Takao et al. Progress in Earth and Planetary Science
https://doi.org/10.1186/s40645-020-00343-1(2020) 7:37

RESEARCH ARTICLE

Progress in Earth andPlanetary Science

Open Access

Spatial and temporal variations indepositional systems in the Kazusa Group:
insights into the origins of deep-water
massive sandstones in a Pleistocene forearc
basin on the Boso Peninsula, Japan
Akihiko Takao1, Keisuke Nakamura2,6, Shinichi Takaoka3, Masaya Fuse4, Yohei Oda5, Yasushi Shimano6,8,
Naohisa Nishida7 and Makoto Ito8*

Abstract
A detailed chronostratigraphic framework established by the mapping of tephra key beds and application ofoxygen isotopic data allows assessment of the synchroneity and diachroneity of depositional systems formed incoastal and deep-water environments. This framework also allows estimation of the timing of active delivery ofcoarse-grained sediments beyond the shelf margin in relation to relative sea-level changes. The depositional
processes of deep-water massive sandstones (DWMSs) are still enigmatic; their formation is a result of active
delivery of sands in association with the supply of organic carbon into deep-water environments. DWMSs are also
important as reservoirs for hydrocarbon explorations. This study investigated the origins of DWMSs in the upper
Umegase, Kokumoto, and Chonan formations (in ascending order) of the Pleistocene Kazusa Group on the Boso
Peninsula, central Japan. Each formation contains several packets of DWMSs that are interpreted to have formed in
response to the progradation of gravelly shelf-margin deltas or fan deltas during the falling and lowstand stages ofrelative sea-level changes controlled primarily by glacioeustasy. The development of DWMSs and associatedsandstone beds is interpreted to have been induced by hyperpycnal flows, in association with sediment gravity
flows that were initiated by breaching and/or collapse of sandy substrates on the shelf-margin deltas or fan deltas.
The timings of the initial and final deposition of the packets vary within and between the formations, and areconsidered to have been controlled by the interaction between allogenic and autogenic processes operating in thegravelly shelf-margin deltas or fan deltas. A muddy horizon that contains the Lower–Middle Pleistocene Subseries
boundary (the base of the Chibanian Stage) in the Kokumoto Formation is also underlain and overlain by the
packets and represents a deposit formed in a condensed section in an upper slope environment. This depositionalsetting may have favored the development of the Global Boundary Stratotype Section and Point (GSSP) for the
Lower–Middle Pleistocene Subseries boundary in the formation.
Keywords: Deep-water massive sandstones, Sediment gravity flows, Shelf-margin delta and fan delta, Slope apron,
Hyperpycnal flow, Glacioeustasy, Falling and lowstand stages, Kazusa Group* Correspondence: mito@faculty.chiba-u.jp
8Department of Earth Sciences, Chiba University, Chiba 263-8522, Japan
Full list of author information is available at the end of the article© The Author(s). 2020 Open Access This article is licensed under a Creative Commons Attribution 4.0 International License,
which permits use, sharing, adaptation, distribution and reproduction in any medium or format, as long as you give
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Takao et al. Progress in Earth and Planetary Science(2020) 7:37

Introduction
Mapping of tephra key beds within heterogeneous siliciclastic successions can provide physically defined timelines for establishing detailed chronostratigraphic correlations of various types of lithofacies assemblages that
document different depositional processes and environments. These physically defined chronostratigraphic
markers enable identification of the synchronous or diachronous development of distinctive lithofacies assemblages, bounding surfaces, and depositional cycles inonshore outcrops (Ito 1995). In particular, the identification of local and regional synchroneity and diachroneityof stratigraphic cyclicities is crucial for elucidating
spatial and temporal variations in the interaction between allogenic and autogenic controls (i.e., external andinternal forces to the system, respectively, in the sense ofBlum et al. 2018) on stratal formation within and between sedimentary basins. Tephra key beds are commonly intercalated in siliciclastic successions that areformed in active margin basins, such as forearc andbackarc basins in the Japanese islands, and have beenused for local and regional chronostratigraphic correlations and dating of both marine and nonmarine sedimentary successions (e.g., Machida and Arai 2011).
These correlations and dating have enabled highresolution analyses of the origins of stratigraphic cyclicities in terms of glacial and interglacial sea-level cycles,
particularly during the Pliocene and Pleistocene (e.g., Ito
and Katsura 1992; Sakai and Masuda 1996; Pickering
et al. 1999).
A thick Pleistocene marine succession that is definedas the Kazusa Group (maximum thickness ~ 3000 m) isdeveloped on the central part of the Boso Peninsula,central Japan (Figs. 1, 2, and S1). The group is characterized by high sedimentation rates (1.5 m/1000 year on
average; Ito and Katsura 1992), and contains many
tephra key beds, which have enabled high-resolution
stratigraphic analyses of the group (e.g., Okada andNiitsuma 1989; Ito and Katsura 1992; Satoguchi 1995;
Pickering et al. 1999; Tamura et al. 2019). Part of the
muddy interval of this group in the Kokumoto Formation (the Chiba composite section) has been assigned as
an international type section for the Lower–Middle
Pleistocene Subseries boundary (the base of the Chibanian Stage (0.774 Ma: International Commission of Stratigraphy, 2020) (Kazaoka et al. 2015; Suganuma et al.
2018) (Figs. 2 and S1). The high-resolution stratigraphic
framework for the Kazusa Group allows detailed analyses of the spatial and temporal variations in depositional systems in terms of relative sea-level changes,
which are considered to have been controlled primarily
by glacioeustasy (Ito and Katsura 1992). However, therelative timing of different types of depositional systems,
which formed in between coastal, shallow-marine, andPage 2 of 21

deep-water environments, has not yet been precisely analyzed in terms of the depositional processes and development of lithofacies architecture from coastal andshallow-marine to deep-water systems.
In particular, the muddy interval that contains the
Chiba composite section is underlain and overlain bydeep-water deposits, which are characterized by thedevelopment of thick- to very thick-bedded (cf.,Campbell 1967) sandstones in the Kokumoto Formation. These sandstones are commonly ungraded or
weakly graded in the uppermost part of the beds and/or convoluted, and can be classified as deep-water
massive sandstones (DWMSs) (in the sense of Stowand Johansson 2000). DWMSs similar to those in the
Kokumoto Formation are also developed in the upper
Umegase and Chonan formations of the Kazusa
Group. Although DWMSs are considered to be important reservoirs in many hydrocarbon systems
worldwide (Stow and Mayall 2000), their origins anddepositional processes still remain controversial (Stevenson and Peakall 2010). In addition, DWMSs, aswell as other types of sandstone beds formed in deepwater environments, are generally interpreted to havebeen formed by sediment gravity flows (sensu lato),
which are an important type of underwater flows andmove huge amounts of sediments on Earth (Talling
et al., 2013b). These flows are also interpreted to havebeen responsible for the fluctuations in the carbon
budget and storage in the deep-water environments
through time, and therefore have played an importantrole in modulating CO2 levels in the atmosphere(Galy et al. 2007; Talling et al., 2013b; Hodgson et al.
2018). A better understanding of the origins of
DWMSs in the Kazusa Group can constrain one typeof variations in the possible processes and timing ofactive delivery of sands, which may also have been associated with the supply of organic carbon, into deepwater environments. Understanding of DWMSs is also
crucial for establishing an accurate depositionalframework for the development of the Chiba composite section that contains the Lower–Middle Pleistocene Subseries boundary (the base of the ChibanianStage) in the group.
The major objective of the present study is to clarify the origins of thick- to very thick-bedded sandstones that can be identified as DWMSs in the
Kazusa Group in relation to coeval coastal andshallow-marine deposits on the basis of the chronostratigraphic framework that is defined by the mapping of tephra key beds. In particular, one aim of thestudy is to clarify the relative timing of the active delivery of coarse-grained sediments beyond the shelf
margin to develop DWMSs within the framework of
the oxygen isotopic sea-level index.

