Izumi et al. Progress in Earth and Planetary Science
https://doi.org/10.1186/s40645-020-00393-5(2021) 8:10

RESEARCH ARTICLE

Progress in Earth andPlanetary Science

Open AccessMultiproxy sedimentological andgeochemical analyses across the Lower–
Middle Pleistocene boundary:chemostratigraphy and paleoenvironmentof the Chiba composite section, central
Japan
Kentaro Izumi1*, Yuki Haneda2, Yusuke Suganuma3,4, Makoto Okada5, Yoshimi Kubota6, Naohisa Nishida7,
Moto Kawamata4 and Takuya Matsuzaki8

Abstract
The Chiba composite section (CbCS) in the Kokumoto Formation, Kazusa Group, central Japan is a thick andcontinuous marine succession that straddles the Lower–Middle Pleistocene boundary and the well-recognized
Matuyama–Brunhes paleomagnetic polarity boundary. Although recent studies extensively investigated the CbCS,
its chemostratigraphy, particularly around the Lower–Middle Pleistocene boundary, is poorly understood. Therefore,
in this study, we performed multiproxy sedimentological and geochemical analyses of the CbCS, including the
Chiba section, which is the Global Boundary Stratotype Section and Point for defining the base of the Middle
Pleistocene Subseries. The aim of these analyses is to establish the high-resolution chemostratigraphy and toreconstruct the paleoenvironments of its sedimentary basin in detail. We used the K/Ti ratio as a broad proxy forthe clastic material grain size of the sediments. Although the K/Ti ratio generally varies throughout the studiedinterval, the K/Ti ratio especially during Marine Isotope Stage (MIS) 19a shows a variation pattern like those of theforaminiferal oxygen isotope (δ18O) records. The records of the C/N ratio of bulk samples and carbon isotope ratioof the organic carbon (δ13Corg) suggest that the organic matter in the CbCS sediments during MIS 19c mostly
originated from marine plankton, whereas the organic matter during MIS 18 and 19a was characterized by a
mixture of marine plankton and terrestrial plants. These records are clearly indicative of changes in mixing ratio of
marine vs. terrestrial organic matter in association with glacial–interglacial cycles from the late MIS 20 to the early
MIS 18. In addition, we calculated the mass accumulation rates (MARs) of organic carbon, biogenic carbonate, andterrigenous material for quantitative interpretations on the paleoenvironmental changes. MAR calculations revealed
that the contribution of marine organic carbon relative to terrestrial organic carbon increased during MIS 19c, andthat the contribution of the terrigenous material relative to biogenic carbonate decreased during MIS 19c.
(Continued on next page)* Correspondence: izumi@chiba-u.jp
1
Faculty & Graduate School of Education, Chiba University, 1-33 Yayoi-Cho,Inage-Ku, Chiba-shi, Chiba 263-8522, Japan
Full list of author information is available at the end of the article© The Author(s). 2021 Open Access This article is licensed under a Creative Commons Attribution 4.0 International License,
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Izumi et al. Progress in Earth and Planetary Science(2021) 8:10

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Furthermore, we observed relatively large variations in the total organic carbon and total nitrogen contents during
MIS 19a. These variations were probably caused by the relative decrease in bottom-water oxygen level, which isalso supported by our trace-fossil data, although it is not certain whether the increase in organic-carbon flux at ~760 ka was due to the synchronous increase in biogenic productivity in surface water. Such a relative decrease inbottom-water oxygen level was partly due to the increased ocean stratification because of the northward
displacement of the Kuroshio Extension Front.
Keywords: Chemostratigraphy, Chiba composite section, GSSP, Kokumoto Formation, Organic matter, Paleoenvironment,
Pleistocene, Trace fossils, XRF

1 Introduction
The Lower to Middle Pleistocene is an important geological period that is characterized by gradual changes in
the climate systems that affected the evolution of biota(Maasch 1988; Mudelsee and Schulz 1997; Mudelsee andStattegger 1997; Raymo et al. 1997; Head and Gibbard
2005; Clark et al. 2006; Head et al. 2008; Elderfield et al.
2012; Head and Gibbard 2015). The Chiba composite section (CbCS) in the Kokumoto Formation of central Japanis a thick and continuous marine succession that straddles
the Lower–Middle Pleistocene boundary and the wellrecognized Matuyama–Brunhes paleomagnetic polarity
boundary (M–B boundary) (Niitsuma 1976; Okada andNiitsuma 1989; Suganuma et al. 2015; Nishida et al. 2016;
Okada et al. 2017; Suganuma et al. 2018; Haneda et al.
2020a). Because the M–B boundary serves as the primary
guide for the Lower–Middle Pleistocene stage boundary(Head and Gibbard 2005), the Chiba section, which is oneof the main sections of the CbCS (see Section 2 for a detailed description) that also contains many key tephra
layers and thus chronostratigraphic tie-points, has recently
been determined as the Global Boundary Stratotype Section and Point (GSSP) for the base of the Middle Pleistocene Subseries (Suganuma et al. 2021). Furthermore, theM–B boundary is underlain by a distinctive and widespread tephra deposit, Byk-E (also known as the OntakeByakubi Tephra: Takeshita et al. 2016), which allows theboundary to be identified also in successions in other partsof Japan.
Because the Kokumoto Formation is well-exposed
along the Yoro River of the Boso Peninsula, central
Japan, there have been numerous studies on the formation based on litho-, bio-, tephro-, magneto-, and oxygen-isotope stratigraphy (as reviewed by Kazaoka et al.
2015; Suganuma et al. 2018). In addition, the sedimentological features of the Kokumoto Formation have also
been extensively investigated (Katsura 1984; Ito 1992; Itoand Katsura 1992; Nakamura et al. 2007; Nishida et al.
2016; Takao et al. 2020). However, the detailed chemostratigraphy (i.e., organic carbon isotope ratio, the contents of organic elements such as carbon and nitrogen,and major inorganic elements) of the CbCS, particularly

near the Lower–Middle Pleistocene boundary, remains
poorly understood.
Therefore, the aims of this study are to (1) establish
the chemostratigraphy of the CbCS in detail and (2) reconstruct the paleoenvironmental changes recorded in
the CbCS. Because the Boso Peninsula faces the Pacific
Ocean where the subtropical Kuroshio and subarctic
Oyashio currents meet at the southern margin of the
Subarctic Front (Fig. 1), these data are especially important and unique for a better understanding of the fundamental changes in the ocean and climate systems of thenorthwestern Pacific Ocean across the Lower–Middle
Pleistocene boundary.

