al. 1998) (Fig. 2a). These fault systems havelikely contributed to the narrow uplift of the Awaji Islandthat extends to the Rokko Mountains (e.g., Huzita 1990;
Kato et al. 2008). Offshore of the Awaji Island, multiple
blind reverse faults have been identified within the Osaka
Bay, including the ~ 70° west-dipping, NE–SW trending
Osaka Bay Fault and a subsidiary fault ~ 15 km and ~ 3 to
5 km east of the island, respectively (Iwabuchi et al. 1995,
2000; Yokota et al. 1997; Yokokura et al. 1998; Sato et al.
1998; Nanayama et al. 2000) (Fig. 2a). These fault systems
have likely made a significant contribution to the longterm subsidence of the basin in the footwall side (e.g.,
Itoh et al. 2000, 2001). Despite the proximity to the faults
around the Awaji Island that ruptured during the 1995
earthquake, the Osaka Bay Fault is not reported to haveruptured during the earthquake, as inferred from mainshock and aftershock distributions, although aftershock
microseismicity was observed offshore northern Awaji
Island, possible along the subsidiary faults between the
Kariya Fault and the Osaka Bay Fault (Hirata et al. 1996;
Katao et al. 1997).
2.2 The Osaka bay fault

The Osaka Bay Fault is responsible for the deformationof the ~ 3-km-thick Late Pliocene to Quaternary sediments of the Plio-Pleistocene (< 3 Ma) Osaka Group andthe Late Pleistocene-Holocene strata (colluvium, alluvium) that extend to the Osaka Basin, and the underlying
Cretaceous granitic basement rocks of the Ryoke Group(e.g., Yokota et al. 1997). The deformation caused by thefault is observed in the form of thrust folding, characterized by an asymmetric anticline (~ 1 to 2 km-wide flexure) of which its forelimb becomes steeper with depth,
and a broad syncline (~ 10 km wide) in the footwall side,
depositing packages of growth strata (Iwabuchi et al.
1995, 2000; Yokota et al. 1997; Yokokura et al. 1998; Satoet al. 1998; Usami 2002; Grothe et al. 2014). The fault tip
is not directly imaged by previous studies, although it isinferred to lie > 1 km below the seafloor based on trishear
modelling (Grothe et al. 2014). The vertical displacementby the Osaka Bay Fault reaches ~ 1 km at the base of thesediments during the last ~ 3 Ma, and reaches to a totalof ~ 2.5 to 3 km relative to the uplifted Awaji Island whencombining with the displacements by other faults along
the east coast and the Kariya and Higashiura Fault systems (e.g., Sato et al. 1998). The lateral displacement by
the Osaka Bay Fault is unknown, and previous studies
have not been able to clarify the strike-slip offsets fromtheir seismic images (e.g., Yokokura et al. 1998). AlthoughHamahashi et al. Progress in Earth and Planetary Science(2024) 11:8

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dip slip is largely supported from the geodetic principal
strain axis (ESE–WNW to ENE–WSW) that is nearly
perpendicular to the strike (NE–SW) of the main fault
traces of the Osaka Bay Fault (Sagiya 2004; Nishimura
et al. 2018, 2020) (Fig. 1b), right-lateral strike-slip faulting is also expected from the steep dips (> 70°) of thesefaults, and the regional earthquake focal mechanismsthat show strike-slip and reverse slip senses (Hirata et al.
1996; Katao et al. 1997; Nakamukae et al. 2003; Matsushita and Imanishi 2015). The P-axes of regional focal
mechanisms generally show ESE–WNW to ENE–WSW
compressions, but with some complexities (e.g., NE–SW,
SE–NW and N–S compressions and normal faulting),
possibly due to the variable geometry of faults and/or
local stress changes (Katao et al. 1997; Nakamukae et al.
2003; Matsushita and Imanishi 2015). In the northern
region of the Osaka Bay, right-lateral strike slip is likely to
be more dominant where the Osaka Bay Fault bends andstrikes WSW–ENE parallel to the Rokko-Awaji Island
Fault zones (e.g., Nakamukae et al. 2003) (Fig. 2a).
The Osaka Bay Fault can be divided into the northern,
central, and southern segments based on the alongstrike variations in fault geometry, that may partly be
due to the structural relationship with the surrounding
Rokko-Awaji Island Fault zone, the Median Tectonic
Line, and the Uemachi Fault zone (Fig. 2a). Understanding how these major faults structurally link witheach other is important for assessing the regional seismic hazard and the complex rupture processes that
may involve movement of multiple fault strands during a single or sequential earthquake (e.g., Honda andYomogida 2003). At the northern segment, the Osaka
Bay Fault is known to splay into three strands (namely
the Wadamisaki, Maya, and Rokko Island Faults), as
the fault approaches the adjacent Rokko-Awaji Island
Fault zone to the north (Yokokura et al. 1998; Nanayama et al. 2000). At the central segment, displacementby the Osaka Bay Fault is mainly occurring at the main
fault, in addition to distribution of subsidiary faults inthe eastern and western portions of the bay (Yokokura
et al. 1998; Sato et al. 1998; Iwabuchi et al. 2000). In
contrast, at the southern segment, the geometry of the
Osaka Bay Fault changes, and is likely partitioning into
several strands, as also inferred from gravity anomalies and isopach models that show complex basement
morphologies and varying sediment thickness (Komazawa et al. 1996; Yokokura et al. 1998; Iwabuchi et al.
2000; Afnimar et al. 2002; Itoh et al. 2013a). However,
due to the lack of seismic imaging in the southern portion of the Osaka Bay, the fault geometry at this segment is unknown. This segment has importance in the

structural relationships with the Median Tectonic Lineto the south, and the Uemachi Fault zone beneath the
Osaka Basin to the east that is gradually bending westward towards the southern Osaka Bay (Fig. 2a; Researchand Development Bureau/MEXT & DPRI 2013).
While previous seismic reflection surveys in the
Osaka Bay have mostly targeted great depths of ~ 2 km
below the seafloor, inferring the overall deformation
structure from the basement to the sedimentary section (Iwasaki et al. 1994; Yokokura et al. 1998; Satoet al. 1998; Iwabuchi et al. 2000), the shallow structuresof the faults have not been studied in detail. Previous
surveys that targeted shallow depths (Hayakawa et al.
1964; Huzita and Maeda 1969; Inouchi 1990; Iwabuchi et al. 1995; Nanayama et al. 2000) have inferred theexistence of multiple blind faults in the bay (e.g., Iwabuchi et al. 1995; Usami 2002), suggesting possible fault
ruptures associated with the 1995 earthquake (Iwabuchi et al. 1995), but detailed studies have been limited
to the northern segment of the fault close to the coast(Nanayama et al. 2000). The recent deformation historyof the Osaka Bay Fault is not well constrained, together
with the lack of historical earthquakes (The Headquarters for Earthquake Research Promotion 2014).
Detailed constraints on the shallow structure and subsurface deformation are crucial for assessing the Holocene fault slip rates and seismic risks that may involve
tsunami generation at the seafloor (e.g., Nakamura
1999; Kawata et al. 2005) and strong ground motions inthe sedimentary basin, and is the primary focus of thisstudy.
2.3 Depositional environment and late quaternary
sedimentary sequence around the Osaka Bay

The Osaka Bay is situated at the eastern edge of the
Seto Inland Sea in western Japan (Figs. 1, 2). The Seto
Inland Sea has generally shallow water depths of ~ 30 to
50 m, and is connected with the Pacific Ocean via the
Kitan and Naruto Straits to the southeast and the Hoyo
Strait to the west, and connects with the Japan Sea via
the Kanmon Strait, consisting of multiple enclosed bays
(Osaka, Harima-nada, Hiuchi-nada, Hiroshima, Iyonada, and Suo-nada Bays) laterally linked by narrowstraits (Akashi, Bisan-seto, Kurushima and Tsurushima
Straits) (e.g., Yasuhara 2008; Saito et al. 2016). While
the paleo-Seto Inland Sea is thought to have originatedfrom repeated, episodic rifting events in the Asian
continent since the Eocene–Oligocene (e.g., Nakajima2018), partially overprinted by Setouchi volcanism in
the Mid-Miocene (~ 15 Ma) (Kimura et al. 2005; Tatsumi 2006), the current Seto Inland Sea is overlain byHamahashi et al. Progress in Earth and Planetary Science(2024) 11:8

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Plio-Pleistocene (< 3 to 6 Ma) sedimentary basins (e.g.,
Mizuno 2018) that have likely interacted with the contemporary compressional regime associated with thesubduction of the Philippine Sea Plate and reactivationof the Median Tectonic Line since the Late Miocene(Kamata and Kodama 1994; Kimura et al. 2005, 2014).
Past fluctuations in relative sea-levels and consequent opening of the straits and formation of bays
have resulted in a dynamic depositional environment
in and around the Osaka Bay (e.g., Irizuki et al. 2001;
Yasuhara et al. 2004; Yasuhara 2008; Masuda and Itomoto 2015). The sedimentary sequence underlying
the Osaka Bay and Osaka Basin consists of alternating
marine and non-marine strata of the Plio-Pleistocene
Osaka Group, overlain by Late Pleistocene colluviumand Holocene alluvium (e.g., Huzita and Maeda 1984;
Itihara eds. 1993; Itoh and Takemura 2019a). While the
Osaka Group and colluvium sediments are partially
exposed at the surface in the surrounding mountains
and elevated fluvial terraces in the Osaka Basin, theHolocene alluvium is distributed more broadly, forming
the coastal alluvial fan and floodplains (e.g., Itoh andTakemura 2019a) (Fig. 2a). The stratigraphy of the sedimentary sequence in this region has been characterizedon the basis of tephrochronology, magnetostratigraphy,
pollen assemblage, and marine clay beds (individuallynumbered “Ma 1”–“Ma13”) correlated to eustatic sea
level high-stands and marine oxygen isotope records(e.g., Itihara 1960, 1993; Yoshikawa et al. 1987; Itohet al. 2000, Itoh and Takemura 2019a), in addition toregional unconformities (e.g., Huzita 1990), establishedfrom outcrops and deep drilling results in the Osaka
Basin and coastal areas in Kobe (e.g., Ikebe et al. 1970;
Mitamura et al. 1998, 2000).
The Late Pleistocene colluvium in this region consistof terrigenous fluvial sands and gravels during the Last
Glacial Period, containing marine clay beds “Ma 12”
and “Ma 11”, correlated to Marine Isotope Stages (MIS)
3 and 5, respectively, whereas the major composition of
the Holocene alluvium is the “Ma13” clay, silt, and sand,correlated to MIS 1, consisting of inner bay deposits andtidal delta sediments (e.g., Masuda 1992; Nanayama et al.
2001). Based on detailed facies analysis and dense 14C
dating from sediment cores offshore Kobe and Kansai
Airport, and along the coastal edge of the Osaka Basin(e.g., Nakaseko 1987; Masuda 1992; Irizuki et al. 2001;
Yasuhara et al. 2004; Kitada et al. 2001), the sediments
beneath the Osaka Bay are interpreted to have transitioned from fluvial, estuary to inner bay deposits during
a rapid rising stage of sea level (transgression) startingat ~ 11 ka when sea level was -51 m below modern sea
level (Masuda and Miyahara 2000; Nanayama et al. 2001;
Masuda and Itomoto 2015). This period was followed bytidal delta deposition consisting of sand and mud, originating from the Akashi Strait that was formed at ~ 9.7 ka,
when sea level was − 30 m below modern sea level (Masuda and Miyahara 2000). The sedimentary sequence wasthen overlain by bay-floor deposits during marine floodand sea level high-stands that occurred at ~ 5.3 to 5 ka,during which sea level was + 1 to 2 m above modern sea
level (Masuda et al. 2000). The most recent sediments
consist of coarse bedload deposits transported by strong
tidal currents from the Akashi Strait, and the uppermost
marine clay which are primarily floodplain suspension
loads transported by the Yodo and Mukogawa River systems from the Osaka Basin, inferred to have deposited
following periods of marine regression that started at ~ 1
to 1.7 ka (Masuda and Miyahara 2000; Nanayama et al.
2001).
The distribution of contemporary seafloor sedimentsin the Osaka Bay is not homogeneous, where marine clay
is restricted to the eastern portion of the bay, whereas
coarser sediments are more dominant in the central towestern part of the bay (Huzita and Maeda 1969; Inouchi
1982; Inouchi 1990; Fujiwara et al. 1989; Yashima 1992;
Masuda and Miyahara 2000; Nanayama et al. 2001).
Sediment distribution is affected by wave and tidal currents operating in the bay, producing various erosionaland depositional processes near the straits and in the
shelf (e.g., Yashima 1992; Ikehara and Kinoshita 1994).
In a longer timescale, the sediments are also affected bycoastal tectonic deformation, forming the sedimentarybasin (e.g., Sato et al. 2017). In this study, we investigate
the shallow structure and distribution of seafloor sediments in the Osaka Bay, and document the geometries ofsequence stratigraphic surfaces and deformation, providing constraints on the dynamic interactions of tectonics,
sea level fluctuations, and tidal currents that led to theformation of the late Quaternary sedimentary basin inthis region.

3 Methods
We conducted our surveys in the central-western portion of the Osaka Bay along E–W trending lines that arenearly perpendicular to the fault traces of the Osaka Bay
Fault (e.g., Iwabuchi et al. 1995; Yokokura et al. 1998;
Usami 2002; National Institute of Advanced Industrial
Science and Technology 2012), and along N–S trending lines that intersect with the E–W trending lines tocharacterize the three-dimensional structure (Fig. 2).

Hamahashi et al. Progress in Earth and Planetary Science(2024) 11:8

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We acquired high-resolution seismic profile data alongeight survey lines using a Sercel Mini-GI airgun withvolume reducers (15 i­n3 generator and 15 ­in3 injector) as
active source, equipped at Kobe University onboard theT/S Fukaemaru (449 t). The shot spacing was 10 s witha record length of 8 s, every 23 m while the vessel ran at
an average speed of 4.5 knot. The towed streamer cable
had 6 channels with 6.25 m hydrophone group intervals
and a total length of 150 m. The received seismic reflection signals were recorded at a 1 ms sampling rate by the
DAQLink 4 seismograph (Seismic Source). The locationsof receivers and shots were determined by GPS. The
acquired data was processed with static correction, bandpass filter, amplitude recovery, 12.5 m binning, normal
moveout correction using constant velocity (1500 m/s),
and Stolt migration. The data are also visualized with anautomatic gain control. The frequency range for bandpass filter was set to 50–450 Hz.
We also acquired high-resolution seismic profile data
along seven survey lines onboard the R/V Onokoro (9.7t), using a Boomer (Applied Acoustics, AA251) as active
source, with an energy output of 100 J or 200 J per shot,
equipped at the Geological Survey of Japan (Suzuki et al.
2021). The shot spacing was 1 s with a record length of
800 ms, every 1.56 m while the vessel ran at an average speed of 3–3.5 knot. The towed streamer cable had
16 channels with 3.125 m hydrophone group intervals,
combined to a total length of 50 m. The received seismic reflection signals were recorded at a 0.125 ms sampling rate by the GeoEel multi-channel digital streamer
system (Geometrics), and the locations of receivers
and shots were determined by GPS. The acquired data
was processed with the sequence of format conversion
and trace edit, geometry application to trace header,band-pass filter, gain recovery, deconvolution, velocity analysis, normal moveout corrections, CMP stacking, and data visualization after a second band-passfilter. The frequency range for band-pass filter was set to300–2000 Hz.
At the vicinity of our seismic reflection profile data,high-resolution multi-beam bathymetry data covering ~ 170 ­km2 around the survey area was acquired
onboard the T/S Fukaemaru, using a Kongsberg EM 712(1.0° × 1.0°) Multi-beam Echo Sounder, operating at sonarfrequencies in the 40–100 kHz range, and swath coverage of 140° with 400 beams by dual swath (Figs. 3, 4). To
account for submetric error on the beam positioning, thesystem was interfaced with Differential Global Positioning System and motion Reference Unit (Kongsberg) for
attitude determination and motion compensation. The

in situ surface water sound speed was acquired for calibrating proper beam steering, by sound velocity probes
equipped at on-keel station located near the multi-beamtransducers, and by CTD profiling of the water column.
Bathymetry and backscatter data were processed with theSeafloor Information System (SIS) and CARIS HIPS &SHIPS software. Noise reduction due to artificial instrumental spikes and poor-quality beam exclusion were carried out by swath editing and de-spiking tools.
We interpret lithostratigraphy and sequence boundaries and assign the ages in our seismic profiles by correlating with published borehole data acquired in thevicinity of our survey lines offshore Kobe city and Kansai
Airport, and along the coastal plain of the Osaka Basin(Nakaseko 1987; Masuda 1992; Masuda et al. 2000; Masuda and Miyahara 2000; Nanayama et al. 2000, 2001;
Kitada et al. 2001; Masuda and Itomoto 2015). We use
observed stratigraphic geometries and age constraints
from boreholes to estimate fault displacements and slip
rate.