Takao et al. Progress in Earth and Planetary Science(2020) 7:37

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Fig. 1 Geologic sketch map of the Boso Peninsula and the adjacent area. Modified from Ito et al. (2016). The inset map shows the present-day
plate-tectonic framework of the Boso Peninsula and its adjacent area (modified from Ito 2019). SRFD in the inset map denotes the location of themodern Sakawa River fan delta

Geologic setting
The Kazusa Group is an infill of the Pleistocene Kazusa
forearc basin. The basin is interpreted to have developed
in response to the west-northwestward subduction of
the Pacific plate beneath the Eurasia plate (later the
Okhotsk/North American plate; Nakajima and Hasegawa
2010) along the Izu–Bonin trench, which was superimposed by the northwestward subduction of the Philippine Sea plate beneath the Eurasia plate along the
Sagami trough between 2.4 and 0.45 Ma (Ito and
Masuda 1988; Ito and Katsura 1992) (Fig. 1). The lower
part of the group is characterized by deep-water basin
plain, submarine-fan, and lower-slope deposits in thenortheastern downslope area, and by upper-slope andouter-shelf deposits in the southwestern upslope area
(Ito and Katsura 1992) (Fig. S1). In contrast, upper-slope

and outer-shelf deposits are widely developed in theupper part of the group, in association with minor
inner-shelf and coastal deposits in the southwestern
upslope area (Fig. S1). The lithofacies succession of theentire Kazusa Group has been classified into 17 highfrequency depositional sequences, which are considered
to have been controlled primarily by glacioeustatic sealevel changes that were superimposed on the forearc tectonics, and which comprise transgressive and highstandsequence sets of a third-order composite sequence witha duration of about 1.95 Myr. (Ito and Katsura 1992).
DWMSs are developed in several stratigraphic horizons of the upper Umegase, Kokumoto, and Chonan formations of the Kazusa Group, in ascending order (Fig.
2). These DWMS-containing stratigraphic horizons are
hereafter defined as packets: PU1 and PU2 in the upperTakao et al. Progress in Earth and Planetary Science(2020) 7:37

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Fig. 2 Stratigraphy of the Kazusa Group on the Boso Peninsula. Modified from Ito (1992). Solid orange lines indicate stratigraphic horizons thatare characterized by the development of deep-water massive sandstones. Brown color indicates the stratigraphic units for which descriptions andinterpretations of lithofacies features are given in the text. The red arrow between B and C indicates the stratigraphic position of the Chiba
composite section. Data sources for this figure are given by Ito (1992). The age of the Matuyama–Brunhes paleomagnetic boundary is corrected
to 0.773 Ma following the International Chronostratigraphic Chart v. 2020/01 (International Commission on Stratigraphy, 2020)

Umegase Formation; PK1 and PK2 in the Kokumoto
Formation; and PC1–PC4 in the Chonan Formation(Figs. 3, 4, and 5a). In general, each packet is defined byabrupt replacement of muddy deposits with thick- tovery thick-bedded sandstones at the base and also by an
abrupt change from thick- to very thick-bedded sandstones to muddy deposits at the top, although somethick- to very thick-bedded sandstones show abruptchanges in thickness in the upslope and/or downslope
directions (Figs. 3, 4, and 5a). Downslope correlation of
the base and top of each packet is also supported by thecorrelation of some tephra key beds that are intercalated
in muddy deposits between the packets. The first occurrence of PU1 in the upper Umegase Formation coincides
with the abandonment of a submarine canyon–submarine
fan system in the Kazusa Group, and occurred after thedeposition of the U3 tephra key bed (Ito 1998; Ito andSaito 2006) (Fig. 3). The submarine canyon is interpreted
to have been about 8 km long, up to 1 km wide, and morethan 100 m deep (Ito and Saito 2006), and is considered
to have acted as a conduit for northeastward-flowing

turbidity currents and other types of sediment gravity
flows from which submarine-fan deposits (the Otadai and
Umegase formations) with a maximum thickness of 1050
m formed between ca 1.2 and ca 0.8 Ma (Hirayama andNakajima 1977) (Figs. 2 and S1).
Even after the abandonment of the submarine canyon in the Kazusa forearc basin, active shedding ofcoarse-grained sediments into the upper-slope or
deeper environments eventually occurred, particularly
during the falling and lowstand stages (Ito 1992,
1998). This sediment shedding might have also facilitated the development of the packets in the threeformations.

Methods/experimental
We conducted detailed outcrop analyses of the upper Umegase, Kokumoto, and Chonan formations and of the coeval
coastal and shallow-marine deposits of the Chonan Formation, such as the Mandano and Nagahama formations on
the central part of the Boso Peninsula, Japan. In particular,
we conducted detailed descriptions and interpretations of

Takao et al. Progress in Earth and Planetary Science(2020) 7:37

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Fig. 3 Stratigraphic cross-section and sequence-stratigraphic classification of the upper Umegase Formation. Modified from Ito (1998). The
oxygen isotopic data are from Pickering et al. (1999) and are adjusted to the stratigraphic section 6. The inset map shows the locations of themeasured sections and the distribution of the submarine-fan system. U8–Ku5 denotes the coding of tephra key beds. SB11 is a sequenceboundary defined by Ito and Katsura (1992). See Fig. S1 for the stratigraphic position of SB11 within the Kazusa Groupthe lithofacies features of each formation in outcrops to establish the depositional processes and depositional environments of each formation on the basis of the facies analysis
methodology. Some sediment samples were also collectedfrom the outcrops to analyze vertical changes in grain sizes
within sandstone beds using a laser diffraction grain size
analyzer (Malvern Mastersizer 3000) at Niigata University,
Japan. In addition, we also conducted sequencestratigraphic analyses of the formations along with the oxygen isotopic sea-level index of the formations to assess

spatial and temporal variations in the timing of active delivery of coarse-grained sediments beyond the shelf margin,
which may have been responsible for the development of
DWMSs.

Results
Relationships between the packets and coeval shallowmarine and coastal deposits

Because the stratigraphic units that contain the packets
record an overall regression within the highstandTakao et al. Progress in Earth and Planetary Science(2020) 7:37

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Fig. 4 Stratigraphic cross-section and sequence-stratigraphic classification of the Kokumoto Formation. The oxygen isotopic data are from
Pickering et al. (1999) and are adjusted to stratigraphic section 7. The inset map shows the locations of the measured sections. Ku6–Ka2.6 denotethe coding of tephra key beds. The filled red circle on the right side of section 2 indicates a stratigraphic horizon approximately 110 cm above
the Byakubi tephra key bed (Byk) (Byk-E in Suganuma et al. 2018) that is inferred to represent the Lower–Middle Pleistocene Subseries boundary.
See Fig. S1 for the stratigraphic positions of SB12 and SB13 within the Kazusa Groupsequence set of the third-order composite sequence ofthe Kazusa Group (Ito and Katsura 1992), PC1–PC4 in
the Chonan Formation have the highest potential forclarifying the sequence-stratigraphic relationships between the development of the packets (PC1–PC4) andthat of coeval shallow-marine and coastal deposits.
PU1–PU2 and PK1–PK2, for which coeval coastal deposits are not exposed in outcrop on the Boso Peninsula(Fig. S1), are not suitable for this purpose. Consequently,
we first examine the relationship between PC1–PC4 andcoeval shallow-marine and coastal deposits (the Mandano and Nagahama formations) in the southwestern
upslope areas (Figs. 2 and 5a).

Lithofacies features of the Chonan, Mandano, andNagahama formations

Descriptions and interpretations of the major lithofacies features of the Chonan, Mandano, and Nagahama formations
have been given elsewhere (Ito 1992; Ito and Katsura, 1992;
Ito et al. 2006a; Fuse et al. 2013). Here, we expand on previous studies and describe the distinctive lithofacies featuresof these formations to clarify the relationship between PC1–
PC4 and their coeval coastal and shallow-marine deposits.
Chonan Formation

The Chonan Formation consists of interbedded veryfine- to medium-grained sandstones, silty sandstones,

