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Progress in Earth and Planetary Science

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  • Research article
  • Open Access

A record of the upper Olduvai geomagnetic polarity transition from a sediment core in southern Yokohama City, Pacific side of central Japan

  • Chie Kusu1Email author,
  • Makoto Okada2,
  • Atsushi Nozaki3,
  • Ryuichi Majima4 and
  • Hideki Wada5
Progress in Earth and Planetary Science20163:26

https://doi.org/10.1186/s40645-016-0104-7

©  The Author(s). 2016

Received: 7 March 2016

Accepted: 24 August 2016

Published: 23 September 2016

Abstract

A detailed paleomagnetic record of the upper Olduvai polarity transition was obtained from a 106.72 m-long sediment core drilled in southern Yokohama City, located on the northern Miura Peninsula, on the Pacific side of central Japan. The core spans the upper part of the Nojima Formation and the lowermost part of the Ofuna Formation, both of which correspond to the middle Kazusa Group (Lower Pleistocene forearc basin fill). The record was reconstructed using discrete specimens taken throughout mudstone and/or sandy mudstone sequences in the Nojima Formation. In this record, the virtual geomagnetic pole (VGP) fluctuation accompanying the polarity transition was determined to occur between depths of 66.99 and 63.60 m. These depths have been dated at 1784.4 and 1779.9 ka, respectively, and the duration of the polarity transition is estimated to be 4.5 kyr using an age model based on a δ18O record from that core. The VGP paths during the transition do not appear to show any preferred longitudinal bands. However, the VGP positions cluster in five areas: (A) eastern Asia near Japan, (B) the Middle East, (C) eastern North America (North Atlantic), (D) off southern Australasia, and (E) the southern South Atlantic off South Africa. The primary locations of the observed VGP clusters coincide with the areas on the Earth’s surface that possess a strong downward flux of the vertical component of the present geomagnetic non-axial dipole field. The relative paleointensity rapidly decreased approximately 1 kyr before the beginning of the polarity transition and gradually recovered to its initial level in 12 kyr.

Keywords

GeomagnetismPaleomagnetismGeomagnetic polarity transitionOlduvai subchronVGP pathVGP clusterReversalRelative paleointensitySediment core

Introduction

Studies of geomagnetic reversals, which are among the most conspicuous phenomena associated with the Earth’s magnetic field, provide valuable information about geodynamo processes. Previous paleomagnetic investigations have documented the details of geomagnetic field behavior during geomagnetic polarity transitions. The virtual geomagnetic poles (VGPs) during polarity transitions, as inferred from sediments, tend to pass within two preferred longitudinal bands over the Americas and antipodally over eastern Asia and through Australia (Tric et al. 1991; Clement 1991; Laj et al. 1991). Conversely, late Cenozoic volcanic records from intervals spanning polarity transitions suggest that VGPs were clustered mainly within two regions near the southern portions of South America and Western Australia (Hoffman 1991, 1992). These regions lie within the two preferred longitudinal bands identified in the sedimentary records. However, the VGPs inferred from the volcanic records appear not to have moved continuously but to have “jumped” between the regions. The two preferred bands coincide with the centers of the radial magnetic flux of the present geomagnetic field on the Earth’s surface when the axial dipole component is removed (Hoffman 1992). Laj et al. (1991) reported that the two preferred bands coincide well with regions of fast seismic wave propagation (regions of cooler temperature) observed in the lowest mantle. Therefore, they suggested that the transitional field was constrained by temperature patterns at the core-mantle boundary, and its configuration might have persisted over the past 10 Myr. Recently, several high-resolution sedimentary records from deep-sea cores have shown that transitional VGPs still passed within the preferred longitudinal bands but also stayed within regions similar to the VGP clusters observed in volcanic records (e.g., Mazaud and Channell 1999; Ohno et al. 2008; Mazaud et al. 2009).

The detailed processes relating to geomagnetic polarity reversals remain unclear, although it is generally accepted that field intensities during polarity transitions decline to ~10 % of their usual values, which is close to the magnitude of the non-dipole component of the current geomagnetic field (Merrill and McFadden 1999). Absolute paleomagnetic intensities are typically only recovered from the thermal remanent magnetization of volcanic rocks, but this does not yield a continuous record. Tauxe (1993) proposed a method to deduce relative paleointensities using sedimentary records, and continuous relative paleointensity records have thus been obtained from deep-sea sediments. In recent years, studies of relative paleointensity have progressed (e.g., Tauxe and Yamazaki 2007), and global stacks of relative paleointensity variations spanning the last 1 to 3 Myr have been generated (e.g., Valet et al. 2005; Yamazaki and Oda 2005; Channell et al. 2009). Those records, which have focused on reconstructing long-term variations in geomagnetic dipole moment, have been successful in revealing the broad features of geomagnetic dipole fluctuations. However, the details of paleointensity variations during polarity transitions are poorly documented because the accumulation rates of deep-sea sediments used to generate the stacked records are quite low (approximately a few cm/kyr). Because the average duration of polarity transitions has been estimated at 7 kyr, based on available sediment records spanning the four most recent transitions (Clement 2004), records with resolutions higher than 1 kyr are needed to observe detailed variations in paleointensity as well as direction during polarity transitions.