4 Results
4.1 Bathymetry and backscatter data

Figures 3, 4, 5 shows the acquired multi-beam bathymetry data (10-m horizontal grid spacing) acquired in thisstudy, plotted with the M7000 digital bathymetry data
published by the Japan Hydrographic Association. The
seafloor of the Osaka Bay has generally shallow water
depths (10–20 m) at the eastern region, abruptly transitioning into deeper water depths (30–65 m) at the central-western part of the bay, across a concave-shaped
inner shelf and slope (Fig. 3a). The deepest portion of thebay (> 70 m water depth) is restricted to the northwest
and northcentral regions offshore northern Awaji Island,
which gradually shallows towards the south (Fig. 3b).
Seafloor morphology is most complex near the Akashi
Strait to the north and the Kitan Strait to the south,
and is characterized by multiple pockets of bathymetric lows and highs, interpreted to be caused by depositional and erosional processes linked to tidal currents(e.g., Inouchi 1990; Yashima 1992; Ikehara and Kinoshita1994) (Fig. 3b). The Akashi Strait has a water depth of up
to ~ 100 m, indicating dominant erosion; previous studies
have documented bedrock granites exposed at the seafloor beneath the strait (Huzita and Maeda 1984). Coincidently, the epicenter of the 1995 Kobe Earthquake isalso at this vicinity, where the Nojima and Kariya Faults(Rokko-Awaji Island Fault zone) run across northern
Awaji Island towards Kobe city (e.g., Kato et al. 2008)(Fig. 2a).

Hamahashi et al. Progress in Earth and Planetary Science(2024) 11:8

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Fig. 3 Multi-beam bathymetry data (10-m horizontal grid spacing) acquired in this study, plotted with the M7000 digital bathymetry data (light
blue contours) by the Japan Hydrographic Association. (a) Location of survey area. (b). Close-up of survey area. Grey lines: seismic reflection surveylines (#1–15). Red line: Trace of the Osaka Bay Fault. Box shows region of distinct sand wave structures closed-up in Fig. 5a. Blue dotted square: areaof negative polarity reflector observed along Lines 3, 5, and 12 (Sect. 4.2.1)

At about 6 km offshore northern Awaji Island, ourbathymetry data clearly image a NW–SE trending
bathymetry high region (27–30 m water depth), extending 6.4 km (long axis) × 2.6 km (short axis) wide withrelative heights of 5–10 m (Fig. 3b). This bathymetrichigh, locally called the “Okinose” sand bank (Inouchi
1990; Yashima 1992; Kojima 2020), is associated with
ESE–WNW trending lineation fabrics interpreted tobe sand wave bedforms comprising a subaqueous dune
complex (Figs. 3b, 5a). The sand waves are oriented
approximately 12–45 degrees to the long axis of the
Okinose sand bank, and are characterized by asymmetric shapes with steeper southern flanks (Fig. 5b). The
largest sand waves are up to ~ 5 m in height with 100–
150 m wavelengths and 1–5 km along-crest length, andare concentrated on the northwestern region, whereas
the smaller waves are < 1 m high with a wavelength

less than 20 m and > 1 km along-crest length (Fig. 5b).
The morphology of the western and southern partof the dune complex are bounded by steep edges andsharp steps in bathymetry, characterized by an NW–SE
trending ~ 5 m-high escarpment in the western side,and WNW-ESE trending ~ 5 m-high escarpment in thesouthern side (Fig. 3b). These escarpments intersect at
34.53° N, 135.10° E south of the summit of the bathymetric high, with an angular contact between the westernand southern flanks (Fig. 3b). The Okinose sand bankis interpreted to be a sand dune complex formed bystrong tidal currents from the Akashi Strait (Inouchi
1990; Yashima 1992; Kojima 2020). To the east and westof the dune complex, the water depth deepens to ~ 60
to 78 m (Fig. 3b). The regional depressions to the eastof the Okinose sand bank is an area locally named“Yokose”, and are likely forming seafloor cauldrons andHamahashi et al. Progress in Earth and Planetary Science(2024) 11:8

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Fig. 4 Backscatter intensities acquired in this study linearly mapped to a greyscale mosaic with a resolution of 1 m and dynamic mosaic rangeof -41.2–15 dB. Light blue contours: M7000 digital bathymetry data. Three characteristic backscatter imagery facies are identified (boundary traced
by green-dotted lines). Facies 1: heterogeneous, moderate reflective (light grey) facies. Facies 2: homogeneous, weak-to-moderate reflective (darkgrey) facies. Facies 3: homogeneous, moderate-to-strong reflective (light grey) facies. Black lines: seismic reflection survey lines (#1–15). Red line:
Trace of the Osaka Bay Fault. Box shows region of distinct sand wave structures closed-up in Fig. 5bscour hollows eroded by tidal currents (Inouchi 1990;
Yashima 1992; Ikehara and Kinoshita 1994).
Figure 4 shows the backscatter intensities acquired inthis study linearly mapped to a greyscale mosaic witha lateral resolution of 1 m and dynamic mosaic rangeof − 41.2 to 15 dB. The compensated backscatter imagery,normalized at reference levels on flat seafloor, provide
relative intensities for qualitative analyses. We define
three characteristic backscatter imagery facies at theseafloor within our survey area: (1) a heterogeneous,moderate reflective (grey; − 20 to 10 dB) facies (“backscatter facies 1”), (2) a homogeneous, weak-to-moderate
reflective (dark grey; − 40 to − 20 dB) facies (“backscatterfacies 2”), and (3) a homogeneous, moderate-to-strong
reflective (light grey; − 20 to 10 dB) facies (“backscatterfacies 3”) (Fig. 4).
The region of backscatter facies 1 is associated with the
Okinose sand bank. Notably, the moderate backscatter(See figure on next page.)
Fig. 5 Close-up of bathymetry and backscatter imageries across the Okinose sand bank. (a) Bathymetry data image the morphology of the sandbank, with characteristic ENE–WSW trending lineation fabrics interpreted to be sand wave bedforms. (b) Elevation of bedforms across i–ii (left),and iii–iv (right) shown in (a). The sand waves generally have steeper southern flanks. Note smaller-scale waves in southern portion of the dune. (c)Backscatter intensities detect the morphology of individual sand waves (blue arrows: example of larger-scale sand waves, yellow arrows: example
of smaller-scale sand waves). The backscatter imagery provides fine details of the small-scale sand waves that extends to the southern flanksof the dune complex. Note occasionally patchy distributions. Black box: example of heterogeneous, high backscatter intensities not related to sandwavesHamahashi et al. Progress in Earth and Planetary Science

34.56°(2024) 11:8

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a

0

500 m

15
m
70

Sand waves

iiii

34.55°

34.54°

ii

iv

28
30
32
34
36
38
40

34.56°

135.09°135.08°

i

ii

300 m

Water depth (m)bWater depth (m)135.07°30
32135.1°

iii

iv

34
36
38

c

Smaller waves
300 m

0

500 m

15
dB
-41.2

34.55°

Larger-scale
sand waves

34.54°

Smaller-scale
sand waves

135.07°
Fig. 5 (See legend on previous page.)

135.08°135.09°135.1°

Hamahashi et al. Progress in Earth and Planetary Science(2024) 11:8

reflectivity extends east and southeast of the sand bank(Fig. 4). The sand wave bedforms and morphology of thedune complex are also clearly detected in the backscatter
imagery (Fig. 5c). The backscatter over the sand waves isgenerally weak (dark grey to grey), whereas low backscatter level (dark grey) is observed on the sand wave crests,
and strong levels (light grey) in the flanks and troughs,
with occasionally patchy distributions (Fig. 5c). Patchesof heterogenous, high backscatter intensities (light grey)
not associated with sand waves are also present (Fig. 5c).
Across the local depressions northeast of the Okinose
sand bank, patches of low backscatter intensities (darkgrey) are observed, indicating a complex distribution ofseafloor sediments and erosion (Figs. 4, 5c).
The crest of the southern escarpment of the Okinose
sand bank is visible from the lineations of low backscatter levels (dark grey) (Fig. 4). At the southern flank of
the Okinose bank, the water depths gradually deepen
from ~ 35 to ~ 40 m, and a sharp transition from backscatter facies 1 to facies 2 is observed at the foot of thedune (Fig. 4). The seafloor south of the Okinose bankcorresponds to the region of backscatter facies 2, which
is distributed broadly and has an overall smooth bathymetry (< 55 m water depth) (Fig. 4).
Westward from the region of facies 1 and 2, a sharptransition from backscatter facies 2 to facies 3 is observed
from east to west at our northern and central survey
areas (Fig. 4). Backscatter facies 3 is associated with thegradual increase of water depths (> 60 m) as the seafloor
forms a local embayment (Fig. 3b). At the southern area,
in contrast, the boundary between backscatter facies
2 and 3 is gradually curved westward along the southern rim of the embayment as the water depths shallow(Fig. 4). Here, backscatter facies 2 is continuously distributed from east to west.
Generally, the intensities of acoustic signals backscattered from the seafloor indicate seafloor characteristics
such as interface roughness, acoustic impedance, andsurficial heterogeneity, related to seafloor compositionand small-scale topographies (e.g., Lamarche et al. 2011;
Yu et al. 2015). The increase in backscatter intensities
with grain size have been noted in previous studies (e.g.,
Goff et al. 2000; Collier and Brown 2005; Ferrini andFlood 2006), and is likely perceptible in this study. We
attribute the high backscatter intensity facies 1 and 3 torepresent coarser sediments such as sands and gravels,and lower backscatter intensity facies 2 to represent finer
sediments such as silt and clay.

Page 13 of 50

4.2 Seismic reflection data

The Osaka Bay Fault is inferred to run NE–SW acrossthe western-central part of the Osaka Bay (Fig. 2a). In
this section, we interpret the structures below the seafloor imaged by seismic reflection data obtained froma Mini-GI airgun (imaged to ~ 1000 ms, ~ 700 m below
the seafloor) and a Boomer (imaged to ~ 150 ms, ~ 50 to
80 m below the seafloor) as active sources (Sect. 3). Dueto challenges in seafloor multiples in the data and scattering of signals across higher-angle strata at greater
depths, we use the upper < 300 to 500 ms (~ 200 to 400 m
below the seafloor) to resolve for confidence in structural interpretations, and assume a constant velocityof 1500 m/s for depth conversion. Based on the bandpath width of 300–2000 Hz for Boomer and 50–450 Hz
for airgun (Sect. 3), we estimate the accuracy of depth
measurement for Boomer to range ~ 0.4 to 2.5 m, and ~ 3
to 17 m for airgun. In this section, we divide our survey
areas into the northern, central, and southern regions,and constrain the geometry of observed structures andstratal dips (Figs. 6, 7, 8, 9, 10, 11, 12, 13, 14, 15, 16, 17,
18, 19, 20, 20, 21). We trace distinct seismic reflectors
and assign seismic units based on stratal terminations(angular disconformities) and characteristics of seismicreflections (Figs. 6, 7, 8, 9, 10, 11, 12, 13, 14, 15, 16, 17,
18, 19). The area of greatest folding, itself a proxy for fault
shear at a particular stratigraphic level in growth strata
above the Osaka Bay Fault is mostly contained between
the anticlinal fold axis and the synclinal fold axis (Figs. 7,
8, 9, 10, 13, 14, 15, 16, 17). The widths of fold limbs in thisstudy are measured from the range of width between the
anticlinal fold axis and synclinal fold axis in the imaged
depths.Fence diagrams are constructed to observe thespatial distributions of seismic units and their geometries
across our study area (Fig. 21). To interpret the surface
expressions of structures and distribution of seafloor sediments, we combine our observations with bathymetry
and backscatter data (Sect. 4.1) (Figs. 3, 4). For the deeper
sections, we identify seismic horizons at the base of representative seismic reflectors that are traceable among
all the sections (Horizons C1–C5; Figs. 6, 7, 8, 9, 10, 11,
12, 13, 14, 15, 16, 17, 18, 19). In Sect. 5.1.2, the ages ofthese deeper horizons are estimated by comparing withthe age-correlated reflectors of marine clay marker beds
by Grothe et al. (2014) along previously acquired deep
seismic lines (Yokokura et al. 1998), OD-B (Iwasaki et al.
1994; Yokokura et al. 1998; Iwabuchi et al. 2000).

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Fig. 6 Seismic profile Line 1 (Boomer). TWT: two-way travel time. Vertical line represents location of cross lines. (a) Uninterpreted image (post-stack,vertical exaggeration = 20). Red triangle: representative seafloor multiple. (b) Trace of seismic reflectors in (a) and interpreted image. Star shows
location of sediment thickness measurement in Fig. 24. Unit A1 exhibits chaotic internal structure with a wavy and irregular base, indicating
an erosional contact (CMP #3000–3800). Continuous tilting of strata by the hanging wall anticline of the Osaka Bay Fault is observed to reachto shallow depths

Fig. 7 Seismic profile Line 2 (Airgun). TWT: two-way travel time. Vertical line represents location of cross lines. (a) Uninterpreted image (post-stack,vertical exaggeration = 20). White triangle: representative seafloor multiple. (b) Trace of seismic reflectors in (a) and interpreted image. The seismic
units at deeper sections (light to dark pink) are assigned at the base of representative seismic reflectors traceable with other sections. Bluedotted line: anticlinal axial surface. Yellow dotted line: synclinal axial surface. Dip measurements of representative layers are shown. Star shows
location of sediment thickness measurement in Fig. 24. The base of Unit A1 is wavy and irregular, indicating an erosional contact (CMP #120–260).
Continuous tilting of strata by the hanging wall anticline of the Osaka Bay Fault is observed to reach to shallow depths

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Fig. 8 Seismic profile Line 3 (Airgun). TWT: two-way travel time. Vertical line represents location of cross lines. (a) Uninterpreted image (post-stack,vertical exaggeration = 20). White triangle: representative seafloor multiple. Note the presence of high-amplitude negative polarity reflector
at shallow depths (CMP #180–270). (b) Trace of seismic reflectors in (a) and interpreted image. The seismic units at deeper sections (light to darkpink) are assigned at the base of representative seismic reflectors traceable with other sections. Blue dotted line: anticlinal axial surface. Yellow
dotted line: synclinal axial surface. Dip measurements of representative layers are shown. Star shows location of sediment thickness measurementin Fig. 24. Continuous tilting of strata by the hanging wall anticline of the Osaka Bay Fault is observed to reach to shallow depths

4.2.1 Northern survey area

At our northern survey area, our E–W, NE–SW, and
NW–SE sections that cross oblique to high-angle to thestructure of the Osaka Bay Fault show a monoclinal panel
consisting of a coupled anticline and syncline (Figs. 3b, 6,
7, 8, 9, 10). We observe a distinct fold axis of the hanging
wall anticline buried beneath recent sediments above the
Osaka Bay Fault. The anticlinal axial surface, or growth
axial surface, dips 12–25° to the west, separating shallowly-dipping reflectors to the west from a 0.4–13° eastdipping panel which comprises the ~ 0.3–2.0-km-wide
forelimb to the east (Figs. 6, 7, 8, 9, 10). The anticline
has a 0.2–0.9° west-dipping backlimb, which dips gentler
than the forelimb (Figs. 6, 7, 8, 9, 10). The synclinal foldaxial surface, in contrast, dips ~ 84–85° to the east along
Lines 2 and 3, and dips ~ 75° to the west along Lines 4 and5, and separates the forelimb to the west from shallowlydipping reflectors to the east (Figs. 7, 8, 9, 10). The seismicreflectors are continuously folded across these fold axes,
thinning the strata across the fold, with the forelimb andbacklimb of the anticline progressively steepening withdepth, and the overall monocline widening at shallower

depths. It is noted that Lines 1–3 cross at high-angle tothe structure of the Osaka Bay Fault, reflecting the geometry of the fault-normal fold and its greater synclinal foldaxial surface (84–85°), compared to Lines 4 and 5, whichcross oblique to the strike of the fault (Fig. 2a).
At our northern sections, the uppermost sequence
consists of a 7–14-m-thick seismically transparent
package with occasional chaotic fabric, associated withthe local topographic high (Okinose sand bank) imaged
from bathymetry data (Sect. 4.1), which we assign
as Unit A1 (Figs. 6, 7, 8, 9, 10). The base of Unit A1 iserosive along the northern edge of the Okinose bank(Lines 1 and 2) (Figs. 6, 7), whereas continuous deposition is observed at the center of the dune (Lines 3–5)
(Figs. 8, 9, 10).
Along Lines 3–5, the section (up to ~ 11 m thick)beneath Unit A1 consists of parallel reflectors contrasting with the less-reflective Unit A1, which we assign asUnit A1’ (Figs. 8, 9, 10). Between this transition, a localized section of high-amplitude negative polarity reflector is observed, indicating the presence of a low velocity

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Fig. 9 Seismic profile Line 4 (Airgun). TWT: two-way travel time. Vertical line represents location of cross lines. (a) Uninterpreted image (post-stack,vertical exaggeration = 20). White triangle: representative seafloor multiple. (b) Trace of seismic reflectors in (a) and interpreted image. The seismic
units at deeper sections (light to dark pink) are assigned at the base of representative seismic reflectors traceable with other sections. Blue dotted
line: anticlinal axial surface. Yellow dotted line: synclinal axial surface. Dip measurements of representative layers are shown. Star shows locationof sediment thickness measurement in Fig. 24. Units A1, A1’ and A2 onlap onto Unit B1, underlain by continuous tilting of strata by the hanging wall
anticline of the Osaka Bay Fault

material relative to the surrounding strata (Lines 3, 5, 12;
Figs. 8, 10, 11).
Unit A1’ is underlain by a 13–16-m-thick sequence ofthin, well-defined parallel seismic reflections, which we
assign as Unit A2 (Figs. 6, 7, 8, 9, 10). Beneath this section, < 5 m-thick sequences of continuous, high-amplitude reflections are present, which we assign as Unit
B1. At the base of Unit B1, sequences of discontinuous
and wavy reflectors are observed, contrasting with thesequence above, which we classify as Unit B2 (Figs. 6, 7,
8, 9, 10). At the deeper section of Unit B2, well-definedcontinuous reflectors are present. Towards the westernside of our sections, Units A1, A2 and B1 gradually thin,and Unit B2 is exposed at the seafloor, as the water depth
increases west of the Okinose bank (Figs. 9, 10).
Along Lines 1 and 2, the base of Unit A1 is wavy andirregular, indicating an erosional contact (Figs. 6, 7).