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Fig. 5 a Stratigraphic cross-section of the Chonan and Mandano formations. Modified from Ito (1992) and Ito et al. (2006a). SB14–SB16 aresequence boundaries defined by Ito and Katsura (1992). See Fig. S1 for the stratigraphic positions of SB14–SB16 within the Kazusa Group. The
coding of tephra key beds, such as Ch1–Ch3 and Ks22–Ks10, is given on the left side of sections. Thin red lines indicate the correlation of thetephra key beds, and those with single arrows show onlap or toplap terminations of time lines. Solid red lines indicate sequence boundaries. b
Tentative correlation of the timing of falling-stage deposition (a–d in a) and some tephra key beds with the oxygen isotopic sea-level index of
Bassinot et al. (1994). Thin red horizontal lines in Ch2 and Ks11 indicate error bars for the numerical ages of these tephra key beds described in
the text. The inset map shows the locations of the measured sections

and sandy siltstones, and is up to 170 m thick (Figs. 5, 6,
and S2a). The sandstone beds generally have a lenticulargeometry and show compensation stacking (Fig. 6). Internally, these beds are either ungraded, ungraded withnormal grading in the uppermost part, or entirely normally graded (Fig. S2b–c). In addition, inverse-to-normal
grading, or alternation of inverse and normal grading is
also observed in some sandstone beds (Fig. S2d–e). Parallel lamination and current-ripple cross-lamination are
also developed in the sandstone beds, except in ungraded or some normally graded sandstones (Fig. S2e–f).
Normally graded sandy siltstones with carbonaceous
fragments commonly rest on the top of normally gradedsandstones, particularly in the northeastern downslope
area (Fig. S2c). Minor hummocky cross-stratificationand dune structures are also observed in some sandstonebeds, particularly those formed in the southwestern upslope area, where the formation unconformably overlies
the Ichijiku Formation (Figs. 2 and S2g–h). Ungraded or
ungraded and normally graded (in the uppermost part)sandstones are commonly one or more meters thick andcontain convolute bedding and mudstone clasts. Locally,
the interbeds also exhibit slump structures (Fig. 6b). In
addition, the Chonan Formation is characterized by theoccurrence of cold-water molluscan shell and benthic
foraminifer assemblages that indicate outer-shelf toupper-slope environments (ca. 50–300 m paleowater
depth; Baba 1990; Kitazato 1997). The formation is moderately to intensely bioturbated with Rosselia-, Thalasinoides-, and Teichichnus-type burrows. Three tephra key
beds (Ch1, Ch2, and Ch3, in descending order) are intercalated in sandy siltstones or silty sandstones (Figs. 5aand S2a–b). Four packets (PC1–PC4) with thick- to very
thick-bedded sandstones are developed below and above
the muddy deposits that contain the three tephra key
beds (Fig. 5a).
Mandano Formation

The Ch1 and Ch2 tephra key beds are also intercalated in the Mandano Formation, the outcrops ofwhich are developed in the farther southwestern upslope area from the outcrops of the Chonan

Takao et al. Progress in Earth and Planetary Science(2020) 7:37

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Fig. 6 a Downflow (= northeastward) variations in thickness and lithofacies features of thick- to very thick-bedded sandstones in PC4 of the
Chonan Formation, Heizo, Ichihara City. b Outcrop photograph of PC4 of the Chonan Formation. A–F denote the locations of the measured
sections (A–F) in a. Arrows 1-a and 1-b, and arrows 2-a and 2-b indicate downflow thickening and thinning of sandstone beds in the northeast,
respectively. This outcrop (35° 21′ 01″ N, 140° 12′ 10″ E) is located about 400 m east of the location of section 6 in Fig. 5

Formation (Machida et al. 1980), where the formationis also underlain by the lower part of the Chonan
Formation, which contains the Ch3 tephra key bed(Figs. 5a, 7, and S3a). The Mandano Formation ischaracterized by medium- to very coarse-grainedsandstones and pebbly sandstones, and pebble-sized
conglomerates (Figs. 5a, 7, and S3b–d). The formation
is informally subdivided into lower, middle, and upperparts, with the middle part being characterized bymuddy deposits (0–17 m thick) (Figs. 7 and S4a).
The lower Mandano Formation, which has a maximum thickness of 50 m, shows an overall coarseningupward succession. The lower section contains veryfine- to medium-grained sandstones that show planar
stratification and trough cross-stratification, andminor hummocky cross-stratification, and local intercalation with silty sandstones. The middle section
comprises fine- to very coarse-grained sandstones thatshow planar stratification and trough cross- and hummocky cross-stratification, in association with waveripple lamination, and intercalation with pebble- togranule-sized conglomerates that also display planar

stratification and trough cross- and wave-ripple crossstratification. Finally, the upper section contains pebble- to granule-sized conglomerates (Figs. 7, S3d, andS4a). The basal contact with the underlying Chonan
Formation is sharp but locally erosional (Figs. 7 andS3a). In addition, the lower section locally shows
inclined bedding exhibiting a downlap termination to
the upper Chonan Formation (i.e., clinoform), discordance of bedding, and some scours up to 5 m deep
(Figs. 7 and S3b–c). The lower part of the conglomerates in the upper section is characterized by troughcross-stratification and local intercalations of planar
stratified and trough cross-stratified, medium- to very
coarse-grained sandstones (Figs. 7 and S4a). The basal
contact of the lower part of the conglomerates withthe underlying sandy deposits of the middle section is
generally sharp or slightly erosional and is marked bya distinct grain size break (Fig. S3d). The upper part ofthe conglomerates in the upper section is characterized byplanar stratification, which originally gently inclined in thenortheastern offshore direction, and by preferred clast imbrication inclined in the northeastern offshore direction.

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Fig. 7 Stratigraphic cross-section of the Mandano Formation. The inset map shows the distribution of the Mandano Formation and the locationsof the measured sections

The conglomerates are locally overlain by fine- tomedium-grained sandstones with wavy stratification, rootlets, and burrows. The Ch2 tephra key bed is intercalated
in the uppermost part of the upper section of the lower
Mandano Formation (Figs. 7 and S4a).
The middle Mandano Formation is represented bymuddy deposits that lie on an erosional basal surface incised into the lower Mandano Formation (locally up to 1
m deep) (Figs. 7 and S4a). The muddy deposits are sandy
siltstones and silty sandstones and commonly containleaf and wood fragments, with locally developed rootlets(Fig. 7). In addition, very fine- to medium-grained sandstones with planar stratification and trough crossstratification, and heterolithic bedding, which locally exhibit parallel lamination and current-ripple crosslamination, also occur within the muddy deposits (Fig.
7). Overall, both the muddy deposits and the intercalated
sandy deposits are moderately to intensely bioturbated.
The Ch1 tephra key bed is intercalated in the middle
part of the muddy deposits (Machida et al. 1980), where
finer-grained muddy deposits are dominant and shallowmarine and embayment molluscan shells, such as Raetellops pulchellus and Macoma incongrua, are found, particularly in a horizon lying mainly above the Ch1 tephra

key beds (Figs. 7 and S4a–b). Locally, intercalations ofcarbonaceous fragments increase in abundance again
upsection in the muddy deposits. The thickness of the
middle Mandano Formation varies from 0 to 17 m, consistent with the amount of erosion at the base of theupper Mandano Formation (Fig. 7).
The upper Mandano Formation shows an overall
convex-up, lenticular geometry, is up to 70 m thick (Fig.
5a), and consists of medium- to very coarse-grainedsandstones and pebbly sandstones that are characterized
by trough cross-stratification (height up to 4 m) andminor planar stratification (Fig. S4c). These deposits
commonly have bioturbation structures with Ophiomorpha-, Thalassinoides-, and Skolithos-type burrows, andcontain shallow-marine molluscan shells (Fig. S4d) thatare interpreted to indicate a paleowater depth of 30–250m (Baba 1990), as well as bone fragments of marine
mammals. The upper Mandano Formation fines upward
to muddy deposits of the Kasamori Formation, whichcontains the Ks12 tephra key bed in a horizon just above
the top of the upper Mandano Formation (Tokuhashi
and Endo 1984) (Fig. 5a), and includes very fine- tomedium-grained sandstones that lack any wave-built
sedimentary structures. The trough cross-stratification

Takao et al. Progress in Earth and Planetary Science(2020) 7:37

of the upper Mandano Formation generally indicateseastward–northeastward-directed paleocurrents, and theupper Mandano Formation pinches out into the middle
part of the Kasamori Formation in the northeastern
downslope area (Fig. 5a), where sandy siltstones and silty
sandstones are dominant and are intercalated with veryfine- to coarse-grained sandstone beds and contain molluscan shells that indicate a paleowater depth of 50–300m (Baba 1990).
Nagahama Formation

