1

Quaternary International

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Volume 397, 18 March 2016, Pages 417–421

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Japanese Quaternary Studies

45Constraints on eustatic sea–level changes during the Mid-Pleistocene Climate Transition:

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evidence from the Japanese shallow-marine sediment record

7

Akihisa Kitamura

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Institute of Geosciences, Shizuoka University, 836 Oya, Suruga-ku, Shizuoka, 422-8529, Japan

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ABSTRACT11

During the Middle Pleistocene, the nature of glacial–interglacial fluctuations changed from12

low-amplitude and a periodicity of 41 kyr to high-amplitude and quasi-periodic of 100 kyr. The

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origin of the Mid-Pleistocene Climate Transition (MPT) is an unsolved mystery. At present,

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there is a debate about whether the initiation of the MPT was a gradual or an abrupt process.

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This study investigated the process of initiation of the MPT from reconstructions of eustatic16

sea–level changes, as a proxy for global ice volume, based on a reexamination of lithofacies and17

fossil occurrences from shallow-marine sediments (the Omma Formation) exposed on the west

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coast of Japan. The Omma Formation comprises 19 depositional sequences spanning marine

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isotope stages (MIS) 56–21.3, reflecting sedimentation under alluvial plain to offshore

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conditions. The data indicate that (1) sea–level was lowest during MIS 22 (~0.9 Ma); (2)21sea–level during MIS 34 (~1.13 Ma) and MIS 26 (~0.96 Ma) was lower than during any other22

glacial stage, except for MIS 22; and (3) sea–level during MIS 22 was at most 20 m lower than23

during MIS 34 and 26. Together, these findings suggest that the initiation of the MPT was a

24

gradual, rather than abrupt, process.
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1. Introduction

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Glacial–interglacial cycles, and corresponding changes in eustatic sea–level, have had a28

considerable impact on both global climate and ecosystems throughout the Quaternary. Between

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2.7 and 1.2 Ma, these cycles occurred with a periodicity of 41 kyr (the “41-kyr world”), and are30

attributed primarily to changes in the Earth’s obliquity (Raymo and Nisancioglu, 2003; Huybers,

312006). In contrast, for the past 0.7 Myr, glacial cycles have followed an approximately 100-kyr

32

periodicity (the “100-kyr world”) (Hays et al., 1976; Imbrie et al., 1992). The timing of the33

glacial–interglacial cycles is explained by the theory postulated by Milankovitch in the early34

20th century, which changes in boreal summer insolation are responsible for changes in the

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volume of boreal glacial ice sheets. However, the transition from the 41-kyr to the 100-kyr

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world, termed the Mid-Pleistocene Climate Transition (MPT) (e.g., Pisias and Moore, 1981),

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appears to bear no relation to orbital forcing. Furthermore, there is little agreement as to when38

the MPT occurred.

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Clark et al. (2006) suggested that the constructed LR04 ‘‘stacked’’ benthic 18O record

40(Lisiecki and Raymo, 2005) (Fig. 1) shows the MPT beginning at ~1.25 Ma with a gradual41

increase in global ice volume and decrease in deep-water temperature. This hypothesis was

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supported by Sosdian and Rosenthal (2009), who reconstructed early Pleistocene eustatic43

sea–level (and hence global ice volume) changes with orbital-scale resolution using changes in44

the 18O of seawater based on analyses of 18O and Mg/Ca ratios of the epifaunal benthic

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foraminifera Cibicidoides wuellerstorfi and Oridorsalis umbonatus from North Atlantic Deep46

Sea Drilling Project (DSDP) site 607 (Fig. 1). However, Yu and Broecker (2010) questioned the47

result of Sosdian and Rosenthal (2009), because of the confounding influence of carbonate ion

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saturation on epifaunal benthic foraminiferal Mg/Ca ratios.
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49

Recently, Elderfield et al. (2012) provided a record of eustatic sea–level for the past 1.5

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Myr from Ocean Drilling Program (ODP) 181, site 1123, off New Zealand (Fig. 1). The record,

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based on 18O and Mg/Ca ratios of the shallow-infaunal benthic foraminifera Uvigerina spp.