It is generally considered that the magnetization of sediments is stabilized through post-depositional remanent magnetization (PDRM) processes (Irving and Major 1964; Kent 1973). The fixing of magnetic particles in a sedimentary column is thought to occur gradually within a zone due to sediment compaction and dewatering, which generates an offset between the depth where the magnetization is fixed and the sediment surface (e.g., Verosub 1977; Hyodo 1984). This offset is referred to as the “lock-in depth” of PDRM and has been the focus of much debate for a long time (e.g., Okada and Niitsuma 1989; deMenocal et al. 1990; Tauxe et al. 1996, 2006; Channell and Guyodo 2004). Suganuma et al. (2010, 2011) convincingly showed that magnetization in a sedimentary column takes place within a zone just below the base of the bioturbation zone. Based on direct comparisons of 10Be flux and paleomagnetic records through the Matuyama-Brunhes transition in deep-sea cores, they deduced by means of a Gaussian lock-in function that the thickness of the zone of magnetization is approximately 17 cm, regardless of the sedimentation rate. This indicates that sedimentary sequences with accumulation rates that exceed 17 cm/kyr are enough to reconstruct records of past geomagnetic field fluctuations with a time resolution of 1 kyr. Although the PDRM process is uncertain and still under debate, it is possible to minimize the length of delay and smoothing width in remanence acquisition due to the PDRM process by using a sediment sequence that has as high a sedimentation rate as possible.

In this paper, we present an ultra-high-resolution paleomagnetic record of the upper Olduvai polarity transition from a sediment core comprising a homogeneous siltstone whose sedimentation rate exceeded 70 cm/kyr. We also present the VGP paths and relative paleointensity variations during the transition.

Methods/Experimental

Geology of the core site

The core used in this study (Core M; 35.354 6° N, 139.606 3° E) was recovered from the middle part of the lower Pleistocene Kazusa Group in September 2010 using a rotary drilling method and an 86 mm diameter bit. The core is 106.72 m long and composed of sediments with almost 100 % recovery (Fig. 1). The Kazusa Group, a forearc basin fill, is exposed in the northern Miura Peninsula, central Japan. Core M is composed of the upper part of the Nojima Formation (alternations of mudstones and sandy mudstones below a core depth of 20.5 m) and overlies the lowermost part of the Ofuna Formation (massive mudstones from the top of the core to 20.5 m) in the middle part of the Kazusa Group.
Figure 1
Fig. 1

Location maps of the study site. a Tectonic plates under the Japanese Islands (modified from Fig. 1 of Seno and Takano 1989). b Distribution of the Kazusa Group in the southern Kanto region (compiled from Mitsunashi and Suda 1980; Mitsunashi et al. 1980; Unozawa et al. 1983; Eto 1986; and Sakamoto et al. 1987). c Locality map showing the core sites (solid stars) plotted on a 1:25,000 scale topographic map published by the Geospatial Information Authority of Japan

The depths at which the sediments of the Nojima and Ofuna Formations were deposited are estimated to 400 to 500 m in the lower horizon of the Nojima Formation (Utsunomiya and Majima 2012) and 200 to 300 m in the upper horizon of the Ofuna Formation (Tate and Majima 1998; Kitazaki and Majima 2003). The strike and dip of the bedding around the Core M site are N63° W and 12° NE, respectively, based on outcrop observations of the NOT-1 tuff bed (Nozaki et al. 2014; the YH02 tuff bed of Takahashi et al. 2005 and Kusu et al. 2014), which is intercalated with sediments between 14.46 and 14.41 m depth in Core M. The detailed lithologies of Core M will be described in a separate paper.

In May 2004, Core I (105.00 m in total length; 35.357 5° N, 139.589 5° E) was drilled and recovered approximately 1.6 km west of the Core M drill site (Fig. 1c), and Kusu et al. (2014) published a paleomagnetic study of the upper Olduvai polarity transition and the vicinity in this core. In Core I, the upper Olduvai polarity transition was determined to lie between 86.77 and 84.64 m depth in the Nojima Formation on the basis of paleomagnetic inclination data at a horizon below the NOT-1 tuff bed that is intercalated between 27.03 and 26.83 m core depth. Core I was shown to provide a detailed record of the paleomagnetic behavior during the upper Olduvai transition; however, the record consists only of in situ paleomagnetic inclinations. Because Core I was stored as horizontally unoriented pieces of sections, information on the core’s orientation and thus the declinations were lost. Conversely, when the newly recovered Core M was drilled, its orientation was recorded and described, thus enabling the reconstruction of the paleomagnetic vectors through the upper Olduvai polarity transition.