2 Geological setting and chronological framework
The Kazusa Group outcrops widely in the middle regionof the Boso Peninsula, central Japan. It unconformably
overlies the Miocene and Pliocene Miura Group and isoverlain by the Middle and Upper Pleistocene ShimosaGroup (Fig. 2). The Kazusa Group, which comprises thefills of the Kazusa fore-arc basin, was deposited in thebasin plain, submarine fan, slope, shelf, and coastal environments (Katsura 1984; Ito and Katsura 1992). The
thickest succession (up to 3000 m) crops out along the
Yoro River, where the outcrops of the Kazusa Group canbe observed continuously. Therefore, numerous studies
based on lithostratigraphy, biostratigraphy, tephrostratigraphy, paleomagnetism, and oxygen-isotope stratigraphy have focused on the type section of the Kazusa
Group (as reviewed by Kazaoka et al. 2015). The Kazusa
Group includes the following 14 main lithostratigraphic units, in ascending order: the Kurotaki,
Katsuura, Namihana, Ohara, Kiwada, Otadai, Umegase,
Kokumoto, Kakinokidai, Ichijiku, Chonan, Mandano,
Kasamori, and Kongochi formations (Tokuhashi and Endo
1984; Kazaoka et al. 2015; Fig. 2).
The Kokumoto Formation, which is the middle part of
the Kazusa Group, is generally 350–400 m thick and ismainly composed of alternating sandstone, siltstone, andsome distinct tephra layers (Kazaoka et al. 2015; Nishida
et al. 2016; Fig. 2). One of the most distinctive and important key tephra layers is the Byk-E tephra (seeIzumi et al. Progress in Earth and Planetary Science(2021) 8:10

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Fig. 1 Oceanographic circulation in summer around Japan. The CbCS (white circle) is located in the middle of the Boso Peninsula, central Japan.
Positions of major ocean currents are based on Brown et al. (2001). Ocean temperature gradients are from World Ocean Atlas 2013 (Locarnini
et al. 2013), which was drawn using Ocean Data View software (Schlitzer 2015). This figure was redrawn after Suganuma et al. (2018)

Kazaoka et al. 2015). The Kokumoto Formation represents an expanded and well-exposed sedimentary succession across the Lower–Middle Pleistocene boundary,especially at the CbCS. The CbCS comprises the Yoro
River section, the Yoro-Tabuchi, Yanagawa, Urajiro, andKokusabata sections (Suganuma et al. 2018; GSSP proposal group 2019; Fig. 2; see also Suganuma et al. 2021).
The Chiba section is an extension of the Yoro River section (Suganuma et al. 2021, for details). These sections
are adjacent and contiguous throughout the Lower–
Middle Pleistocene boundary interval (Fig. 2). The lithology of the CbCS is dominated by bioturbated siltstones, as described by previous studies (e.g., Suganuma
et al. 2018). Although there are minor sandy beds, particularly in the lower part of the succession, the CbCS isgenerally interpreted to be a continuous depositional
unit (Nishida et al. 2016). For more detailed descriptions
on the lithology of the Kokumoto Formation, seeKazaoka et al. (2015).
The Yoro-Tabuchi and Yanagawa sections of the
CbCS show no evidence of slump scars or unconformities and are composed mainly of hemipelagites(Nishida et al. 2016). In addition, Nishida et al. (2016)
reported ichnogenera, mainly from these sections, which
are typical of deep-sea siliciclastic systems (Hubbard
et al. 2012; Uchman and Wetzel 2012; Wetzel and Uchman 2012). Therefore, it is generally assumed that the
CbCS was deposited on the continental slope under generally stable conditions (Nishida et al. 2016), which were
free of the wave/current influences that are common inshallow shelf environments (Nishida et al. 2016). The
trace fossil assemblage of the CbCS is the characteristicof the Zoophycos ichnofacies (Nishida et al. 2016). Basedon a compilation of the marine sediment cores from theworld’s oceans, it has been suggested that Zoophycos occurs only at water depths of more than 800–1000 m, although there are significant variations in its bathymetric
distribution (Löwemark and Werner 2001). Importantly,
burrows assignable to Zoophycos have never been foundin shallow water settings (Löwemark and Werner 2001),
further supporting the conclusion that the CbCS was deposited on the continental slope.
Highly detailed bio-, tephro-, magneto-, and oxygenisotope stratigraphic investigations have been performedon the Kokumoto Formation so far, allowing the integrated stratigraphy of the CbCS to be established (e.g.,
Suganuma et al. 2015, 2018, 2021; Okada et al. 2017; Simon et al. 2019; Haneda et al. 2020a, b; Kameo et al.
2020). An age model of the CbCS proposed in a previous
study based on the detailed oxygen isotope (δ18O) stratigraphy yielded a depositional age spanning the lateMarine Isotope Stage (MIS) 20 through the early MIS 18(see Fig. 6 in Suganuma et al. 2018). This model alsosuggests that the minimum sedimentation rate, which
occurred during an interval of maximum flooding, was
44 cm/kyr. The M–B boundary is located ~ 110 cm
above the Byk-E tephra layer in the Chiba section(Okada et al. 2017) and lies in an interval with a sedimentation rate of 89 cm/kyr (Suganuma et al. 2018).
MIS substages (MIS 19c, 19b, and 19a) have also been
assigned based on the climatic variables proposed byNomade et al. (2019); MIS 19c corresponds to the climatically stable part of Stage 19, MIS 19b to the first climatic cooling stage, and MIS 19a to the most unstable

Izumi et al. Progress in Earth and Planetary Science(2021) 8:10

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Fig. 2 a A simplified geological map of the entire Kanto area, which is redrawn after Kazaoka et al. (2015) and Suganuma et al. (2018). Box
indicates location of a detailed geological map in b. b A detailed geological map of the Boso Peninsula, highlighting stratigraphy of the Kazusa
Group and location of CbCS. Panel b is redrawn after Suganuma et al. (2018) and GSSP proposal group (2019)

Izumi et al. Progress in Earth and Planetary Science(2021) 8:10

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Fig. 3 Lithological logs of the Yoro River and Yoro-Tabuchi sections. The horizons of the samples for geochemical analyses (CN analysis, and XRF
analysis) are highlighted. Logs are modified from Suganuma et al. (2018). The Lower-Middle Pleistocene boundary corresponds with the base of
the Byk-E tephra (at 0 m in the lithological logs)

Izumi et al. Progress in Earth and Planetary Science(2021) 8:10

part of Stage 19 (Haneda et al. 2020b). The astronomicalages for the stage and substage boundaries have been estimated to be 787.5 ka (MIS 20/19c), 773.9 ka (MIS
19c–19b), 770.1 ka (MIS 19b–19a), and 756.9 ka (MIS
19a/18), based on the age model for the CbCS by Suganuma et al. (2018) (Haneda et al. 2020b).