The upper Mandano Formation has been correlatedwith the Nagahama Formation (Machida et al. 1980;
Mitsunashi 1990; Ito 1992; Nakajima and Watanabe2005). The Nagahama Formation is developed mainlyin the southwestern upslope area of the upper Mandano Formation (Fig. 2). The outcrops of the Nagahama Formation are about 1.8 km wide and about 6
km long in the northeastern downslope direction, andthe formation has a distinct erosional basal surface
that incises into the underlying Chonan and Ichijiku
formations and also some older muddy deposits thatare equivalent to the Kokumoto and Umegase formations (Nakajima and Watanabe 2005) (Fig. S5a–b).
The Nagahama Formation contains a basal pebblesized conglomerate that locally contains mudstone
rip-up clasts up to 2 m long. The conglomerates fine
upward to fine- to medium-grained, medium- to very
thick-bedded, amalgamated sandstones and pebbly
sandstones that commonly have convolute bedding,and finally to interbedded very fine- to very coarsegrained sandstones and sandy siltstones that are locally associated with slumped deposits, yielding a total
thickness of up to 30 m (Fig. S5a–b). The conglomerates in the basal part of the formation show gently
undulating waveforms in outcrops oriented nearlyparallel to the northeastward-directed downslope direction, and also locally have gently inclined beddingin the northeast (Fig. S5a). In contrast, the conglomerates in outcrops oriented obliquely or orthogonallyto the downslope direction show trough crossstratification (Fig. S5b). Medium- to very coarsegrained sandstones and pebbly sandstones locally reston the waveforms and internally contain gently undulating spaced stratification and/or ungraded or normally graded bedding (Fig. S5a). The formation fines
upward to sandy siltstones of the Sanuki Formation,
which are intercalated with the Sn15 tephra key bed
that is equivalent to the Ks11.5 tephra key bed in the
Kasamori Formation (Machida et al. 1980) (Figs. 5aand S1). The Nagahama Formation is overlain by thesouthwestern extension of the upper Mandano Formation; the contact between the two formations is aslight erosional surface (Fig. S5c).

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Lithofacies organization of the packets

To clarify the origins of the packets in the upper Umegase, Kokumoto, and Chonan formations that containthick- to very thick-bedded DWMSs, we investigated thecommon lithofacies features of the sandstone beds inthe packets, incorporating the descriptions of lithofacies
features of the Chonan Formation provided above, andtheir inferred formation processes. The lithofacies assemblages of the sandstone beds that constitute thepackets contain six common lithofacies types (types 1–6), as summarized in Table 1. Types 1 and 2 characterize
the packets and can be identified as DWMSs, in association with types 3–6. Types 3–6 are also found in
muddy intervals; these intervals are characterized by bioturbated sandy siltstones and/or silty sandstones intercalated with very thin- to medium-bedded types 3–6
sandstone beds in the three formations.
Types 1 and 2 sandstone beds commonly show abrupt
thinning into the other types in both depositional dip
and strike directions (Figs. 8, 9, and 10). Some types 1
and 2 sandstone beds also locally contain several flow
units that are separated by muddy deposits (Ito et al.
2006b) (Fig. 9). In a section oriented oblique to their
paleocurrent directions in the northeast, some thinned
sandstone beds can be farther correlated with thicker
and coarser sandstone beds that formed in a more
northeastern downslope area (Fig. 10). Types 3 and 4
sandstone beds show vertical changes in grain sizes andsedimentary structures that differ from those in sandstone beds, which have been interpreted as deposits
from waning turbidity currents (i.e., the Bouma model:
Bouma 1962; the Lowe model: Lowe 1982; Talling et al.
2012). Locally, type 4 sandstone beds contain internal
erosional surfaces and siltstone streaks. Types 5 and 6
sandstone beds are mainly medium- to very thin-bedded(in the sense of Campbell 1967) (Table 1), and are also
common in muddy deposits below and above thepackets. Overall, the assemblages of the six types ofsandstone beds in each packet do not show any distinct
thinning- and fining-upward or thickening- andcoarsening-upward patterns.

Discussion
Depositional environments of the Chonan, Mandano, andNagahama formations
Chonan Formation

The Chonan Formation is interpreted to have formed inouter-shelf to upper-slope environments, which intermittently received an active supply of sandy sedimentfrom the southwestern hinterlands via sediment gravity
flows (Ito 1992; Ito and Katsura 1992; Fuse et al. 2013)(Fig. 11). Some sandstone beds show vertical changes ingrain sizes and sedimentary structures that differ fromthose of the Bouma and Lowe models (Bouma 1962;

Takao et al. Progress in Earth and Planetary Science(2020) 7:37

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Table 1 Lithofacies types and their inferred depositional
processes. The outcrop features of each lithofacies type areshown in Fig. S6.

Table 1 Lithofacies types and their inferred depositional
processes. The outcrop features of each lithofacies type areshown in Fig. S6. (Continued)

Type 1 Ungraded sandstones with convolution
Thickness: 50 cm–up to more than 250 cm
Description: Medium- to coarse-grained sandstones with sharp and flator locally distinct erosional bases. No distinct sedimentary structures aredeveloped and convolute structures are common. Locally, internal erosion surfaces in association with mudstone clasts are developed, andscattered coarse- to very coarse-grained sands and granules are also locally observed.
Interpretation: Rapid fall out of suspended particles from locallyerosional high-density turbulent flows, and/or recurrent fall out of suspended particles from sustained high-density gravity flows (Lowe 1982;
Kneller and Branney 1995; Talling et al. 2012), in association with bypassing of sedimentation from wakes. Locally developed internal erosion surfaces and mudstone clasts suggest amalgamation of thinnercomponent sandstone beds.

grading in local association with parallel- and/or current ripplelaminations. Locally, this type sandstones fine upward to sandy siltstones
that commonly contain carbonaceous fragments. Sharp and flat basal
surfaces are common with locally developed distinct erosional basal surfaces. Mudstone clasts are included in the middle and upper parts of asingle bed.
Interpretation: Locally erosive, waning, lower density turbidity currents(Talling et al. 2012). Carbonaceous fragments suggest an influence ofhyperpycnal flows (Mulder et al. 2003; Zavala et al. 2006).

Type 2 Ungraded sandstones with normal grading in the uppermost
part
Thickness: 50 cm–up to more than 250 cm
Description: Major lithofacies features are very similar to those of type 1
except for the development of normally graded sandstones up to 10
cm in the uppermost part. Graded sandstones are medium- to very finegrained and locally contain parallel- and/or current ripple-laminations
and convolute laminations.
Interpretation: Rapid fall out of suspended particles from locallyerosional high-density turbulent flows, and/or recurrent fall out of suspended particles from sustained high-density gravity flows (Lowe 1982;
Kneller and Branney 1995; Talling et al. 2012), in association with fall outof finer grained particles from wakes.
Type 3 Inverse-to-normally graded sandstones with current ripple- andparallel-laminations
Thickness: 20–40 cm
Description: Inverse-to-normally graded, medium- to very fine-grainedsandstones, in association with current ripple- and/or parallellaminations. The basal and upper surfaces of sandstone beds are sharp
and flat, and locally upper contacts are gradational to interbedded
sandy siltstones.
Interpretation: Invers-to-normal grading indicates waxing and waning oflower density gravity flows, possibly ignited by hyperpycnal flows(Mulder et al. 2003).
Type 4 Alternation of current ripple- and parallel-laminated sandstones
Thickness: 20–150 cm
Description: Multiple repetition of inverse-to-normally graded, mediumto very fine-grained sandstones in association with current ripple- andparallel-laminations. Structureless and/or normally graded, medium- tovery fine-grained sandstones are also locally intercalated.
Internal erosion surfaces and/or muddy streaks are also locallyintercalated.
Interpretation: Alternation of current ripple- and parallel-laminations, inassociation with internal erosion surfaces and muddy streaks indicate
multicycles of waxing and waning gravity flows, possibly induced byhyperpycnal flows and/or breaching of sandy substrates (Mastbergen
and Van den Bergen 2003; Mulder et al. 2003)
Type 5 Structureless sandstones
Thickness: 3–30 cm
Description: Medium- to fine-grained sandstones without any distinct
normal and/or inverse grading, and sedimentary structures, except for
locally developed convolute laminations. Sharp and flat basal and uppersurfaces are common, and some beds have distinct erosional bases andcontain mudstone clasts in the basal parts.
Interpretation: Rapid sediment fall out of turbulent suspension in locallyerosional gravity flows (Talling et al. 2012).
Type 6 Normally graded or parallel- and/or current ripple-laminatedsandstones
Thickness: 3–50 cm
Description: Medium- to very fine-grained sandstones with normal