Paleomagnetic and rock-magnetic measurements

Core M was not initially oriented in a horizontal plane. Thus, the absolute declination values could not be directly obtained; however, the relative declination has been preserved within a continuously recovered interval without any core break. Because the core barrel used for the drilling was 2 m long, a 2 m-long continuous core segment was recovered at each stroke in the drilling operation. To determine the relative declinations for as long a continuous interval as possible, for each core segment, we confirmed whether its end surface could be connected to the next core segment and carefully checked the connection between the two segments without any horizontal rotation to ensure the continuity of the core. Additionally, when the dip of a bedding plane was identified from the sedimentary structures in the core, we drew a vertical orientation line at the position of the maximum dip direction on the core surface. Even when the dip could not be identified, we drew a vertical line on the core surface to ensure the directional continuity of the core segments.

For paleomagnetic and rock-magnetic measurements, 25.4 mm diameter mini-cores were collected horizontally from the side surface of the core along the vertical line marked on the core surface. The mini-cores were taken using a core-picker every 1 m over the entire core interval, every 10 cm between 85.00 and 75.20 m and between 60.00 and 50.00 m and every 5 cm between 75.20 and 69.84 m. Each mini-core was cut into approximately 2 cm-long sections. Between depths of 69.84 and 60.00 m, the core was quartered vertically using a rock cutter, and one quarter of the core was cut continuously into approximately 2 cm-long cubic blocks from the center portion of the core. In total, specimens were obtained from 766 horizons in Core M.

Before demagnetization experiments were carried out, low-field magnetic susceptibility (volumetric) and anisotropy of magnetic susceptibility (AMS) measurements were performed on all specimens using a KLY-3 Kappabridge susceptibility meter (AGICO, Czech Republic). The natural remanent magnetization (NRM) was measured using an RF-SQUID Cryogenic magnetometer model 750R (2G Enterprises, USA). Progressive alternating field demagnetization (AFD) with a static 3-axis was performed in 5 mT increments up to 60 mT using a separate AF demagnetizer (DEM-8601C, Natsuhara-Giken, Japan), and progressive thermal demagnetization (THD) was performed in 50 °C increments up to 600 °C in air using a thermal demagnetizer (TD-48, ASC Scientific, USA). The measurements were performed at Ibaraki University, Japan. To deduce relative paleointensity, progressive AFD must be employed as the main demagnetization technique because after the demagnetization of the NRMs, the same specimen is repeatedly used for the acquisition of artificial remanent magnetizations to estimate relative paleointensities. The results of the progressive AFD and THD were displayed using an orthogonal vector diagram (Zijderveld 1967). Characteristic remanent magnetization (ChRM) directions were determined by principal component analysis (Kirschvink 1980) using more than five vector endpoints.

Relative paleointensity was estimated by normalizing for the magnetic grain content using anhysteretic remanent magnetization (ARM). Subsequent to the AFD of the NRM, the ARM was imparted to each specimen using a 30 μT DC field in an alternating field that decreased from a peak of 60 mT. After the ARM acquisition, AFD was then performed at 20, 30, 40, 50, and 60 mT to calculate the paleointensity proxy (NRM/ARM). Thermomagnetic analyses were performed on two samples (from depths of 70.01 and 58.58 m, near the polarity transition) using a Curie Balance NMB-89 (Natsuhara-Giken, Japan). The samples were heated to 700 °C and cooled in air and in a vacuum with an applied field of 0.3 T. Magnetic hysteresis was measured on eight samples at 2–3 m intervals between 70 and 50 m in Core M using a MicroMag 2900 alternating-force gradient magnetometer (Princeton Measurements, USA) to peak fields of ±1 T. The thermomagnetic experiments and hysteresis measurements were performed at the Kochi Core Center, Japan.

Oxygen isotope measurements

Stable oxygen and carbon isotopic ratios were measured to establish an age model based on correlation with the LR04 curve, which is a stacked δ18O curve that uses oxygen isotope ratios measured in benthic foraminifers from 57 globally distributed deep-sea cores (Lisiecki and Raymo 2005). The isotopic measurements were performed on tests of the planktonic foraminifer Globorotalia inflata. Because G. inflata precipitates its test at approximately 300–500 m below the sea surface where water temperatures are relatively constant, the G. inflata δ18O profile tends to be parallel to that of benthic foraminifers (Oba et al. 2006). To extract the G. inflata tests, a total of 104 sediment samples were collected from the core at approximately 1 m intervals. Each dried sediment sample (approximately 80 g) was crushed and disaggregated using a sodium sulfate solution. From each sediment sample, we picked tests of G. inflata that were unbroken, untarnished, “glassy,” and larger than 250 μm in size. Each test was immersed in ethanol, broken using a needle, and then ultrasonically cleaned to remove filling impurities.

The isotopic measurements were performed at Shizuoka University, Japan, using the procedure of Wada et al. (1984). The 20–30 individual tests used for each measurement were reacted in saturated phosphoric acid at 60 °C, and the evolved CO2 gas was analyzed with a MAT-250 mass spectrometer. The obtained value was converted into a value against a PDB standard employing the US National Bureau of Standards NBS-20 (Vienna PDB). The standard deviations of the in-house standard (total of 19 measurements) were 0.04 ‰ for δ18O and 0.02 ‰ for δ13C, which represented the precisions of the isotopic measurements.