Here, Unit A2 is eroded at the base of Unit A1, and Unit
A1’ is absent or only partially distributed. Erosion is mostsignificant across the bathymetric depression east of the
Okinose bank, and the uppermost sequence (10–11 m) at
the eastern side of Line 1 exhibits chaotic internal structure, overlying the < 20-m-thick sub-horizontal Units A1’
and A2 (Line 1; Fig. 6).
Our N-S sections (Lines 12, 13) show that Unit A1 terminates south of the Okinose bank, and a 13–14-m-thicksequence of thin, well-defined parallel seismic reflections,which we assign as Unit A1’, is exposed at the seafloor(Figs. 11, 12). The sub-horizontal base of Unit A1’ underlain by Unit A2 is well-imaged as clear seismic reflectors
from north to south. Here, Units A2, B1, and B2 are continuously distributed below Units A1 and A1’ (Figs. 11,
12). From the fence diagram crossing our N-S and E–W
sections, these units are spatially correlated (Fig. 21;

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Fig. 10 Seismic profile Line 5 (Airgun). TWT: two-way travel time. Vertical line represents location of cross lines. (a) Uninterpreted image (post-stack,vertical exaggeration = 20). White triangle: representative seafloor multiple. Note the presence of high-amplitude negative polarity reflector
at shallow depths (CMP #1–20). (b) Trace of seismic reflectors in (a) and interpreted image. The seismic units at deeper sections (light to dark pink)
are assigned at the base of representative seismic reflectors traceable with other sections. Blue dotted line: anticlinal axial surface. Yellow dotted
line: synclinal axial surface. Dip measurements of representative layers are shown. Star shows location of sediment thickness measurement in Fig. 24.
Units A1’ and A2 onlap onto Unit B1, underlain by continuous tilting of strata by the hanging wall anticline of the Osaka Bay Fault

Additional file 1: Fig. S1). The top of Unit A1’ beneath the
Okinose bank (the boundary between A1 and A1’), however, is discontinuous and likely erosive (Figs. 6, 7, 8, 9, 10).
Our NW–SE sections (Lines 1, 2, 3) cross at highangle to the structure of the Osaka Bay Fault, andnearly parallel to the long axis of the Okinose bankand to the strike of the shelf northeast of our sections(Figs. 3b, 6, 7, 8). Here, continuous tilting of strata ofUnits A1, A1’, A2, B1, and B2 are observed from near
the seafloor, revealing the shallow deformation by uplifton the hanging wall anticline forelimb above the Osaka
Bay Fault. We measure 0.5–3.2° east-dipping strata andshallowly-dipping strata separated by a synclinal axial
surface within these units, associated with distinct thinning of sediments in the hanging wall and thickening in

the footwall (Figs. 6, 7, 8). Here, the uplifted stratigraphy in the hanging wall is preserved from erosion, andthe continuous deformation by the anticline to the shallow subsurface is clearly observed.
Our E–W and NE–SW sections (Lines 4, 5), by contrast, cross at an oblique angle to the structure of the
Osaka Bay Fault, and also oblique to the shelf and to
the long axis of the Okinose bank (Figs.3b, 9, 10). Fromthese sections, we observe progressive onlap and overlap of Units A1, A1’, and A2 onto Unit B1. Here, Units
A1 and A1’ dip sub-horizontally to the west, sub-parallel to the dip of the current seafloor (Figs. 9, 10). In contrast, tilting of strata are observed in Units A2, B1, andB2, which we interpret to represent the shallow deformation by the hanging wall anticline forelimb above

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Fig. 11 Seismic profile Line 12 (Airgun). TWT: two-way travel time. Vertical line represents location of cross lines. (a) Uninterpreted image(post-stack, vertical exaggeration = 20). White triangle: representative seafloor multiple. Note the presence of high-amplitude negative polarity
reflector at shallow depths (CMP #120–180). (b) Trace of seismic reflectors in (a) and interpreted image. The seismic units at deeper sections (lightto dark pink) are assigned at the base of representative seismic reflectors traceable with other sections

Fig. 12 Seismic profile Line 13 (Boomer). TWT: two-way travel time. Vertical line represents location of cross lines. (a) Uninterpreted image(post-stack, vertical exaggeration = 20). Red triangle: representative seafloor multiple. (b) Trace of seismic reflectors in (a) and interpreted image(See figure on next page.)
Fig. 13 Seismic profile Line 6 (Airgun). TWT: two-way travel time. Vertical line represents location of cross lines. (a) Uninterpreted image (post-stack,vertical exaggeration = 20). White triangle: representative seafloor multiple. (b) Trace of seismic reflectors in (a) and interpreted image. The seismic
units at deeper sections (light to dark pink) are assigned at the base of representative seismic reflectors traceable with other sections. Bluedotted line: anticlinal axial surface. Yellow dotted line: synclinal axial surface. Dip measurements of representative layers are shown. Star shows
location of sediment thickness measurement in Fig. 24. Note the onlap of Units A1’ and A2 onto Unit B1, underlain by continuous tilting of strata
by the hanging wall anticline of the Osaka Bay Fault. Unit A1’ is also likely truncated by Unit A2 (CMP #100–160)

Hamahashi et al. Progress in Earth and Planetary ScienceFig. 13 (See legend on previous page.)(2024) 11:8

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Fig. 14 Seismic profile Line 7 (Boomer). TWT: two-way travel time. Vertical line represents location of cross lines. (a) Uninterpreted image(post-stack, vertical exaggeration = 20). Red triangle: representative seafloor multiple. (b) Trace of seismic reflectors in (a) and interpreted
image. Blue dotted line: estimated anticlinal axial surface. Yellow dotted line: estimated synclinal axial surface. Star shows location of sediment
thickness measurement in Fig. 24. Note the onlap of Units A1’ and A2 onto Unit B1, underlain by continuous tilting of strata by the hanging
wall anticline of the Osaka Bay Fault. An additional synclinal fold structure is observed west of the hanging wall anticline (CMP #400–1200). The
widths of the syncline and anticline reported in the main text are estimated from the width of the two points where the folded strata intersectwith the inferred undeformed stratal dip (Additional file 1: Figs. S2, S3)

the Osaka Bay Fault (Figs. 9, 10). We measure 0.4°–3.4°
east-dipping strata and shallowly-dipping strata separated by a synclinal axial surface within these units,
associated with distinct thickening of sediments fromthe hanging wall to the footwall. The onlap of Units A1,
A1’ and A2 onto Unit B1 infer major sequence stratigraphic surfaces. At the deeper section of Unit B2,
folded strata are continuously observed in the hanging
wall and footwall.
4.2.2 Central survey area

At our central survey area, our NW–SE sections (Lines
6–10) cross at oblique to high-angle to the structure of
the Osaka Bay Fault, and nearly perpendicular to theshelf east of our sections (Figs. 3b, 13, 14, 15, 16, 17).
Our seismic sections show a monoclinal panel boundedby synclinal and anticlinal axial surfaces beneath the seafloor. We observe a distinct fold axis of the hanging wall
anticline above the Osaka Bay Fault, separating a ~ 0.2 to15° east-dipping panel which comprise the ~ 0.1–3.7-kmwide forelimb to the east, from shallowly-dipping reflectors to the west (Figs. 13, 15, 16). This anticlinal axial
surface dips 13–25° to the west. Notably, the dip of theforelimb is gentler in the south (0.2–5.2°) (along Lines 8
and 9) (Figs. 15, 16), compared to the north (1–15°) (Line6) (Figs. 13). Towards the eastern side on the footwall,
our sections (Lines 7, 10) image the strata gradually tilting 0.2°–0.8° to the west, forming a broad syncline witha gentler dip than the hanging wall anticline (Figs. 14,
17). This synclinal fold axis dips ~ 89° to the west along
Line 6, whereas it dips 85°–87° to the east along Lines 8
and 9, separating shallowly-dipping reflectors to the east

from the forelimb to the west (Figs. 15, 16). Our sections also image the 0.4°–1.4° west-dipping backlimbof the hanging wall anticline above the Osaka Bay Fault(Figs. 15, 16). The seismic reflectors are continuously
folded across these fold axes, thinning the hanging wall
strata across the fold and with the dips of tilted strata in
the forelimb and backlimb progressively steepening withdepth (Figs. 13, 14, 15, 16, 17). The deeper sections westof the backlimb dip east, forming an additional synclinal
fold structure buried by overlying sediments along Line
7 (Fig. 14). This structure may relate to the subsidiaryfault between the Kariya and Osaka Bay Faults reported
by previous studies (Yokokura et al. 1998; Iwabuchi et al.
2000) (Fig. 2a). This eastern synclinal fold axis dips ~ 8.6°
to the west, and separates shallowly-dipping reflectors
to the east from ~ 0.8° east-dipping panel to the west
(Fig. 14). Overall, our sections show that the monocline
above the Osaka Bay Fault at shallow depths consistsof a coupled ~ 4.4 to 4.5 km wide anticline bounded bya broad ~ 11 km wide syncline to the east, and a narrow ~ 2.6 km wide syncline to the west (Figs. 14, 17; Additional file 1: Figs. S2, S3).
The uppermost ~ 31-m-thick sequences in the footwall are on-lapping the underlying tilted strata; thesesequences are deposited only on the eastern side ofour sections (Figs. 13, 14, 15, 16, 17). Here, Unit A1 isabsent, and the observed sequences are characterized
by thin, parallel reflectors similar to Units A1’ and A2
at the northern survey area, and the bases of Units A1’
and A2 are correlated with our N-S sections (Figs. 12,
13, 14). However, the top of Unit A1’ is not correlated from north to south, and Unit A1’ at the central

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Fig. 15 Seismic profile Line 8 (Airgun). TWT: two-way travel time. Vertical line represents location of cross lines. (a) Uninterpreted image (post-stack,vertical exaggeration = 20). White triangle: representative seafloor multiple. (b) Trace of seismic reflectors in (a) and interpreted image. The seismic
units at deeper sections (light to dark pink) are assigned at the base of representative seismic reflectors traceable with other sections. Bluedotted line: anticlinal axial surface. Yellow dotted line: synclinal axial surface. Dip measurements of representative layers are shown. Star shows
location of sediment thickness measurement in Fig. 24. Note the onlap of Units A1’ and A2 onto Unit B1, underlain by continuous tilting of strata
by the hanging wall anticline of the Osaka Bay Fault

survey area is 16–22-m thick, deposited to the surface (Figs. 12, 13, 14, 15, 16, 17). Unit A1’ is relatively
undeformed, onlapping and truncated by the underlying 9–15-m-thick sequence of Unit A2, indicating someerosion (Figs. 13, 17). The strata of Unit A1’ dip ~ 0.1°

to the west, sub-parallel to the dip of the current seafloor (Figs. 13, 14, 15, 16, 17). In contrast, we measure
0.06–0.2° east-dipping strata and ~ 0.1° west-dipping
strata separated by a synclinal axial surface within Unit
A2 along Lines 6, 8, 9, and 10, which we attribute toHamahashi et al. Progress in Earth and Planetary Science(2024) 11:8

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the shallow deformation via the hanging wall anticline
forelimb above the Osaka Bay Fault (Figs. 13, 15, 16,17). We interpret the observed stratal termination andtruncation of Unit A2 below Unit A1’ to represent anerosional sequence stratigraphic surface. Near the eastern edge of our sections, an additional downlap surface
may also be present at the shallow portion of Unit A1’,
although the imaging is not clear (Figs. 14, 17).
At the western side of our sections in the hanging
wall, the 4–11-m-thick tilted sequence below Unit A2
is exposed at the seafloor, and is on-lapped by Units A1’
and A2 (Figs. 13, 14, 15, 16, 17). This sequence is characterized by well-defined high-amplitude reflections, which
we assign as Unit B1. We measure 0.2–1° east-dipping
strata and ~ 0.14° west-dipping strata separated by a synclinal axial surface within Unit B1 on the hanging wall
anticline forelimb (Figs. 13, 14, 15, 16, 17). The stratal termination and onlap of Units A1’ and A2 above Unit B1
infer a major sequence boundary. Notably, the geometryof the boundary between Units A2 and B1 is wavy andirregular, indicating an erosional contact (Figs. 13, 14, 15,
16, 17). Here, thinning of strata in the hanging wall andthickening in the footwall within Unit B1 are obscured
due to erosion in the footwall.
Our E–W sections and bathymetry data show a gradualwestward increase in water depth, in a direction nearly
perpendicular to the shelf to the east, with a deeper
bathymetry in the hanging wall and shallower bathymetry in the footwall of the Osaka Bay Fault (Figs. 3b, 13,
14, 15, 16, 17). The shallower water depth in the easternregions coincides with the deposition of Units A1’ and A2
on the footwall side, which are absent or thinner in thewestern regions on the hanging wall. The observed variations in sediment thickness may partly be explained byenhanced erosion due to uplift by the fault in the hanging
wall and larger sediment accommodation in the footwall
due to subsidence. Furthermore, the clear difference instratal dips observed between the folded strata (e.g. Units
A2 and B1) and the shallow strata (Unit A1’) sub-parallelto the current dip of the seafloor, suggest additional processes at shallow depths (Figs. 13, 17). Here, additional,
non-tectonic surface processes such as tidal currents
operating in the bay and sea-level fluctuations across theshelf, are also likely playing a major role in creating thevariations in sediment thickness (Sect. 5.2).
Beneath Unit B1, sequence of wavy and discontinuous, irregular reflections is present, which we assign asUnit B2, underlain by well-defined continuous reflections at the deeper section (Figs. 13, 14, 15, 16, 17).
Along Lines 6–8, Unit B1 is eroded at the seafloor,
exposing Unit B2 on the hanging wall (Figs. 13, 14, 15).
By contrast, south of these sections along Lines 9 and
10, Unit B1 is preserved at the seafloor, overlying Unit
B2 (Figs. 16, 17). The contact between Units B1 and B2
is also observed in our N-S sections (Lines 15 and 16) at
the hanging wall side, where the southern sub-horizontal sequence of Unit B1 is underlain by ~ 0.2° southwarddipping sequence of Unit B2, associated with the overall southward thickening of strata (Figs. 18, Additional
file 1: Fig. S4). At the deeper sections of Unit B2, thesouthward tilting of strata (~ 0.6°) becomes progressively steeper with depth, indicating that these structures are tectonically driven, showing an increase in
uplift to the north (Figs. 18, Additional file 1: Fig. S4).
The on-lap of Unit B1 onto Unit B2 observed from the
N-S section may suggest a significant age gap between
Units B1 and B2 (Fig. 18). The southward thickening of strata is coincident with the gradual southward
shallowing in water depth shown by bathymetry; thenorthward dip of the seafloor is contrary to the southward dip of the underlying strata, indicating additionalvariations in surface processes (e.g., erosion by tidal
currents) (Figs. 3b, 18). The deepest regions coincide
where Unit B1 is eroded at the seafloor, exposing Unit
B2, whereas the shallower region preserves Unit B1 at
the seafloor (Figs. 3b, 18).
4.2.3 Southern survey area

Our E–W section (Line 11) at our southernmost survey
area is located where the bathymetry shallows southwardof the northern embayment, at an oblique angle to theinferred structure of the Osaka Bay Fault (Figs. 3b, 19).
Units A1’, A2, B1, and B2 have similar seismic characteristics as those in the central and northern survey areas,
and are traceable along our N-S section (Lines 13 and 14)(Figs. 12, 18). Notably, the boundary between Units B1(See figure on next page.)
Fig. 16 Seismic profile Line 9 (Airgun). TWT: two-way travel time. Vertical line represents location of cross lines. (a) Uninterpreted image (post-stack,vertical exaggeration = 20). White triangle: representative seafloor multiple. (b) Trace of seismic reflectors in (a) and interpreted image. The seismic
units at deeper sections (light to dark pink) are assigned at the base of representative seismic reflectors traceable with other sections. Bluedotted line: anticlinal axial surface. Yellow dotted line: synclinal axial surface. Dip measurements of representative layers are shown. Star shows
location of sediment thickness measurement in Fig. 24. Note the onlap of Units A1’ and A2 onto Unit B1, underlain by continuous tilting of strata
by the hanging wall anticline of the Osaka Bay Fault

Hamahashi et al. Progress in Earth and Planetary ScienceFig. 16 (See legend on previous page.)(2024) 11:8

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Hamahashi et al. Progress in Earth and Planetary Science(2024) 11:8

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Fig. 17 Seismic profile Line 10 (Boomer). TWT: two-way travel time. Vertical line represents location of cross lines. (a) Uninterpreted image(post-stack, vertical exaggeration = 20). (b) Trace of seismic reflectors in (a) and interpreted image. The seismic units at deeper sections (light to darkpink) are assigned at the base of representative seismic reflectors traceable with other sections. Blue dotted line: estimated anticlinal axial surface.
Yellow dotted line: estimated synclinal axial surface. Star shows location of sediment thickness measurement in Fig. 24. Note the onlap of Units
A1’ and A2 onto Unit B1, underlain by continuous tilting of strata by the hanging wall anticline of the Osaka Bay Fault. An additional synclinal foldstructure is observed west of the hanging wall anticline (CMP #3500–3700). The widths of the syncline and anticline reported in the main text are
estimated from the width of the two points where the folded strata intersect with the inferred undeformed stratal dip (Additional file 1: Figs. S2, S3)