Lowe 1982), and are interpreted to have formed fromunsteady turbidity currents, which exhibit temporal fluctuations in terms of waxing and waning of flow speeds(Kneller 1995). Hyperpycnal flows (in the sense ofMulder et al. 2003; Zavala et al. 2006) (Fuse et al. 2013)
and/or breaching-induced turbidity currents (Mastbergen and Van Den Berg 2003) are considered to havebeen responsible for the development of unsteady turbidity currents. Thick- to very thick-bedded sandstones,
which are ungraded or ungraded with normal grading in
the uppermost part, are interpreted to have formed fromrapid suspension fall out of sand particles from highdensity sediment gravity flows (e.g., Lowe 1982; Talling
et al. 2012) and/or continued suspension fall out of sandparticles from quasi-steady and depletive turbidity currents (Kneller 1995). Normally graded sandy siltstones,
which gradationally overlie graded sandstone beds, areinterpreted to represent turbiditic muddy deposits (Talling et al. 2012), although laminated and ungradedmuddy deposits, which also characterize muddy turbidites (Stow and Shanmugam 1980; Talling et al. 2012),
are uncommon in the present examples. Locally observed hummocky cross-stratification in sandstone interbeds in the southwestern upslope area indicates theepisodic influences of oscillatory current-dominatedcombined flows in an outer-shelf environment (e.g., Dumas et al. 2005), although the origin of hummocky
cross-stratification remains disputed (Quin 2011; Morsilli et al. 2012).
Mandano Formation

The lower section of the lower Mandano Formation
locally contains storm-wave-built sedimentary structures and is interpreted to have formed in a sandy
inner-shelf environment, which is equivalent to, forexample, the modern shelves around California andthe Boso Peninsula (Clifton 2006; Nishida et al. 2019).
The basal surface of the lower Mandano Formation isconsidered to have formed as a storm-wave-induced
erosional surface in response to progradation of asandy shelf system in association with offshoredirected migration of the storm wave base (Petersand Loss 2012) (Fig. 11a). Sandstones intercalated

Takao et al. Progress in Earth and Planetary Science(2020) 7:37

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Fig. 8 Downflow changes in thickness, lithofacies types, and grain sizes of a sandstone bed developed just above the U2 tephra key bed in the
Umegase Formation. The inset map shows the locations of the measured sections. Black filled circles in grain-size diagrams denote volume
percentages of muddy grains less than 20 μm. See Table 1 for descriptions and interpretations of lithofacies types

with conglomerates in the middle section of the lower
Mandano Formation also contain distinctive stormwave-built sedimentary structures and are interpreted
to be lower shoreface deposits (Massari and Parea
1988; Hart and Plint 1989; Ito 1992; Ito et al. 2006b).
The lower part of the conglomerates in the upper
section of the lower Mandano Formation is considered to record the migration and aggradation of gravelly bars in an upper shoreface environment, whereas
the upper part of the conglomerates may have formedin response to swash and backwash processes in agravelly foreshore environment (Hart and Plint 1989;
Clifton 2006; Bluck 2011). Locally observed fine- tomedium-grained sandstones with wavy stratification,
rootlets, and burrows in the uppermost part of theupper section are interpreted to have formed as eolian dune deposits in a backshore environment (Clifton 2006). The sharp and locally erosional basal
contact of the upper shoreface deposits is considered
to have formed by the migration of a gravel bar andtrough system in an upper shoreface environment andis equivalent to a surf diastem (in the sense of Swiftet al. 2003).
The distinct erosional basal surface of the middle
Mandano Formation suggests an unconformity thatwas developed in response to a fall in relative sea
level. This relative sea-level fall may have induced

preceding forced regression to form the lower Mandano Formation before the development of the basal
unconformity of the middle Mandano Formation. The
muddy deposits with carbonaceous fragments androotlets overlying the erosional surface are considered
to have formed in response to the reduction incoarse-grained sediment supply during the ensuingtransgression. The muddy deposits in the middle
Mandano Formation are interpreted to have formed
initially in an estuary associated with mudflats andsalt-marsh environments, which might have been affected by subsequent inundation of seawater to develop a shallow-marine muddy embayment (e.g., Allen
1997; Davidson-Arnott 2010) (Fig. 11c). The development of heterolithic bedding in the muddy deposits
suggests reduced wave processes or the influence oftidal processes in the estuary (e.g., Boyd et al. 2006;
Ahokas et al. 2014). Because the coastal gravelly deposits are considered to have been formed by active
wave processes, longshore currents may have played
an important role in alongshore transportation ofcoarse sediments to form spits at a river mouth (Burningham 2015) (Fig. 11c). The shallow-marine muddy
embayment that was documented in the middle Mandano Formation appears to have formed in responseto the development of gravelly spits (see Hiroki and
Masuda 2000; Nielsen and Johannessen 2008;

Takao et al. Progress in Earth and Planetary Science(2020) 7:37

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Fig. 9 Downflow changes in lithofacies types of sandstone beds above the Ku1 tephra key bed in PK2 of the Kokumoto Formation along the Yoro River(Location 4 in Fig. 4). The inset map shows the locations of the measured sections. See Table 1 for descriptions and interpretations of lithofacies types

Takao et al. Progress in Earth and Planetary Science(2020) 7:37

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Fig. 10 Downflow changes in thickness and lithofacies types of two sandstone beds below the Ku1 tephra key bed. The inset map shows thelocations of the measured sections. See Table 1 for descriptions and interpretations of lithofacies types. The trend of this stratigraphic crosssection (74°) is oblique to the paleocurrent directions (47–70°)

Fig. 11 Schematic illustration of temporal variations in the depositional systems in which the packets of thick- to very thick-bedded sandstonesin the upper Umegase, Kokumoto, and Chonan formations developed

Takao et al. Progress in Earth and Planetary Science(2020) 7:37

Bartholdy et al. 2017), although distinct platform foresets that indicate the migration of gravelly spits are
not observed in the present outcrop belt.
An overall convex-up lenticular geometry, unidirectional paleocurrents, and downslope termination into
the Kasamori Formation, which is considered to haveformed in an outer-shelf environment, indicate thatthe upper Mandano Formation represents shelf sandridge deposits formed by an overall transgression (Ito
1992). The basal erosional surface of the upper Mandano Formation represents a ravinement surface, andthe development of the sand-ridge deposits is interpreted to have been influenced by the northeastwarddirected intrusion of the paleo-Kuroshio Current into
the Kazusa forearc basin in response to glacioeustatic
sea-level rise (Ito 1992).
Nagahama Formation

The basal erosional surface of the Nagahama Formationand the overlying conglomerates indicate that the deposition of this formation was influenced by preceding erosion by sediment gravity flows in shelf and shelf-margin
environments and the subsequent deposition in response
to downslope migration and aggradation of gravelly bedforms from high-density sediment gravity flows (cf., Ito2019). The shallow incision of its base and its thin stratigraphic thickness may indicate that the Nagahama Formation was formed as an infill of a shallow valley on anouter-shelf and/or shelf margin (Ito 1992; Ito et al.
2006a) (Fig. 11b). An overall fining-upward pattern inthe entire Nagahama Formation is considered to havebeen formed by backfilling of the valley in response toan overall rise in relative sea level.
Sequence stratigraphy of the Chonan, Mandano, andNagahama formations

The Chonan Formation is interpreted to have formed
mainly during the marine isotopic stage (MIS) 16(Okada and Niitsuma 1989; Pickering et al. 1999), andthe Ch2 tephra key bed has yielded a zircon U–Pb ageof 0.61 ± 0.02 Ma (Ito et al. 2017). In addition, the Ks11
tephra key bed in the Kasamori Formation is interpreted
to have formed during MIS 13.3 (in the sense of Bassinot
et al. 1994) (0.52 ± 0.02 Ma) (Shirai 2001). A stratigraphic horizon between the Ks11 and Ks 12 tephra key
beds, which are intercalated in muddy deposits that developed over the sand-ridge deposits of the upper Mandano Formation, was assigned as a sequence boundary(SB16 in Ito and Katsura 1992). The present interpretation of lithofacies architecture and bounding surface
features (such as the sharp base of the lower Mandano
Formation, the erosional bases of both the middle andupper Mandano Formation, and the reinterpretation of
the stratigraphic relationship between the Nagahama

