ARTICLE IN PRESS

Quaternary Science Reviews 25 (2006) 323–335

Eustatic sea-level change at the Mid-Pleistocene climate transition:
new evidence from the shallow-marine sediment record of Japan
Akihisa Kitamuraa,, Takeshi KawagoebaInstitute of Geosciences, Faculty of Science, Shizuoka University, Shizuoka 422-8529, Japanb
Kumagaigumi, Co. Ltd., 2-1, Tsukudo-Cho, Shinjuyuku, Tokyo 162-8557, Japan
Received 23 May 2004; accepted 23 February 2005

Abstract
Based on the excellent sedimentary record, we have identiﬁed ﬁve depositional sequences in the upper part of the Omma
Formation, exposed along the Japan Sea coast of central Japan. The base of this stratigraphic section is placed just below the top of
the Jaramillo Subchron, and the top is located below the Brunhes-Matuyama boundary. Biostratigraphy supports the interpretationthat the lower three sequences correlate with marine oxygen isotope stages (MIS) 28–26, 26–24 and 24–22, respectively. We havereconstructed a eustatic sea-level curve from analyses of sedimentary facies, fossil molluscs, basin subsidence and sedimentaccumulation. The resulting sea-level curve shows that mean eustatic sea-level dropped between 1.0 and 0.9 Ma, and that sea-level
was lowest at MIS 22, which has the heaviest d18O value between MIS 28 and 22. The total fall in mean sea-level is inferred to be
20–30 m. This independent sea-level record demonstrates that ice-sheet volume rapidly increased at the Mid-Pleistocene Climate
Transition (MPT) and culminated in MIS 22, as inferred from the d18O record. The upper two depositional sequences imply thecontribution of precession to changes in ice sheet volume would have arisen in MIS 21, just after the MPT.
r 2005 Elsevier Ltd. All rights reserved.

1. Introduction
The switch from predominant 41-ka to predominant
100-ka glaciation cycles at around 1.0–0.9 Ma (the MidPleistocene Climate Transition (MPT)) occurred without a corresponding change in orbital forcing (e.g.,
Pisias and Moore, 1981). High-resolution oxygen isotope records imply that an ice volume-related increase inmean d18O took place near the top of the Jaramillo
Subchron, with a value of 0.36% or 0.18% (Prell, 1982;
Maasch, 1988). According to Mudelsee and Schulz(1997), the MPT was centered at 922712 ka and lasted
about 4079 ka, but the amplitude of the 100-ka cycle
abruptly increased much later on, at 64179 ka. The
MPT consisted of a d18O increase of 0.2970.05%,
which was caused by an expanding ice mass and minor
Corresponding author. Tel.: +81 54 2384798;
fax: +81 54 238 0491.
E-mail address: seakita@ipc.shizuoka.ac.jp (A. Kitamura).

0277-3791/$ - see front matter r 2005 Elsevier Ltd. All rights reserved.
doi:10.1016/j.quascirev.2005.02.009

temperature effects. Raymo et al. (1997) discovered that
this transition was accompanied by a 0.3% decrease inmean ocean d13C at the MPT and suggested that thisone-time addition of isotopically depleted terrestrial
carbon may possibly have been related to an increase inglobal aridity. Internal climatic mechanisms such as icebedrock interaction, atmospheric CO2 levels, tectonic
processes, and a nonlinear response to precession havebeen among the hypothesized causes (Raymo et al.,
1988; DeBlonde and Peltier, 1991; Raymo and Ruddiman, 1992; Clark et al., 1999; Elkibbi and Rial, 2001).
To solve this problem requires precise information
about changes in global ice volume at the MPT.
Although the history of continental ice volume comesfrom oxygen isotope records of foraminifera in deepmarine sediments, these records are also inﬂuenced byregional climate changes. In fact, there are differences in
the pattern and amplitude of ﬂuctuations amongstudies of the MPT (e.g., Mudelsee and Schulz, 1997).
Direct calibration of the ice-volume component of the

ARTICLE IN PRESS
324

A. Kitamura, T. Kawagoe / Quaternary Science Reviews 25 (2006) 323–335

deep-marine oxygen isotope records with actual eustatic
sea-level changes has been accomplished in studies oflate Pleistocene coral-terrace sequences (e.g., Chappell
et al., 1996) and submerged speleothems (e.g., Richards
et al., 1994). Whether the same calibration is valid for
the MPT isotopic records is unknown, because an
independent calibration of the ice volume does not exist.
Shelfal and nearshore sedimentary records are an
independent archive of eustatic sea-level changes (Naish,
1997; Kamp and Naish, 1998). Exposures of midPleistocene shallow-water sediment are known fromthe Wanganui Basin in New Zealand (Abbott andCarter, 1994), the Croton Basin in southern Italy (Rio etal., 1996; Massaria et al., 1999), and the San Francisco
Peninsula in northern California (Carter et al., 2002).
However, these authors did not identify oxygen isotope
stage 23 in their lithostratigraphic schemes.
The Omma Formation is exposed in the vicinity ofKanazawa City on the Sea of Japan coast of central
Japan and consists mainly of shallow-marine deposits.
Its upper part occupies the interval between the top of
the Jaramillo Subchron and the Matuyama/Brunhes
transition (Kitamura, 1994). On the basis of litho- andbiofacies, the upper part of the Omma Formation at its
type section (the Okuwa section) was deposited in avariety of coastal plain, shoreface, and marine-shelfenvironments. Although three depositional sequences
have been recognized in the upper part of the Omma
Formation, and correlated with oxygen isotope stages
28–22, there is a discrepancy between these correlations
and biostratigraphic data (Kitamura, 1994). However,
we have gained important insights into the sequencestratigraphy of these beds by examining new exposuresand evaluating drilling data associated with highway
construction. As a result, we deﬁne the upper part of
the Omma Formation to contain ﬁve depositionalsequences and, as noted below, resolve the discrepancy
in correlation between biostratigraphic and oxygenisotope ages. Further, by comparing a sea-level curve
inferred from this revised sequence stratigraphic
interpretation with deep-sea oxygen isotope records,
we provide constraints on global sea-level changes at
the MPT.

2. Geological setting of Omma Formation depositionalsequences
The sedimentary basin in which the Omma Formation
was deposited is a back-arc basin with respect to thesubducting plate boundaries of the Paciﬁc, Philippine
and Eurasia plates (Fig. 1). The Omma Formation
unconformably overlies the Middle Miocene Saikawa
Formation that consists of massive siltstone (Ogasawara, 1977). Because the unconformity is penetrated by
the marine rock-boring bivalve Zirfaea sp., the Saikawa

Fig. 1. Map showing tunnel, locations of cores and type section of theOmma Formation. Geologic map is modiﬁed from Imai (1959).