Fig. 18 Seismic profile Line 14 (Boomer). TWT: two-way travel time. Vertical line represents location of cross lines. (a) Uninterpreted image(post-stack, vertical exaggeration = 20). Red triangle: representative seafloor multiple. (b) Trace of seismic reflectors in (a) and interpreted image. The
seismic units at deeper sections (light to dark pink) are assigned at the base of representative seismic reflectors traceable with other sections. Notethat the strata dips and thickens southward, with an increase of tilting at greater depths, suggesting tectonic deformation

and A2 is wavy, indicating an erosional contact (Fig. 19).
Here, sediments are overall thinner compared to thenorth, but Units A1’ and A2 (13–23-m thick in total) arecontinuously deposited on top of Unit B1 (6–11-m thick).
Unlike other sections to the north, Unit B1 is not exposedat the seafloor, and the tilting of strata is less distinct atshallow depths (Fig. 19). We interpret the tilting of strata

in the deeper section beneath Unit B2 to represent the
deformation by the hanging wall anticline.
4.2.4 Comparison of seismic facies with backscatter data

When combining seismic reflection and backscatter data,
the seismic unit boundaries at the seafloor constrained
from the seismic sections are coincident with theHamahashi et al. Progress in Earth and Planetary Science(2024) 11:8

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Fig. 19 Seismic profile Line 11 (Boomer). TWT: two-way travel time. Vertical line represents location of cross lines. (a) Uninterpreted image(post-stack, vertical exaggeration = 20). (b) Trace of seismic reflectors in (a) and interpreted image. Star shows location of sediment thickness
measurement in Fig. 24

observed transitions in backscatter intensities (Sect. 4.1)
(Fig. 4). Units A1 correspond to the heterogenous, moderate backscatter intensity facies 1, whereas Units A1’
and B1 correspond to the homogeneous, weak-to-moderate backscatter intensity facies 2, and Unit B2 to thehomogeneous moderate-to-high backscatter intensity
facies 3 (Sect. 4.1, Fig. 4). These comparisons suggest that
Units A1 and B2 consist of coarser-grained sediments
such as sands and gravels, and Units A1’ and B1 likely
consist of finer-grained sediments such as silt and clay.
The lithostratigraphy of seismic units is further discussed
in Sect. 5.1.
4.2.5 Summary of observed deformation structuresof the Osaka Bay Fault

Our seismic sections are all consistent with the presenceof an asymmetric anticline above the Osaka Bay Fault
imaged at shallow depths (50–400 m below the seafloor),
characterized by a ~ 0.1 to 3.7-km-wide, steep forelimbthat is dipping 0.2°–15° east and a gentler and broader
backlimb that is dipping 0.2°–1.4° west, associated withan additional ~ 2.6-km-wide syncline at the westernextent of the anticline (Figs. 6, 7, 8, 9, 10, 11, 12, 13, 14,
15, 16, 17, 18, 19, 20, 20, 21). Towards the eastern sideof our sections, the tilted strata dips 0.2–0.8° gently west,
forming a broad ~ 11-km-wide syncline in the footwall.
Our northern, central, and southern sections clearlyshow that the dip of the anticline forelimb is steeper in
the northern area (0.4°–15°) with larger offsets across thefold, compared to that of the south (0.2°–5.2°) (Fig. 20).
The southward tilting and thickening of strata observedin our N-S sections on the hanging wall to the west, also

indicate larger amounts of uplift to the north (Fig. 18).
The observed north–south variations in the width andgeometry of the folds indicate along-strike variationsin the amount of displacement of the underlying Osaka
Bay Fault and its slip rate, and are further discussed in
Sect. 5.2.
Our seismic data show no presence of fault plane reflections at our imaged depths, and confirm that the Osaka
Bay Fault is blind. The deformation, however, reaches toand near surface by an asymmetric fold that grows progressively steeper with depth, showing clear evidenceof limb rotation and fanning of limb dips (Fig. 20). The
lengths of the limbs increase with depth, forming growth
triangles between the anticlinal and synclinal fold axes,
characterized by upward narrowing kink bands, whereshear occurs within the triangular zone (“trishear”) (e.g.,Erslev 1991; Hardy and Ford 1997; Cardozo et al. 2011;
Allmendinger 1998) (Fig. 20). The shear along the fault is
not localized onto a narrow zone that is meters to tensof meters thick, but rather, shear is spread out acrossthe entire forelimb in the trishear envelope, reaching tonear the surface. An asymmetric fold commonly develops in the hanging wall of a thrust fault that is propagating towards the surface (e.g. Jamison 1987; McClay 1992;
Almeida et al. 2018). We observe slip on the fold decreasing towards the surface, which is typical for a trishear
fault-propagation fold that is continuously buried by sediments during its uplift history (e.g., Erslev 1991; Hardy
and Ford 1997; Cardozo et al. 2011; Allmendinger 1998).
At the bottom of the sedimentary basin (~ 2 km depths),
previous studies image the Osaka Bay Fault offsetting the
pre-growth strata of the granitic basement of the RyokeHamahashi et al. Progress in Earth and Planetary Science(2024) 11:8

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Fig. 20 Overview of E–W sections acquired in this study. Left: uninterpreted image. Right: interpreted image. Vertical exaggeration = 20. Notethe northward increase in vertical displacement across the Osaka Bay Fault

Hamahashi et al. Progress in Earth and Planetary Science(2024) 11:8

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Fig. 21 Fence diagram of all seismic sections acquired in this study. Vertical exaggeration = 20. (a) Uninterpreted image. (b) Interpreted image. (c)Interpreted image excluding Lines 7, 10, and 11

Hamahashi et al. Progress in Earth and Planetary Science(2024) 11:8

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Group (e.g., Sato et al. 1998; Yokokura et al. 1998), indicating the presence of a fault plane extending to deeper
depths, with the fault tip estimated to exist at 1.5–4.5 km
based on trishear modelling (Grothe et al. 2014). We
interpret that the observed fold is a fault-propagation
fold, and that the Osaka Bay Fault has accrued enoughslip to generate this entire fold panel at its fault tip.
The seismic data from both our study and previous
studies do not reveal clear stratigraphic cutoffs acrossthe Osaka Bay Fault and measurable fault-dip angles.
We measure fold axial surfaces that characterize theshear occurring within the growth triangle bounded by
the anticlinal axial surface (growth axial surface) dipping ~ 12 to 25° to the west, and the synclinal axial surface dipping ~ 75 to 89° (to the west along Lines 4, 5, 6,
and to the east along Lines 2, 3, 8, 9), along which shear ismost concentrated throughout the stratigraphic column(Fig. 20; e.g., Hardy and Ford 1997; Cardozo et al. 2011).
The more gently inclined anticlinal fold axis compared tothe synclinal fold axis is consistent with the growth axial
surface in the growth triangle above the blind thrust but
at shallower levels (Grothe et al. 2014). It likely records
the interplay between the migrating edge of the sheared
trishear envelope and deposition of sediments depositedas the fold grows upward through the stratigraphic section (Grothe et al. 2014). The anticlinal fold axis is thus
not necessarily a planar surface, as rates of sedimentation can vary its geometry; for example, faster sedimentation would tend to steepen the anticlinal axial surface
and vice versa. It also marks the change in thickness ofthe growth strata in the trishear envelope, as predictedby fault related fold theory for trishear fault propagation
folds (e.g., Allmendinger 1998; Grothe et al. 2014).
The synclinal axial surface marks the edge of the trishear envelope on the footwall side of the fault, but does
not indicate the actual dip of the thrust per se (Grothe
et al. 2014). The synclinal axial surface at shallow depths
measured in this study ranges 75°–85° (84°–85° along
Lines 1–3, and 75° along Lines 4 and 5) and 85°–89° in
our northern (Figs. 6–10) and central (Figs. 13, 14, 15,
16, 17) survey areas, respectively, steeper than the synclinal axial surface measured in previous studies (~ 72°)
at greater depths (e.g. Sato et al. 1998; Grothe et al. 2014).
The steeper dips of the synclinal axial surface at shallow depths are associated with an upward propagating
growth triangle with ongoing sedimentation (e.g., Hardy
and Ford 1997; Cardozo et al. 2011). The presence of abacklimb on the fault-related fold of the Osaka Bay Fault
may relate to the steepening of the trishear envelope(e.g., Seeber and Sorlien 2000). The fact that the backlimb steepens progressively with depth is also indicativeof a curved thrust fault (e.g., Seeber and Sorlien 2000).

The shallow part of the fault-structure documented byour study in previously unavailable detail, provides new
insights on the geometry of subsurface deformation
caused by the Osaka Bay Fault.
Our sections show that the strata thin across the folds
and thicken in the footwall, indicating that these units
represent syn-tectonic growth stratigraphy (i.e., deposition during deformation) (e.g., Hardy and Ford 1997;
Cardozo et al. 2011; Grothe et al. 2014) (Figs. 6, 7, 8, 9, 10,
11, 12, 13, 14, 15, 16, 17, 18, 19, 20, 20, 21). The sedimentsin the growth strata have continuously buried the fold
scarps formed by slip on the Osaka Bay Fault, recording
the amount of relative uplift as the difference in stratigraphic thickness between the uplifted fold crest in the
hanging wall and the subsided trough in the footwall.
The ages of growth strata therefore define the timing of
deformations. In Sect. 5.2, we measure the differences instratigraphic thicknesses and combine age constraints
from nearby borehole data to discuss the recent slip history of the Osaka Bay Fault.
The geometries of growth structures are controlled
primarily by folding mechanism and the relative rates ofsedimentation and tectonic uplift (e.g., Hardy and Ford
1997; Cardozo et al. 2011). From our seismic sections,
we observe two contrasting features of growth strata; (1)progressive folding and thinning of recent strata acrossthe fold, and (2) distinct on-lap of recent strata concentrated on the footwall above tilted strata (Figs. 6, 7, 8,
9, 10, 11, 12, 13, 14, 15, 16, 17, 18, 19, 20, 20, 21). These
structural contrasts are either showing variations in rates
of sedimentation and/or uplift within the bay. We investigate these processes and interpret the sequence stratigraphy in Sects. 6 and 7.
Our seismic sections and bathymetry data overall show
no clear surface expression of the active fold scarps of the
Osaka Bay Fault (Figs. 6, 7, 8, 9, 10, 11, 12, 13, 14, 15, 16,
17, 18, 19, 20, 20, 21). The slope of the seafloor is likely
modified by various non-tectonic surface processes (erosional and depositional) such as by tidal currents. We
observe along-strike variation of sediment thicknessof recent sediments on the hanging wall and footwall,
revealing horizons where the erosional surface of the foldis exposed at the seafloor, and horizons where the fold isburied by recent sediments (Figs. 20, 21). This trend isassociated with the eastward and southward shallowingof water depth, and deeper water depths at the northwest
region (Fig. 3b). The along-strike variation of sediment
thickness is inconsistent with the northward increase
in uplift on the hanging wall (Fig. 20, Sect. 5.2), which
would predict a northward decrease in sediment thickness, clearly illustrating the imbalance between uplift anderosion.

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Fig. 22 Seismic Lines 7 and 8 combined with lithology previously reported at borehole sites #57-30, 56-9, and 56-2 near Kansai Airport, ~ 2.3 km
east of Line 7. Boreholes are placed in their reported distance and depths. Lithology bar is modified after Nakaseko (1987). Ages (ka) are basedon tephra and pollen assemblages (Nakaseko 1987). Ma: marine clay. The bases of Ma 13 and 11 correspond to the bases of the Holocene alluviumand Late Pleistocene colluvium, respectively. Units A1’, A2, and B1 defined in this study are interpreted to represent the Holocene alluvium (Ma13),and Unit B2, C1, and C2 to comprise the Late Pleistocene colluvium. Note that the correlation of the base of C3 with Ma 11 is speculative, and may
not necessarily coincide (Sect. 5.1.2). The red-dotted line is an extrapolation of the bedding geometry indicated from the seismic section

5 Discussion
5.1 Interpretation of stratigraphy and age

To investigate the lithostratigraphy and provide age constraints for the structures imaged by our seismic reflection data, we compare our seismic survey with publishedborehole data acquired from offshore Kobe city andKansai Airport, and along the coastal plain of the Osaka
Basin (Nakaseko 1987; Masuda 1992; Masuda et al. 2000;
Masuda and Miyahara 2000; Nanayama et al. 2000, 2001;
Kitada et al. 2001; Masuda and Itomoto 2015) (Figs. 2a,22, 23). Although our survey lines do not directly crossover the boreholes, we are able to compare geologicalinformation and infer sequence stratigraphy (e.g., Van
Wagoner et al. 1988; Catuneanu et al. 2009) with our seismic reflection and backscatter facies units and observed
stratal geometries.
We start by extending our sections to the nearest
boreholes (Sites 57–30, 56–23, 56–24, 56–25) offshore
Kansai Airport ~ 2.3 km east of our survey Lines 7 and
8 (Figs. 2a, 22; Nakaseko 1987; Masuda 1992). Based on
previous studies, the sedimentary sequence in the upper
30–35 m at these sites is primarily marine clay “Ma 13”of the Holocene alluvium (bearing the K-Ah tephra
layer; ~ 7.3 ka) interpreted to be estuary and inner bay
deposits, underlain by non-marine, sandy to gravely fluvial sequences of the Late Pleistocene colluvium (bearing the AT tephra layer; ~ 25 ka) (Fig. 22; Nakaseko

1987; Masuda 1992). Our E–W sections show that thesesediments thicken seaward and thin towards the coast(Fig. 22). We correlate Units A1’ and A2 observed fromour sections with the Holocene marine clay, and Unit B2
with the fluvial sands and gravels of the Late Pleistocene
colluvium (Fig. 22). We interpret Unit B1 to represent
a sandy to clayey section transitioning from fluvial tomarine facies across the Holocene–Late Pleistocene. Unit
B1 thins to the east and is not well preserved in boreholes from offshore Kansai Airport, but is recovered andconstrained by abundant microfossil 14C ages and pollen
assemblages from boreholes offshore Kobe (Sites OB-1
and OB-2) and coastal Osaka (Sites Suminoe-A and B) to
the north (Figs. 2a, 23; Nanayama et al. 2000, 2001; Masuda et al. 2000; Kitada et al. 2001; Masuda and Miyahara
2000).
At Sites OB-1 and OB-2, ~ 6.8 km north of our survey Line 1 (Fig. 2a), the upper fluvial sequence at thebase of marine clay, that we correlate with Unit B1, consists of normally-graded, interlayered silt and medium
to coarse sand with pebbles, and has reported 14C agesof 11.6–11.9 ka at Site OB-2 (Fig. 23; Nanayama et al.
2000). This sequence of Unit B1 is deposited up to ~ 10 mthick above the AT tephra layer (~ 25 ka), overlain by a
15–40 m thick coarser gravel sequence (Nanayama et al.
2000; Kitada et al. 2001), which we correlate to Unit B2(Fig. 23). The AT tephra layer is also observed to cap theHamahashi et al. Progress in Earth and Planetary Science(2024) 11:8

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Fig. 23 Lithology bar at boreholes sites OB-1 and OB-2 (offshore Kobe) and sites Suminoe-A and B (Osaka Basin), modified after Masuda et al.
(2000), Masuda and Miyahara (2000), Nanayama et al. (2000, 2001), and Kitada et al. (2001). Ages (yr) are based on these studies (red: tephra, black:radiocarbon dating, blue: pollen assemblage). The Holocene alluvium consist of a lower sequence of clay, silt, and sand interpreted to be marine
clay deposited in an estuary and inner bay environment during marine transgression via the Kitan Strait, overlain by middle and upper sequencesof coarser-grained, silt, sand, and clay in marine and tidal delta environment, following the opening of the Akashi Strait (~ 9.7 ka) (Masudaand Miyahara 2000; Nanayama et al. 2001). Green, blue, yellow, and pink lines mark the base of seismic units defined in this study. Bases of Units
A1 and A1’ in the northern area in this study are correlated with the bases of the middle and upper sequences (tidal delta facies), and the baseof Unit A2 with the that of lower sequence (inner bay and estuary facies). Unit A1 comprises the contemporary Okinose sand bank, which havelikely initiated during the sea-level highstand at ~ 5 ka (also see Fig. 24). Unit B1 is correlated to the finer-grained fluvial facies of the early Holocene
alluvium, and Unit B2 to the coarser-grained fluvial facies of the Late Pleistocene colluvium. Note that in the central to southern area, Unit A1
is absent, and Unit A1’ is deposited to the surface (~ 9.7 ka to present)

gravely sequence of Unit B2 at the boreholes offshore
Kansai airport (Fig. 22; Nakaseko 1987; Masuda 1992).
At Site Suminoe ~ 25 km northeast of our survey Line
1, the sandy and gravely sequence of Unit B1 (4–8 mthick) is dated 10–14 ka based on pollen assemblage at
Suminoe-A, with a radiocarbon age of ~ 12.7 ka at Suminoe-B, and Unit B2 (7–12 m thick) of similar lithology isdated > 47 ka based on radiocarbon at Suminoe-A and B,
separated by an unconformity at > 13 ka (Fig. 23; Nanayama et al. 2001). The youngest pollen assemblage (P2zone, subgroup b) obtained from Unit B2 at Site Suminoe-A is 50–55 ka (Fig. 23; Nanayama et al. 2001). Here,the AT tephra layer is not well preserved, and were likely

removed or disrupted by erosion at the base of Unit B1 orat the top of B2 (Fig. 23; Nanayama et al. 2001). Based onthese data, we assign the age of Unit B1 to be ~ 11–14 ka,
with a possibly older age (~ 25 ka or older) at the base (asinferred from the presence of AT tephra and the youngest pollen age (50–55 ka) at the top and upper section of
Unit B2, respectively) (Fig. 23).
The stratal termination and onlap of Units A1’ andA2 above Units B1 and B2 observed from our E–W
sections are consistent with a major sequence boundary (Figs. 20, 21). The younger strata of the Holocene
alluvium containing more distal facies of marine clay