Page 15 of 21

and upper Mandano formations), together with the revised age framework of the Chonan, Mandano, andNagahama formations, enable the three formations to be
subdivided into four distinctive depositional stages interms of higher-frequency sequence stratigraphy compared to the previous interpretation of Ito and Katsura(1992) (Figs. 5b and S1).
The development of four packets (PC1–PC4) is considered to have been induced by active shedding ofcoarse-grained clastic particles by both low-density
and high-density sediment gravity flows into shelfmargin to upper-slope environments during the falling
and lowstand stages of relative sea-level changes (Ito
1992, 1995; Ito and Katsura 1992) (Fig. 11a). Sandstone beds coarser than medium grain size, such as
the types 1 and 2 beds, are considered to haveformed from high-density gravity currents (Mulder
and Alexander 2001; Talling et al. 2012).
The muddy horizon that is intercalated with the Ch3
tephra key bed is interpreted to have formed during MIS
16 (Nanayama et al. 2016). This muddy horizon isencased between PC1 and PC2 (Fig. 5a). In addition,
PC2 is overlain by an overall coarsening-upward succession in the upper Chonan Formation that is replaced by
the lower Mandano Formation with a sharp basal contact (Fig. 5a). From this stratigraphic architecture, weinfer that PC1 and PC2 may have formed during an
overall glacioeustatic sea-level fall from MIS 17 to MIS
16. This sea-level fall was interrupted by a short-term
still-stand at MIS 16.3, which is marked by the development of a muddy interval that contains the Ch3 tephra
key bed (Fig. 5b).
An abrupt glacioeustatic sea-level rise and the following fall from MIS 16.2 to MIS 15.4 formed an overall
coarsening-upward succession, which is interpreted asconsisting of transgressive and highstand deposits of theupper Chonan Formation and falling-stage deposits of
the lower Mandano Formation in the southwestern upslope area (Fig. 5). The sharp basal surface of the lower
Mandano Formation can be assigned to a regressive surface of marine erosion (in the sense of Plint and Nummedal 2000). Because the Ch2 tephra key bed isintercalated in the uppermost part of the lower Mandano Formation and the lowest part of PC3 (Machida
et al. 1980; Tokuhashi and Endo 1984) (Fig. 5a), shedding of coarse-grained sediments beyond the shelf margin may have commenced during the late falling stage,
just before the uppermost part of the lower Mandano
Formation was subjected to slight incision that enabled
the bypassing of coarse-grained sediments to form PC3
in the northeastern downslope area. The ensuing transgression and regression are considered to have formed
the marsh and shallow-marine embayment deposits ofthe middle Mandano Formation from MIS 15.4 to MIS

Takao et al. Progress in Earth and Planetary Science(2020) 7:37

15.3 (Fig. 5b). During this period, the muddy intervalthat contains the Ch1 tephra key bed replaced PC3 in
the northeastern downslope area (Fig. 5a).
The subsequent glacioeustatic sea-level fall fromMIS 15.3 to MIS 15.2 is interpreted to have caused
development of a shallow valley that represents thebase of the Nagahama Formation and may have acted
as a conduit to form PC4 in the northeastern downslope area (Fig. 11b). Backfilling of the valley is considered to have occurred during the late falling andearly transgressive stages, and the active supply ofgravel and sand may have subsequently declined. Because the coarse-grained coastal and shallow-marine
deposits of the lower Mandano Formation show anabrupt change downslope into slope deposits of thecoeval Chonan Formation from locations 4 to 5 (seesections 4 and 5 in Fig. 5a), in association with a distinct clinoform that shows a downlap termination to
the underlying upper Chonan Formation (Fig. S3b),
the lower Mandano Formation is considered to haveformed as a coarse-grained, shelf-margin delta or fandelta system (in the sense of Porębski and Steel 2003)(Fig. 11). The clinoform and discordance in bedding
within the lower section of the lower Mandano Formation (Fig. S3b–c) are interpreted to reflect progradation and switching of depocenters of the shelfmargin delta or fan delta.
PC4 was first onlapped by the outer-shelf depositsof the lower Kasamori Formation in response to theglacioeustatic sea-level rise from MIS 15.2 to MIS
15.1 (Fig. 5). Subsequently, the middle Mandano andNagahama formations were further transgressed by
the sand-ridge deposits of the upper Mandano Formation; these sand-ridge deposits were succeeded byouter-shelf muddy deposits of the middle Kasamori
Formation that contains the Ks12 tephra key bed
and is interpreted to comprise transgressive andhighstand systems tracts (Ito 1992) (Fig. 5a).

Origins of the packets

In terms of the sequence-stratigraphic architectures of
the Chonan, Mandano, and Nagahama formations,
PC1–PC4 are interpreted to have formed during thefalling and lowstand stages of glacioeustatic sea-levelchanges as a response to progradation of a coarsegrained shelf-margin delta or fan delta (Figs. 5 and11). PU1–PU2 in the upper Umegase Formation and
PK1–PK2 in the Kokumoto Formation are also interpreted to have developed in depositional andsequence-stratigraphic settings similar to those of the
Chonan Formation, although their coeval coastal depositional systems are not exposed in outcrop (Figs. 3
and 4).

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Depositional processes of lithofacies types in the packets

Thick- to very thick-bedded sandstones of types 1 and 2
are interpreted to have formed by the amalgamation ofthinner sandstone beds as a result of either recurrent fallout of suspended loads from sustained high-density sediment gravity flows (Kneller and Branney 1995) or discontinuous fall out of suspended loads fromintermittently occurring waning turbidity currents.
Abrupt downflow thinning of thick- to very thickbedded sandstones is also interpreted to be a result of en
masse deposition from sandy debris flows (Shanmugam
1996; Talling et al. 2013a). Because type 1 and 2 sandstone beds commonly contain convolute bedding andpossess erosional bases, and laterally show a transition
into type 5 and 6 sandstone beds over a short distance(Fig. 10), their deposition is considered to have been induced by rapid fall out of sand particles from highdensity sediment gravity flows rather than by en masse
processes in debris flows, although flow transformationfrom sandy debris flows to turbulent gravity flows (inthe sense of Fisher 1983) is an alternative possible
process for the lateral variation in lithofacies types.
Abrupt fining and thinning of thick- to very thickbedded sandstones can alternatively be interpreted to reflect the deposition of sand particles, forming thick- tovery thick-bedded sandstones that may have resulted ina decrease in the density of the residual sediment gravity
flows, which initially contained lower-density fresh water
and were finally transformed into buoyant flows (Gladstone and Pritchard 2010). Buoyant flows are commonly
generated in hyperpycnal flows that are interpreted to be
more widely occurring in river-fed turbidite systems
than previously thought (Stevenson and Peakall 2010).
Consequently, thick- to very thick-bedded sandstones
are expected to have been deposited by hyperpycnal
flows (Plink-Björklund and Steel 2004).
Sandstone beds that have single or multiple zones of
inverse-to-normal grading in association with alternating
parallel lamination and current-ripple cross-laminationand internal erosion surfaces (types 3 and 4) are interpreted to have formed in response to single or multiple
episodes of waxing and waning of unsteady sedimentgravity flows. Hyperpycnal flows (Mulder et al. 2003;
Zavala et al. 2006) are possible candidate gravity flows
for the development of type 3 and 4 sandstone beds.
Breaching of sandy sediments in response to retrogressive failure of a sandy substrate is an alternative explanation for the initiation of single and multiple episodes ofwaxing and waning of sediment gravity flows (Mastbergen and Van Den Berg 2003). Because the developmentof the four packets (PC1–PC4) in the Chonan Formation
is interpreted to have been induced by the progradationof a gravelly shelf-margin delta or fan delta, which arecharacterized by the development of wave-dominatedTakao et al. Progress in Earth and Planetary Science(2020) 7:37