Formation sediments must have lithiﬁed prior todeposition of the Omma Formation (Kitamura, 1997).
The sediments of the Omma Formation are poorly
unconsolidated and can be easily scored with a knife.
The 210-m thick Omma Formation has been divided
into lower, middle and upper parts, based on its lithoand biofacies (Kitamura and Kondo, 1990). Its lower
and middle parts are composed of fourteen sixth-order(41-ka) depositional sequences that were deposited in
inner- to outer-shelf depths during oxygen isotope stages
56–28 (Fig. 2) (Kitamura et al., 1994, 2001). These
depositional sequences include the following architectural elements, in ascending stratigraphic order: (1) abasal sequence boundary that is superposed on theravinement surface; (2) a transgressive systems tracts(TST) (2–5 m thick) consisting of a basal shell bed (0.3 m
thick) (a condensed onlap shell bed) and overlying ﬁneto very-ﬁne-grained sandstone; (3) a maximum ﬂooding
horizon that coincides with the horizon that has the
maximum concentration of sand-size carbonate grains;(4) a highstand systems tracts (HST) (2–3 m thick)consisting of ﬁne-grained sandstone and sandy siltstone;and (5) a regressive systems tracts (RST) (o1 m thick)comprising ﬁne-grained sandstone with a coarseningupward trend (Kitamura et al., 2000). These parts of theOmma Formation show no progressive shift in lithoand biofacies toward deeper or shallower deposits. In
addition, if oxygen isotope records (e.g., Ruddiman
et al., 1989; Shackleton et al., 1990; Berger et al., 1994)

ARTICLE IN PRESS
A. Kitamura, T. Kawagoe / Quaternary Science Reviews 25 (2006) 323–335

325

Fig. 2. Columnar section of the Omma Formation at its type section. Biostratigraphic datum horizons are after Takayama et al. (1988) and Sato and
Takayama (1992); magnetostratigraphic data from Ohmura et al. (1989) and Kitamura et al. (1994); time scale for the oxygen isotope record at
DSDP Site 607 and ages of biostratigraphic datum horizons and magnetic polarity changes are based on chronology of Berger et al. (1994). SB:
Sequence boundary. L-1 to 3, 1–11, I–III: depositional sequence numbers.

ARTICLE IN PRESS
326

A. Kitamura, T. Kawagoe / Quaternary Science Reviews 25 (2006) 323–335

are regarded as a proxy of the glacio-eustatic sea-level
record, there were no signiﬁcant changes in averagesea-level during deposition of this formation. These
characteristics imply that the sedimentation rate
matched the subsidence rate.
The upper part of the Omma Formation is 110-m
thick, has an E–W strike and dips 101 north. Based on
the stack of seven sedimentary facies, this upper part isthought to consist of three depositional sequences(Kitamura, 1994), numbered from DS I to DS III(Fig. 3). Depositional sequence I (DS I) was partiallycovered by ﬂuvial deposit (Fig. 2). The ﬁrst-appearance
datum of Gephyrocapsa parallela (1045725 ka; Berger
et al., 1994) and the last-appearance datum of Reticulofenestra asanoi (889725 ka; Berger et al., 1994) havebeen placed within DS I (Takayama et al., 1988). A
magnetostratigraphic study by Ohmura et al. (1989)
inferred that the Omma Formation is assignable to the
Matuyama Reversed Polarity Chron. The top of the
Jaramillo Subchron is placed 65715 cm above the lower
sequence boundary of DS I (Kitamura et al., 1994).
Kitamura (1994) inferred that the depositional ages of
DS I, II and III correspond with oxygen isotope stages
28–26, 26–24 and 24–22, respectively. However, based on
the new information herein, this preliminary correlation
is not in agreement with the stratigraphic position of thelast-appearance datum of R. asanoi (Fig. 3).

3. Sedimentary facies
A new outcrop of the upper part of the Omma
Formation was exposed at its type section during
construction of bridge piers, and allows us to observe
all of depositional sequence I. A large tunnel (663 m inlength) connected to a bridge was dug into a terrace on
the right bank of the Saikawa River in a E–W direction(Fig. 1). Prior to construction of the tunnel, eight cores
were drilled at horizontal intervals of 50–100 m along
the proposed path of the tunnel (Fig. 1). These cores
correspond with the upper portion of DS I and the lower
portion of DS II. In addition, we directly observed thesemi-elliptic face of the tunnel (7.8 m height and 7.0 mwide) at each 1 m of drilling advance. The stratigraphic
interval we studied in the tunnel corresponds with theupper portion of DS I. By relying on a combination ofthese new sedimentary records and published studies, wehave identiﬁed seven sedimentary facies. The stratigraphic distribution of these facies is illustrated in
Figs. 3 and 4.

life positions, and are mostly aligned parallel to bedding.
This preservation implies that these deposits do notcontain species transported post-mortem from otherhabitats, and so evidently are indigenous. These deposits
occupy the lowest portion of DS I and the middle
portion of DS II (Fig. 3).
Within the 4-m thick facies 1 of DS I, molluscan fossil
associations change, in ascending order, from a transitional association to a Clinocardium–Turritella association (Kitamura et al., 1994). The former is characterized
by a mix of cold-water species (e.g., Acila insignis,
Clinocardium fastosum and Yoldia notabilis) and warmwater species (e.g., Stellaria exutum and Paphia schnelliana) (Fig. 3). This association probably lived at a depthof 20–30 to 50–60 m in a transitional environment
between cold and warm water masses (Kitamura et al.,1994). The Clinocardium–Turritella association probably
dwelled in the upper sublittoral zone from the tidemark to 50–60 m depth in a cold-water environment(Kitamura et al., 1994).
The thickness of facies 1 in DS II is about 16 m
(Fig. 3), and the lower part of this facies contains
numerous well-preserved fossil molluscs. This fossiliferous interval is about 7 m thick and changes upwards into
unfossiliferous sandstone. The molluscan faunal associations within this fossiliferous interval change, in
ascending order, from a Raetellops–Macoma association, to a Macoma–Jupiteria association (Fig. 3)(Kitamura et al., 1994). The former is characterized byRaetellops pulchellus and Macoma tokyoensis, which
commonly live at depths shallower than 20–30 m in
inner bays around Japan (Matsushima, 1984). The
characteristic species of the latter association are Ringicula doliaris, M. tokyoensis, Jupiteria confuse and Oliva
mustelina, while Conus sieboldii, Pecten albicans, Nitidotellina hokkaidoensis and Panopea japonica are alsopresent. Based on the modern habitat of these species
around Japan, this association dwells between 30 and
50 m deep in warm water (Kitamura, 1994). Molds ofthe large bivalves of M. stimpsoni, A. insignis andMegangulus zyonoensis are present in a generally
unfossiliferous sandstone bed 15 m above the base ofthis unit (Kitamura, 1994). These species are classiﬁed
as Megangulus association and live between the low-tide
mark and 30 m depth in cold-water (Fig. 3) (Kitamura
et al., 1994).
Judging from molluscan fossil associations and thenear absence of physical sedimentary structures such as
laminae, facies 1 is inferred to be an inner shelf deposit.