3 Methods
In this study, high-resolution multiproxy analyses were
performed on siltstone samples collected from the Yoro
River and Yoro-Tabuchi sections. Samples used for Xray fluorescence analysis, dry bulk density analysis, andtrace fossil analysis are from the 1 inch-diameter cores(paleomagnetic cores), which were previously taken fromthe outcrops for paleomagnetic measurements (Okada
et al. 2017; Simon et al. 2019). The lithological logs of
the Yoro River and Yoro-Tabuchi sections are shown in
Fig. 3 (see also Fig. 5 in the GSSP proposal group(2019)). The stratigraphic level for each sample (for all
analyses) was taken to be the stratigraphic distance fromthe Byk-E tephra at the Yoro-River section (Suganuma
et al. 2018, 2021; GSSP proposal group 2019). The
Lower–Middle Pleistocene boundary corresponds to thebase of the Byk-E tephra (0 m in the lithological logs in
Fig. 3).
Samples analyzed by this study cover a ~ 55-m-thickinterval that ranges from 794 ka (late MIS 20) to 756 ka(early MIS 18).
3.1 X-ray fluorescence analysis of major elements

A total of 90 paleomagnetic cores with ca. 0.2 to 1.6(average = 0.43) kyr resolution were subjected to X-ray
fluorescence (XRF) core-scanning analysis. The measurement samples were collected from the same orclosely similar horizons as those for the samples used forthe organic geochemical analysis (see detailed methods
in the next section). The XRF core scanning analysis was
performed using the nondestructive XRF core-imaging
scanner TATSCAN-F2 installed at the Center for Advanced Marine Core Research, Kochi University, Japan(Sakamoto et al. 2006). The TATSCAN-F2 is designed
to allow for rapid elemental scanning through energydispersive-type XRF (EDXRF) spectroscopy (Sakamoto
et al. 2006). Although it can measure the elemental
fluorescence intensity of the surface of the sediment
samples (Sakamoto et al. 2006), the analytical data aresemi-quantitative. The TATSCAN-F2 employs a new Xray detector with a rhodium (Rh) target that can generate X-rays with a five-times higher intensity comparedwith standard EDXRF instruments (Sakamoto et al.
2006).
During the analysis using the TATSCAN-F2, the Xrays were projected with an analytical spot diameter of 7mm using an X-ray tube voltage of 30 kV and an

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appropriate X-ray tube current of 0.01-1.0 mA; this was
performed to ensure that the X-ray intensity was sufficient. The obtained dataset of the concentrations of the
major elements is tentatively represented as the weightpercentages of their oxides (e.g., Na2O, MgO, Al2O3,
SiO2, P2O5, S, K2O, CaO, TiO2, MnO, and Fe2O3), except for sulfur. Reproducibility of the measurements,
which was calculated as the standard deviation of themeasured values of repeated measurements divided by
the average of the repeatedly measured values, was ±0.48% (Na2O), 0.24% (MgO), 0.14% (Al2O3), 0.23%(SiO2), 0.05% (P2O5), 0.01% (S), 0.04% (K2O), 0.29%(CaO), 0.02% (TiO2), 0.01% (MnO), and 0.22% (Fe2O3)(Sakamoto et al. 2006).
In addition, to further evaluate the abundance of eachelement obtained by the TATSCAN-F2 analysis (i.e.,
semi-quantitative analysis), a conventional (i.e., quantitative) XRF analysis was conducted for ten samples using
the Magix PRO+PW2440/00 (Malvern Panalytical) installed at the Center for Advanced Marine Core Research. During measurements, an X-ray tube voltage was
32-60 kV and the appropriate X-ray tube current was
66-125 mA. For conventional XRF analysis, dry sediment
samples were calcined at 950 °C for 6 h before making
glass beads. Subsequently, glass bead samples were prepared by fusing 0.5 g of the calcined sediment samples
with 5.0 g lithium tetraborate in a Pt crucible at 1250 °C.
The weight of the sediment sample and lithium tetraborate was measured using an electronic balance (measurement error = 0.1 mg). Each sample was analyzed threetimes, and the average value was used as the abundanceof each element.
Finally, to compare the results obtained from the
TATSCAN-F2 and those obtained from Magix PRO appropriately, we performed a correlation analysis.
3.2 Carbon and nitrogen analyses

A total of 89 samples with ca. 0.2 to 1.7 (average = 0.43)
kyr resolution were collected from the Yoro River andYoro-Tabuchi sections, and their total organic carbon(TOC), total nitrogen (TN), and stable organic carbon
isotope (δ13Corg) values were analyzed. In addition, theC/N ratio (= TOC/TN) of each sample was calculatedusing the obtained data. The samples were treated with
1 N HCl several times at room temperature to remove
the carbonate fraction. Afterward, the samples were
washed with distilled water until neutralized and thendried at 110 °C. The TOC, TN, and δ13Corg values were
measured using an elemental analyzer (Flash2000, ThermoFisher Scientific Ltd.) and a stable isotope mass spectrometer (MAT253, ThermoFisher Scientific Ltd.) at the
Department of Geology and Paleontology, National Museum of Nature and Science, Japan. A few milligrams ofeach sample were sealed in Sn foil for the measurements

Izumi et al. Progress in Earth and Planetary Science(2021) 8:10

of TOC, TN, and δ13Corg. During the measurements, thetest sample was heated to 1000 °C in the furnace of theelemental analyzer, and the resulting purified N2 andCO2 gas was fed directly into the mass spectrometerusing pure helium as the carrier gas. We performed themeasurements on a working standard (L-Alanine;δ13Corg = − 18.50‰) with every sixth sample to calibrate
the measured isotopic values with respect to the Vienna
Pee Dee Belemnite (VPDB).
3.3 Measurement of dry bulk density

To calculate the mass accumulation rate of the specificmaterial using the results of our XRF and organic geochemical (C, N) analyses, the dry bulk density of the CbCS
sediments was calculated. Eighty-three paleomagnetic
cores with ca. 0.2 to 2.8 (average = 0.46) kyr resolution
were measured for size (diameter, length) and weight. The
size and weight were measured by using an electronic caliper and an electronic balance, respectively. The volume ofeach paleomagnetic core was calculated using its diameterand length, and subsequently, the dry bulk density of each
paleomagnetic core was calculated using the volume andweight. The measurement errors of the electronic caliper
and electronic balance are 0.1 mm and 0.01 g, respectively.
However, due to some knotty-shaped paleomagnetic
cores, the actual measurement error of the size (diameter,
length) of paleomagnetic cores may be ~ 1 mm. Therefore,
the relative error of the dry bulk density may be quite
large (probably up to ~ 15%), especially in the case ofsome knotty-shaped paleomagnetic cores.
3.4 Analysis of trace fossils

Eighty-three paleomagnetic cores with ca. 0.2 to 2.8(average = 0.46) kyr resolution were subjected to the observation and analysis of trace fossils. Thesepaleomagnetic cores were cut perpendicular to the

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bedding planes. Trace fossils of the polished surface ofthe paleomagnetic cores were observed. Afterward, theburrow size of the most abundant trace fossil (the ichnogenus Phycosiphon in the case of this study; Fig. 4) was
measured using an image-analysis software ImageJ.
When analyzing the temporal change in the size of Phycosiphon, the burrow diameter on the cross section wasused in this study because it is commonly interpreted
that the burrow diameter reflects the body diameter of
the trace-maker. In this study, the burrow diameter of
the analyzed trace fossil (i.e., Phycosiphon) is defined as
the short axis of the cross section of the burrow, assuming that the burrow shape is cylindrical (e.g., Bednarz
and McIlroy 2009) and the effect of compaction of thesediments is small (e.g., Izumi et al. 2017). In addition,
the maximum burrow diameter represents a biologicalparameter of the trace-maker. Namely, the maximum
burrow diameter preserves a record of the largest infaunal organism present at a given time, which is linked
with relative oxygen levels correlated with body size
based on physiological oxygen demands (Boyer and Droser 2011 and references therein).
In the case of Phycosiphon, this ichnogenus is a burrow tube filled with dark-colored, fine-grained material,
which is generally defined as a “core,” and the core issurrounded by pale-colored, coarser-grained material,
which is generally defined as a “mantle” (e.g., Kern 1978;
Goldring et al. 1991; Wetzel and Bromley 1994; Izumi
2014; see also Fig. 4). As for the paleoecological information of Phycosiphon, this ichnogenus has been regarded
as a trace produced by a subsurface deposit feeder (e.g.,
Kern 1978; Wetzel and Bromley 1994), and the infill of
the core has been interpreted as fecal material (e.g.,
Ekdale and Lewis 1991; Bednarz and McIlroy 2009).
Based on the formation model of Phycosiphon (e.g., Wetzel and Bromley 1994; Bromley 1996; Seilacher 2007),Fig. 4 Scan images of the selected samples of the polished paleomagnetic core. a Sample ID: TB3-37. b Sample ID: TB3-53. The paleomagnetic
cores were cut perpendicular to bedding planes, thus, the upward direction of these images represent stratigraphically higher direction. Noteabundant Phycosiphon in the surface. In addition, the core and mantle of Phycosiphon can clearly be observed