gravelly coasts, active shedding of coarse-grained sediments beyond the shelf margin is considered to havebeen promoted by hyperpycnal flows (Fuse et al. 2013),
although collapse of delta-front sediments cannot necessarily be excluded as a possible mechanism for the origins of the component sandstone beds of the packets.
Concurrent occurrence of hyperpycnal flows andcollapse-induced gravity flows has been documented in
the modern Sakawa River fan delta in the Sagami Bay,
about 75 km west of the Boso Peninsula (location
marked in the inset of Fig. 1). This fan delta underwent
a large flood event in 1972, and abundant sand andgravel, together with wood and plant fragments, weretransported from the river mouth into the fan-delta
front and farther downslope area by gravity flows. This
flood event also induced the collapse of the fan-deltaslope, generating submarine landslides and gravity flows
and associated submarine cable breaks (Otsuka et al.
1973). Subsequently, younger surface sediment samples
were collected from this fan-delta slope; some of thesesediments have been interpreted as hyperpycnites (Ikehara et al. 2018). The packets in the upper Umegase andKokumoto formations contain lithofacies assemblages
that also characterize the four packets of the Chonan
Formation, and are considered to have developed via
processes similar to those involved in the formation of
the packets in the Chonan Formation.
In general, the sandstone beds in the packets of the threeformations have a lenticular geometry, compensation stacking patterns, and laterally discontinuous distribution patterns. These architectural features suggest that thesandstone beds were formed as lenticular deposits, withlimited downslope and along-slope extents, and the depressions between sand deposits, which were formed by thepreceding depositional events, provided accommodation
volume for the later sand deposition (Fig. 11a). Of the sandstone beds that show laterally discontinuous distribution
patterns, some thinned sandstone beds can be farther correlated with thicker and coarser sandstone beds in a more
downslope area (Fig. 10). These sandstone-bed distribution
patterns suggest that these sandstone beds were sourced
from different staging points simultaneously in association
with the progradation of a shelf-margin delta or fan delta in
upslope areas (Fig. 11a). In addition, PC1 and PC2 also
show laterally discontinuous distribution patterns, andthicker sandstone beds are developed in a more downslope
area in PC1 (Fig. 5). These distribution patterns may also
indicate the development of different staging points. Consequently, the depositional setting of the packets in the threeformations is interpreted to be classified as a slope apron ora ramp, characterized by a linear source system (Heller andDickinson 1985; Stow and Johansson 2000; Stow andMayall, 2000) that enabled coeval multiple sandbodies todevelop beyond the shelf margin (Fig. 11a).

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Relative timing of formation of the packets

Active delivery of coarse-grained sediments beyond theshelf-margin into deep-water environments has been
considered to occur during the falling and lowstandstages of relative sea level in terms of the sequencestratigraphic model (e.g., Posamentier and Allen 1999).
Subsequently, the joint operation of several factors, including the positions of canyon heads relative to coastal
and river-mouth areas, the shelf width, and the intensities of longshore currents and fluvial discharges, hasbeen interpreted to vary the timing of the onset of active
delivery of coarse-grained sediments beyond the shelfmargin into deep-water environments, particularly during the Quaternary (Covault and Graham 2010; Sweet
and Blum 2016). However, these studies have focused
mainly on submarine canyon–submarine fan systems,
and the timing in other deep-water systems, such as
slope apron and ramp systems, is not yet clearly understood. Because the packets of the upper Umegase andKokumoto formations, as well as those of the Chonan
Formation, are interpreted to have formed as slope
aprons or ramps (Fig. 11), the relative timing of formation of the packets can explain some of the variation in
the active delivery of coarse-grained sediments beyond
the shelf-margin into deep-water environments in an active margin setting.
On the basis of the age framework of the upper Umegase and Kokumoto formations, we investigated the relative timing of the initial and final delivery of coarsegrained sediments in the development of PU1–PU2 in
the upper Umegase Formation and of PK1–PK2 in the
Kokumoto Formation with respect to the oxygen isotopic sea-level index as an allogenic control.
Upper Umegase Formation

The age framework of the upper Umegase Formation isbased on magnetostratigraphy, oxygen isotopic stratigraphy, and nannofossil datums (Sato et al. 1988; Okada
and Niitsuma 1989; Pickering et al. 1999). Except for
those of Pickering et al. (1999), the age data indicate thatthe stratigraphic horizons containing the U4 and U3
tephra key beds can be assigned to MIS 24 and MIS 23,
respectively (Kazaoka et al. 2015). PU1 and PU2 areinterpreted to have formed after the abandonment ofsubmarine-fan channel deposits and were included in
the falling- and lowstand deposits of DS 11 of Ito and
Katsura (1992). On the basis of the age framework described above, the packets are considered to have formed
mainly during the falling and lowstand stages from MIS
23 to MIS 22 (Fig. 3). The separation of PU1 and PU2
by a muddy horizon, which contains the U2 tephra key
bed (cf. Mitsunashi et al. 1959; Geological Survey ofJapan 1961), is considered to have developed in responseto an ephemeral lateral shift of a river mouth on theTakao et al. Progress in Earth and Planetary Science(2020) 7:37

shelf-margin delta or fan delta, or alternatively to havebeen induced by a minor glacioeustatic sea-level rise between MIS 23 and MIS 24 that interrupted an overall
sea-level fall, as proposed by Lisiecki and Raymo (2005).
Kokumoto Formation

A detailed age framework for the Kokumoto Formation
has been proposed in previous studies (Pickering et al.
1999; Kazaoka et al. 2015; Nishida et al. 2016). Tentative
correlations of the bases and tops of the two packets(PK1–PK2) were made by Nishida et al. (2016) on thebasis of the sequence-stratigraphic interpretation of Ito(1992) and Ito and Katsura (1992). PK1 and PK2 areconsidered to have been deposited between MIS 21.1
and MIS 20.2 and between MIS 18.4 and MIS 18.3, respectively (Fig. 4). Additional oxygen isotopic data froma composite section for a stratigraphic horizon between
the upper part of PK1 and the lower part of PK1 (Suganuma et al. 2018) indicate that deposition of PK2 commenced during the late stage of MIS18 when theeustatic sea level fell approximately 60 m from the highest stage of MIS 19, although a minor eustatic sea-level
rise occurred during this overall falling stage (Elderfield
et al. 2012), similar to the ephemeral rise during MIS
18.3 reported by Bassinot et al. (1994). The minor glacioeustatic sea-level rise seems to have been documented
in the formation of muddy deposits in the upper part of
PK2 in location 7 in Fig. 4. The muddy deposits, whichcontain the Lower–Middle Pleistocene Subseries boundary, show an overall fining-upward pattern above PK1(Fig. 4), and are interpreted to have formed close to thehighest stage of eustatic sea level of Elderfield et al.
(2012). This proposed boundary horizon is also considered to be in the uppermost transgressive systems tract(= downlap surface) (Fig. 4), and to have a higher potential for preserving tephra beds, although the age resolution may be limited by the development of acondensed section along the downlap surface horizon.
Variation in the relative timing of packet deposition

The timing of the initial deposition of the packets in theupper Umegase, Kokumoto, and Chonan formations varies within a single formation and between the three formations, although the packets developed mainly during
the falling and lowstand stages of glacioeustatic sea-levelchanges. The variation is considered to have been controlled by the width of the shelf, the positions of rivermouths relative to the shelf margin, the amounts ofcoarse-grained sediment discharges, and the magnitude
and rate of falls in relative sea level.
The youngest packets of each formation were abruptly
replaced by muddy deposits in the basal parts of theoverlying formations, such as the basal parts of the
Kokumoto, Kakinokidai, and Kasamori formations (Figs.

Page 18 of 21

3, 4, and 5), although the replacement of PK2 by thebasal muddy deposits of the Kakinokidai Formation occurred during the latest stage of glacioeustatic sea-level
fall (Fig. 4). An abrupt reduction in the active delivery ofcoarse-grained sediments beyond the shelf margin may
have been a response to the storage of coarse-grained
sediments at river mouths on the shelf-margin deltas or
fan deltas. The earlier replacement of PK2 with the
muddy deposits during a late stage of glacioeustatic sealevel fall is considered to have been induced by theactivation of basin subsidence, giving rise to an earlycommencement of relative sea-level rise before MIS18.2.
Autogenic processes, such as the lateral migration of astaging point for the coarse-grained sediment supply anda short-lived decrease in the sediment discharge fromthe river system, may also have caused the abrupt abandonment of packet formation for PK2 and the otherpackets.