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(the shells of which are barely affected by carbonate-ion saturation), suggests that the MPT was

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initiated by an abrupt increase in Antarctic ice volume at MIS 22 (~0.9 Ma). According to

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Elderfield et al. (2012), the uncertainty of the sea–level changes are ± 20 m. More recently,

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Rohling et al. (2014) reconstructed eustatic sea–level changes over the past 5.3 Myr using

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eastern Mediterranean planktonic foraminiferal 18O records. The authors reported a strong

57

agreement between their reconstruction and the sea–level estimates of Elderfield et al. (2012),

58

although the 95% probability interval of the Rohling et al. (2014) is ±6.3 m.

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Prior to these studies, Bintanja et al. (2005) reconstructed the sea–level curve during the60

past 1.1 Myr from an ice-sheet–ocean-temperature model and the LR04 ‘‘stacked’’ benthic 18O

61

record (Lisiecki and Raymo, 2005), although they did not discuss the pattern of initiation of the62

MPT. According to Bintanja et al. (2005), the uncertainty of their sea–level changes is ± 10 m.

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Consequently, independent constraints on sea–level during the early Pleistocene glacial

64

periods are required to help determine whether the initiation of the MPT was a gradual or an65

abrupt process. As the upper limits of U/Th dating and polar ice cores records are 0.5 Ma and66

0.8 Ma, respectively, shelfal and nearshore sedimentary records provide useful constraints on

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eustatic sea–level changes during the early Pleistocene.

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Existing shallow-water sediment records include those of the Wanganui Basin in New

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Zealand (Beu and Edwards, 1984; Abbott and Carter, 1994; Carter and Naish, 1998; Kondo et70al., 1998), the Croton Basin in southern Italy (Rio et al., 1996; Massari et al., 1999, 2011), the71

Merced Formation in northern California (Carter et al., 2002), the Seoguipo Formation in72

southern Korea (Kim et al., 2010), and collisional marine foreland basin of southern Taiwan
3

73(Chen et al., 2001) (Fig. 1). However, because these marine basins are too deep to detect small

74

fluctuations in sea–level, none of these previous studies identified depositional sequences that75

correspond to MIS 22–24 in their lithostratigraphic schemes. In contrast, the depositional76

sequence in the Omma Formation, Central Japan, is firmly correlated with MIS 55–21

77(Kitamura and Kawagoe, 2006) (Fig. 2) and thus encompasses this transitional period. To

78

investigate the process of initiation of the MPT, this study reconstructed eustatic sea–level

79

changes in glacial period, as a proxy for global ice volume, based on a reexamination of80

lithofacies and fossil occurrences from the Omma Formation.

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2. Geological setting of depositional sequences in the Omma Formation

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The Omma Formation is exposed around Kanazawa City, along the west coast of Central

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Japan (Fig. 1). The formation is up to 220 m thick in the type section along the Saikawa River at

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Okuwa. The Omma Formation overlies the middle Miocene Saikawa Formation (Ogasawara,

861977) and is in turn overlain unconformably by the Utatsuyama Formation (Ichihara et al.,871950) (Fig. 2). The marine Saikawa Formation is mainly composed of massive siltstone

88(Ogasawara, 1977). The Utatsuyama Formation is about 100 m thick and comprises fan-delta

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deposits of alternating beds of mudstone, coarse-grained sandstone, and conglomerate (Nirei,

901969).