Izumi et al. Progress in Earth and Planetary Science(2021) 8:10

the diameter of the core directly reflects the body diameter of the trace-maker.

4 Results
In this study, the main results are shown using graphs toillustrate the temporal patterns. The x-axis of each graphis based on the established age model (Suganuma et al.
2018). For the stratigraphic distance of each sample, see
Supplementary Tables in Additional files.
4.1 Concentrations of major inorganic elements

Data obtained by TATSCAN-F2 and Magix PRO analyses are summarized in Table S1 (see Additional files).
The absolute values of each element concentration are
different in TATSCAN-F2 and Magix PRO partly due tothe different sample conditions and analytical settings.
However, the values of CaO obtained by TATSCAN-F2
and Magix PRO show strong positive correlation (r =0.87; see Fig. 5), while the values of low-concentration

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elements such as K2O and TiO2 obtained by these twoanalytical instruments do not show strong correlation(Fig. 5; see also Table S1 in Additional files).
The temporal trends of sedimentologically and paleoceanographically important proxies (K/Ti ratio, and Ca/
Ti ratio, respectively) are shown in Figs. 6 and 7, respectively; these are discussed in detail later in Section 5.
The results of the correlation analysis for these proxies
are also presented in Fig. 5. The K/Ti and Ca/Ti values
obtained by TATSCAN-F2 are highly consistent withthose by Magix PRO, which is supported by the highcorrelation coefficient (r = 0.91 and 0.97, respectively;
see Fig. 5, and Table S1 in Additional files). Thus, thetemporal trends of K/Ti and Ca/Ti ratios obtained by
TATSCAN-F2 are discussed in Section 5.
4.2 TOC, TN, C/N ratio, and δ13Corg records

The TOC and TN records show distinct temporal variations throughout the analyzed interval. The TOC value

Fig. 5 The relationships between the major element abundances between TATSCAN-F2 and Magix PRO analyses for important proxies discussed
in this paperIzumi et al. Progress in Earth and Planetary Science(2021) 8:10

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Fig. 6 Temporal changes in sedimentological parameters from MIS 20 to MIS 18 at the CbCS. a K/Ti record obtained by TATSCAN-F2, which isinterpreted as a grain-size proxy (see details in the text). Note the smaller K/Ti value indicates coarser sediment. b High-resolution δ18O records
from planktonic (Globigerina bulloides) and c benthic foraminifers at the CbCS (Haneda et al. 2020b). Thick lines are 1000 yr moving average
profiles. d The LR04 benthic stack (Lisiecki and Raymo 2005). e Calculated sedimentation rate variation at the CbCS (after Suganuma et al. 2018)with (light gray bars) and without (dark gray bars) sandstone beds. Detailed information on the age model is given in Suganuma et al. (2018).
Dashed lines are Marine Isotope Stage and substage boundaries, based on Haneda et al. (2020b)

remained relatively constant at ~ 0.4–0.5 wt% from 795
to 780 ka but generally increased to ~ 0.65 wt% from
780 to 770 ka with some variation (Fig. 7). Subsequently,
it decreased gradually from 770 to 755 ka, with the exception at ~ 764 and ~ 760 ka. The TN content from
795 to 790 ka was ~ 0.04–0.05 wt% with minor variation,and gradually increased to ~ 0.08 wt% from 790 to 770
ka. From 770 to 755 ka, it was generally 0.04-0.06 wt%
with the exception at ~ 764 and ~ 760 ka. Relatively
large variations in the form of at least two maxima, at ~764 and ~ 760 ka, were observed in the cases of TOC
and TN (Fig. 7). Overall, the TOC and TN contents generally increased during MIS 19; however, the exact timings of these increases are not synchronous, resulting ina pronounced temporal variation in the C/N ratio.
Throughout the analyzed interval, the C/N ratio record changes in association with glacial–interglacial cycles from the late MIS 20 to the early MIS 18. In otherwords, the value of the C/N ratio from 795 to 790 ka(MIS 20) and from 770 to 755 ka (late MIS 19 to early
MIS 18), corresponding to glacial periods, was generally
higher (~ 7 to ~ 11) than that from 790 to 770 ka (lateMIS 20 to early MIS 19 corresponding to deglacial tointerglacial periods when the C/N ratio = ~ 5.5 to ~ 9.5)
(Fig. 7). However, the ratio showed relatively large millennial to multimillennial-scale variations especially during 795 to 790 ka and 765 to 755 ka within early glacial
periods (Fig. 7).
The δ13Corg record generally showed a similar trend asthat for the C/N ratio; this was also indicative of the glacial–interglacial cycles from the late MIS 20 to the early
MIS 18, having relatively large variations especially during 795 to 790 ka and 765 to 755 ka. The δ13Corg values
from 795 to 790 ka (MIS 20) and from 770 to 755 ka(late MIS 19 to early MIS 18) corresponding to glacial
periods generally ranged from ~ − 23.0 to − 24.5‰,
whereas they are ~ − 21.5 to − 23.0‰ from 790 to 770
ka (late MIS 20 to early MIS 19) corresponding toIzumi et al. Progress in Earth and Planetary ScienceFig. 7 (See legend on next page.)(2021) 8:10

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Izumi et al. Progress in Earth and Planetary Science(2021) 8:10