3.1. Facies 1-inner shelf3.2. Facies 2 sand-dominated lower shoreface-inner shelftransition

This facies is composed of massive, bluish gray, ﬁnegrained sandstone with abundant well-preserved fossil
molluscs. Most of bivalves are conjoined, but are not inThis facies is subdivided into two parts. Facies 2a ischaracterized by brown, well-sorted, massive, mediumgrained sandstone with sporadic pebbles and ichnotaxa

ARTICLE IN PRESS
A. Kitamura, T. Kawagoe / Quaternary Science Reviews 25 (2006) 323–335

327

Fig. 3. Columnar section for the upper part of the Omma Formation. Note that there is a discrepancy between biostratigraphic data and oxygenisotope stages for the correlation of depositional sequences.

such as Rosselia. Small-scale trough lamination and lowangle hummocky cross-stratiﬁcation (HCS) are common
features. Fossil molluscs are rare, and are not identiﬁable due to dissolution of the shells. Facies 2a deposits

occupy the middle portion of DS I and the upper
portion of DS II. The thicknesses of the former andlatter facies are about 17 and 35 m, respectively (Figs. 3
and 4). A few individuals of the echinoid Scaphechinus

ARTICLE IN PRESS
328

A. Kitamura, T. Kawagoe / Quaternary Science Reviews 25 (2006) 323–335

Fig. 4. Correlation panel for drill cores. See legend in Fig. 3 and text.

sp. occur in this facies in DS II. A coarsening-upward
trend is evident in both horizons (Figs. 3 and 4).
Facies 2b is characterized by rhythmic 40-cm-thickgraded sandstone beds with sharp erosional bases. The
lower part of each bed consists of HCS or parallel- andtrough cross-laminated, coarse to medium sand. Symmetrical ripple laminae are present but rare. The upper
part of each bed consists of massive sandstone with
Ophiomorpha and echinocardium burrows. Facies 2b is
2 m thick, is recognized in the uppermost part of DS II
and overlies facies 2a (Fig. 3).
According to Hamblin and Walker (1979), parallel orswaley cross-stratiﬁcation (SCS) and wave ripple-laminated sands are storm-generated sand sheets that were
deposited between mean fair-weather wave base andmean storm wave base. Facies 2 is interpreted as an
open-ocean, sand-dominated, shoreface-inner shelf transition deposit. Although storm-generated sand sheets
are present in both facies 2a and 2b, their preservation in
the latter is very good compared with the former. As aresult, we believe that the depositional depth for facies
2a was greater than for facies 2b.
3.3. Facies 3 mud-dominated lower shoreface-inner shelftransition
This unit is composed of interbedded sandstone andmudstone. The sandstone beds, commonly 10–40 cm
thick, are graded and have sharp bases. They also arecommonly parallel-laminated and wave-ripple laminated, and rarely show SCS or minor planar cross-Fig. 5. Alternation of sandstone and mudstone (Facies 3) observed at
the tunnel face: (a) Overview of alternation of sandstone andmudstone. Arrows show top of mudstone layers. (b) Close view
showing the alternation of sandstone and mudstone.

stratiﬁcation. The thickness of the massive mudstone
beds is less than 10 cm. These sediments were notaffected by the burrowing of benthic organisms, and thisfacies is present in two horizons within the upper part of
the Omma Formation. The upper horizon is 2 m thickand makes up the lower portion of DS III (Fig. 3). As
seen in the tunnel exposures, this facies has a lower
horizon that lies within the middle portion of DS I, andit pinches out westward (o3 m thick) so that it is notpresent in the type section (Figs. 3–5). Coarsening- andthickening-upward trend are identiﬁed in the lower
horizon (Fig. 4).

ARTICLE IN PRESS
A. Kitamura, T. Kawagoe / Quaternary Science Reviews 25 (2006) 323–335

329

It is widely accepted that the parallel or SCS and wave
ripple-laminated sands are storm-generated sand sheets
deposited between mean fair-weather wave base andmean storm-wave base. In most cited examples, sandsheets are interbedded with bioturbated mudstone.
However, in the present study, burrows are not present
in mudstones or at the upper boundaries of sandstone
sheets. This implies that the recurrence intervals ofstorm deposits were relatively short. Saito (1989)
examined modern storm deposits in the nearshore zone
(20 m depth) close to the mouths of rivers near Sendai
Bay, northeast Japan, and discovered that they consistof alternations of silt and sand layers without burrowing. Saito (1989) estimated the recurrence intervals ofthese storm deposits at 20–100 years. Since recent
sediments show the same sedimentlogical characteristics
as facies 3, we infer that facies 3 was deposited at a
shoreface-inner shelf transition off a river mouth
in a bay.

and deers have been found on mudstone bedding planes(Matsuura, 1996). A 3-m-thick conglomerate bed ispresent at the base of back-marsh deposits on the eastside of the study section (Fig. 4), and consists of clastimbricated pebble gravels. The dip of the imbrication
indicates that direction of paleo-ﬂow was to the northwest (present seaward).
The occurrence of erect stumps and land mammal
tracks indicates that the sediment of this unit was
subaerially exposed. The sandstone beds are thin andgraded, with sharp bases. These sedimentary features are
simply the products of episodic decelerating ﬂows. In acontinental context, the sedimentary environment of thisalternation may be interpreted as back marsh deposits,
whereas the sandstone beds can be interpreted as a ﬂood
deposit in crevasse splays. This facies is present in thelower portion of DS II and is 10 m thick.

3.4. Facies 4 lower shoreface

The facies is characterized by 20-cm-thick shell beds
with erosional surfaces, and a ﬁning-upward trend in thematrix. This facies is found from the bases of facies 1
and 2a within DS I, respectively (Fig. 3). The upper shell
bed is covered by a 30-cm-thick massive siltstone.
These shell beds consist of M. zyonoensis, A. insignis,
Glycymeris yessoensis and Mizuhopecten yessoensis
yokoyamae (Kitamura, 1994). These shells are absent
in sediment immediately below the shell beds, so areunlikely to have been reworked from older deposits. In
the upper shell bed, the underlying amalgamated andSCS very ﬁne sandstone do not contain fossils ofmolluscs or other macroinvertebrates. These molluscs
inhabit well-sorted sandy bottoms between the low-tide
mark and a depth of 30 m in cold water (Kitamura et al.,1994). Most molluscs in this facies do not show
abrasion, damage due to Cliona boring, or encrustation
by calcareous algae, barnacles or bryozoans. These
characteristics are not consistent with a general model ofshell accumulation resulting from low sedimentation
rates. Because the molluscan fossil associations change
from a cold-water association in the lower shell bed to atransitional association in overlying facies 1, this shell
bed may have formed during a deglacial period. As
noted below, the upper shell bed also exhibit evidence ofa deepening-upward trend. We concluded that this facies
is transgressive lag formed by coastal shoreface erosion.