Hamahashi et al. Progress in Earth and Planetary Science(2024) 11:8

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indicate transgression, whereas the underlying proximal
facies of the coarser-grained colluvium indicate retrogradation (e.g., Van Wagoner et al. 1988; Masuda 1992;
Catuneanu et al. 2009). We interpret the base of Unit B1
to represent a maximum regressive surface at the beginning of transgression (sea-level rise), when depositional
trends switched from coastal progradation (regression)
to retrogradation (transgression). Units B1 and B2 areinterpreted to represent fluvial sediments in a lowstandsystems tract across the Last Glacial Period, correlated
to the incised lower terrace levels of the Tenman and the
Uemachi Sub-Groups in the Osaka Basin, with Unit B1
incorporated to the early Holocene alluvium due to its
relatively young age (Fig. 23; Masuda and Miyahara 2000;
Nanayama et al. 2001). The boreholes indicate that theage of Unit B2 derived from AT tephra and pollen assemblage is significantly older (> 25–55 ka) than the overlying
sediments, thus the base of Unit B1 is likely erosive with apossible hiatus, due to enhanced fluvial erosion during asea level drop at MIS 2 (Fig. 23; Nanayama et al. 2001). At
Sites OB-1, OB-2, Suminoe-A and B, the deepest sectionof the coarse gravel sequence of Unit B2 is marked by a
sharp transition from finer-grained sediments (transitioning to marine clay Ma 12 at greater depth), floored by
the Ata tephra (~ 100 ka) at OB-1 and OB-2, with a pollen assemblage suggesting ~ 70 ka (P3 zone, subgroup e)
at Suminoe-A and B (Fig. 23; Nanayama et al. 2000, 2001;
Kitada et al. 2001). The base of Unit B2 (~ 70 to 100 ka) isinterpreted to represent the onset of forced regression at
the beginning of sea-level drop at MIS 4 (Nanayama et al.
2001). In the current study, however, we cannot correlate
the location of the base of Unit B2 from our seismic sections with the previously reported boreholes, due to thelack of traceable reflectors for the base of this horizon(Sect. 4.2).
Notably, the geometry of the boundary between Units
B1 and A2 observed in our sections is wavy, suggesting
fluvial erosion and/or wavy bedforms (possibly similar tothe contemporary Okinose sand bank) of the sedimentary sequence (Figs. 13, 14, 15, 16, 17). We interpret thisbase of Unit A2 to represent an erosional surface thatformed by means of waves during transgression (i.e., atransgressive ravinement surface; Catuneanu et al. 2009).
The overlying sediments of Unit A2 are observed todrape and pond in the erosional spaces, rather than eroding the underlying sediments, consistent with the interpretation of a marine clay deposition in an estuary andinner bay environment (Figs. 13, 14, 15, 16, 17; Nanayama et al. 2000; Masuda and Miyahara 2000). The abrupt
landward shift of marine systems (as a result of a landward shift of fluvial delta systems) suggests that the risein sea level occurred rapidly and at rates higher than therate of sedimentation (e.g., Catuneanu et al. 2009). Here,the marine clay of Units A1’ and A2 have likely onlappedin the direction of coastal retreat/advancement of thesea, perpendicular to the shelf or parallel to the shoreline
trajectory, as sediment accommodation space increased.
The ages of the Holocene alluvium are well-constrainedat Sites OB-1, OB-2, and Suminoe (Figs. 23, 24; Masudaand Miyahara 2000; Nanayama et al. 2000, 2001). The
Holocene alluvium recovered from Sites OB-1 and OB-2
consist of a lower sequence of clay, silt, and sand interpreted to be marine clay deposited in an estuary andinner bay environment during transgression, dated at thebase to be ~ 11 ka by 14C (Figs. 23, 24; Nanayama et al.
2000; Masuda et al. 2000). This lower sequence is overlain by middle and upper sequences of silt, sand, and clay,
interpreted to represent tidal delta deposition driven bystrong currents emerging from the Akashi Strait, in addition to marine clay deposition from floodplains of the
Osaka Basin, dated by 14C to be ~ 9.7 ka (radiocarbon,OB-1) at the base of the middle sequence, which is also
supported by a pollen assemblage suggesting ~ 9.5 ka at
OB-2, with a slightly older radiocarbon age of 10.9 ka
below (Figs. 23, 24; Nanayama et al. 2000; Masuda et al.
2000). Here, the upper sequence of the Holocene alluvium is distinguished with the middle sequence by adecrease in sedimentation rate starting at ~ 8 ka, andostracod assemblages suggestive of deepening of waterin the upper section at ~ 5 to 5.3 ka, and is interpreted
to represent highstand systems tract across a maximum flooding surface at ~ 5 ka, which is supported by(See figure on next page.)
Fig. 24 a Depositional curves at representative sites along Lines 1–11 (location shown in Figs. 1–10, 13–17, 19) and at Site OB-1 offshore Kobe
(Masuda et al. 2000), plotted with the relative sea-level curve derived from ostracod paleo-depth assemblage from a borehole at the Osaka Basin,
Kitatsumori (Miyahara et al. 1999; Masuda et al. 2000; Masuda and Itomoto 2015). Elevation is measured from present sea-level. Note that the middle
sequence of the Holocene alluvium (Unit A1’) is distinguished with the lower sequence (Unit A2) by a decrease in sedimentation rate startingat ~ 8 ka and deepening of water at ~ 5 ka representing the sea-level highstand. The decrease in sedimentation rate is likely due to the further
landward shift of marine and fluvial delta system near the sea-level highstand. During the sea-level highstand, the tidal currents have strengthened,
and the Okinose sand bank (Unit A1) was likely initiated during this period (Masuda et al. 2000; Nanayama et al. 2001). Note the increase
in sedimentation rate of Unit A1 compared to Unit A1’ at Lines 1–5. (b) Comparison of stratigraphic records with Northern Hemisphere summer
solar insolation (W/m2) (red line; Berger and Loutre 1991), orbital precession cycle (amplitude/ no unit) (orange line; Berger and Loutre 1991)
and Greenland ice core records (δ18O /‰) (gray line; Wolff et al. 2010). MIS: marine isotope stage. YD: Younger Dryas. H1-H6: Heinrich events

Hamahashi et al. Progress in Earth and Planetary Science(2024) 11:8

Page 32 of 50

radiocarbon ages of 4.7–5.4 ka at the onset of tidal delta
sequence identified at Sites OB-1 and OB-2 (Figs. 23,
24; Masuda and Miyahara 2000; Masuda et al. 2000;

Nanayama et al. 2000, 2001). During the sea-level highstand, the tidal currents have likely strengthened, depositing coarser sediments; the Okinose sand bank was likely

Fig. 24 (See legend on previous page.)Hamahashi et al. Progress in Earth and Planetary Science(2024) 11:8

Page 33 of 50

initiated during this period (Masuda et al. 2000; Nanayama et al. 2001). While the fluvial delta shifted further
landward, decreasing the sedimentation rate of marine
clay during the sea-level highstand, this period is followed by a gradual increase in sedimentation rate as thesea-level went down again (Fig. 24; Masuda et al. 2000;
Nanayama et al. 2001). We correlate the lower sequence(inner bay and estuary facies) with Unit A2, and the
middle sequence (marine clay and tidal delta facies) andupper sequences (marine clay and tidal delta facies; sealevel highstand) with Units A1’ and A1 in our northern
sections, respectively (Figs. 23, 24). The erosional surface
at the base of Unit A1 observed in our section is interpreted to have formed by wave scouring and high-energy
tidal currents, indicating a transgressive ravinement surface during the sea-level highstand (Lines 1, 2; Figs. 6,
7, 24). In the northern part of our survey area (Fig. 23),
based on the information above, we assign the following
units the following ages; base of Unit A2 ~ 11 ka, base of
Unit A1’ ~ 9.7 ka, and base of Unit A1 ~ 5 ka.
Our seismic sections show that Unit A1 (~ 5 ka to present) is present only in the northern area, and is replaced
by Unit A1’ to the surface (~ 9.7 ka to present) in thecentral area (Sect. 4.2.2, Figs. 13, 14, 15, 16, 17). Our sections in the central survey area show Unit A1’ onlapping
onto Units A2 and B1 (Figs. 13, 14, 15, 16, 17), consistent with marine transgression (retrogradation). The
observed truncation of Unit A2 below Unit A1’ may
indicate erosion of Unit A2, suggesting that the base of
Unit A1’ is also an earlier transgressive ravinement surface produced by waves and tides during sea-level rise.
At the shallower portion of Unit A1’, we infer a possible
additional downlap surface at the eastern side of ourE–W sections (Figs. 14, 17), which may correlate with aswitch to progradation of marine clay and further seaward shift of floodplains during recent marine regressionat ~ 1 to 1.7 ka (Masuda et al. 2000). Similarly, the shallowest section of Unit A1’ closer to the fault (to the west)
may also be downlapping onto the underlying sediments.
Although we interpret an overall onlapping relationshipthat is consistent with marine transgression, Unit A1’
may additionally be prograding towards the east, due tothe growth of the hanging wall anticline creating more
sediment accommodation space eastward, forming adownlap surface onto Unit A2. The geometry of growthstrata is further discussed in Sect. 5.1.1.
The Holocene alluvium in the northern regions offshore Kobe and coastal Osaka are likely to be coarsergrained compared to the clay-rich alluvium recovered
offshore Kansai Airport, possibly due to the stronger
effects of tidal currents around the Akashi Strait (Masuda

et al. 2000; Masuda and Miyahara 2000; Nanayama et al.
2000, 2001; Kitada et al. 2001; Masuda and Itomoto 2015)(Fig. 23). We infer that Units A1’ and A2 are a combination of both marine clay and tidal delta facies, but arelikely to be finer-grained in our central and southern survey areas compared to the north, although these units
could be seismically traced laterally across our survey
area (Figs. 21, 23).
The timing of marine transgression (base of Unit A2) isdated ~ 11 to 11.4 ka from Sites OB-1 and OB-2, slightly
older than the timings dated at the boreholes in coastal
Osaka such as in Suminoe (~ 8.5 to 9.2 ka) (Masuda andMiyahara 2000; Nanayama et al. 2001) (Fig. 23). The timing of transgression is likely diachronous and becomes
younger towards the basin margin. Similarly, sedimentdeposition arising from the opening of the Akashi Strait(Units A1 and A1’) is also expected to be diachronous,
likely becoming younger with increasing distance fromthe sediment source. However, we are unable to constrain
the spatial variations of age and their full error range atthe base of these units, as we do not have sufficient age
data from offshore across the Osaka Bay. In this study, we
therefore refer to the ages from OB-1 and OB-2 which
are nearer to our survey lines, and use ~ 11 ka for the baseof Unit A2, and ~ 9.7 ka for the base of Unit A1’ in all our
sections.
In summary (Table 1), we interpret Units A1, A1’ andA2 to represent combined facies of marine clay andtidal delta sediments of the Holocene alluvium. Unit A1
comprises the contemporary Okinose sand bank and iscoarser-grained, compared to the marine clay-rich Units
A1’ and A2. We correlate Unit B1 with the finer-grained
fluvial facies of the early Holocene alluvium, and Units
B2 to the coarser-grained fluvial facies of the Late Pleistocene colluvium. These interpretations are consistent
with facies assignments based on backscatter intensity(Sect. 4.2.4), from which it can be inferred that coarsergrained facies comprise the seafloor of Units A1 and B2,
whereas Units A1’ and B1 comprise finer-grained facies(Fig. 4; Table 1). Internal seismic reflections also represent differences in lithology; the sequences of discontinuous and wavy reflectors that characterize Unit B2 in thisstudy indicate the distribution of buried channels anddeposition of coarser-grained, fluvial sediments forming wavy beds and scour surfaces (Figs. 6, 7, 8, 9, 10, 11,
12, 13, 14, 15, 16, 17, 18, 19). The seismically-transparent
and chaotic facies of Unit A1 is consistent with coarsegrained tidal facies, whereas the thin, parallel seismicreflections of Units A1’ and A2 likely represent interbeds
of finer-grained marine clay (Figs. 6, 7, 8, 9, 10, 11, 12, 13,
14, 15, 16, 17, 18, 19).

Thick sequences of continuous, high-amplitudereflections

Sequence of discontinous
and wavy reflectors

B2

B1

Thick sequences of continuous, high-amplitudereflections

B1

Continous, thin, parallel
reflectors

Continous, thin, parallel
reflectors

A2

A2

Continous, thin, parallel
reflectors

A1’

Continous, thin, parallel
reflectors

Seismically transparent
and occasionally chaotic
reflectors

A1

North (Lines 1–5)

Central (Lines 6–10) A1’

Seismic unit Characteristics ofinternal seismicreflections

Survey area

Tilting of strata by folding

Fluvial sandy/gravely
to clayey section transitioning to marine facies

Marine clay, silt, sand(estuary and inner bay
deposits), transitioning
to tidal delta facies

Marine clay, silt, sandand tidal delta facies

Marine clay, silt, sand,and tidal delta facies(Okinose sand bank)

Inferred sediment facies

Homogeneous, weakto-moderate backscatter
intensity facies 2

Homogeneous, weakto-moderate backscatter
intensity facies 2

Fluvial sandy/gravely
to clayey section transitioning to marine facies

Marine clay, silt, sand(estuary and inner bay
deposits), transitioning
to tidal delta facies

Marine clay, silt, sand,and tidal delta facies

Homogeneous, moderate- Fluvial sands and gravels
to-high backscatter
intensity facies 3

Homogeneous, weakto-moderate backscatter
intensity facies 2

N/A

Homogeneous, weakto-moderate backscatter
intensity facies 2

Heterogenous, moderate backscatter intensity
facies 1

Backscatter intensity
faciesOnlap and truncation,
N/A
wavy and irregular contact(erosional) at the base

Onlap

Tilting of strata by folding

Tilting of strata by folding

Tilting of strata by folding(Lines 1, 2, 3), onlap (Lines4, 5), wavy and irregular contact (erosional)
at the base

Tilting of strata by folding(Lines 1, 2, 3), onlap (Lines4, 5)

Tilting of strata by folding(Lines 1, 2, 3), onlap (Lines4, 5), wavy and irregular contact (erosional)
at the base (Lines 1, 2)

Stratal geometries

Table 1 Summary of interpreted stratigraphy in this study

Lowstand systems tract,
Maximum regressive
surface at the beginning of transgression(base of Unit B1), switchfrom coastal progradation(regression) to retrogradation (transgression)

Marine transgression,transgressive ravinement
surface (base of Unit A2)

Marine transgression
and sea-level highstand,possible transgressiveravinement surface (baseof Unit A1’)

Lowstand systems tract

Lowstand systems tract,
Maximum regressive
surface at the beginning of transgression(base of Unit B1), switchfrom coastal progradation(regression) to retrogradation (transgression)

Marine transgression,transgressive ravinement
surface (base of Unit A2)

Marine transgression

Sea-level highstand, Maximum flooding surface(base of Unit A1), possible
transgressive ravinement
surface (base of Unit A1;
Lines 1, 2)

Intepreted sequencestratigraphic surfaces~ 11 to 14 ka~ 11 ka~ 9.7 ka

~ 70 to 100 ka~ 11 to 14 ka~ 11 ka~ 9.7 ka

~ 5 kaInferred age (ka)at the base ofunit

Hamahashi et al. Progress in Earth and Planetary Science(2024) 11:8
Page 34 of 50

South (Line 11)

Survey area

Continous, thin, parallel
reflectors

Continous, thin, parallel
reflectors

Thick sequences of continuous, high-amplitudereflections

Sequence of discontinous
and wavy reflectors

A1’

A2

B1

B2

Continous deposition

Continous deposition

Inferred sediment facies

Homogeneous, weakto-moderate backscatter
intensity facies 2

Fluvial sandy/gravely
to clayey section transitioning to marine facies

Homogeneous, moderate- Fluvial sands and gravels
to-high backscatter
intensity facies 3

Homogeneous, weakto-moderate backscatter
intensity facies 2

Marine clay, silt, sand(estuary and inner bay
deposits), transitioning
to tidal delta facies

Marine clay, silt, sand,and tidal delta facies

Homogeneous, moderate- Fluvial sands and gravels
to-high backscatter
intensity facies 3

Backscatter intensity
faciesContinous deposition,
N/A
wavy and irregular contact(erosional) at the baseContinous deposition

Tilting of strata by folding

B2

Sequence of discontinous
and wavy reflectors

Stratal geometries

Seismic unit Characteristics ofinternal seismicreflections

Table 1 (continued)

Lowstand systems tract

Lowstand systems tract,
Maximum regressive
surface at the beginning of transgression(base of Unit B1), switchfrom coastal progradation(regression) to retrogradation (transgression)

Marine transgression,transgressive ravinement
surface (base of Unit A2)

Marine transgression
and sea-level highstand

Lowstand systems tract

Intepreted sequencestratigraphic surfaces~ 70 to 100 ka~ 11 to 14 ka~ 11 ka~ 9.7 ka