Page 11 of 19(See figure on previous page.)
Fig. 7 Temporal change in paleoceanographic parameters from MIS 20 to MIS 18 from CbCS. a Total organic carbon (TOC) record. b Total
nitrogen (TN) record. c Carbon isotope ratio of organic carbon (δ13Corg) record. d C/N record. e Ca/Ti record obtained by TATSCAN-F2. f
Maximum diameter of the core of Phycosiphon. g The difference in water temperature between surface and bottom water (ΔT) from Haneda
et al. (2020b). Thick lines are 1000 yr moving average profiles. h Radiolarian concentrations (Suganuma et al. 2018). i S-ratio-0.3T record (Haneda
et al. 2020a). j Magnetic susceptibility record (Haneda et al. 2020a). k High-resolution δ18O records from planktonic (G. bulloides) and l benthic
foraminifers from CbCS (Haneda et al. 2020b). Thick lines are 1000 yr moving average profiles. Dashed lines are marine isotope stage andsubstage boundaries, based on Haneda et al. (2020b). Yellow shades at ~ 764 and 760 ka highlight Millennial Isotopic Oscillation (MIO)-Interstadial
2 and MIO-Interstadial 3, respectively (Haneda et al. 2020b), which correspond to relatively large variations in TOC and TN values. Although Sratio-0.3T, magnetic susceptibility, and δ18O records of the CbCS has been the subject of a series of investigations over the past 5 years (Suganuma
et al. 2015, 2018; Okada et al. 2017; Simon et al. 2019; Haneda et al. 2020a, b) during which successive refinements have been incorporated, thispaper refers to the latest and refined data set of Haneda et al. (2020a, b)

deglacial to interglacial periods (Fig. 7). Similar to thetemporal change in the C/N ratio, there may be millennial to multimillennial-scale changes in the δ13Corg
values, although the amplitudes of such shorter-scalechanges are low (Fig. 7).
The obtained results are also listed in Table S2 (seeAdditional files).

from the late MIS 20 to the early MIS 18, unlike thetemporal trends of the C/N ratio and δ13Corg records
that are described in the previous section (Fig. 7). Rather, the diameter of the core of Phycosiphon varied
largely without any distinct long-term trend throughout
the analyzed interval (Fig. 7).

5 Discussion
4.3 Dry bulk density

5.1 Variations in grain size as inferred from the K/Ti ratio

Dry bulk density of the CbCS sediments ranges from1.34 to 1.73 g/cm3 (Table S3 in Additional files). Amongthe analyzed samples (83 samples), approximately 90%of the samples (75 samples) show a dry bulk density
value ranging from 1.45 to 1.60 g/cm3.
As for the temporal trend, the values of the dry bulk
density during MIS 19c are generally lower than those inother periods (Table S3 in Additional files). This feature
may partly be due to the difference in lithology. In particular, the lithology especially during MIS 19c issiltstone-dominated and is also characterized with theless presence of sand layers (Fig. 3).

For the interpretation of the high-resolution temporalchange in grain size, it is useful to evaluate a proxy forthe grain size as well as physical grain-size analysis. In
particular, the K/Ti ratio is used as the grain-size proxyin this study. On the one hand, relative changes in Ti
abundance have been used previously as grain-size proxies owing to the expected enrichment of this element insilty to sandy sediment relative to its abundance in clay
minerals (despite the presence of clay-bound Ti, for example) (Calvert and Pedersen 2007). On the other hand,
K is most strongly associated with aluminosilicates suchas clay minerals (Calvert and Pedersen 2007). Therefore,
the elemental ratio of coarse grain-enriched element (Ti)relative to clay-enriched element (K) can be used as agrain-size proxy. In fact, the K/Ti ratio has been used asa proxy for the grain size of clastic material in various
studies (e.g., Kemp and Izumi 2014; Xie et al. 2014; Yanget al. 2015).
As for the temporal trend, the K/Ti ratio from 794 to
792 ka (MIS 20) shows an increasing trend, which issuggestive of a fining trend (Fig. 6). Subsequently, the ratios from 792 to 786 ka (late MIS 20 to early MIS 19c)
are essentially constant, although there are slight variations (Fig. 6). The K/Ti ratio from 786 to 784 ka (MIS19c) shows a decreasing trend (= coarsening trend), andthen the ratio keeps constant from 784 to 772 ka (MIS
19c to early MIS 19b), although there are two distinct
peaks at ~ 774 and 772 ka (Fig. 6). The K/Ti ratio after
772 ka (late MIS 19b to MIS 19a) shows large change
patterns (i.e., multimillennial-scale changes), which areconcordant with the temporal trends in the δ18O recordsof the benthic, subsurface, and surface planktonic foraminifers (Haneda et al. 2020b) (Fig. 6). As is discussed in

4.4 Trace fossil records

Several ichnogenera were recognized in the surface ofthe paleomagnetic cores, including Chondrites, Phycosiphon, and the burrows of indeterminate ichnogenus.
Among them, Phycosiphon is the most abundant ichnogenus (Fig. 4), and thus the maximum diameter of this
ichnogenus was measured (see Section 3.4). As mentioned earlier, it is commonly interpreted that the burrow diameter reflects the body diameter of the tracemaker; thus, the maximum burrow diameter represents
biological parameter of the trace-maker, which is linked
with relative oxygen levels (Boyer and Droser 2011 andreferences therein).
The obtained dataset is summarized in Table S4 (seeAdditional files). As a result of our observation, the maximum diameter of the core of Phycosiphon observed in
the surface of each paleomagnetic core ranges from0.254 to 1.905 mm (Table S4). As for the temporal
trend, the maximum diameter of Phycosiphon does notchange in association with glacial–interglacial cycles

Izumi et al. Progress in Earth and Planetary Science(2021) 8:10

later section (see Section 5.4), the peaks in δ18O records
during MIS 19a reflect the northward displacement ofKuroshio Extension Front (Haneda et al. 2020b). The
northward displacements of Kuroshio Extension Front
during MIS 19a may have been coincident with the decreases in the K/Ti ratio (i.e., coarsening trends). In particular, based on the sedimentological study using the
Holocene marine core samples collected from the continental slope setting (water depths: 449-1728 m), off the
Boso Peninsula, the intensification of the Kuroshio influence may result in the coarsening of the grain size because finer grains are easily transported when the
Kuroshio Current is intensified (Ajioka et al. 2019).
However, to further evaluate the mechanism of the likely
multimillennial-scale changes in the grain size of clastic
material, high-resolution grain-size analysis with the pretreatments of the removal of organic matter, opal, andbiogenic carbonate is required in the future.
5.2 Source and mass accumulation rate of organic carbon

In general, the C/N ratio of marine sediments varies depending on the source of the organic matter. The C/N ratio
is high (typically > 10) in the case of terrestrial plants andlow (< 10) for marine plankton/algae (Meyers 1994). Thus,
the C/N ratio is used as a proxy for organic matter source.
However, the TN values occasionally include inorganic nitrogen absorbed in clay minerals. Therefore, it is necessary
to cross-check with the other proxies to more precisely
evaluate the source of the organic matter in marine sediments. The δ13Corg value of marine sediments can also be
used as a proxy for organic matter source. The δ13Corg

Page 12 of 19

value of terrestrial C3 plants is generally − 27‰, whereasthat of marine plankton ranges from − 22 to − 18‰ (Fryand Sherr 1984). The scatter plot of the δ13Corg versus theC/N values shows a moderate negative correlation (r = −0.67; Fig. 8). However, considering that the correlation coefficient is − 0.67, it is suggested that, in the case of this study,
the influence of the inorganic nitrogen is not negligible.
Therefore, to interpret the source of organic matter (especially organic carbon), it is better to use the δ13Corg values.
Based on the results of the δ13Corg measurements, it isinterpreted that the organic carbon of the CbCS sediments was mostly composed of marine plankton during
MIS 19c and 19b, whereas that during MIS 20, 19a and18 was characterized by a mixture of both marine plankton and terrigenous organic carbon (Fig. 7). To further
interpret this quantitatively, we estimated the contributions of marine plankton organic carbon and terrestrialorganic carbon by using a two-end-member isotope mixing model (cf. Belicka and Harvey 2009), assuming thatthe end member values of marine and terrestrial organic
matter are δ13Corg = − 21 and − 27‰, respectively (seeTable S2 in Additional files). As a result, the contributions of marine organic carbon during MIS 19c and 19b
are approximately estimated as from 70 to 90%, whereas
those during MIS 20, 19a and 18 were generally lower
than 60% (see Table S2 in Additional Files).
As for the temporal changes in the source and amountof organic carbon, it is interpreted that the input of marine and terrestrial organic carbon and their balance control the contribution of marine organic carbon versus
terrestrial organic carbon within the sediments. To