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Biostratigraphic and magnetostratigraphic data (Takayama et al., 1988; Ohmura et al.,

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1989; Sato and Takayama, 1992; Kitamura et al., 1994) indicate that the basement of the Omma

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Formation at the type section is located between the first occurrence (FO) of Gephyrocapsa

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oceanica (1.664 ± 0.025 Ma; Berger et al., 1994) and the FO of Gephyrocapsa (large) (1.515 ±95

0.025 Ma; Berger et al., 1994) (Fig. 2). These data show that the top of the formation is located96

below the Brunhes–Matuyama magnetic polarity reversal (Fig. 2).
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The Omma Formation has been divided into lower, middle, and upper parts (Fig. 2)

98(Kitamura et al., 1994, 2001). The lower and middle parts consist of 14 depositional sequences

99(Fig. 2, L1–L3, M1–11) that include the following architectural elements, in ascending

100

stratigraphic order: (1) a basal sequence boundary that is superposed on the ravinement surface;

101(2) a transgressive systems tracts (TST) (2–5 m thick) consisting of a basal shell bed of 0.3 m

102

thick (a condensed onlap shell bed) overlain by fine- to very-fine-grained sandstone; (3) a103

maximum flooding horizon coinciding with the horizon with the highest concentration of104

sand-size carbonate grains; (4) a highstand systems tracts (HST) (2–3 m thick) consisting of

105

fine-grained sandstone and sandy siltstone; and (5) a regressive systems tracts (RST) (<1 m106thick) comprising fine-grained sandstone with a coarsening-upward trend (Kitamura et al.,1072000). The upper part of the formation consists of five depositional sequences (Fig. 2, U1–U5)108

associated with back-marsh to inner-shelf environments (Kitamura and Kawagoe, 2006). Erect

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stumps and tracks of elephants and deers have been found from back-marsh deposits (Kitamura

110

and Kawagoe, 2006). These parts of the Omma Formation show no progressive shift in litho-111

and biofacies toward deeper or shallower deposits.

112

Except for four depositional sequences in the upper part, during the deposition of each113

sequence, the molluscan fauna changed from cold-water, upper-sublittoral species to114

warm-water, lower-sublittoral species, followed by a return to cold-water, upper-sublittoral

115

species (Kitamura et al., 1994). The term “cold-water species” is applied to fauna living in the116

area north of southern Hokkaido and/or in water deeper than 150–160 m off Sanin and117

Hokuriku. These species are present in the Pacific coast area north of the Boso Peninsula at

118

35°N, where the warm Kuroshio Current diverges away from the Japanese Islands. The term

119

“warm-water species” is applied to fauna living in the area south of southern Hokkaido, and120

living at the area shallower than 150–160 m depth off Sanin and Hokuriku. The area off Sanin5

121

and Hokuriku is strongly influenced by the warm Tsushima Current, which is a branch of the122warm Kuroshio Current. These species are present south of 35°N on the Pacific coast of Japan.

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The cyclic changes in molluscan content in these depositional sequences indicate that ocean

124

conditions and water depth fluctuated concurrently. Specifically, increased water depths

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correspond to periods of warmer marine conditions associated with the inflow of the warm

126

Tsushima Current. Thus, the depositional sequences of the Omma Formation are correlated with127

obliquity-driven glacio-eustatic changes, with a periodicity of 41 kyr (Kitamura et al., 1994).

128

Kitamura and Kimoto (2006) correlated 19 depositional sequences in the Omma Formation with

129

oxygen isotope stages 56 to 21.3 using a combination of sequence stratigraphic, biostratigraphic,

130

and magneto stratigraphic data (Fig. 2). As noted above, the Omma Formation contains only131

shallow-marine facies that can be perfectly correlated between oxygen isotope stages and132

depositional sequences during MIS 55–21.

133
134

3. Sea–level reconstruction

135

Water depth in sedimentary basins can be influenced by several factors, including

136

compaction, basement subsidence, sediment supply, hydro-isostatic effects, and eustatic137

sea–level changes. The basement rock on which the Omma Formation was deposited (i.e., the138

Saikawa Formation sediments) was consolidated prior to deposition of the Omma Formation,

139

since the lower unconformity is penetrated by marine rock-boring bivalves (Kitamura, 1997).