Results

Rock magnetism

The magnetic fabrics were analyzed via AMS. The directions of the maximum susceptibility (Kmax), intermediate susceptibility (Kint), and minimum susceptibility axes (Kmin) of the AMS ellipsoids were plotted on lower hemisphere equal area projections. Figure 2 shows the AMS data for depths of 67.85–66.11 m within the same core segment for which the directional continuity was confirmed. The data are not corrected for tilt. The directions of Kmax and Kint indicate a girdle distribution perpendicular to Kmin, which indicates a flattening fabric, and all results of the AMS analysis showed similar flattening fabrics (foliations). The mean inclination of Kmin between depths of 67.85 and 66.11 m was 74°, and the dip of the foliation plane formed by Kmax and Kint is therefore estimated to be 16°. The estimated dip angle of the bedding plane at the core site is approximately 12°, which suggests that the foliation plane reasonably coincides with the bedding plane. In the four sections of the core, bedding planes comprised of intercalated thin sandstone beds were directly examined. In each segment, the difference between the observed maximum dip direction and the direction of the foliation plane was less than 10°. Accordingly, we treat the foliation plane, as deduced from the AMS data, as a bedding plane at the site.
Figure 2
Fig. 2

Lower hemisphere equal area projections of the anisotropy of magnetic susceptibility (AMS). AMS data from the 67.85–66.11 m depth interval in Core M. The squares, triangles, and circles represent the directions of the maximum, intermediate, and minimum susceptibilities, respectively

Thermomagnetic curves from samples at depths of 70.01 and 58.58 m are shown in Fig. 3. Both curves show an identical Curie temperature of approximately 570 °C, which indicates that magnetite (or Ti-poor titanomagnetite) is the predominant magnetic carrier. The heating curves in air as well as in vacuum show a slight increase above 450 °C, but the cooling curves do not show such inflections. When iron sulfide minerals such as pyrite are present in sediments, they are changed into magnetite and maghemite and, finally, into hematite through heating by oxidation that occurs primarily above 450 °C, which produces a heating curve that at first shows an increase in the magnetization and then a decrease (Passier et al. 2001). Therefore, although the thermomagnetic curves suggest a minor amount of iron sulfide minerals, (titano)magnetite is considered to be the main magnetic mineral in the samples.
Figure 3
Fig. 3

Strong-field thermomagnetic behaviors. a sample from the depth of 58.58 m, b sample from the depth of 70.01 m in Core M. Air.: Under an air condition. Vac.: Under a vacuum condition

A plot of anhysteretic remanent magnetization susceptibility (κARM) versus magnetic susceptibility (κ) for Core M and Core I is shown in Fig. 4a. According to the grain size estimates of King et al. (1983), the mean grain sizes of the magnetic grains in the specimens from Core M and Core I should primarily range from 1 to 5 μm. Figure 4b shows a Day plot (Day et al. 1977) using selected specimens from Core M and Core I. Both plots indicate that pseudo-single domain (PSD) or mixture of single domain (SD) and multidomain (MD) magnetite grains are the predominant magnetic carriers. The Core I data are from Kusu et al. (2014).
Figure 4
Fig. 4

Rock-magnetic results. a Anhysteretic remanent magnetization susceptibility (κARM) plotted against volume magnetic susceptibility (κ). b The Day plot used selected specimens from Core M and Core I (Day et al. 1977). Mrs saturation remanent magnetization, Ms saturation magnetization, Hcr remanent coercivity, Hc coercivity. SD, PSD and MD indicate single domain, pseudo-single domain, and multidomain fields, respectively, with the limits given by Dunlop (2002)

A plot of the magnetic susceptibility (κ) and anhysteretic remanent magnetization susceptibility (κARM) on a logarithmic horizontal scale is shown in Fig. 6g. The κ and κARM show similar gradual fluctuations, which range within an order of magnitude. Additionally, the fluctuation patterns of κ and κARM strongly resemble each other. Therefore, it is possible that the changes in the values of κ and κARM are related to changes in the magnetic mineral content in the specimen, and the changes in the magnetic mineral content with depth are relatively small.

Remanent magnetization

Typical orthogonal vector diagrams for NRMs during progressive AFD and THD are shown in Fig. 5. Most of the results show that the secondary magnetization components were removed by AF demagnetization at a peak field strength of 15 mT. Large fluctuations in the inclinations can be recognized between depths of 67.03 and 64.38 m (Fig. 5e), and steep upward inclinations are observed between depths of 64.36 and 63.20 m (Fig. 5d). In the lower part, below a depth of 67.05 m, the inclinations are positive (Fig. 5b, f), and the maximum angular dispersion (MAD) values for the ChRMs are generally less than 5°. However, values slightly over 10° are seen in a few specimens just below a depth of 67.05 m (Fig. 6d). In the upper part, above a depth of 63.18 m, the inclinations are negative (Figs. 5a, c), and the MAD values for the ChRMs are generally less than 10°, approximately 80 % of which are less than 5° (Fig. 6d). ChRMs were not obtained from five specimens (from depths of 66.61, 66.59, 66.55, 66.49, and 65.81 m), which had very weak remanences (Fig. 5g) and/or MAD values exceeding 25°. From these inclinations, we infer that the Olduvai normal polarity subchronozone occupies the lower part below a depth of 67.05 m, and the post-Olduvai Matuyama reversed polarity chronozone occupies the upper part above a depth of 63.18 m.
Figure 5
Fig. 5