Fig. 8 Scatter plot of the δ13Corg versus the C/N value. Note the moderate negative correlation (r = − 0.67)

Izumi et al. Progress in Earth and Planetary ScienceFig. 9 (See legend on next page.)(2021) 8:10

Page 13 of 19

Izumi et al. Progress in Earth and Planetary Science(2021) 8:10

Page 14 of 19(See figure on previous page.)
Fig. 9 Mass accumulation rates (MARs) of some selected materials. a Calculated MAR of marine organic carbon. Gray plots are calculation for
individual samples, and black bold lines indicate the average values of the interval between two age controlling points. b Calculated MAR of
terrestrial organic carbon. Gray plots are calculation for individual samples, and black bold lines indicate the average values of the interval
between two age controlling points. c Calculated MAR of biogenic carbonate. Gray plots are calculation for individual samples, and black boldlines indicate the average values of the interval between two age controlling points. d Calculated MAR of terrigenous material. Gray plots arecalculation for individual samples, and black bold lines indicate the average values of the interval between two age controlling points. e Highresolution δ18O records from planktonic (G. bulloides) and f benthic foraminifers from CbCS (Haneda et al. 2020b). Thick lines are 1000 yr moving
average profiles. Dashed lines are Marine Isotope Stage and substage boundaries, based on Haneda et al. (2020b). Yellow shades at ~ 764 and760 ka highlight Millennial Isotopic Oscillation (MIO)-Interstadial 2 and MIO-Interstadial 3, respectively (Haneda et al. 2020b), which correspondwith relatively large variations in TOC and TN values. Black arrows indicate the age controlling points (cf. Suganuma
et al. 2018, Haneda et al. 2020b)

further discuss this issue quantitatively, we calculated
the mass accumulation rate (MAR) of both marine andterrestrial organic carbon by using the data on organiccarbon, sedimentation rate, and dry bulk density. For thedetails on the MAR calculation, see Table S2 (see Additional files). The MARs of marine and terrestrial organiccarbon and their temporal changes are illustrated in Fig.
9. Because of the large variations in the sedimentation
rate (Fig. 6), the MARs of marine and terrestrial organiccarbon throughout the studied interval varied considerably (Fig. 9). In particular, during MIS 19c, the MARs ofboth marine and terrestrial organic carbon were extremely low (Fig. 9), due to the low sedimentation rate.
The MAR of marine organic carbon during late MIS 19a
ranges from ~ 10 to 20 g/m2 year (average = 14.4 g/m2
year; see Table S2), whereas that from late MIS 20 toMIS 19c ranges from ~ 2 to 3 g/m2 year (average = 2.6
g/m2 year; see Table S2) (Fig. 9). The MAR of terrestrialorganic carbon during late MIS 19a ranges from ~ 9 to
13 g/m2 year (average = 10.7 g/m2 year; see Table S2),
whereas that from late MIS 20 to MIS 19c ranges from~ 0.2 to 1 g/m2 year (average = 0.5 g/m2 year; see Table
S2) (Fig. 9). These results mean that the MAR of terrestrial organic carbon during late MIS 19a was ~ 20 times
higher relative to that from late MIS 20 to MIS 19c,
while the MAR of marine organic carbon during late
MIS 19a was only ~ 5.5 times higher relative to thatfrom late MIS 20 to MIS 19c (Fig. 9). Taken together,
the results of the MAR calculation are indicative of thecharacteristics of the temporal change in the organiccarbon source; specifically, the contribution of marineorganic carbon relative to terrestrial organic carbon significantly increased during MIS 19c (Fig. 9).
5.3 Biogenic carbonate and terrigenous material

The Ca/Ti ratio, which reflects biogenic carbonate content, tends to be higher especially during MIS 19c (Fig.
7). It is occasionally interpreted that the Ca/Ti ratio generally reflects biogenic productivity (e.g., Riethdorf et al.
2013; Hyodo et al. 2017). The radiolarian (i.e., secondary
producer) concentration was also higher especially during MIS 19c (Suganuma et al. 2018; Fig. 7). Furthermore,

based on the observation of the TB2 core that was collected at a site ca. 200 m east of the Chiba section, thetemporal trend in the diatom (i.e., primary producer)
concentration is highly similar with that in the radiolarian, showing that the diatom concentration was higher
especially during MIS 19c (Tanaka et al. 2017).
However, it is not reasonable to consider that the temporal variations in Ca/Ti ratio and siliceous microfossil
abundances (radiolarian and diatom concentrations) of
the sediments directly reflect the past biogenic productivity in the case of this study, because the sedimentation
rate was lower especially during the earlier part of MIS
19 (Fig. 6). Therefore, it is also possible that the temporal changes in the Ca/Ti ratio and radiolarian concentrations were controlled by of the changes in the
terrigenous input rate.
To further discuss this issue quantitatively, the MARsof both biogenic carbonate and terrigenous material
were calculated using the data from the XRF analysis,
sedimentation rate, and dry bulk density (see Table S1for detailed methods of calculation). The assumption of
MAR calculation, in particular, the equations for calculating the percentages of biogenic carbonate and terrigenous material of the sediments are based on Kido
et al. (2007). The detailed steps and assumptions on the
MAR calculation have also been described in Table S1(see Additional files).
The results are shown in Fig. 9. Because of the large
variations in sedimentation rate (Fig. 6), the MARs ofbiogenic carbonate and terrigenous material throughout
the studied interval varied considerably (Fig. 9; see Table
S1 for detailed methods of calculation). The MAR ofbiogenic carbonate during late MIS 19a ranges from ~250 to 600 g/m2 year (average = 393.0 g/m2 year; see
Table S1), whereas that from late MIS 20 to MIS 19c
ranges from ~ 30 to 95 g/m2 year (average = 68.1 g/m2
year; see Table S1) (Fig. 9). For the terrigenous material,
its MAR during late MIS 19a ranges from ~ 3500 to4500 g/m2 year (average = 3921.7 g/m2 year; see Table
S1), whereas that from late MIS 20 to MIS 19c ranges
from ~ 350 to 500 g/m2 year (average = 405.9 g/m2 year;
see Table S1) (Fig. 9). Such results mean that the MAR