This facies is characterized by amalgamated, swaley
cross-stratiﬁed, well-sorted, ﬁne to very ﬁne-grainedsandstone, and is present in the middle portion of DS I
(13 m thick) and the upper part of DS III (2 m thick)(Fig. 3). The facies is inferred to represent a lowershoreface environment, and was deposited above mean
fair-weather wave base (Bourgeois, 1980; Murakoshi
and Masuda, 1992; Walker and Plint, 1992).
3.5. Facies 5 upper shoreface
Facies 5 is characterized by composite sets of highangle, planar or trough cross-stratiﬁed, well-sorted,
medium to very coarse sandstone beds. The thicknessof a set is less than 1 m. Fossils were not found. The dipsof the cross beds are landward, seaward or parallel tothe shoreline (Fig. 3) (Kitamura, 1994). Judging fromthese features, Kitamura (1994) interpreted this facies as
an upper shoreface. This facies occupies the uppermost
part of DS I and is less than 12 m thick (Figs. 3 and 4).
3.6. Facies 6 back marsh
This unit is composed of alternating massive mudstone beds and coarse-grained sandstone beds, withindividual beds ranging from 10 to 100 cm thick (Fig. 3).
The sandstone beds are graded, with erosional bases.
Small-scale trough lamination is often recognized within
the sandstone beds. Although most mudstone beds are
massive, parallel lamination is present in a few beds.
Plant debris is common in the mudstone, and erect
stumps have been found in the lowermost (Kitamura
and Kondo, 1990) and the uppermost (Takayama et al.,1988) mudstone beds (Fig. 3). The tracks of elephants

3.7. Facies 7 transgressive lag

4. Depositional sequences
Kitamura (1994) showed that the upper part of theOmma Formation consists of three depositional sequences. In the present study, we have divided thelowest of these sequences (DS I) into three depositionalsequences, based on new information about lithofacies

ARTICLE IN PRESS
330

A. Kitamura, T. Kawagoe / Quaternary Science Reviews 25 (2006) 323–335

distribution (Fig. 4). As a result, the upper part of theOmma Formation is now interpreted as consisting ofﬁve sequences (Fig. 6).
Depositional sequence U1 (DS U1) is 17 m thick andconsists of the following facies, in ascending stratigraphic order: basal shell bed (Facies 7, transgressivelag), massive ﬁne-grained sandstone with abundantmolluscan fossils (Facies 1, inner shelf) and amalgamated SCS ﬁne to very ﬁne-grained sandstone (Facies 4,
lower shoreface) (Fig. 6). The erosional base of the basal
shell bed is also the lower sequence boundary andcoincides with a ravinement surface formed by coastal
shoreface erosion during a transgression (Bruun, 1962).
Warm-water molluscs in the transitional association
imply that the lower part of the massive ﬁne-grainedsandstone was deposited during an interglacial period.

Thus, we regard the massive sandstone with thisassociation and the underlying shell bed as a TST.
Inner-shelf, massive sandstone is directly overlain by alower-shoreface SCS sandstone. This indicates that
sediments of the lower shoreface-inner shelf transition
were eliminated. The base of the SCS sandstone is sharp
and overlain by a 20-cm-thick granule-gravel bed(Kitamura, 1994). Such a sharp-based shoreface succession has been reported from many areas (e.g., Plint,
1988, 1991; Plint and Norris, 1991). The base of the SCS
sandstone is believed to have formed by wave scour
during a sea-level fall (Posamentier et al., 1992). The
sediment above the erosinal base is deﬁned as a fallingstage systems tract (FSST) (Plint and Nummedal, 2000).
Thus, we think that the SCS sandstone and theunderlying massive sandstone with the cold-water

Fig. 6. Sequence stratigraphic interpretation and comparison of eustatic sea-level curve derived from sedimentary record in the upper part of theOmma Formation, with oxygen isotope records from sites 607, 677 and 806 (Ruddiman et al., 1989; Shackleton et al., 1990; Berger et al., 1994). Seelegend in Fig. 3.

ARTICLE IN PRESS
A. Kitamura, T. Kawagoe / Quaternary Science Reviews 25 (2006) 323–335

Clinocardium–Turritella association correspond with an
FSST and an HST, respectively (Fig. 6).
Depositional sequence U2 (DS U2) has a maximum
thickness of 17 m and commences with a 20-cm-thickshell bed (Facies 7, transgressive-lag) resting on a sharp
erosional surface and being overlain by a 20-cm-thickmudstone sheet (Fig. 6). The presence of mudstone sheet
indicates that shell bed deposited below fair-weather
wave base. Since the sediment below the shell bed islower-shoreface SCS sandstone, the shell bed can beinterpreted as transgressive lag formed by coastal
erosion. Its base is also the lower sequence boundaryand superposed the ravinement surface. A well-sorted,
massive, medium-grained sandstone lies above the basal
shell bed and exhibits a coarsening-upward trend (Facies
2a, sand-dominated lower shoreface-inner shelf transition) (Fig. 4). Therefore, we infer that the sandstonecorresponds to an HST. In the case of all fourteen
depositional sequences in the lower and middle parts of
the Omma Formation, it is not possible to identify amaximum ﬂooding surface without examining thestratigraphic distribution of planktonic foraminiferal
abundance and/or fossil molluscs (Kitamura, 1998;
Kitamura et al., 2001). Shells of both micro- andmacrofossils in DS U2 have been dissolved, so we wereunable to identify the HST base. We regard the basal
shell bed and the overlying deposits as the TST and
HST, respectively.
The well-preserved depositional sequence U3 (DS U3)
is exposed within the tunnel. The sequence is 13 m thick,and comprises interbedded sandstone and mudstone(Facies 3, mud-dominated lower shoreface-inner shelftransition) and cross-stratiﬁed, well-sorted, medium tovery coarse sandstone (Facies 5, upper shoreface) (Figs.
4 and 6). The stacking pattern from the well-sorted,
massive sandstone (HST) in DS U2 to the presentinterbedded sandstone and mudstone implies that thedepositional environment drastically changed from asand-dominated to a mud-dominated shelf. Because thisstacking pattern cannot be explained as a regressive
setting, we believe that the sharp erosional base of theinterbedded sandstone and mudstone is an unconformity and sequence boundary. The Facies 3 sediment
represents coarsening- and thickening-upward trends
for the sandstone beds, which we interpret as a TST
and HST. This unit is directly overlain by theupper-shoreface sandstone. This means that the lowershoreface zone has been eliminated, and that the uppershoreface sandstone corresponds with the FSST, whosebase represents a regressive surface of marine erosion(RSME, Plint, 1988; Hart and Long, 1996). Thus, thissequence mainly consists of an upward-shoaling facies
succession and is regarded as a parasequence. Our
observations suggest that the degree of erosion duringformation of the RSME was grater to the westward(Fig. 6), where the lower shoreface-inner shelf transition