Examples of demagnetization diagrams. Typical results from the orthogonal vector diagrams for data obtained during progressive thermal demagnetizations (THD) and alternating-field demagnetizations (AFD) for samples from Core M. a, c Samples from the post-reversal zone. d, e Samples from the transitional zone. b, f Samples from the pre-reversal zone. g Sample from a section yielding weak remanence intensity. The solid and open symbols indicate the projection of the vector endpoints onto the horizontal and vertical planes, respectively

Figure 6
Fig. 6

Rock-magnetic and paleomagnetic results from a depth interval of 80–50 m in Core M. a Paleomagnetic declinations. b Section breaks. c Paleomagnetic inclinations. d Maximum angular dispersions (MAD) of principal component analysis. e VGP latitudes, in which the letters A–E are the VGP cluster locations see in Fig. 7a. f Relative paleointensities (NRM20–40 mT/ARM20–40 mT). g Volume magnetic susceptibilities (κ) and anhysteretic remanent magnetization susceptibilities (κARM). The vertical dotted lines are the expected declinations and inclinations for the two polarities at the core site from the geocentric axial dipole field

Orientation of the core

The core was not oriented, and absolute paleomagnetic declinations therefore could not be derived directly from the measured ChRMs. To obtain the declinations of the ChRMs, we reconstructed the core orientation via the following process. First, we calculated the dip azimuth of the foliation plane, which was deduced from the AMS in each core segment where the continuity of the cores was confirmed. We then adjusted the dip azimuth to the north direction. Using this process, we connected core segments between depths of 80 and 50 m. However, significantly different declinations of ChRMs were observed at three depth intervals: 79.78 to 76.59 m, 73.71 to 72.05 m, and 51.48 to 51.17 m (intervals are indicated by the gray bars in Fig. 6b). We then readjusted the azimuth directions at the three intervals to minimize the differences in the declinations between the three intervals and other depths. Next, the dip azimuth observed in the core was adjusted to the dip azimuth of the bedding plane observed around the core site (strike and dip: N63° W and 12° NE), and the tilt was then corrected. Finally, the mean declination of the normal polarity (below 67.05 m) was adjusted to 0°.

Paleomagnetic directional changes and relative paleointensity

The resultant declinations and inclinations of the ChRMs are shown in Fig. 6a, c. The reversed polarity declinations were generally aligned to 180°, and the inclinations after the tilting correction were aligned to ±53°, which would be expected given a geocentric axial dipole (GAD) at the core site. Large fluctuations in the declinations and inclinations were observed between depths of 67.03 and 64.38 m, and steep upward inclinations were observed at approximately 63.5 m depth. The VGP latitudes calculated from the ChRMs are shown in Fig. 6e. The VGP moves across the equator from the Northern Hemisphere to the Southern Hemisphere between 65.69 and 65.67 m depth. The VGP begins to move rapidly at a depth of 67.03 m, and the VGP settles at approximately 63.44 m after moving to the Southern Hemisphere.

The rock-magnetic measurement results indicate that the predominant carrier of magnetic remanences in the sediments is PSD, or a mixture of SD and MD-sized (titano)magnetite (Figs. 3 and 4), and the fluctuations in κ and κARM range within an order of magnitude (Fig. 6g). From these results, the sediment satisfies the criteria for relative paleointensity reconstruction proposed by Tauxe (1993). To obtain an appropriate proxy for relative paleointensity, we compared the NRM/ARM ratios from three different coercivity fractions, which were between 20 and 40 mT, 30 and 50 mT, and 30 and 40 mT. The variations of the three coercivity fractions were almost identical, and we therefore used the NRM/ARM ratio between 20 and 40 mT as a proxy for relative paleointensity in this study (Fig. 6f).