140

The sediments of the Omma Formation are unconsolidated and consist primarily of very-fine- to141fine-grained sandstone, with a burial depth equal to the total formation thickness (210 m). From142

the porosity–depth relationship (Sclater and Christie, 1980; Ramm and Bjiørlykke, 1994),

143

sandstone porosity within the formation decreases by only 5% from the surface to a burial depth

144

of 250 m, indicating that the water–depth curve is only weakly affected by compaction.
6

145With the exception of depositional sequence U4 (DS U4), which corresponds to MIS

146

22–21.4 (Kitamura and Kawagoe, 2006), the sequence boundaries for the depositional147

sequences of the Omma Formation coincide with ravinement surfaces formed by shoreface

148

erosion during marine transgression (Bruun, 1962). As the depth of shoreface erosion is149

generally less than 40 m (e.g., Saito, 1989), water depths at sequence boundaries probably

150

remained at least than 40 m for the 800-kyr period between MIS 56 and 21.3, indicating that151

basin subsidence kept pace with sediment supply.

152

Being located far from the former continental ice-sheets, the marine basin of the Omma

153

Formation was influenced by hydro-isostasy alone, in which rising (falling) sea–level causes

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basin subsidence (uplift) due to increasing (decreasing) water load. Consequently, relative

155

sea–level during glacial periods can be compared directly, allowing water-depth changes

156

represented by the Omma Formation to be used as a proxy for eustatic sea–level changes.

157

Water depth change, as recorded by the Omma Formation, has been reconstructed by

158

analyses of lithofacies and fossil occurrences (Kitamura, 1991; Kitamura et al., 1994; Kitamura

159

and Kawagoe, 2006; Kitamura and Kimoto, 2006) (Fig. 2). Non-marine sediments, comprising160

back-marsh deposits found in the lowest levels of DS U4 (MIS 22), suggest that sea–level

161

during MIS 22 was lower than at any other time between MIS 56 and 21 (Kitamura and162

Kawagoe, 2006). Kitamura and Kimoto (2004) and Kitamura and Kawagoe (2006) described163

well-sorted, fine-sand units that contain parallel laminations or hummocky cross-stratification in164

the upper levels of depositional sequences 8 and U1, corresponding to MIS 36–34 and 28–26,

165

respectively (Fig. 2). These sediments do not contain fossils such as molluscs and foraminifera,

166

due to the dissolution of calcareous shell materials. Sedimentary structures indicate that the167

sediments were deposited in a lower shoreface environment. Because the fair-weather wave

168

base lies at depths of about 5–15 m (Walker and Plint, 1992), the depth of the lower shoreface is7169inferred to be 5–15 m. The reexamination herein indicates that, with the exception of MIS 22,

170

sea–level was lower during MIS 34 (~1.13 Ma) and 26 (~0.96 Ma) than during any other glacial

171

stage between MIS 56 and 21.

172
173

4. Discussion and conclusion

174

Based on a reexamination of lithofacies and fossil occurrences from the shallow-marine

175

Omma Formation, sea–level was lower during MIS 22 than at any other time between MIS 56

176

and 21. This finding is in close agreement with the findings of Bintanja et al. (2005), Elderfield

177

et al. (2012), and Rohling et al. (2014), and supports the argument for a significant increase in178

global ice volume at MIS 22, referred to by Clark et al. (2006) as the 900-ka event.

179

This reexamination also indicates that sea–level was lower during MIS 34 and 26 than180

during other glacial stages (except for MIS 22) of the same time period, in agreement with the181

conclusion of Bintanja et al. (2005). According to Elderfield et al. (2012) and Rohling et al.

182(2014), sea–level in MIS 34 and 26 was not significantly lower than in the other glacial periods

183

between MIS 56 and 21. This is inconsistent with the new evidence of the present study.

184

The differences in sea–level between MIS 22 and the two glacial stages MIS 34 and 26 are

185

estimated to have been 20 m (Bintanja et al., 2005), 40 m (Rohling et al., 2014) or 50 m

186

(Elderfield et al., 2012) (Fig. 2). As noted above, the water depth is thought to have been187

between 5 and 15 m during MIS 34 and 26. Conversely, the sea–level during MIS 22 is

188

uncertain, because the study area emerged and was a back-marsh environment during this period

189(Fig. 2). When the difference in sea–level of 50 m (20 m) is applied, the elevation of the study

190

area is estimated to have been 35–45 m (5–15 m) during MIS 22. As back-marsh commonly

191

develops in low coastal plains, it is likely that the difference in sea–level was 20 m, as

192

determined by Bintanja et al. (2005), rather than the 40–50 m estimated by Elderfield et al.
8193(2012) and Rohling et al. (2014).