331

sediments were completely truncated at the bed along
Saikawa River. Thus, DS U3 in this section contains
only the FSST that comprises a progradational uppershoreface facies (Fig. 4). This caused the misidentiﬁcation of depositional sequences in DS I as presented byKitamura (1994). The degree of preservation in DS U3suggests that water depth increased eastward. This isconsistent with reconstructions based on fossil molluscs
in the middle part of the Omma Formation (Kitamura
et al., 1997).
Depositional sequence U4 (DS U4) correspondswith depositional sequence II in Kitamura (1994). In
this 55-m-thick sequence, the lithofacies are, in ascending order, an alternation of massive mudstone beds andcoarse-grained sandstone beds (Facies 6, back marsh),massive bluish gray ﬁne-grained sandstone (Facies 1Inner shelf), well-sorted, massive, medium-grainedsandstone (Facies 2a, lower shoreface-inner shelf transition), and rhythmic graded sandstone beds (Facies 2b,
lower shoreface-inner shelf transition) (Fig. 6). Basedon relative sea-level change inferred from the stacking
pattern of these lithofacies, Kitamura (1994) thoughtthat the base and top of back-marsh deposits correspond with a lower sequence boundary and a ravinement surface, respectively. In such a case, backmarsh deposits are interpreted to represent a low
system tract and early TST deposits. The changes ofmolluscan fossil associations in the inner-shelf deposits
indicate that the marine climate became warm and thencooled. In glacial-eustatic terms, sea-level is relative highduring warm interglacial periods. Therefore, Kitamura(1994) placed the maximum ﬂooding surface, whichis the boundary between the TST and HST, at
the midpoint of the horizon containing warm-watermolluscs.
Depositional sequence U5 (DS U5) coincides withdepositional sequence III of Kitamura (1994), and iscomposed of interbedded sandstone and mudstone(Facies 3, mud-dominated lower shoreface-inner shelftransition) and amalgamated SCS ﬁne to very ﬁnesandstone (Facies 4, lower shoreface) (Fig. 6). Thisfacies succession has been reported from wave-dominated and gradationally based shoreface successions(e.g., Heward, 1981; MuCubbin, 1982). Such successions
typically begin with outer-middle shelf, bioturbatedmudstone. Therefore, the stacking pattern from thesandstone in the uppermost part of DS U4 to theoverlying interbedded sandstone and mudstone cannot
be explained by a progradational inner-shelf to shorefacies assemblage. We conclude that the sharp boundary
between them corresponds to a sequence boundary.
Kitamura (1994) tentatively placed the base of the HSTat the midpoint of interbedded sandstone and mudstone,
because these sediments do not exhibit systematicchanges such as ﬁning-upward or thickening-upward
trends.

ARTICLE IN PRESS
332

A. Kitamura, T. Kawagoe / Quaternary Science Reviews 25 (2006) 323–335

5. Discussion
We ﬁrst correlated depositional sequences in theupper part of the Omma Formation with oxygen isotope
records (Fig. 6). As noted above, the top of the
Jaramillo Subchron is placed just above the basal shell
bed in DS U1. Thus, the TST sediment in this sequence
was deposited during a deglacial sea-level rise in MIS 28
to 27. According to Takayama et al. (1988) and Satoand Takayama (1992), Gephyrocapsa parallela is only
found from inner-shelf sandstone in DS U4, whereas
R. asanoi does not occur in the sandstone in DS U4.
This implies that the sea-level highstand during thedeposition of DS U4 is younger than the last-appearance datum of R. asanoi (889725 ka). Furthermore, thetop of the Omma Formation is younger than the
Matuyama/Brunhes transition (79075 ka; Berger etal., 1994) (Ohmura et al, 1989). From these observations, the depositional age of the inner-shelf sandstone
in DS U4 correlates with either MIS 23 or 21. Deep-sea
d18O records for sites 607, 677 and 806 (Ruddiman et al.,
1989; Shackleton et al., 1990; Berger et al., 1994) showthat the values for MIS 22 is signiﬁcantly heavier thanthose for MIS 26 and 24. In addition, the duration of
MIS 22–20 seems to have been 70–78 ka (Shackleton etal., 1990; Bassinot et al., 1994; Berger et al., 1994). Since
DS U4 is anomalously thick and only includes nonmarine sediments in its lowest portion, we believe that
this sequence corresponds to MIS 22–20, and that DS
U2 and U3 correlate with MIS 26–24 and MIS 24–22,
respectively (Fig. 6). If this correlation is correct, then
DS U5 may represent ‘‘interstadial’’ oscillations in
sea level during the glaciation from MIS 21 to 20(see below).
We evaluated water depth based on facies successions
and molluscan fossil associations in the ﬁve depositionalsequences within the upper part of the Omma Formation. Because fair-weather wave base lies at depthsof about 5–15 m (Walker and Plint, 1992), the maximum
depths of the lower and upper shoreface are inferred
to be 15 and 5 m, respectively. Based on the bathymetricranges of molluscan fossil associations, the minimum
water depth for the inner-shelf deposits is inferred to
be 30 m. A water-depth curve plotted against time,
derived from the above-noted correlation between
depositional sequences and oxygen isotope stages, isshown in Fig. 6.
Kitamura et al. (1994) showed that water depthchanged from 20–30 to 100–120 m during depositionalsequences 10 and 11 (MIS 32–28) in the middle part of
the Omma Formation. These sequences do not include
sediments that were deposited shallower than the innershelf. However, the upper part of the Omma Formation
mainly consists of deposits shallower than 30 m (96. 4 mof 116.4 m in thickness, 83%). This implies that meanwater depth decreased from 1.0 Ma (just above the topof the Jaramillo Subchron) to 0.9 Ma (MIS 22). The
total decrease in water depth is inferred to be 20–30 m.
This value is a minimum estimate, because of remaining
uncertainty about sea level during the deposition of
back-marsh sediments in MIS 22, and about thelowstand strata lost at the bounding unconformities.
Water-depth changes were inﬂuenced by basement
subsidence, sediment supply and eustatic sea-levelchange. The aggradational stacking pattern of deposits
in the lower and middle parts of the Omma Formationsuggests that sedimentation kept pace with basin
subsidence at 600 ka (MIS 56–28). There are no reportsof local tectonic movement in the study area during
deposition of the Omma Formation. In addition, wewere unable to identify differences in the mineralcomposition or in relative abundances of fossil molluscs
within the massive sandstone that lies between the
middle and upper parts of this formation (Facies 1). It
is, therefore, reasonable to think that the sedimentation
rate also matched the subsidence rate during depositionin the upper part of the Omma Formation. In addition,
the effects of local isostasy may not have changed. Basedon correlations between depositional sequences withoxygen isotope stages, the sedimentary ratios of depositional sequences 10 and 11 are estimated to be 18.8 and27.6 cm/ka, respectively. On the other hand, theapparent rates of sediment accumulation in depositionalsequences U1, U2 and U3 are 36 to 42 cm/ka. Such a
difference in sedimentation rates between the middle
and upper parts of the Omma Formation resulted froma drop in mean water depth.
The water-depth curve can also be affected bycompaction due to water and sediment loading. The
basement rock for the Omma Formation had alreadybeen consolidated prior to deposition of this formation(Kitamura, 1997). The Omma Formation mainly consists of very ﬁne-to-ﬁne grained sandstone, with theirmaximum burial depth equal to the thickness (210 m) of
the Omma Formation. Based on the porosity-depth
relationship (Sclater and Christie, 1980; Ramm and
Bjiørlykke, 1994), sandstone porosity decreases only 5%
between the surface and a burial depth of 250 m. As aresult, the water-depth curve was hardly inﬂuenced by
the effects of compaction. The sum of these factors
suggests that a decrease in mean water depth of 20–30 m
was caused by global sea-level fall (Fig. 6).
Prell (1982) and Maasch (1988) documented a rapid
increase in mean d18O composition of 0.36% or 0.18%
near the top of the Jaramillo Subchron. These values
equate to a sea-level fall of about 33 m and 16 m,
respectively. More recently, Mudelsee and Schulz (1997)
suggested that the MPT represents a d18O increase of
0.2970.05% ( ¼ 2375 m, sea-level fall) and was causedby an expanding ice mass. These values are very close to
the change in mean water depth recorded in the Omma
Formation.