VGP and paleointensity

The VGP paths between 70.0 and 60.0 m are shown in Fig. 7a. In the upper Olduvai polarity transition, the VGP apparently did not move within a preferred longitudinal band but settled in several VGP cluster areas. The VGP can be observed in five areas (Fig. 7a): (A) eastern Asia near Japan, (B) the Middle East, (C) eastern North America (North Atlantic), (D) off southern Australasia, and (E) the southern South Atlantic off South Africa. The VGP apparently moved rapidly between the clusters.
Figure 7
Fig. 7

Paleomagnetic results in the vicinity of the upper Olduvai polarity reversal. Comparison between Core M and ODP Site 983. a VGP path. b VGP latitude. c Relative paleointensity (NRM20–40 mT/ARM20–40 mT) from Core M and Core I. d VGP path. e VGP latitude. f Relative paleointensity (NRM20 mT/ARM20 mT) from ODP Site 983. For Site 983, the VGP positions and age model are based on Channell et al. (2002), and the relative paleointensity is based on Mazaud and Channell (1999). The declination and inclination data from ODP Site 983 are plotted using the PANGEA database (Channell et al. 2013). The dotted circles in a and d and the shaded circles in b and e indicate VGP clusters, and the bold letters represent the temporal order. The blue shaded areas in a indicate that the vertical component of the NAD field at the Earth’s surface is less than −15 μT (Constable 2007). The stars indicate the core sites

In the relative paleointensity of the normal polarity, a maximum value is observed at 75.50 m depth (Fig. 6f) and it then decreases with the short-term oscillations. Before the onset of the paleomagnetic directional change, the relative paleointensity has a peak at 67.75 m depth (with a value approximately 56 % of the maximum value), and values then decrease at 67.15 m to approximately 11 % of the maximum value. Subsequently, the relative paleointensity slightly recovers up to 66.89 m. In this interval, the directional change begins at a depth of 67.03 m, and the VGP near the North Pole rapidly moves to cluster A (Fig. 6e). The relative paleointensity subsequently drops at 66.5 m depth. Between 66.53 and 65.81 m, the relative paleointensity is approximately 9 % of the maximum value. In the same interval, the VGP settles in cluster C. The relative paleointensity then remains low up to 64.56 m. In this interval, the VGP in the Northern Hemisphere moves into the Southern Hemisphere between depths of 65.69 and 65.67 m, and the VGP moves to cluster D at ~65.17 m. Afterwards, the relative paleointensity gradually increases. From 64.65 to 63.46 m, the VGP settles in cluster E, and then, the VGP moves to the area near the Antarctic.

During the polarity reversal, especially between depths of 67.03 and 63.46 m, the relative paleointensity drops to approximately 12 % of the maximum value. After the large movements in the VGP, the relative paleointensity remains low for a time and then gradually recovers. At 58.38 m depth, the relative paleointensity returns to the peak value observed at 67.75 m, just before the paleomagnetic directional change began, and at about 55 m, the relative paleointensity becomes comparable to the maximum value observed pre-reversal.

Age model

To construct an age model for the interval spanning the upper Olduvai polarity reversal (1.8–1.7 Ma) in Core M, we correlated our δ18O profile generated from tests of G. inflata in the core (Fig. 8g) with the LR04 global δ18O stack (Lisiecki and Raymo 2005; Fig. 8k) and determined the peaks of marine isotope stages (MIS) 64 and 63.
Figure 8
Fig. 8

Comparison between Cores I and M in the vicinity of the upper Olduvai polarity reversal. Lithofacies, volume magnetic susceptibility (κ), and oxygen isotope fluctuations based on measurements in Globorotalia inflata (δ18O) and in situ paleomagnetic inclinations are shown in a–d for Core I and e–h for Core M, respectively. i Correlation of the peaks of the magnetic susceptibility between the two cores, using the depths from Core M. j Correlation of in situ paleomagnetic inclinations between the two cores, using the same depth as i. k The LR04 benthic δ18O stack (Lisiecki and Raymo 2005). Abbreviations shown on the lithology columns are the following: m mudstone (mud content ≥70 %), sm sandy mudstone (70 % > mud content ≥ 45 %), fs fine-grained sandstone, ft fine tuff, ct coarse tuff, pt pumice-rich lapilli tuff. For the red lines, see the text

The peak observed at 75.00 m (1.61 ‰) is comparable to the peak of MIS 64 because it is just below the Olduvai termination, and the broad low at approximately 58.00–52.00 m (0.78–0.83 ‰) can be correlated to MIS 63 because it is just above the Olduvai termination. The maximum amplitude of δ18O in the core is 0.83 ‰. This is quite comparable to the amplitude of 0.71 ‰ observed in the LR04 stack between MIS 64 and MIS 63. We therefore assigned the peak at 75.00 m and the broad low at approximately 58.00–52.00 m to MIS 64 and to MIS 63, respectively.

It is, however, difficult to determine a particular position within the broad trough at approximately 58.00–52.00 m for correlation with the lowest point of MIS 63. To determine which position is suitable as the MIS 63 trough, we used Core I, which was drilled in the same study area (Fig. 1c). From Core I, the Olduvai termination boundary, magnetic susceptibility, and relative paleointensity were obtained, and a significant trough comparable to MIS 63 was recognized in the δ18O curve (Kusu et al. 2014). The relative paleointensity and δ18O curve for Core I were determined using the same methods as in this study for Core M. Figure 8 shows a comparison between the two cores and includes the horizons of MISs 64 and 63. The vertical axes are expressed as core depths. Because a coarse-grained tuff bed, a pumice-rich lapilli tuff bed, and a fine-grained tuff bed (A, B, C, respectively, in Fig. 8) are commonly recognized in both cores, these three tephra beds were used to correlate the two cores. In the interval between the coarse-grained tuff bed (A) and the fine-grained tuff bed (C), the fluctuating patterns in the two cores’ magnetic susceptibility records coincide well (Fig. 8i), as do the records of in situ paleomagnetic inclinations (Fig. 8j). We therefore used the magnetic susceptibility curves to more closely correlate the two cores. The red lines in Fig. 8 indicate the correlation between the two cores using representative tie horizons in the vicinities of MIS 64 and MIS 63.