~ 70 to 100 ka

Inferred age (ka)at the base ofunit

Hamahashi et al. Progress in Earth and Planetary Science(2024) 11:8
Page 35 of 50

Hamahashi et al. Progress in Earth and Planetary Science(2024) 11:8

5.1.1 Variations in sedimentation rates and geometriesof growth strata

Figure 24 shows the depositional curves of Units A1’, A1,
A2, and B1 using the assigned ages at representative locations above the footwall syncline along our northern,
central and southern sections, plotted with the publisheddepositional curve at Site OB-1 offshore Kobe (Masuda
et al. 2000; Masuda and Miyahara 2000) and the relative
sea-level curve of the Osaka Bay constrained from a borehole in Kitatsumori, coastal Osaka (Miyahara et al. 1999;
Masuda et al. 2000; Masuda & Itomoto 2015) (Additional
file 2: Table S1). Based on the inferred age and sediment
thickness, we identify a sharp decrease in sedimentation
rate of Unit A1’ compared to Units A2 and B1 in all our
sections (Fig. 24). Units A2 and B1 generally show higher
sedimentation rates of 1.0–12 m/ka and 2.2–7.6 m/ka,
respectively, whereas Unit A1’ show lower sedimentation
rates of 0.9–3.7 m/ka (Fig. 24). This is likely reflecting
the decrease in sedimentation rate of marine clay due tothe further landward shift of the fluvial delta during thesea-level highstand (Masuda et al. 2000; Nanayama et al.
2001) (Fig. 24).
The sediments are generally thicker in the northern
region around the Okinose sand bank, characterized byhigher sedimentation rates of 1.4–4.0 m/ka for Unit A1
and 2.1–3.7 m/ka for Unit A1’ compared to the central
and southern regions (Fig. 24). The progressive thinningof Units A1, A1’, A2, and B1 across the fold observed in
our northern survey areas indicate that the sediments
were deposited while deformation was occurring (Figs. 6,7, 8). This was likely achieved by the high sedimentation
rate of coarser deposits that possibly exceeded the uplift
rate, preserving the overall growth stratigraphy from surface erosion beneath the Okinose sand bank (Figs. 6, 7,
8).
The high sedimentation rate around the Okinose sandbank is also inferred from the high-amplitude negative
polarity reflector that is partially distributed at the baseof Unit A1. The high-amplitude negative polarity reflector may imply a localized horizon of abnormally highsediment porosity, suggesting that compaction and dewatering may not have kept pace with high rates of sedimentation. Rapid rates of sedimentation may produce
elevated fluid pressure, possibly inducing submarineland-sliding at the base of the sand dune complex (e.g.,
Terrinha et al. 2019; Zhang and Askarinejad 2019). The
steep flanks and multiple escarpments of the Okinose
sand bank observed from our bathymetry data also imply
that past gravitational collapse and syn-sedimentary normal faulting may have shaped the current morphology(Sect. 4.1, Figs. 3, 4, 5). The distribution of the negative
polarity reflector along Lines 3, 5, and 12 is located across
one of the steep flanks of the sand bank (Figs. 3, 8, 10,

Page 36 of 5011), and therefore may be related to past landslides thatformed this topography. It is notable that this negative
polarity reflector is only partially distributed and rather
localized, which may indicate that this horizon could
itself be a past landslide deposit. While we do not havedirect evidence suggesting the occurrence of landslides(e.g., collapsed materials around the escarpments), multiple escarpments can be also be formed by other processes such as tidal erosion, rapid deposition of sand,and oversteepening of slope by sediment deposition atdune crest (e.g., Inouchi 1990; Yashima 1992; Ikehara andKinoshita 1994).
In our central survey area, we observe distinct on-lap
of Unit A1’ onto Units A2 and B1, with no tilting of strata
within Unit A1’ (Figs. 13, 14, 15, 16, 17). This is likely due
to the relatively low sedimentation rate of Unit A1’ (~ 1.1
to 1.5 m/ka), possibly slower than or nearly equal to the
uplift rate, suggesting that the marine clays were not continuously deposited during deformation (Fig. 24). The
onlap may in fact be a result of several independent factors: marine transgression across the shelf, low sedimentation rate, and decrease in tectonic uplift to the south. The
stratal dips of Unit A1’ are sub-parallel to the current dipof the seafloor, thinning towards the west above the hanging wall (Figs. 13, 14, 15, 16, 17). The erosional contact
between Units A1 and A2 suggest a major shift in depositional styles with possible age gaps. We therefore cannot
use the upper section of Unit A1 as continuous record ofgrowth strata to infer tectonic rates (Figs. 13, 14, 15, 16,17). The underlying Units A2 and B1, however, show tiltingof strata, preserving growth stratigraphy due to its higher
sedimentation rate compared to Unit A1.
5.1.2 Implications for stratigraphy at deeper sections(Horizons C1–C5)

Past cycles of sequence stratigraphy are continuously
recorded in the Pleistocene sediments in the Osaka Bay(e.g., Muslim et al. 2002; Inoue et al. 2003) (Fig. 22, Additional file 1: Fig. S5), providing a reference for inferring
stratigraphy at greater depths. Marine clay layers that
correlate with global eustatic sea-level highstands possess distinctive physical property contrasts with the surrounding non-marine strata that create high-amplitudereflections, and have been used as prominent stratigraphic markers by previous studies (Muslim et al. 2002;
Inoue et al. 2003; Grothe et al. 2014). Notable of themarine clays are Ma 12 (~ 127 ka), Ma 11 (~ 242 ka), Ma
10 (~ 334 ka), Ma 9 (~ 427 ka), Ma 6 (~ 621 ka), Ma 3(~ 865 ka), Ma 1 (~ 1070 ka), Ma 0 (~ 1120 ka), and Ma-1 (~ 1200 ka), that correlate with isotope stages MIS 5,
7, 9, 15, 21, 25, and 27, respectively (Additional file 1: Fig.
S5). Ma 11 marks the base of the Late Pleistocene colluvium and the top of the Upper Osaka Group, whereas theHamahashi et al. Progress in Earth and Planetary Science(2024) 11:8

Page 37 of 50

bases of Ma 3 and Ma -1 mark the top and bottom of the
Middle Osaka Group, respectively (e.g. Itoh et al. 2000,
2001). In Additional file 1: Figures S6–S20, we compare
the bases of Horizons C1–C5 in our seismic profiles withthe age-correlated reflectors of marine clays reported
by Grothe et al. (2014) near the intersection points with
our survey lines, along previously acquired deep seismic
Lines GS-8ME (Yokokura et al. 1998), OD-B (Iwasaki
et al. 1994), and HD-3 (Iwabuchi et al. 2000), correlatedwith the 1700-m deep borehole (Site GS-K1) at Higashinada, Kobe city (e.g., Itoh et al. 2001; Muslim et al. 2002;
Inoue et al. 2003). Here, we refer to the published depthconverted profiles by Grothe et al. (2014), in which thetwo-way travel time section to depth conversion wasmade based on published stacking velocity functions by
Yokokura et al. (1998). Due to the various uncertainties

in correlation of seismic reflectors arising from velocitymodel-dependent depth-conversions between this study
and previous studies, however, we are unable to constrain
reliable ages for the bases of Horizons C1–C5 in the current study Additional file 3: Table S2).
5.2 Estimation of fault displacement and slip rateof the Osaka Bay Fault: Implications for seismic
hazards
5.2.1 Late Pleistocene–Holocene slip rate of the Osaka Bay
fault

To document the strain accommodated by the Osaka
Bay Fault, we measure the vertical offsets across theforelimb (i.e. the amount of relative uplift between the
uplifted fold crest in the hanging wall and the subsided
footwall trough) for the following stratigraphic horizonsFig. 25 Vertical uplift measured in this study plotted with inferred age and error bars. Vertical offsets across the forelimb (i.e. the amount of relative
uplift between the uplifted fold crest in the hanging wall and the subsided footwall trough) were measured at the base of Units A1, A1’, A2, and B1.
The range of uplift rates estimated along each seismic line incorporating the full error of the ages (based on the minimum and maximum agesof each horizon) are shown in Additional file 1: Figs. S30–S33. (a) Uplift versus age measured from Lines 1–5, 9, and 10 in this study. (b) Uplift
versus age measured from the northern sections Lines 1–5. The slope of the linear regression represents the estimated uplift rate using averageages. (c)Vertical uplift measured in this study and from previous deep seismic lines plotted with average age of marine clays and bases of Units A1,
A1’, A2, and B1. Uplift rate is estimated from the range of the slope of linear regression. Note the northward increase in uplift rate. Box on upper rightshows data plotted from this study closed-up in a

Hamahashi et al. Progress in Earth and Planetary Science(2024) 11:8

Page 38 of 50(Sect. 5.1); the bases of Units A1 (~ 5 ka), A1’ (~ 9.7 ka),
A2 (~ 11 ka), and B1 (12.5 ± 1.5 ka), and the bases of
Horizons C1, C2, C3, C4, and C5 (Additional file 1: Figs.
S21–S29, Additional file 4: Table S3). For comparison, we
also measure the vertical offsets of the marine clays Ma
12, Ma 11, Ma 10, and Ma 9 using the depth-converted
images along the deep seismic Lines GS-8ME (Yokokura
et al. 1998), OD-B (Iwasaki et al. 1994), and HD-3
(Iwabuchi et al. 2000) from previous studies (Grothe
et al. 2014) (Additional file 1: Figs. S6–S19; Additional
file 4: Table S3). For this analysis, we picked the highest position on the hanging wall fold crest and the mostsubsided position on the footwall syncline, and avoided
undulating surfaces that are not planar (Additional file 1:
Figs. S21–S29). We then plot the measured vertical offsets with known age to obtain uplift rates from the slopeof the linear regression (Figs. 25, Additional file 1: Figs.
S30–S33). To incorporate the full error of the ages, weestimate the range of uplift rates based on the minimum
and maximum ages of each horizon (Additional file 1:
Figs. S30–S33).
The most recent time period defined in this study isfrom the bases of Units A1 (~ 5 ka) and A1’ (~ 9.7 ka) tothe present. The total vertical uplift over the last ~ 5 ka
ranges ~ 5 to 11 m, and ~ 6 to 14 m over the last ~ 9.7 ka
in our northern survey area (Fig. 25a, b; Additional file 1:
Figs. S30–S33). At the base of Unit A2, the vertical uplift
ranges ~ 8 to 14 m and > 2–9 m over the last ~ 11 ka at
the northern and central regions, respectively, and adds
up to a total of ~ 9 to 19 m (northern) and > 6 m (central)over the last ~ 11–14 ka at the base of Unit B1 (Fig. 25a,b; Additional file 1: Figs. S30–S33). Units A2 and B1 at
the central survey area are partly eroded at the seafloor,
giving a minimum value for uplift (Figs. 13, 14, 15, 16,
17, 25a; Additional file 1: Figs. S31–S33). From the linear regression of the measured vertical offsets and age,the Late Pleistocene–Holocene uplift rate is estimated torange ~ 1.0 to 1.7 m/ka (Additional file 1: Figs. S31–S33)(average ~ 1.2 to 1.4 m/ka; Fig. 25b, Additional file 1: Fig.
S30) and > 0.6 m/ka at the northern and central regions,
respectively (Additional file 1: Figs. S31–S33). Whenusing fault dip of ~ 70° estimated from greater depths by
previous studies (Grothe et al. 2014), the uplift rates correspond to fault slip rates of ~ 1.1–1.8 m/ka and > 0.6 m/
ka.
From previous studies along deep seismic Lines GS8ME (Yokokura et al. 1998), OD-B (Iwasaki et al. 1994),
and HD-3 and HD-4 (Iwabuchi et al. 2000), average
uplift and fault slip rates of ~ 0.5 m/ka (with a rangeof ~ 0.5 to 1.0 m/ka), ~ 0.3 m/ka (with a range of ~ 0.2 to
0.7 m/ka), ~ 0.3 m/ka (with a range of ~ 0.03 to 0.7 m/ka), and ~ 0.1 m/ka (with a range of ~ 0.05 to 0.4 m/ka),
are inferred in the northern, central, south-central, andsouthern regions along the Osaka Bay Fault, respectively,
based on measured displacements at the bases of marine
clay bed markers (Ma 12 to Ma − 1) during the Middle
Pleistocene (< ~ 1.2 Ma), using the velocity functions by
Yokokura et al. (1998) (Grothe et al. 2014) (Fig. 25c; Additional file 4: Table S3). This is generally consistent with
the Pleistocene uplift rates estimated by previous studies;
Yokokura et al. (1998) reported vertical uplift rate of 0.5–
0.6 m/ka along the Osaka Bay Fault during the early–latePleistocene, whereas Iwabuchi et al. (2000) reported vertical uplift rate of 0.5 m/ka over the last 200 ka, 0.6 m/ka
over the last 600 ka, and 0.6–0.7 ka over the last 1200 ka.
It is noted that the measured depths and amountof uplift from the seismic profiles are velocity-model
dependent. The depth conversion of the seismic profiles
acquired in this study uses constant velocity at 1500 m/s(Section 4.2). If we use the downward increasing velocity functions by Yokokura et al. (1998) which predict thatthe velocity increase from ~ 1500 m/s up to ~ 1700 m/s
at ~ 500 ms along Line GS-8ME close to the fault andclose to our survey line, the total depths (up to ~ 500ms) along our seismic profiles would deviate up to 7%
increase at the trough of the footwall syncline, and up to
4% increase at the crest of the hanging wall anticline. If
using this velocity function at the hanging wall and footwall (Yokokura et al. 1998), we estimate the deviation of
uplift from that measured using constant velocity (1500m/s) in the northern survey area at the bases of Units
A1–C1 to be 0%, and ~ 1.3% increase at the base of Unit
C2, indicating that the consequent errors for Late Pleistocene–Holocene uplift rates documented from Units
A1–B1 in this study are minimal. In comparison, we also
estimate the deviation of uplift in the central survey area
to be ~1.6% increase at base of Unit C1, ~ 2.3% increaseat the base of Unit C2, and ~1% increase at the base of
Unit C3. As a result, the estimated uncertainties of measured uplift from our seismic profiles would range + 0.2–
0.8 m at the base of Unit C1, + 0.9 to 1.3 m at the base of
Unit C2, and + 0.8–1.1 m at the base of Unit C3.
The uplift rates estimated across the Late Pleistocene–
Holocene in this study (~ 1.0 to 1.7 m/ka) (Fig. 25b; Additional file 1: Figs. S30–33) is larger than estimated acrossthe Pleistocene (~ 0.5 to 0.7 m/ka; Yokokura et al. 1998;
Iwabuchi et al. 2000), suggesting that the slip rate on the
Osaka Bay Fault may have increased since the Late Pleistocene. We note, however, that the Pleistocene uplift
rates are averaged over longer time periods than our Holocene uplift rates measured in higher resolution, and theassumption of an uplift rate remaining constant between
the time intervals of marine clays (~ 100 ky) in the Pleistocene may not be applicable. The fault slips along the
Osaka Bay Fault have likely been variable in space andtime within the ranges stated above (Fig. 25c).