ARTICLE IN PRESS
A. Kitamura, T. Kawagoe / Quaternary Science Reviews 25 (2006) 323–335

Based on a comparison of maximum water depth ineach depositional sequence, sea-level during MIS 23 waslower than during MIS 27, 25 or 21 (Fig. 6), which isconsistent with many published d18O records (Site 607,
Ruddiman et al., 1989; Site 677, Shackleton et al., 1990;
Site 806, Berger et al., 1994). On the other hand, there isa difference in the d18O values between stages 25 and 27.
The value for stage 25 is 0.3–0.4% lower than for stage
27 in d18O records for Sites 607 and 677, while both
values are nearly the same at site 806. If the entire
difference in d18O values contribute to ice-volume
change, sea level during stage 25 was 27–36 m higher
than during stage 27, based on 0.011% d18O for eachmetre of sea-level change. This water-depth value is notconsistent with our curve, which incorporates uncertainty about water depth by using analyses of sedimentary facies and mollusc fossils (Fig. 6). In addition, thed18O value for stage 21 is slightly heavy relative to stage
25 at Sites 607 and 677, while the former is 0.1% lighter
than that of the latter in Site 806. Our reconstructed amaximum water depth in each depositional sequence isconsistent with the oxygen isotope record at Site 806.
We therefore believe that oxygen isotope record of
G. sacculifer from the Ontong Java Plateau (ODP Site
806; Berger et al., 1994) is most faithful record of sealevel ﬂuctuations at the MPT. However, our generated
curve indicates that the amplitude of sea-level ﬂuctuation for each depositional sequence is signiﬁcantly
smaller than at Site 806 (calibration of d18O records
using 0.011% per metre). There are at least two possible
causes for this discordance. First, we do not know what
the sea level was during deposition of the back-marsh
sediments and lowstand strata lost at the bounding
unconformities. Our estimates of water depths are
therefore regarded as minimum values. Second, thed18O record at Site 806 was also inﬂuenced by regional
climate change. Many studies have noted that averagesea-surface temperature during the last glacial maximum
was only 1–2 1C lower than in the modern western
Paciﬁc (Broecker, 1986; Trend-Staid and Prell, 2002).
However, the magnesium/calcium ratio of G. ruber at
Site 806 indicates a surface temperature was 3 1C colder
than present during the last glacial maximum (Lea et al.,2000). In this latter case, the glacial–interglacial d18O
amplitude at the site would be larger than that due to
ice-volume changes.
It is noteworthy that DS U5 represents a sea-level
oscillation in stage 21 (Fig. 6). The DS U4 horizon
exhibiting maximum water depth is a massive, innershelf sandstone, whereas in DS U5 this horizon consistsof lower shoreface-inner shelf transition sediments. Thismeans that the peak sea-level highstand during DS U5
was lower than in DS U4. According to Bassinot et al.
(1994), a high-resolution oxygen isotope record fromcore MD900963, in which precession-related oscillationsof d18O are particularly well expressed, was caused by

333

superposition of a local salinity signal over the global ice
volume signal, and shows three precessional oscillations.
The ice-volume curve of Bassinot et al. (1994) showsthat ice volume was reduced in the latter interstadial
periods. Thus, we conclude that the horizons withmaximum water depth in DS U4 and U5 correspondwith stages 21.5 and 21.3, respectively. This implies thatthe precession contribution to sea-level variability wasstronger MIS 21.

6. Conclusions
We identiﬁed ﬁve depositional sequences in the upper
part of the Pleistocene Omma Formation, based on the
excellent sedimentary record at new exposures and in
drilling cores associated with highway construction.
When these results are combined with publishedbiostratigraphic and magnetostratigraphic data, weconclude that the lower three depositional sequences
correlate with MIS 28–22, which clearly experienced
obliquity-related climate cycles. On the other hand, theupper two depositional sequences in the Omma Formation correspond with MIS 21.5 and 21.3, respectively.
Our resulting eustatic sea-level curve, based on analysesof lithofacies, fossil molluscs and subsidence, indicatesthat a fall in eustatic sea-level of 20–30 m took placefrom MIS 28 to 22 (1.0–0.9 Ma). This result is inexcellent agreement with the sea-level fall inferred fromthe d18O record. Thus, the rapid growth of ice sheets at
the mid-Pleistocene transition is conﬁrmed by the shallowmarine record in the Omma Formation. In addition, theupper two depositional sequences in the Omma Formation imply that the precessional contribution to theformation and demise of continental ice sheets was
already in evidence at MIS 21, just after MPT.