From this correlation, we determined that the most likely position corresponding to the trough of MIS 63 in Core M is at a depth of 58.00 m. We assigned MIS 64 (1795.0 ka) to the peak at 75.00 m and MIS 63 (1772.5 ka) to the trough at 58.00 m, which serves as an age model for Core M. Based on the age model, the average sedimentation rate between MISs 64 and 63 is 73.9 cm/kyr.

Discussion

Our results from Cores M and I show that the fluctuating patterns in the records of magnetic susceptibility (Fig. 8i), in situ paleomagnetic inclinations (Fig. 8j), and relative paleointensities (Fig. 7c) coincide well with each other, and the ChRM inclinations of Core M are in close agreement with the value expected for a geocentric axial dipole (Fig. 6c). This supports the reliability of our results for the upper Olduvai polarity transition.

The VGP latitudes and relative paleointensities are plotted against age in Fig. 7b, c, respectively. If the polarity transition interval is defined as the interval when the VGPs cross the equator and are more than 45° apart from the geographical poles, the transitional interval for Core M lies between depths of 66.99 and 63.60 m. This interval is dated at between 1784.4 and 1779.9 ka, and the duration of the transition is thus estimated to be 4.5 kyr. The relative paleointensity evidently decreased rapidly approximately 1 kyr before the beginning of the polarity transition and gradually recovered to its original value after almost 12 kyr.

The VGP cluster over the southern South Atlantic has a similar location to VGP clusters seen in lava-derived records of polarity transitions spanning the last 10 Myr (Hoffman 1991, 1992). Previous studies suggest that the VGP cluster areas during polarity transitions are associated with the vertical flux patches in the present non-axial dipole (NAD) field, although the VGP records from different polarity transitions merge (e.g., Channell et al. 2003, 2004; Hoffman et al. 2008). In Fig. 7a, the blue-shaded areas indicate the downward flux patches of the NAD, averaged for the last 400 years, where the vertical component is less than −15 μT (Constable 2007). This figure shows that our results of the VGPs from the upper Olduvai polarity transition match well with the NAD vertical flux patches.

The results for the upper Olduvai polarity transition obtained from this study were compared with records from ODP Site 983 (Mazaud and Channell 1999; Channell et al. 2002). At Site 983 (60.40° N, 23.64° W), detailed records of the upper Olduvai polarity transition were obtained by u-channel measurements at 1 cm intervals. Figure 7 compares the VGP path, VGP latitude, and relative paleointensity results of the two studies. The VGP position data and age model from Site 983 are based on Channell et al. (2002), and the relative paleointensity data from this site is based on Mazaud and Channell (1999). Note that the age models used in the two studies are different. The VGP path from the present study does not exhibit the large loop seen in the Site 983 record (Fig. 7d). However, the VGP cluster areas in eastern North America, the South Atlantic, and the Middle East obtained from Site 983 are similar to the areas reported in this study, whereas the VGP paths between the clusters are different. The bold letters in Fig. 7 represent the temporal order of the VGP clusters at each study site. As shown in Fig. 7, the temporal orders of the VGP clusters are different for the two studies, but the positions of the clusters are similar. The two records indicate that the VGP cluster moved to a location close to the study site at the beginning of the transition, relocated to eastern North America, and then relocated to the Southern Hemisphere. This suggests that the VGPs at the beginning of the transition could have been affected by one of the vertical flux patches of the NAD field located near the study site, and that, just before the polarity reversal, the VGPs were strongly affected by the NAD field near eastern North America.

Based on the ten most detailed records of reversals from volcanic materials, Valet et al. (2012) proposed that a geomagnetic reversal can be characterized by three successive phases: a precursory event (phase 1), a 180° polarity switch (phase 2), and a rebound (phase 3). Furthermore, they proposed that the maximum durations of phases 1 and 3 are ~2.5 kyr and the maximum duration of phase 2 is approximately ~1.0 kyr. Our records for VGP cluster A could correspond to phase 1, the VGP path between clusters C and D could correspond to phase 2, and VGP cluster E could correspond to phase 3.

A previous study that analyzed variations in field intensity across the five polarity reversals over the past 2 Myr suggested that the intensity records were characterized by a rapid recovery following the transitions (Valet et al. 2005). However, the relative paleointensity variations during the upper Olduvai polarity transition reconstructed in this study (Fig. 7c) as well as from Site 983 (Fig. 7f) show a slow and gradual recovery after the reversal. Thus, the results from this study and Site 983 suggest that the paleointensity variations associated with geomagnetic reversals could differ between each reversal.