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For a given fault slip, a steeper fault dip will produce
larger vertical uplift. We infer steeper synclinal fold axis(~ 75 to 90°) in the recent strata from the dips of the eastern synclinal fold axis, compared to that measured fromgreater depths by previous studies (~ 72°) (Sect. 4.2). The
presence of a backlimb also implies that the fault steepens upward (Seeber and Sorlien 2000). However, theoverall range of these steep dips (~ 70° or near vertical)
do not contribute largely to the increase in fault slip,
showing that the slip rates may have likely increased inrecent strata.
Larger fault slip rates may imply a higher probability of frequent earthquake recurrences, although fault
slip can be accommodated by both coseismic, interseismic, and aseismic events (e.g., Scholz 2019). The lack of
localized microseismicity and aseismic creep along the
Osaka Bay Fault (e.g., Katao et al. 1997; Nakamukae et al.
2003; Matsushita and Imanishi 2015; Nishimura et al.
2020), however, indicate that fault slip is largely accommodated by large earthquakes and long-term inelastic
deformation such as folding. Coseismic events, especiallylarge magnitude earthquakes, produce large and rapid
fault slips compared to aseismic deformation. During
the 1995 M­ w6.8 earthquake, for example, vertical offsets of up to ~ 1.2 m was inferred along the Nojima Fault(Awata et al. 1995; Nakata and Yomogida 1995; Nakata
et al. 1996). The total vertical uplift of ~ 5–11 m overthe last ~ 5 ka observed along the Osaka Bay Fault may
therefore indicate the occurrence of multiple earthquake
events.
Previous shallow seismic surveys using a single-channel sonoprobe conducted by Nanayama et al. (2000)
along the northern extent of the Osaka Bay Fault, ~ 9 km
north of our survey Line 1, reported vertical upliftsof ~ 5.5 to 8 m during the last ~ 7.3 to 8.6 ka, and ~ 2 to
3.5 m during the last ~ 0.9 to 1.2 ka, which is the mostrecent time period of fault activity that has been so far
documented along the Osaka Bay Fault. Their results arecomparable with our uplift estimates of ~ 5 to 11 m overthe last ~ 5 ka, and ~ 6 to 14 m over the last ~ 9.7 ka in ournorthern survey area (Fig. 25a).
Although there is no direct paleoseismological record
reported at the Osaka Bay Fault, previous on-land trench
excavation studies at other faults have reported ruptures during the 1596 A.D. Keicho-Fushimi Earthquake(~ M7.5), prior to the 1995 Kobe Earthquake, along multiple fault strands of the surrounding Arima-Takatsuki
and Rokko-Awaji Island Fault zones (e.g., Nojima andGosukebashi Faults), in addition to the Median Tectonic
Line in Shikoku (Nakata et al. 1996; Lin et al. 1998; Sangawa 2001; Ikeda et al. 2019). According to trench studies at the Nojima Fault, the 1596 and 1995 events werepreceded by events that occurred at ~ 2 and ~ 4 ka (Awata

et al. 1996; Nakata et al.1996). At several trenches along
the Arima-Takatsuki Fault zone, rupturing events that
occurred at ~ 2.8 to 3 ka were also identified as predating the 1596 event (Sangawa 2001). Studies of the Gosukebashi Fault, for example, report that the 1596 event
was predated by an event at ~ 416 A.D. (Yamato Earthquake), and an event occurring between ~ 6 ka and 1596
AD (Maruyama et al. 1997; Lin et al. 1998). Although
the recurrence intervals of earthquakes along individual faults are likely variable, the 1995 and 1596 events
have demonstrated that multiple fault segments in thisregion can potentially slip in a single or multiple eventsat around the same time period, leading to disastrous,
large-magnitude earthquakes (Hirata et al. 1996; Katoet al. 2008; Ikeda et al. 2019). The Osaka Bay Fault may
have ruptured in the form of fold scarps during these
events (1596 AD, 416 AD, ~ 2 ka, ~ 4 ka), together withthe surrounding faults in this region, producing the total
vertical uplift of ~ 5–11 m over the last ~ 5 ka, as inferredfrom our study (Fig. 25a,b). It is important to note that,
although blind, the fault displacement along the Osaka
Bay Fault can potentially extend to the surface during
earthquakes, as demonstrated from the shallow deformation across its entire forelimb (Sect. 4.2.5).
The possible increase in fault slip rates inferred inthis study emphasize the importance of documenting
recent deformation of strata and averaging slip rates overshorter time periods for probabilistic seismic hazard
assessments. Our results suggest that slip rates and sedimentation rates averaged over longer time intervals may
yield lower estimates of earthquake recurrence. Variable
fault slips in space and time along the Osaka Bay Fault
have also been documented by Grothe et al. (2014), inferring steadier fault slip rates at longer timescales of ~ 100 sof ky, which have also been noted for faults in other tectonic regions (e.g., Mouslopoulou et al. 2009; Bergen
et al. 2017; Mueller 2017).
5.2.2 Northward increase in slip on the Osaka Bay Fault
and implications for slip rates

Our study and previous seismic reflection studies document a northward increase in vertical uplift and slip rates
along the Osaka Bay Fault (Figs. 20, 25; Additional file 1:
Figs. S30–S33), indicating higher fault activity and/or
structural maturity in the north (Yokokura et al. 1998;
Iwabuchi et al. 2000; Grothe et al. 2014). This is supported by our observation that the dip of the anticline
forelimb is steeper in the northern area (~ 0.4 to 15°) withlarger offsets across the fold, compared to the dip of theforelimb in the central regions (~ 0.2 to 5.2°), indicating
southward decrease of fault-related folding (Sect. 4.2).
The southward tilting and thickening of strata observedin the hanging wall are also consistent with northward

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increase in uplift (Fig. 18). Despite the larger offsets andsteeper forelimbs in the northern region, however, thedip of the eastern synclinal axial surface and inferredfault dip are steeper in the central regions (~ 85 to 89°)
compared to the north (~ 75 to 85°) (Sect. 4.2), suggestingthat the total fault slip indeed increases northward.
Grothe et al. (2014) inverted fault tip depths from thetips of the triangular zones inferred from the geometryof folded growth strata of the Osaka Bay Fault based on
2D asymmetric trishear modeling. Their models suggest greater initiating depths of the fault tip in the north(~ 3.5 to 5 km) and shallower depths in the south (~ 1.6 to
2.5 km) (Grothe et al. 2014). The fault tips initiate fromolder strata (~ 1.5 Ma) in the north than in the south,
implying that the fault may have propagated laterally
from north to south with time (Grothe et al. 2014). The
greater depths of fault tips in the north may have been
influenced by the depth of the basement and thickness ofoverall sediments that increase from south to north, as
inferred by isopach models and gravity anomalies shownby previous studies (Komazawa et al. 1996; Yokokura
et al. 1998; Iwabuchi et al. 2000). Sediment thickness inthe basin is likely related to the effect of flexural loading by surrounding faults, although other far-field processes occurring deeper in the lithosphere may also
play a role. The greater thickness of the depocenter at
depths in the northern region is likely recording subsidence on the footwall of the Kariya Fault offshore northern Awaji Island, counteracted by uplift on the hanging
wall of the Osaka Bay Fault (Fig. 2a). The southward tilting and thickening of recent strata observed in this study(Fig. 18), however, is contrary to the northward thickening at greater depths (> ~ 1000 m depths; > 1.2 Ma)(Iwabuchi et al. 2000), and suggests increased effects of
uplift on the Osaka Bay Fault compared to the Kariya
Fault in relatively recent strata (< 1.2 Ma).
Lateral propagation and growth of the fault may indicate that the fault lengths and areas of the Osaka Bay
Fault have increased with time (Grothe et al. 2014). Generally, for a newly-initiating fault, an increase in fault
length would correspond to greater displacement, leading to estimations for larger maximum potential earthquake magnitudes and seismic hazards (e.g., Bergen andShaw 2010). Accelerated Holocene slip rates along active
faults have also been reported for blind thrust fault systems in other regions such as beneath the Los Angeles
basin (Bergen et al. 2017; Mueller 2017). The possible
increase in slip rates during the Late Pleistocene–Holocene inferred in this study may be a result of enhanced
lateral fault propagation since the Late Pleistocene, that
led to an increase in average displacement per earthquake and occurrence of larger earthquakes.

The trishear modelling by Grothe et al. (2014), however, suggested that the Osaka Bay Fault is reactivated,based on the shallow depth of the fault tip early in its
history, instead of a fault that initiated from the base of
the seismogenic crust (e.g. Figure 1c). A reactivated fault
may not necessarily involve lateral fault propagation in its
recent history. Furthermore, changes in fault slip rate andpropagation may not necessarily be directly correlated;
fault propagation may depend on the strength of rockand faults (Hardy and Allmendinger 2011), whereas fault
slip rates may depend on stressing rates and interaction
among other neighboring faults.
Alternatively, the increase in slip rates may occur as aconsequence of an overall increase in frequency of smallto-large earthquakes along the Osaka Bay Fault, which
may not necessarily relate with increase in fault lengths.
Other external forcing not related to tectonics may also
affect fault activity (Mueller 2017), such as by unloading or loading by sea-level fluctuations and erosion (e.g.,
Calais et al. 2010; Luttrell and Sandwell 2010; Steer et al.
2014), or shorter timescale stress transients caused byearthquakes in the vicinity (e.g., Freed 2005; Ishitsuka
et al. 2020). In the current study, we cannot rule out thecause of the inferred transient slip, although short-term
forcing is less likely based on the time duration (> ~ 10 ka
to present) of the inferred increase in slip rates in thisstudy (Fig. 25).
5.2.3 Comparison of slip rates with regional faults

The Osaka Bay Fault is surrounded by the NE–SW trending Rokko-Awaji Island Fault zone (e.g. Nojima, Kariya,
Kusumoto, and Higashiura Faults) to the west and north,
the Uemachi Fault zone to the east, and the E–W trending Median Tectonic Line to the south (Fig. 2a). The
average slip rate along the Nojima Fault that ruptured
during the 1995 earthquake is 0.4–0.5 m/ka verticallyand 0.9–1.0 m/ka right-laterally (Mizuno et al. 1990).
The right-lateral average slip rate of the Kusumoto Fault
is estimated to range 0.5–1.0 m/ka (Mizuno et al. 1990).
The average slip rates of the E–W faults that extend to
the Rokko Mountains are inferred to have larger lateralslip component, ranging ~ 2.0 m/ka right-laterally and
0.3–0.4 m/ka vertically (e.g. Itoh et al. 2000). For the
Arima-Takatsuki Fault zone to the north, 0.06–0.2 m/
ka vertical and 0.5–1.5 m/ka right-lateral slip rates havebeen reported (Sangawa 1978). The Uemachi and Ikoma
Fault zones to the east have reported average verticalslip rates of ~ 0.4 m/ka and 0.5–1.0 m/ka, respectively(Research and Development Bureau/MEXT & DPRI
2013; The Headquarters for Earthquake Research Promotion/Earthquake Research Committee 2001). The overallsimilar slip rates in the region may suggest that the faults

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behave as a linked system. We note however that ourslip rates are larger than most of the uplift rates reported
by previous studies. By contrast, the long-term uplift
rate estimated from the height of the mountain ranges;
2–3 km of vertical offset between the Osaka Basin andthe Rokko Mountains (and Awaji Island) during the last
3 million years (e.g. Sato et al. 1998), yield a long-term
uplift rate of 0.7–1.0 m/ka. This value is approximately
the average between the Middle–Late Pleistocene slip
rates inferred from previous studies (0.5–0.7 m/ky; see
Yokokura et al. 1998; Iwabuchi et al. 2000) and the Late
Pleistocene–Holocene slip rates (1.0–1.7 m/ka) inferredin this study (Fig. 25b; Additional file 1: Figs. S30–S33)
at the Osaka Bay Fault, supporting that these rates areplausible.
The northern Osaka Bay Fault is known to splay into
three strands (Wadamisaki, Rokko Island, and Maya
Faults), and the average uplift rate of the Wadamisaki
Fault is reported to range ~ 0.2–0.3 m/ka (Yokokura et al.
1998). The observed northward increase in slip along the
Osaka Bay Fault may be translating to these splays. Nanayama et al. (2000) reported large Holocene uplift of ~ 6 m
during the last ~ 6–7.3 ka on the Wadamisaki Fault,
although they did not find that significant uplift occurred
during the Holocene along the Rokko Island and Maya
Faults.
By contrast, the average slip rates along the Median
Tectonic Line to the south range ~ 0.8–1.0 m/ka vertically beneath the Kitan Strait (Nanayama 1999), withright-lateral slip rates of ~ 1.8–3.5 m/ka and ~ 6 m/ka
onland to the east and west of the strait, respectively(Okada 1970; Okada and Sangawa 1978; Saito et al. 1997;
Okada and Tsutsumi 1997). Near the Kitan Strait, E–W
trending fault strands that may possibly be the seaward
extension of the Uemachi Fault zone have been reported
beneath the southern Osaka Bay, revealing vertical
uplifts (Yokokura et al. 1998). These fault strands run
parallel to the Median Tectonic Line, indicating a structural link with the faults beneath the Osaka Bay and theMedian Tectonic Line. We do not have sufficient data
for the southern region to fully address the structurallink between the N-S trending Osaka Bay Fault with theE–W trending fault zones of the Median Tectonic Lineand the Uemachi Fault zone. Structural complexities inthe southern Osaka Bay are inferred from changes in theoverall fault geometry and partitioning of multiple fault
strands (Yokokura et al. 1998; Iwabuchi et al. 2000) andvariable basement morphologies that are overall uplifted
towards the Kitan Strait above the Cretaceous Izumi
Group (Komazawa et al. 1996; Afnimar et al. 2002; Itohet al. 2013a) (Fig. 2a, Sect. 2.2).
The southward decrease in slip of the Osaka Bay
Fault along-strike observed in this study, and the lateralsouthward fault tip migration suggested by Grothe et al.
(2014) may indicate that the southern segment of the
Osaka Bay Fault has accommodated less slip after its
reactivation compared to the northern segment. The
recent increase in slip to the north may be related to its
prior slip history; for example, the northern fault may be
weaker in strength compared to the south, due to more
slip accumulating in prior periods. Additionally, thenorthern segment of the Osaka Bay Fault may be more
active than the south due to the interactions with otherlarger fault systems such as the Rokko-Awaji Island andArima-Takatsuki Fault zones to the north. In the southern Osaka Bay, slip may be partitioning into other fault
strands to the west (Iwabuchi et al. 2000), which may also
relate to the southward decrease in slip of the Osaka Bay
Fault. Alternatively, a change in stress state within thebay (Katao et al. 1997; Nakamukae et al. 2003; Matsushita
and Imanishi 2015) may also contribute to the change invertical component of slip.
If it is assumed that the total amount of slip adding
all the faults in the system has been constant with time,the inferred increase in slip rate on the Osaka Bay Fault
may indicate that the slip rate on other faults may havedecreased. Alternatively, the total amount of slip in thesystem may have overall increased, with increased slip
rates on other faults in addition to the Osaka Bay Fault.
Based on the recent occurrence of the 1995 ­Mw6.8 Kobe
(Southern Hyogo prefecture) Earthquake on the RokkoAwaji Island Fault zone, it is possible that the fault activity of the whole system may have increased since the Late
Pleistocene, leading to greater seismic risks in this region.
It is crucial to study recent slip rates on other faults inorder to fully evaluate the regional fault activities.
Fault slip rates in this region could reasonably be
expected to have been transient during recent and earlier
periods. Thus, while previous studies have inferred thatthe average slip rates have been fairly constant throughtime (e.g., Yokokura et al. 1998), other studies have demonstrated fluctuations in past slip rates (possibly fromas low as ~ 0.1 mm/yr to as high as ~ 1.0 mm/yr) on the
Uemachi, Osaka Bay, and the Rokko-Awaji Island Fault
zones beneath the northern Osaka and Kobe Basins
and in the Osaka Bay during the Middle Pleistocene(Itoh et al. 2000; Grothe et al. 2014). Grothe et al. (2014)
reported variable fault slip in space and time, in which
the Osaka Bay Fault propagated at greater rates in thenorth compared to the south, after its reactivation andonset of slip (~ 1.5 Ma), later building more slip aroundprofile GS-8ME. Itoh et al. (2000) detected three main
episodic changes occurring within the Osaka and Kobe
Basins, where the subsidence rates show a sharp decline
during the Ma 5/6 interval (~ 712–621 ka), followed by asecond episode of differential rise and fall of subsidenceHamahashi et al. Progress in Earth and Planetary Science(2024) 11:8

Page 42 of 50

rate occurred during Ma 7/8 (~ 578–528 ka), and a third
episode of reduced subsidence on the western and eastern edge of the basin during Ma 8/9 (~ 528–427 ka) (Itohet al. 2000). The first episode at Ma 5/6 corresponds to
the Mantidani Unconformity (Huzita 1990), an upwardcoarsening facies boundary at the base of the Upper
Osaka Group found in the Awaji Island, Kobe and the
Osaka Basin, implying a regional event (Inoue et al. 2003).
Previous studies have attributed this episodic decline ofsubsidence in the northern Osaka Basin to be linked withthe shift of the western part of the Median Tectonic Lineat about ~ 700 ka (Itoh et al. 1998, 2000). The second andthird episodic changes are inferred to link with the buildup of the Uemachi Basement High and the N-S warpingof the sedimentary basin that caused accelerated northwestward tilting and subsidence in the south-eastern
Osaka Basin during the Middle to Late Pleistocene (Itohet al. 2000, 2001).
Fault slip rates are important because they are usedin seismic hazard assessments, as a proxy for moment
release along a fault over time, where they are mergedwith records of earthquake magnitude and recurrence
intervals inferred from paleoseismic investigations andhistorical earthquakes (e.g., Matsuda 1975, Kanaori 2000;
The Headquarters for Earthquake Research Promotion
2001; Mueller 2017). Our study investigates slip rates
over the last several to tens of ky, owing to the continuous fold growth strata preserved in the sedimentary basinand high-resolution seismic imaging techniques coupledwith drilling. These intermediate time-scales allow theidentification of transient fault slip rates and bridge theobservational time gaps between geologic cross sections
that average longer time scales (~ few mys) and shortterm observations based on paleoseismology (~ few hundred yrs to several kys), geodesy and seismology (~ few
seconds to few yrs). Transient fault slip rates indicate
changes in potential earthquake magnitudes, frequencies, slip distributions, and/or interactions with surrounding faults with time, and this time-variant naturechallenges characteristic earthquake cycle models thattypically assume repeating rupture size and displacement
on individual fault segments (e.g., Grant 1996; Bergen
et al. 2017; Nakatani 2020). In this study, we highlightthe importance of using observations of slip rates averaged over intermediate (Late Pleistocene-Holocene) timescales, and incorporating these aspects to improve earthquake models for potential seismic hazard assessments.