Acknowledgments
We thank the late Dr. Y. Kaseno, emeritus Professor
at Kanazawa University, for his helpful comments on
our study. We acknowledge Pref. Ishikawa Ken-O
General Public Works Ofﬁce for opportunity ofobservation of the tunnel.
References
Abbott, S.T., Carter, R.M., 1994. The sequence architecture of midPleistocene (0.35–0.95 Ma) cyclothems from New Zealand: facies
development during a period of orbital control on sea-level
cyclicity. In: de Boer, P.L., Smith, D.G. (Eds.), Orbital Forcing
and Cyclic Sequences. International Association of Sedimentologist Special Publication, vol. 19. pp. 367–394.
Bassinot, F.C., Labeyrie, L.D., Vincent, E., Quidelleur, X., Shackleton,
N.J., Lancelot, Y., 1994. The astronomical theory of climate and theARTICLE IN PRESS
334

A. Kitamura, T. Kawagoe / Quaternary Science Reviews 25 (2006) 323–335

age of the Brunhes–Matuyama magnetic reversal. Earth andPlanetary Science Letters 126, 91–108.
Berger, W.H., Yasuda, M.K., Bickert, T., Wefer, G., Takayama, T.,1994. Quaternary time scale for the Ontong Java Plateau:
Milankovitch template for Ocean Drilling Program Site 806.
Geology 22, 463–467.
Bourgeois, J., 1980. A transgressive shelf sequence exhibiting
hummocky cross stratiﬁcation: the Cape Sebastian Sandstone(Upper Cretaceous), southwestern Oregon. Journal of Sedimentary
Petrology 50, 681–702.
Broecker, W.S., 1986. Oxygen isotope constraints on surface ocean
temperature. Quaternary Research 26, 121–134.
Bruun, P., 1962. Sea level rise as a cause of shoreface erosion. Journalof the Waterways and Harbors Division: proceedings of theAmerican Society of Civil Engineers 88, 117–130.
Carter, R.M., Abbott, S.T., Graham, I.J., Naish, T.R., Gammon,
P.R., 2002. The middle Pleistocene Merced-2 and -3 sequencesfrom Ocean Beach, San Francisco. Sedimentary Geology 153,
23–41.
Chappell, J., Omura, A., Esat, T., McCulloch, M., Pandolﬁ, J., Ota,
Y., Pillans, B., 1996. Reconciliation of late Quaternary sea levels
from coral terraces at Huon Peninsula with deep sea oxygenisotope records. Earth and Planetary Science Letters 141,
227–236.
Clark, P.U., Alley, R.B., Pollard, D., 1999. Northern Hemisphere ice
sheet inﬂuences on global climate change. Science 286, 1104–1111.
DeBlonde, G., Peltier, W.R., 1991. A one-dimensional model ofcontinental ice volume ﬂuctuations through the Pleistocene:
implications for the origin of the mid-Pleistocene climate transition. Journal of Climate 4, 318–344.
Elkibbi, M., Rial, J.A., 2001. An outsider’s review of the astronomical
theory of the climate: is the eccentricity-driven insolation the main
driver of the ice ages? Earth-Science Reviews 56, 161–177.
Hamblin, A.P., Walker, R.G., 1979. Storm-dominated shallow marine
deposits; the Fernie-Kootenay (Jurassic) transition, southern
Rocky Mountains. Canadian Journal of Earth Sciences 16,
1673–1690.
Hart, B.S., Long, B.F., 1996. Forced regressions and lowstand deltas:
Holocene Canadian examples. Journal of Sedimentary Research, A
66, 820–829.
Heward, A.P., 1981. A review of wave-dominated clastic shoreline
deposits. Earth-Science Reviews 17, 223–276.
Imai, I., 1959. Geological map of Kanazawa District (1:50,000):
Geological Survey of Japan [Japanese with English abstract].
Kamp, P.J.J., Naish, T., 1998. Forward modeling of the sequencestratigraphic architecture of shelf cyclothems: application to late
Pliocene sequences, Wanganui Basin (New Zealand). Sedimentary
Geology 116, 57–80.
Kitamura, A., 1994. Depositional sequences caused by glacio-eustatic
sea-level changes in the upper part of the early Pleistocene Omma
Formation. The Journal of the Geological Society of Japan 100,463–476 (Japanese with English abstract).
Kitamura, A., 1997. The lower unconformity and molluscan fossils of
the basal part of the early Pleistocene Omma Formation at its typesection, Ishikawa Prefecture, central Japan. The Journal of theGeological Society of Japan 103, 763–769 (Japanese with English
abstract).
Kitamura, A., 1998. Glaucony and carbonate grains as indicators of
the condensed section: Omma Formation, Japan. Sedimentary
Geology 122, 151–163.
Kitamura, A., Kondo, Y., 1990. Cyclic change of sediments andmolluscan fossil associations caused by glacio-eustatic sea-levelchanges during the early Pleistocene -a case study of the middle
part of the Omma Formation at the type locality. The Journal of
the Geological Society of Japan 96, 19–36 (Japanese with English
abstract).

Kitamura, A., Kondo, Y., Sakai, H., Horii, M., 1994. Fourty-one
thousand-year orbital obliquity expressed as cyclic changes inlithofacies and molluscan content, early Pleistocene Omma
Formation, central Japan. Palaeogeography, Palaeoclimatology,
Palaeoecology 112, 345–361.
Kitamura, A., Kimoto, K., Takayama, T., 1997. Reconstruction of the
thickness of the Tsushima Current in the Japan Sea during the
Quaternary from molluscan fossils. Palaeogeography, Palaeoclimatology, Palaeoecology 135, 51–69.
Kitamura, A., Matsui, H., Oda, M., 2000. Constraints on the timing ofsystems tract development with respect to sixth-order (41 ka) sealevel changes: an example from the Pleistocene Omma Formation,
Sea of Japan. Sedimentary Geology 131, 67–76.
Kitamura, A., Takano, O., Takada, H., Omote, H., 2001. Late
Pliocene-early Pleistocene paleoceanographic evolution of the Sea
of Japan. Palaeogeography, Palaeoclimatology, Palaeoecology 172,
81–98.
Lea, D.W., Pak, D.W., Spere, H.J., 2000. Climate impact of Late
Quaternary equatorial Paciﬁc Sea surface temperature variations.
Science 289, 1719–1724.
Maasch, K.A., 1988. Statistical detection of Pleistocene climate.
Climate Dynamics 2, 133–143.
Massaria, F., Sgavettib, M., Rioa, D., D’Alessandroc, A., Prosserd,G., 1999. Composite sedimentary record of falling stages ofPleistocene glacio-eustatic cycles in a shelf setting (Crotone basin,south Italy). Sedimentary Geology 127, 85–110.
Matsushima, Y., 1984. Shallow marine molluscan assemblages of
postglacial period in the Japanese Islands––its historical andgeographical changes induced by the environmental changes.
Bulletin of the Kanagawa Prefectural Museum 15, 37–109(Japanese with English and German abstract).
Matsuura, N., 1996. Vertebrate fossils from the Omma Formation of
Kanazawa area, Ishikawa Prefecture, Japan. The Report of Hokuriku
Geology Institute 5, 55–88 (Japanese with English abstract).
MuCubbin, D.G., 1982. Barrier island and strand plain facies. In:
Scholle, P.A., Spearing, D. (Eds.), Sandstone Depositional
Environments. American Association of Petroleum Geologists,
Memoir 31, pp. 247–279.
Mudelsee, M., Schulz, M., 1997. The mid-Pleistocene climatetransition: onset of 100 ka cycle lags ice volume build-up by
280 ka. Earth and Planetary Science Letters 151, 117–123.
Murakoshi, N., Masuda, F., 1992. Estuarine, barrier-island to strandplain sequence and related ravinement surface developed during thelast interglacial in the Paleo-Tokyo Bay, Japan. Sedimentary
Geology 80, 167–184.
Naish, T., 1997. Constraints on the amplitude of late Pliocene eustatic
sea-level ﬂuctuations: new evidence from the New Zealand shallowmarine sediment record. Geology 25, 1139–1142.
Ogasawara, K., 1977. Paleontological analysis of Omma Fauna from
Toyama-Ishikawa area, Hokuriku Province, Japan. Science Reports of the Tohoku University, Second Series. Geology 47,
43–156.
Ohmura, I., Ito, T., Masaeda, H., Danbara, T., 1989. Magnetostratigraphy of the Omma Formation distributed in Kanazawa City,
Ishikawa Prefecture, Central Japan. Prof. H. Matsuo Mem.
pp. 111–124 (Japanese with English abstract).
Pisias, N.G., Moore Jr., T.C., 1981. The evolution of Pleistocene
climate: a time series approach. Earth and Planetary Science
Letters 52, 450–458.
Plint, A.G., 1988. Sharp-based shoreface parasequences and ‘offshore
bars’ in the Cardium formation of Alberta; their relationship torelative changes in sea level. In: Wilgus, C.K., Hastings, B.S.,
Kendall, C.G.St.C., Posamentier, H.W., Ross, C.A., Van
Wagoner, J.C. (Eds.), Sea-Level Changes: An Integrated Approach. Society of Economic Paleontologists and Mineralogists
Special Publication 42, pp. 357–370.