Izumi et al. Progress in Earth and Planetary Science(2021) 8:10

of terrigenous material during late MIS 19a was ~ 10times higher relative to that from late MIS 20 to MIS
19c, while the MAR of biogenic carbonate during late
MIS 19a was only ~ 6 times higher relative to that fromlate MIS 20 to MIS 19c (Fig. 9). Overall, it is concluded
that the increase in the Ca/Ti ratio (= the proxy for biogenic carbonate) during MIS 19c (Fig. 7) was not causedby the increase in the MAR of biogenic carbonate that
actually decreased during the interval. Instead, the increase in the Ca/Ti ratio during MIS 19c was due to thesignificant decrease in the MAR of terrigenous material
that exceeded the decrease in the MAR of biogenic carbonate during MIS 19c.
5.4 Multimillennial-scale variability in geochemical
proxies during MIS 19a

Another important aspect of our dataset is the relatively
large variations in the TOC and TN values during MIS
19a. In particular, at least two maxima can be observedin the TOC and TN values, at ~ 764 and ~ 760 ka, respectively (Fig. 7). The variation amplitude of the TOC
values during the analyzed interval is 0.32 wt% (0.32–
0.64 wt%), whereas the amplitudes of the peaks at 764
and 760 ka are ~ 0.23 (~ 71.9% of the whole range) and~ 0.20 wt% (~ 62.5% of the whole range), respectively.
Similarly, for the TN values, the amplitudes of the peaks
at 764 and 760 ka are ~ 0.027 (~ 62.8% of the whole
range) and ~ 0.024 wt% (~ 55.8% of the whole range),
respectively.
The factors that are important to control organic matter content in marine sediments are sediment surface
area, sedimentation rate, and preservation efficiency(e.g., Burdige 2006). In general, sediment-specific surface
area controls the amount of organic matter in sediments,and finer-grained sediments (especially clay minerals)
have a larger surface area (Keil and Hedges 1993; Kennedy et al. 2002). However, the temporal trend in the K/
Ti ratio (= proxy for grain size; also see Section 5.1),
which shows relatively small values at both ~ 764 and ~760 ka (Fig. 6), suggests that the sediments at ~ 764 and~ 760 ka are relatively coarse. Regarding sedimentation
rate, higher sedimentation rate prevents the degradationof organic matter by reducing the oxygen exposure time(e.g., Hartnett et al. 1998). However, because the sedimentation rate of the CbCS has been based only on a
few age controlling points (Suganuma et al. 2018), it is
not possible to discuss the actual effect of sedimentation
rate on the changes in organic matter content of the
CbCS sediments on millennial to multimillennial timescale.
Alternatively, as mentioned earlier, because theamount of organic matter in marine sediments is also
controlled by the preservation efficiency (Burdige 2006),
it is possible that preservation efficiency of organic

Page 15 of 19

matter was higher at 764 and 760 ka. In general, the organic carbon content in marine sediments in either
high-productivity (thus high organic-carbon flux) regions or under low-oxygen bottom waters is relatively
high (Hedges 2002; Burdige 2006). In the case of the
CbCS, it is likely that both the organic-carbon flux andthe bottom-water oxygen level affected the organic matter content at 764 and 760 ka (see the later discussion
and Fig. 9).
Especially in terms of the bottom-water oxygen level,
there are several evidences suggesting that the bottomwater oxygen level lowered at ~ 764 and ~ 760 ka. First,
this interpretation is consistent with the lower S-ratio
and magnetic susceptibility in the same intervals (Okada
et al. 2017; Simon et al. 2019; Haneda et al. 2020a) (Fig.
7). Because remarkable decreases in S-ratio and magnetic susceptibility tend to be accompanied by the dissolution of magnetite grains (e.g., Suganuma et al. 2008),
the reductive condition would have occurred at 764 and760 ka. Another supporting evidence is the trace-fossil
data. During MIS 19a, the distinct decrease in the diameter of the core of Phycosiphon occurred at 760 ka (Fig.
7). Although the decrease in the maximum diameter of
Phycosiphon at around 764 ka is less clear because of the
low temporal-resolution data especially during 770 to
765 ka, 764 ka corresponds with the interval dominatedby the smaller Phycosiphon (Fig. 7). In the case of deepmarine sediments characterized with a sedimentation
rate greater than 2 cm/kyr and with the organic-carbon
content greater than 0.1 wt.%, the most significant controlling factor of the burrow size is the bottom-water
oxygen content (e.g., Wetzel and Uchman 2012). In particular, the lower the bottom-water oxygen level is, thesmaller the burrow size is (Wetzel and Uchman 2012).
Therefore, our Phycosiphon records also support the interpretation that the bottom-water oxygen level lowered
at 764 and 760 ka. With respect to the bottom-water
redox conditions of the CbCS, Hyodo et al. (2017) suggested that a persistent anoxic depositional environment
existed and intercalated with short-term oxic events.
However, the siltstones of the CbCS are highly bioturbated, and large (> 1 cm) trace fossils produced by marine
invertebrates are often observed in the CbCS sediments(Nishida et al. 2016). These sedimentary characteristics
are indicative of dysoxic (dissolved oxygen = 0.2–2.0 ml/l) to oxic (dissolved oxygen = 2.0–8.0 ml/l) conditions
instead of an anoxic (dissolved oxygen = 0 ml/l) bottomwater condition (cf. see Tyson and Pearson 1991 for theclassification of redox conditions). As for the modern
Pacific off Boso Peninsula (water depth ranging from 0
to 1500 m), dissolved oxygen values range from ~ 1.5 to
4.5 ml/l (Yang et al. 1993), suggesting that our new interpretation on the redox conditions (dysoxic to oxic) ismore plausible.

Izumi et al. Progress in Earth and Planetary Science(2021) 8:10

In addition, our interpretation that the bottom water
oxygen content lowered at ~ 764 and ~ 760 ka is also
consistent with the available paleoceanographic proxies
from the CbCS. The latest high-resolution (centurialresolution) analysis of the δ18O records of the benthic,
subsurface, and surface planktonic foraminifers from the
CbCS suggests multiple millennial isotopic oscillations(MIOs), which are composed of the MIO-Stadial 1 to
the MIO-Stadial 4 and the MIO-Interstadial 1 to the
MIO-Interstadial 4 during MIS 19 (Haneda et al. 2020b).
The increases in the organic-matter content at 764 and760 ka observed in this study (Fig. 7) correspond to the
MIO-Interstadial 2 and 3, respectively. During these MIO
Interstadials, the difference in the water temperature
between surface and bottom water (ΔT) also fluctuated(Fig. 7); in particular, a decrease in the δ18O value of thesurface water and increase in the ΔT represent an increasein the sea surface temperature (SST) and increased stratification owing to the northward displacement of theKuroshio Extension Front (for more detailed information,
see Haneda et al. 2020b).
Based on the discussion above, we propose the hypothesis that the two peaks in the TOC and TN values
at 764 and 760 ka were caused by enhanced organicmatter preservation due to relative decrease in thebottom-water oxygen level, which were partly attributable to increased ocean stratifications at 764 and 760 ka
because of the northward displacement of the Kuroshio
Extension Front. However, the above hypothesis has notbeen fully tested, because the interpretation based on
the proxies we employed are not unique and there arerooms for other interpretations. For example, increased
surface productivity or decreased terrigenous input at ~764 and 760 ka may be other causes for higher organiccarbon content. In fact, based on our data, it is not certain whether the increases in organic-carbon content at~ 764 and 760 ka were due to the synchronous increasein biogenic productivity in surface water, as mentioned
earlier. If the ocean stratification occurred at 764 and760 ka due to the northward displacement of the Kuroshio Extension Front, it may not have caused the increase in biogenic productivity because the Kuroshio
water is oligotrophic due to low biogenic productivity inthe downwelling-dominant subtropical gyre. Regardless,to verify the hypothesis, further studies should prioritize
the reconstruction of the variation in both the bottomwater oxygen level and the biogenic productivity byhigh-resolution analysis of multiple proxies. Then, analyses of various proxies for paleo-ocean redox conditions, such as pyrite framboids and redox-sensitive traceelements (e.g., Wilkin et al. 1996; Algeo and Maynard2004), and proxies for paleo-ocean productivity, such as
Barium concentrations (Dymond et al. 1992), are important for subsequent studies.