Conclusions
Mapping of tephra key beds within heterogeneous lithofacies assemblages can provide a detailed chronostratigraphic framework for elucidating spatial and temporalvariations in depositional systems from coastal toshallow-marine and deep-water environments. This
study sought to clarify (1) the origins of DWMSs in relations to coeval coastal depositional systems, and (2) the
relative timing of the active delivery of coarse-grained
sediments beyond the shelf margin within a chronostratigraphic framework based on the mapping of tephra
key beds and oxygen isotopic data. The present casestudy was based on detailed outcrop investigations of (inascending order) the upper Umegase, Kokumoto, andChonan formations of the Pleistocene Kazusa Group on
the Boso Peninsula, central Japan. The main conclusionsof the study are as follows.
1. Packets of thick- to very thick-bedded sandstones,
which are assigned to DWMSs, are interpreted tohave developed in response to the progradation ofgravelly shelf-margin deltas or fan deltas during thefalling and lowstand stages of relative sea-levelchanges, which were controlled primarily byglacioeustasy.
2. DWMSs and associated sandstone beds generally
show a lenticular geometry and compensation
stacking patterns. Their lateral changes in bed
thickness and grains sizes suggest that they formed
in multi-sourced systems on the shelf-margin deltas
or fan deltas, such as a slope apron or a ramp.
3. Initiation of sediment gravity flows, which formed
DWMSs and associated sandstone beds, isconsidered to have been caused by hyperpycnal
flows, in association with breaching and/or collapse

Takao et al. Progress in Earth and Planetary Science(2020) 7:37

of coarse-grained substrates on the shelf-margin
deltas or fan deltas in response to flood events.
4. The relative timing of the initial deposition andabandonment of the packets varies within andbetween the three studied formations. The variationis considered to have been caused by interaction
between glacial and interglacial sea-level changes
and basin subsidence superimposed by processes
operated within the shelf-margin delta or fan-deltasystems, such as the lateral migration of a staging
point for the supply of coarse-grained sediment anda short-lived decrease in sediment discharge fromthe fluvial system.
5. Two packets are developed in the Kokumoto
Formation, below and above a muddy interval that
contains the Lower–Middle Pleistocene Subseries
boundary (GSSP). On the basis of sequencestratigraphic classification of these deposits, the
GSSP horizon is interpreted to have formed in theuppermost part of a transgressive systems tract andcan be assigned to a condensed section. This depositional setting likely promoted the development of
the international type section for the Lower–Middle
Pleistocene Subseries boundary in the Kazusa forearc basin.

Supplementary information
Supplementary information accompanies this paper at https://doi.org/10.
1186/s40645-020-00343-1.
Additional file 1: Figure S1. Sequence-stratigraphic classification andspatial and temporal variations in depositional systems of the Kazusa
Group on the Boso Peninsula, central Japan. Modified from Ito and
Katsura (1992) and Ito et al. (2016). The red arrow on the right side ofsection 13 indicates the stratigraphic position of the Chiba composite
section. The inset map shows the locations of the measured sections.
Additional file 2: Figure S2. Outcrop photographs of the Chonan
Formation. a Muddy deposits containing the Ch1 tephra key bed,
underlain and overlain by PC3 and PC4, respectively (Oto, Ichihara City). b
Very thick-bedded sandstone bed occurring just below the Ch2 tephra
key bed and containing convolute bedding and minor normal grading in
the uppermost part (Itabu, Ichihara City). c Normally graded, medium- tovery fine-grained sandstones and graded siltstones containing abundant
plant fragments (Koshikiya, Ichihara City). d Inversely and normally gradedsandstone bed with current-ripple cross-lamination and parallel
lamination in PC3 (Heizo, Ichihara City). e Alternation of current-ripple
cross-lamination and parallel-lamination within a sandstone bed that also
contains massive and normally graded bedding and an internal erosion
surface (Koshikiya, Ichihara City). f Ungraded, very thin- to thin-bedded,
medium- to fine-grained sandstone beds (arrowed) (Koshikiya, Ichihara
City). g Hummocky cross-stratified, fine-grained sandstone bed (Kururioyatsu,
Kimitsu City). h Unconformable contact between the Ichijiku andChonan formations (Okuradai, Kimitsu City).
Additional file 3: Figure S3. Outcrop photographs of the Mandano
Formation. a Sharp basal contact between the Chonan and lower
Mandano formations (Minowa, Kimitsu City). b Distinct clinoform in thelower part of the Mandano Formation (Mandano, Ichihara City). c Scourand bedding discordance in the lower section of the lower Mandano
Formation (Minowa, Kimitsu City). c Medium- to very coarse-grainedsandstones and intercalated pebble-sized conglomerates in the middle

Page 19 of 21

section of the lower Mandano Formation (Minowa, Kimitsu City). Troughcross-stratified conglomerates rest on a sharp contact.
Additional file 4: Figure S4. a Lithofacies succession from the lower tothe middle and upper parts of the Mandano Formation (Yoshino, Kimitsu
City). The basal parts of the middle and upper Mandano Formation show
slightly erosional contacts. Ch1 and Ch2 are tephra key beds that are alsointercalated in the Chonan Formation (see a and b in Fig S2). b Blocksample of muddy deposits of the middle Mandano Formation containing
shallow-marine molluscan shells (Minowa, Kimitsu City). c Lithofacies succession from the lower to upper parts of the Mandano Formation(Yoshino, Kimitsu City). Muddy deposits of the middle Mandano Formation were removed by the erosional base of the upper Mandano Formation that is characterized by trough cross-stratification and shows an
overall fining-upward pattern. d Shallow-marine molluscan fossils intercalated in trough cross-stratified medium- to very coarse-grained sandstones of the upper Mandano Formation (Yoshino, Kimitsu City).
Additional file 5: Figure S5. Outcrop photographs of the Nagahama
Formation. a Gently undulating waveforms and foreset bedding in
conglomerates and overlying pebbly sandstones in the basal part of the
Nagahama Formation (Tsukuriki, Kimitsu City). Solid and dotted red lines
indicates an unconformable contact. b Trough cross-bedded conglomerates in the basal part of the Nagahama Formation (Sasage, Futtsu City).
The solid red line indicates an unconformable contact. c Lithofacies succession from the Ichijiku to the Nagahama and the upper Mandano formations (Manobori, Kimitsu City). A person is circled for scale.
Additional file 6: Figure S6. Outcrop photographs of the lithofacies
types summarized in Table 1. Solid white and black arrows indicate thebase and top of each sandstone bed, respectively. a Very thick-bedded,
ungraded and convoluted, medium-grained sandstone bed of the Umegase Formation (Type 1). b Very thick-bedded, ungraded and convoluted
sandstone bed, with minor normal grading in the uppermost 10 cm of
the Umegase Formation (Type 2). c Climbing-ripple-laminated, fine- tovery fine-grained sandstone bed with inverse and normal grading of the
Kokumoto Formation (Type 3). d Type 4 sandstone bed with alternating
parallel laminations and current-ripple cross-laminations of the Kokumoto
Formation. e Medium-bedded, ungraded, medium- to fine-grained sandstone bed of the Umegase Formation (Type 5). f Normally graded sandstone bed with parallel laminations with lignite (L) and current-ripple
cross-laminations in the uppermost 3 cm of the Umegase Formation(Type 6).
Acknowledgements
We thank M. Okada and Y. Suganuma for inviting us to contribute our study
to this special volume of PEPS. We also thank Y. Takashimizu and hiscolleagues at Niigata University for providing access to their laboratory forgrain-size analysis. A.D. Miall is thanked for critical insights on timestratigraphy of clastic successions. Early versions of the manuscript benefitedof constructive comments from three anonymous reviewers and the editor.
Their reviews greatly improved the paper.
Authors’ contributions
All authors conducted the field researches, and collected and analyzed
outcrop data. The corresponding author wrote the manuscript withsignificant contribution of from all co-authors. All authors read and approved
the final manuscript.
Funding
This study was supported in part by JSPS KAKEN Grant Number 16K055573.
Availability of data and materials
Please contact the corresponding author for data request.
Competing interests
The authors declare that they have no competing interests.
Author details
1
Mitsubishi Corporation, 2-3-1 Marunouchi, Chiyoda-ku, Tokyo 100-8086,
Japan. 2JX Nippon Oil & Gas Exploration Co., Ltd., 1-1-2 Otemachi,
Chiyoda-ku, Tokyo 100-8163, Japan. 3Mitsui Oil Exploration Co., Ltd., 1-2-9

(2020) 7:37

Page 20 of 21

Nishi Shimbashi, Minato-ku, Tokyo 105-0003, Japan. 4Chiba Prefecture, 1-1-1
Ichiba-cho, Chiba 260-8667, Japan. 5Godo Shigen, Co. Ltd., 1365 Nanaido,
Chiba 299-4333, Japan. 6Japan Oil, Gas and Metals National Corporation,
2-10-1 Toranomon, Minato-ku, Tokyo 105-0001, Japan. 7Department ofEnvironmental Sciences, Tokyo Gakugei University, 4-1-1 Nukuikita, Tokyo
184-8501, Japan. 8Department of Earth Sciences, Chiba University, Chiba
263-8522, Japan.

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Received: 22 August 2019 Accepted: 28 May 2020

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