ARTICLE IN PRESS
A. Kitamura, T. Kawagoe / Quaternary Science Reviews 25 (2006) 323–335
Plint, A.G., 1991. High-frequency relative sea-level oscillations in
the Upper Cretaceous shelf clastics of the Alberta forelandbasin: possible evidence for glacio-eustatic control? In: Macdonald,
D.I.M. (Ed.) Sedimentation, Tectonics and Eustasy. International Association of Sedimentologist Special Publication 12,
pp. 409–428.
Plint, A.G., Norris, B., 1991. Anatomy of a ramp region sequence:facies successions, paleogeography and sediment dispersal patterns
in the Muskiki and Marshybank Formations, Alberta forelandbasin. Bulletin of Canadian Petroleum Geology 39, 18–42.
Plint, A.G., Nummedal, D., 2000. The falling stage systems tract:
recognition and importance in sequence stratigraphic analysis.
Geological Society Special Publication 172, 1–18.
Posamentier, H.W., Allen, G.P., James, D.P., Tesson, M., 1992.
Forced regressions in a sequence stratigraphic framework:
concepts, examples and exploration signiﬁcance. American
Association of Petroleum Geologists Bulletin 76, 1687–1709.
Prell, W.L., 1982. Oxygen and carbon isotopic stratigraphy for the
Quaternary of Hole 502B: evidence for two modes of isotopic
variability. Initial Reports of the Deep Sea Drilling Project 68,
455–464.
Ramm, M., Bjiørlykke, K., 1994. Porosity/depth trends in reservoir
sandstones: assessing the quantitative effects of varying porepressure, temperature history and mineralogy, Norwegian Shelf
Data. Clay Minerals 24, 475–490.
Raymo, M.E., Oppo, D.W., Curry, W., 1997. The mid-Pleistocene
climate transition: a deep sea carbon perspective. Paleoceanography 12, 546–559.
Raymo, M.E., Ruddiman, W.F., 1992. Tectonic forcing of lateCenozoic climate. Nature, 117–122.
Raymo, M.E., Ruddiman, W.F., Froelich, P.N., 1988. Inﬂuence of lateCenozoic mountain building on ocean geochemical cycles. Geology
16, 649–653.
Richards, D.A., Smart, P.L., Edwards, R.L., 1994. Maximum sea
levels for the last glacial period from U-series ages of submerged
speleothems. Nature 367, 357–360.

335

Rio, D., Channell, J.E.T., Massari, F., Poli, M.S., Sgavetti, M.,
D’Alessandro, A., Prosser, G., 1996. Reading Pleistocene eustasy
in a tectonically active siliciclastic shelf setting (Crotone peninsula,
southern Italy). Geology 24, 743–746.
Ruddiman, W.F., Raymo, M.E., Martinson, D.G., Clement, B.M.,
Backman, J., 1989. Pleistocene evolution: Northern hemisphere
ice sheets and North Atlantic Ocean. Paleoceanography 4,
353–412.
Saito, Y., 1989. Modern storm deposits in the inner shelf and theirrecurrence intervals, Sendai Bay, northeast Japan. In: Taira, A.,
Masuda, F. (Eds.), Sedimentary Facies in the Active Plate Margin.
Terra Scientiﬁc Publishing Company, Tokyo, pp. 331–344.
Sato, T., Takayama, T., 1992. A stratigraphically signiﬁcant new
species of the calcareous nannofossil Reticulofenestra asanoi. In:
Ishizaki, K., Saito, T. (Eds.), Centenary of Japanese Micropaleontology. Terra Scientiﬁc Publishing Company, Tokyo,
pp. 457–460.
Sclater, J.G., Christie, P.A.F., 1980. Continental stretching: an
explanation of the post Mid-Cretaceous subsidence of the central
North Sea basin. Journal of Geophysical Research 85, 3711–3739.
Shackleton, N.J., Bergeret, A., Peltier, W.R., 1990. An alternative
astronomical calibration of the lower Pleistocene timescale basedon ODP Site 677. The Transactions of the Royal Society of
Edinburgh-Earth Sciences 81, 251–261.
Takayama, T., Kato, M., Kudo, T., Sato, T., Kameo, K., 1988.
Calcareous microfossil biostratigraphy of the uppermost Cenozoic
formations distributed in the coast of the Japan Sea, Part 2:
Hokuriku sedimentary basin. Journal of the Japanese Association
for Petroleum Technology 53, 9–27.
Trend-Staid, M., Prell, W.L., 2002. Sea surface temperature at the Last
Glacial Maximum: a reconstruction using the modern analog
technique. Paleoceanography 17 (17–1–17–18).
Walker, R.G., Plint, A.G., 1992. Wave- and storm-dominated shallow
marine systems. In: Walker, R.G., James, N.P. (Eds.), Facies
Models: Response to Sea Level Change. Geological Association of
Canada, pp. 219–238.