194

As noted above, the uncertainties in estimates of sea–level from geochemical data range

195

from ±6.3 m to ±20 m (Bintanja et al., 2005; Elderfield et al., 2012; Rohling et al., 2014). In

196

contrast, as hydrographic conditions change drastically from shoreface to offshore, lithofacies

197

are highly sensitive to changes in sea–level during sea–level low stands. It is therefore likely

198

that the inconsistency between this study and the geochemical studies of Elderfield et al. (2012)199

and Rohling et al. (2014) reflect uncertainties in the estimates from geochemical data.

200

This means that the early Pleistocene sea–level curve of Bintanja et al. (2005) is more

201

suitable than those of Elderfield et al. (2012) and Rohling et al. (2014), and that the initiation of202

the MPT was a gradual (e.g., Clark et al., 2006), rather than abrupt, process (Elderfield et al.,203

2012).

204
205

Acknowledgments

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I thank three anonymous reviewers, whose comments and suggestions improved the original207

manuscript.

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332
333
334
335
336
337
338
339
340

Figure 1 Locations of deep-sea cores (diamond symbols) used in the LR04 ‘‘stacked’’ benthic

341

18O record (Lisiecki and Raymo, 2005), location of ODP 1123 (Elderfield et al., 2012), and342

locations of early Pleistocene shelfal and nearshore sedimentary records (circles).

343
344
345
346
347
348
349
350
351
352
353
354
355

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100

water regime

inner shelfback marsh
upper shoreface
lower shoreface
transition

Utatsuyama Formation
Fan-delta deposits

sea-level changes
relative to the presentday sea-level (m)Non body fossils

U5molluscan fossils
Megangulus zyonoensis
Mercenaria stimpsoni
Acila insignis
Yoldia notabilis
Jupiteria confusa
Oliva mustelina

Magnestostratigraphy

0
5
15
30

water depth
(m)

10 m

Bintanja et al. Elderfield et al.
(2012)(2005)-150

0 50 -150

0.8● ● ● ● ● ●21
22younger than LAD
Reticulofenestra asanoi
0.899±0.025Ma
Non body fossils

Megangulus zyonoensis
Felaniella usta
Acila insignis
Clinocardium fastosum
Turritella saishuensis
Yoldia notabilis
Cycladicama cumingi
Stellaria exutum
Paphia schnelliana
Jupiteria gordonis

U3

U2

● ● ●

28
30

middle partlower part

34

1.1

33
35

M10
M9

36

M8

40

M5
M4
M3

42

M2
M1

44

L-3

46

L-2

48

L-1

Age
Ma
1.2

37

38

M7
M6

Miocene Saikawa
Formation

1.0

29
31

32

M11

FAD Gephyrocapsa(large) 1.515±0.025Ma

27● ● ● ● ●U1

LAD Helicosphaera
sellii 1.451±0.025Ma

25

26

Top of Jaramillo
0.992±0.01Ma

LAD Gephyrocapsa(large)1.243±0.03Ma

0.9

23

24
molluscan fossils

upper partOmma Formation

Jaramillo

Matuyama

U4

Base of Jaramillo
1.073±0.01Ma

0 50

39
41

1.3

43
45

1.4

47
49

1.5

S VF F M C VC G

sandshell bed
shell fossils
conglomerates
tuff

parallel lamination
cross lamination
hummocky cross stratification
sequence boundary356●

cold-water species
cold-water and extinct species
warm-water species
cold water
warm water

357

Figure 2 Columnar section of the Omma Formation at its type section along the Saikawa River358

at Okuwa. Comparison of water-depth changes reconstructed from the Omma Formation and359

eustaic sea–level changes inferred from geochemical signals of benthic foraminifera in deep-sea

360

sediments (Bintanja et al., 2005; Elderfield et al., 2012).

18