Kobe, and Nara Basins, bounded by the Rokko, Ikoma,
and Suzuka Mountains within the Kinki Triangle (Huzita 1962). The vertical slip component is accommodated
largely by fold-and-thrust faulting, with folding being
dominant in the sedimentary section (Ishiyama 2003;
Sato et al. 2009), as observed in this study. The basesof the sedimentary basins are dated ~ 3 Ma, which iscoincident with the timing of reactivation of regional
faults such as the Median Tectonic Line (Maruyama& Lin 2004; Famin et al. 2014). However, on the basisof regional unconformities and marine clay distributions within the Osaka Group, previous studies haveinferred that the subsidence of basins and uplift of ranges
occurred rapidly after ~ 1 Ma, indicating that the activityof major faults in the region accelerated <  ~ 1 Ma (Huzita
1990; Itoh et al. 2000; Kato et al. 2008). A broad distribution of marine clay Ma 1, 2 and 3 (~ 1070 ka, 960 ka,
and ~ 865 ka, respectively) of the Middle Osaka Group is
identified within the Kyoto and Nara Basins that extend
from the Osaka Basin across the Rokko and Ikoma
Mountains, and the preservation of these strata acrossthese elevated ranges indicates that uplift was not significant during deposition of Ma 1–4, and that significant
uplift occurred later (e.g. Huzita 1990). During and possibly prior to this period (~ 3–1 Ma), the basement rocks of
the Ryoke Group were likely deformed via a detachment
fold without localized uplift (Huzita 1990; Poblet andMcClay 1996). Whereas, the marine clays of the Upper
Osaka Group (Ma 6 ~) (< 621 ka) are not deposited on themountains and are only preserved in the basins, with thestrata clearly offset by dominant faulting and fault-related
folding that have likely accelerated after ~ 1 Ma (e.g.,
Huzita 1990). The current arrangement of the basins
and mountain ranges are therefore likely a topographic
expression of the faulted blocks in the basement rocks
that have propagated to the sedimentary section (Huzita
1990; Itoh et al. 2000).
The onset of the current tectonic framework in SW
Japan is governed by the oblique subduction of the Philippine Sea Plate along the Nankai, Ryukyu and Philippine Trenches, which is estimated to have initiatedaround ~ 6 Ma, as inferred from ocean drilling results in
the Nankai subduction zone (Moore et al. 2015; Kimura
et al. 2014, 2018), geologic surveys on the Median Tectonic Line (Maruyama and Lin 2004; Famin et al. 2014),
and plate reconstruction models (Wu et al. 2016). The slip
sense of the Median Tectonic Line changed from dominantly reverse faulting to right-lateral strike-slip faulting
at around ~ 3 Ma (Maruyama and Lin 2004; Famin et al.
2014), suggesting that the oblique subduction strengthened during this period. The contemporary left-stepping
dextral Median Tectonic Line is likely a result of translation of the forearc sliver, in addition to reactivation of

5.3 Implications for regional tectonics and comparisonswith geodetic slip rates
5.3.1 The onset of active faulting related to forearc sliver

The differential right-lateral and vertical slip motions ofregional faults resulted in the formation of the Osaka,Hamahashi et al. Progress in Earth and Planetary Science(2024) 11:8

other major E–W trending dextral faults to the north(e.g., Arima-Takatsuki Fault zone) (e.g., Kanaori 1990;
Tsukuda 1992). The NE–SW to N-S trending faults northof the Median Tectonic Line, such as the Nojima, Osaka
Bay, Uemachi, and Ikoma Faults, have likely initiated as
en echelon right-lateral- and reverse faults and folds
associated with the forearc sliver (e.g., Tsukuda 1992),
evolving from a detachment fold system to a fault-propagation fold system as the structures matured and as more
sediments accumulated.
The onset of contemporary inland active faulting andthe growth of the accretionary prism at the Nankai subduction zone were likely punctuated events occurring inseveral time steps; ~ 6 Ma, ~ 3 Ma, ~ 2 Ma, and ~ 1 Ma,
most likely related to the kinematics of the plate subduction (e.g., Moore et al. 2001, 2015; Kimura et al. 2018).
Seismic reflection studies coupled with drillings along
the Nankai Trough have shown a major unconformity
at ~ 6 Ma that implied resurgence of Philippine Sea Plate
subduction after a long period of cessation since ~ 12 Ma(Kimura et al. 2018). Drilling results have also indicatedrapid formation of the frontal prism (imbricated thrustzone) in the Nankai Trough after ~ 2 Ma (e.g., Moore
et al. 2001), and rapid <  ~ 2 Ma uplift on the megasplay
fault and landward tilting in the Kumano Basin (e.g.,
Moore et al. 2015), in addition to strike-slip faulting
inferred along the Kumano Basin Edge Fault zone (e.g.,
Martin et al. 2010). The punctuated growths of the offshore accretionary prisms and forearc basins are inferred
to be coincident with the period when the slip sense of
the Median Tectonic Line changed to right-lateral strikeslip faulting (e.g., Maruyama and Lin 2004; Famin et al.
2014), suggesting a large-scale tectonic reorganizationthat involved vast areas of the upper plate. The volcanic
activities around the Beppu-Shimabara grabens andthe Higo volcanic zone are also reported to have onset
at ~ 5–6 Ma and ~ 2 Ma (e.g., Kamata and Kodama 1994).
The opening of the Okinawa Trough is also coincident
with these periods (e.g., Sibuet et al. 1987). Plio-Pleistocene sedimentary basins were formed in many parts of
the region, likely related to tectonic events (e.g., Nakajima 2018).
The recent tectonic phase may also relate to the initiation of convergence of the Amurian Plate with the
Okhotsk Plate along the eastern margin of the Japan Sea(Kimura et al. 2018; Nishimura et al. 2018). Previous
studies have inferred similarities in geodetic strains andtectonic block movements east of Awaji Island with theblocks in the Kanto region that is known as part of the
Okhotsk Plate, implying the presence of a N-S trending
tectonic boundary that may explain the regional E–W
compression (Nishimura et al. 2018).

Page 43 of 50

5.3.2 Comparisons with geodetic strain rates

The contemporary crustal deformation in southwest
Japan has been monitored by dense GNSS (Global
Navigation Satellite System) networks represented by
GEONET (GNSS Earth Observation Network System)
onland stations and GPS-A (Global Positioning SystemAcoustic) seafloor stations (e.g. Sagiya 2004; Yokota
et al. 2016; Nishimura et al. 2018) (Fig. 1b). The inland
velocity fields in most regions show west-northwestward
movement almost parallel to the motion of the subducting Philippine Sea Plate along the Nankai Trough, suggesting strong interplate coupling at the subduction
zone (Nishimura et al. 2018) (Fig. 1b). The existence ofnumerous active faults and high seismicity on the islandarc and continental crust of Japan are consistent with significant slip partitioning and diffuse deformation in theupper plate (Fig. 1a). While the strain rate distribution in
the inland regions are not homogeneous, several zones
of strain concentration are recognized from GNSS data
(Loveless and Meade 2010; Nishimura et al. 2018). Oneof the most active boundaries with geodetic slip ratesof ≥ 9 mm/yr is the Median Tectonic Line, accommodating part of the margin-parallel component of the current
oblique convergence by right-lateral shear motion (Loveless and Meade 2010; Nishimura et al. 2018) (Fig. 1b).
Another significant high-strain rate zone is the Niigata–
Kobe Tectonic Zone, a ~ 50–200 km wide NE–SW trending zone north of the Median Tectonic Line which haslarge geodetic slip rates of > 15 mm/yr (e.g., Sagiya et al.
2000; Loveless and Meade 2010; Nishimura et al. 2018)(Fig. 1b). The active faults around the Osaka Bay and the
Kinki Triangle constitutes the southern portion of the
Niigata–Kobe Tectonic Zone (Fig. 1ab).
Loveless and Meade (2010) and Nishimura et al.
(2018) defined a tectonic block around the Kinki Triangle bounded by the Rokko-Awaji Island Fault zone,
the Arima-Takatsuki Fault zone, and the Median Tectonic Line in their models, and inferred large internalstrain rates of − 52.4 to − 82.8 nanostrain/yr within theblock surrounding Osaka. The geodetic strain rate (surface deformation) within the Kinki Triangle expressed
as a sum of rigid block rotations and interseismic elastic
deformation due to locked faults on the block boundaries(Nishimura et al. 2018) is estimated to be ~ 15 mm/yr (T.
Nishimura, personal communication, December 2023).
The compressional strain within the block dominates ina direction of relative plate convergence, and is nearly
perpendicular to the NE–SW trending Osaka Bay Fault
and other faults such as the Uemachi and Ikoma Faults,
which may indicate relatively dominant dip-slip thrusting for these faults, whereas right-lateral slip is relatively
dominant along E–W trending faults of the Rokko-Awaji

Hamahashi et al. Progress in Earth and Planetary Science(2024) 11:8

Page 44 of 50

Island and Arima-Takatsuki Fault zones (Loveless and
Meade 2010; Nishimura et al. 2018). Across the central–northern Osaka Bay (within a box extending ~ 140 km
oriented perpendicular to the strike of the Osaka Bay
Fault with its center aligned parallel to the fault), GNSS
velocity data by Nishimura et al. (2018) suggest relative
geodetic strain and shortening rates of ~ 10 to 11 mm/
yr (T. Nishimura, personal communication, December2023). The Late Pleistocene–Holocene slip rate (~ 1.1
to 1.8 mm/yr) on the Osaka Bay Fault estimated in ourstudy (Sect. 5.2.1) may account for ~ 10 to 18% of theregional geodetic strain accumulation across the central–northern Osaka Bay (~ 10 to 11 mm/yr), and ~ 7.3 to
12% of the regional geodetic strain accumulation within
the Kinki Triangle (~ 15 mm/yr). The remaining strain
accumulation may be accommodated by other faults,
which may indicate the possibility for more previouslyunknown hidden faults in the region, and/or larger slip
rates on other faults. By contrast, the interblock rightlateral motions along the Rokko-Awaji Island and ArimaTakatsuki Fault zones are inferred to range ~ 3 to 9 mm/
yr, and ~ 3 to 6 mm/yr along the Median Tectonic Line
(Nishimura et al. 2018), which is likely relatively morecompatible with geologic slip rates of known faults.
Geodetic strain rates generally include a sum of the slip
rates and slip senses of individual faults existing in thesystem (e.g., Bourne et al. 1998). However, discrepancies
between geodetic and geologic slip rates may arise fromvarious processes such as by elastic deformation due tointerplate coupling on the plate boundary included ingeodetic strains, long-term (geologic) transient slip rates
that are not accounted for in geodetic rates, and/or when
there is significant amount of deep, inelastic deformationthat is not expressed in geomorphological features andhence excluded in geologic slip rates (e.g., Hashimoto
1990; Meade and Hager 2005; Ohzono et al. 2011; Johnson 2013). We demonstrate that a detailed study of Holocene slip rates of regional faults is necessary to assess theseismic hazards and the internal strain budgets within
the Kinki Triangle and the Niigata–Kobe Tectonic Zone.

syncline at the western extent of the anticline.
Towards the eastern side of our sections, the tilted
strata dip gently west, forming a broad ~ 11-kmwide syncline in the footwall. Our seismic data
confirm that the Osaka Bay Fault is blind, and thatthe deformation reaches near surface by a faultpropagation fold. The forelimb and backlimb of thefault-propagated anticline progressively steepens
with depth, and the overall monocline widens atshallower depths.
(2) The synclinal axial surface at shallow depths measured in this study ranges from 75 to 89°. This dip
observed at shallower depths is steeper than measured in previous studies (~ 72°) from deep seismiclines. The steeper dips of the synclinal axial surface
at shallow depths are associated with an upward
propagating growth triangle with ongoing sedimentation. The presence of a backlimb on the faultrelated fold may also relate to the steepening of thetrishear envelope. The shallow structure of the fault
documented by our study in previously unavailable
detail provide insights on the subsurface deformation by the Osaka Bay Fault.
(3) Our northern, central, and southern sections clearlyshow that the dip of the anticline forelimb is steeper
in the northern area (~ 0.4 to 15°) with larger offsets
across the fold, compared to that of the south (~ 0.2
to 5.2°). The sediment thickness on the hanging
wall, however, is variable (north >  > central < south),suggesting that erosion does not balance uplift. We
attribute this variation to varied mobilization of
sediment by tidal currents active in the bay during the Holocene. Major sequence boundaries suchas onlap of recent sediments onto tilted strata andbasal erosional surfaces during marine transgression and sea level highstands are inferred to affect
the geometry of growth strata.
(4) In the northern Osaka Bay, the most recent deformation by the Osaka Bay Fault reaches to near theseafloor by active folding, with large vertical offsets of ~ 8–14 m over the last ~ 11 ka, and ~ 5–11 mover the last ~ 5 ka. By combining with previously
reported borehole age data, the average uplift rateon the Osaka Bay Fault during the Late Pleistocene
to Holocene in this study is estimated to be ~ 1.0 to
1.7 m/ka, larger than the uplift rates previously documented (~ 0.5 to 0.7 m/ka) from deeper sections
spanning the Pleistocene (Yokokura et al. 1998;
Iwabuchi et al. 2000).
(5) The inferred slip on the Osaka Bay Fault during theLate Pleistocene–Holocene may account for ~ 7.3
to 12% of the regional geodetic strain accumulation
within the Kinki Triangle. Further studies to evalu-

6 Conclusion
To investigate the shallow structure and recent deformation history of active faults in the Osaka Bay, we acquired
high-resolution seismic profiles and multi-beam bathymetry data across the central segment of the Osaka Bay
Fault. Our main results are:(1) The Osaka Bay Fault is characterized by an asymmetric anticline at shallow depths, composed ofa ~ 0.1–3.7-km-wide, steeper forelimb that dipseast and a gentler and broader backlimb that dips
west, associated with an additional ~ 2.6-km-wideHamahashi et al. Progress in Earth and Planetary Science(2024) 11:8

ate the Holocene slip rates of regional faults arenecessary for the assessment of seismic hazards andthe internal strain budgets within the Kinki Triangle
and the Niigata–Kobe Tectonic Zone.

Page 45 of 50

Availability of data and materials
All data used in this study will be uploaded to a FAIR-aligned data repository
upon publication.

Declarations
Competing interests
The authors declare that they have no competing interest.

Abbreviations“Ma 1”–“Ma13”	Marine clay beds 1–13
“TWT”	Two-way travel time

Supplementary Information
The online version contains supplementary material available at https://​doi.​
org/​10.​1186/​s40645-​024-​00607-0.
Additional file 1. Figures S1 to S33.
Additional file 2. Table S1. Measured depths of seismic horizons shownin Figure 24 (main text). 			
Additional file 3. Table S2. Measured depths of seismic horizons andmarine clays.			

Author details
1
Kobe Ocean‑Bottom Exploration Center, Kobe University, 5‑1‑1 Fukaeminamimachi, Higashinada, Kobe 658‑0022, Japan. 2 Geological Survey of Japan,
National Institute of Advanced Industrial Science and Technology, 1‑1‑1
Higashi, Tsukuba, Ibaraki 305‑8567, Japan. 3 Department of Planetology, Graduate School of Science, Kobe University, 1‑1 Rokkodai, Nada, Kobe 657‑8501,
Japan. 4 School of Geosciences, University of Aberdeen, King’s College,
Aberdeen AB24 3FX, UK. 5 Nippon Marine Enterprises, Ltd., 14‑1 Ogawacho,
Yokosuka, Kanagawa 238‑0004, Japan. 6 Seisgadget Ltd, 3‑31‑11, Ohmori‑Kita,
Ohta‑Ku, Tokyo 143‑0016, Japan. 7 School of Marine Resources and Environment, Tokyo University of Marine Science and Technology, 4‑5‑7 Konan,
Minato‑Ku, Tokyo 108‑8477, Japan. 8 Present Address: Faculty of Global and Science Studies, Yamaguchi University, 1677‑1 Yoshida, Yamaguchi 753‑8541,
Japan.
Received: 10 April 2023 Accepted: 2 February 2024

Additional file 4. Table S3. Measured uplift and inferred age along
seismic horizons. 		
Acknowledgements
We are grateful to Editor Ken Ikehara for handling our manuscript, and to
Karl Mueller and an anonymous reviewer for their constructive and insightfulreviews that have helped improve this contribution. We express our gratitude
to Yoshiji Yano, Masaki Fuchi, and crew members of the T/S Fukaemaru for theoperation of the ship during our surveys. We are grateful to Morihito Ishu forthe operation of the R/V Onokoro, and Masata Suzuki and Satomi Takagi for
their assistance on the ship. We are thankful to Shinya Uwai and staff members
from the Research Center for Inland Seas, Kobe University (KURCIS) for their
hospitality during our stay at the Marine Site (Joint Usage/Education Centerof MEXT). We are grateful to Takahiko Inoue, Ayanori Misawa, Tomoyuki Sato,
Fumitoshi Murakami, Kohsaku Arai, and Rei Nakashima for useful discussions.
We are thankful to Takuya Nishimura for the valuable discussions and insights
on geodetic strain rates around the Osaka Bay. We express our gratitude to
Tomoichi Enomoto, Shin Terado, Mitsutec Corp., Hiruzen Institute for Geology
and Chronology Co., Ltd., Masayoshi Kondo, Nakagawa Textile Co., Ltd., andSeisgadget Ltd., for their contribution in the academic crowdfunding toinvestigate the active faults in the Osaka Bay (Academist Project #111). We
thank all the students and staffs that have participated in the seismic surveys
during the marine geophysics field class held during the period of 2017–2021
onboard the T/S Fukaemaru at Kobe University.
Author contributions
MH, HO, YS, JA, TM, NS, SS, HI, MS, KS, KK and KN contributed in data acquisition. HI, HO, JA and YS conducted seismic data processing. MS carried out
processing of bathymetry and backscatter data. MH analyzed the data andconstructed the manuscript. YT and HS acquired funding. MH, HO, YS, JA,
NS, YY, SAB, HS and YT contributed in investigation, data interpretation, andconceptualization. All authors participated in reviewing the manuscript andapproving the final version.
Funding
This research was supported by the Kobe-Ocean Bottom Exploration Center,
Geological Survey of Japan, and academic crowdfunding to investigate theactive faults in the Osaka Bay (Academist Project #111). MH received travel
support to present at the International Joint Workshop on Slow-to-Fast
Earthquakes 2023 supported by JSPS KAKENHI Grant Number JP21H05200
in Grant-in-Aid for Transformative Research Areas (A) “Science of Slow-to-Fast
Earthquakes”.

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