Page 16 of 19

6 Conclusions
In this study, we performed multiproxy sedimentologicaland geochemical analyses of the Chiba composite section(CbCS), including the Chiba section, to establish its highresolution chemostratigraphy and reconstruct its paleoenvironments in detail. Although the temporal trends in the
K/Ti ratio (= a proxy for grain size of clastic material) varied largely throughout the analyzed interval and thus are
not clearly indicative of glacial–interglacial changes, the K/
Ti profile during MIS 19a shows variation pattern similar
to multimillennial-scale variations in δ18O records. The C/
N ratio and carbon isotope ratio of organic carbon (δ13Corg)
records of the CbCS sediments indicate that the sedimentary organic matter was mostly composed of marine organic
matter especially during the interglacial maximum (MIS19c). However, during glacial periods (MIS 18, 19a, and20), the organic matter was characterized by a mixture ofboth marine and terrigenous organic matter. Furthermore,mass accumulation rates (MARs) of organic carbon, biogenic carbonate, and terrigenous material were calculated
to quantitatively interpret the paleoenvironmental changes.
MAR calculations supported the interpretation that thecontribution of marine organic carbon relative to terrestrialorganic carbon increased during MIS 19c, and that the contribution of the terrigenous material relative to biogenic
carbonate decreased during MIS 19c. In addition, althoughtemporal records of the total organic carbon (TOC) andtotal nitrogen (TN) were in association with glacial–interglacial cycles, relatively large variations in TOC and TN
values were observed especially during MIS 19a. It is plausible that these variations can be ascribed to the change inbottom-water oxygen levels, based on the interpretation of
our trace-fossil data, although it is not certain whether theincreases in organic-carbon flux at ~ 764 and 760 ka were
due to the synchronous increases in biogenic productivity
in surface water. The relative decrease in bottom-water
oxygen level was probably due to the increased ocean stratification because of the northward displacement of theKuroshio Extension Front.
7 Supplementary Information
The online version contains supplementary material available at https://doi.
org/10.1186/s40645-020-00393-5.
Additional file 1: Table S1. Full dataset for XRF analysis by TATSCANF2 and Magix PRO, with detailed descriptions of the MAR calculation.
Additional file 2: Table S2. Full dataset for organic carbon andnitrogen analysis, with detailed descriptions of the MAR calculation.
Additional file 3: Table S3. Full dataset for the analysis of dry bulk
density.
Additional file 4: Table S4. Full dataset for trace-fossil analysis.
Abbreviations
CbCS: Chiba composite section; GSSP: Global Boundary Stratotype Section
and Point; M–B boundary: Matuyama–Brunhes paleomagnetic polarity
boundary; MAR: Mass accumulation rate; MIO: Millennial Isotopic Oscillation;

(2021) 8:10

Page 17 of 19

MIS: Marine Isotope Stage; TN: Total nitrogen; TOC: Total organic carbon;
δ13Corg: Carbon isotope ratio of organic carbon

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Acknowledgements
We thank all those who involved in this research project (for the full list ofmembers, see the GSSP proposal group 2019). We also greatly appreciate
the support provided by Ichihara City, Chiba Prefecture, and the Ministry of
Education. We thank Editage (www.editage.com) for English language
editing. We conducted a part of this study through a joint-use system at theCenter for Advanced Marine Core Research, Kochi University (Grants 16A023/
16B021). Our manuscript was greatly improved by the constructive comments from the editor and reviewers.
Authors’ contributions
KI performed the TOC, TN, and δ13Corg measurements, observed andanalyzed trace fossils, analyzed the XRF data, and wrote the main body ofthis manuscript. YH collected the sediment samples and helped with thedata interpretation. YS collected the sediment samples and helped with the
TOC, TN, and δ13Corg analyses as well as the data interpretation. MO
collected the sediment samples, performed the TATSCAN analysis along with
TM, and helped with the data interpretation. YK helped with the TOC, TN,and δ13Corg measurements as well as with the data interpretation. NN
greatly helped with the data interpretation. MK helped with the TOC, TN,and δ13Corg analyses. TM performed the TATSCAN-F2 and Magix PRO analyses along with MO. All authors read and approved the final manuscript.
Authors’ information
KI is an assistant professor at Chiba University, Japan; YH is a post-doctoral
fellow at the Geological Survey of Japan, AIST; YS is an associate professor at
NIPR and the Graduate University for Advanced Studies (SOKENDAI), Japan;
MO is a professor at Ibaraki University, Japan; YK is a researcher at the National Museum of Nature and Science, Japan; NN is an associate professor atTokyo Gakugei University, Japan; MK is a doctoral student at the National Institute of Polar Research and the Graduate University for Advanced Studies(SOKENDAI); TM is a technician at the Center for Advanced Marine Core Research, Kochi University.
Funding
This work was supported by Grants-in-Aid for Scientific Research (KAKENHI)(Grant Numbers 17H06561, 16H04068, and 19H00710) from the Japan Society
for the Promotion of Science (JSPS).
Availability of data and materials
The datasets supporting the conclusions of this article are available asadditional files (Tables S1, S2, S3 and S4).
Competing interests
The authors declare that they have no competing interest.
Author details
1
Faculty & Graduate School of Education, Chiba University, 1-33 Yayoi-Cho,Inage-Ku, Chiba-shi, Chiba 263-8522, Japan. 2Geological Survey of Japan,
AIST, 1-1-1 Central 7, Higashi, Tsukuba, Ibaraki 305-8567, Japan. 3National
Institute of Polar Research, 10-3 Midori-Cho, Tachikawa, Tokyo 190-8518,
Japan. 4Department of Polar Science, School of Multidisciplinary Sciences,
The Graduate University for Advanced Studies (SOKENDAI), Midori-Cho 10-3,
Tachikawa, Tokyo 190-8518, Japan. 5Department of Environmental Sciences,
Ibaraki University, 2-2-1 Bunkyo, Mito, Ibaraki 310-8512, Japan. 6Departmentof Geology and Paleontology, National Museum of Nature and Science, 4-1-1
Amakubo, Tsukuba, Ibaraki 305-0005, Japan. 7Department of Environmental
Sciences, Tokyo Gakugei University, 4-1-1 Nukuikita, Koganei, Tokyo 184-8501,
Japan. 8Center for Advanced Marine Core Research, Kochi University, B200
Monobe, Nankoku, Kochi 783-8502, Japan.
Received: 26 March 2020 Accepted: 19 November 2020

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