Clement (2004) estimated that the average duration of a polarity transition is approximately 7 kyr, based on data from 30 selected sedimentary records spanning the four most recent polarity reversals (the Matuyama-Brunhes, the upper and the lower Jaramillo, and the upper Olduvai), and that this duration varies with site latitude: shorter durations were observed at low-latitude sites, and longer durations were observed at mid- to high-latitude sites. Our estimated duration of 4.5 kyr for the polarity transition closely matches the durations observed at mid-latitude sites, which range from 3 to 10 kyr, by Clement (2004).

At Site 983, the upper Olduvai polarity transition was recorded between 248 and 247 mcd (Mazaud and Channell 1999). The 248 mcd was dated at 1778 ka using the age model constructed by Channell et al. (2002) based on the astrochronology of Shackleton et al. (1990). For Core M, the upper Olduvai polarity transition was determined to be between depths of 66.99 and 63.60 m. The depth of 66.99 m is dated at 1773.4 ka using an age model also based on the astrochronology of Shackleton et al. (1990). This estimated age from Core M is close to but slightly younger (ca. 4 kyr) than the estimated age from Site 983. The reason for this discrepancy could be partly due to a lock-in depth inherent in the Site 983 sediments. However, we acknowledge that our age model, which consists of only two tie points, has to be improved to allow more precise comparison with other records in terms of reversal timing.

Conclusions

We present a nearly continuous paleomagnetic record of the upper Olduvai polarity reversal from a sediment core (Core M) drilled in southern Yokohama City, on the Pacific side of central Japan. The results of our study are summarized as follows.
  1. (1)

    In Core M, the reversal boundary was observed between 65.69 and 65.67 m depth (dated to 1782.7 ka), the polarity transition was observed between 66.99 and 63.60 m depth (dated to between 1784.4 and 1779.9 ka), and the duration of the polarity transition is estimated to be 4.5 kyr, based on an age model derived from δ18O measurements in the planktonic foraminifer G. inflata from the same core.

     
  2. (2)

    In the polarity transition period, VGP clusters were observed in five areas: eastern Asia near Japan, the Middle East, eastern North America (the North Atlantic), off southern Australasia, and the southern South Atlantic off South Africa. The primary locations of the observed VGP clusters are closely related to the vertical flux patches of the present NAD field.

     
  3. (3)

    The relative paleointensity decreased rapidly approximately 1 kyr before the beginning of the polarity reversal and recovered to its original value after almost 12 kyr. Between depths of 67.03 and 63.46 m, where a large movement in the VGPs was observed, the relative paleointensity dropped to approximately 12 % of the maximum value observed pre-reversal.

     

Abbreviations

AFD: 

Alternating field demagnetization

AMS: 

Anisotropy of magnetic susceptibility

ARM: 

Anhysteretic remanent magnetization

ChRM: 

Characteristic remanent magnetization

MAD: 

Maximum angular dispersion

MD: 

Multidomain

MIS: 

Marine isotope stage

NAD: 

Non-axial dipole

NRM: 

Natural remanent magnetization

PSD: 

Pseudo-single domain

SD: 

Single domain

THD: 

Thermal demagnetization

VGP: 

Virtual geomagnetic pole

κ: 

Magnetic susceptibility

κARM : 

Anhysteretic remanent magnetization susceptibility

Declarations

Acknowledgements

We would like to thank Y. Yamamoto who supported the rock-magnetic measurements at the Center for Advanced Marine Core Research (CMCR), Kochi University under the cooperative research program (grants 15A046, 15B041, 14A017, and 14A018). We also thank T. Gokan (Moriya City Office) and Y. Haneda (Ibaraki University) for the cooperation on magnetic measurements. We would like to thank K. Kamoshida (Nitori Co., Ltd.) for the isotopic measurements. We would like to thank M. Arima, M. Ishikawa, R. Wani, and S. Kawagata (Yokohama National University) for the discussions in our seminars. We also would like to thank M. Utsunomiya (Geological Survey of Japan, AIST) for the helpful discussions. We would like to thank two reviewers and the editor of Progress in Earth and Planetary Science for the constructive comments that greatly improved the manuscript.

Funding

This study was supported by the joint research project of the Faculty of Environment and Information Sciences, Yokohama National University in fiscal 2010, by a grant from the Fujiwara Natural History Foundation in 2015 and in part by Grants-in-Aid for Scientific Research (A: no. 16204041 and B: no. 20403015) from the Japan Society for the Promotion of Science (JSPS).

Authors’ contributions

MO proposed the paleomagnetic study of polarity transition. CK and MO wrote the manuscript and coordinated the structure of the paper. CK collected the paleomagnetic and rock-magnetic samples, carried out the experiments and measurements, and analyzed the data. AN carried out the geological surveys. RM conducted the drill core research project and contributed to the construction of the manuscript. HW contributed to the isotope analyses. All the authors read and approved the final manuscript.

Competing interests

The authors declare that they have no competing interests.

Open AccessThis article is distributed under the terms of the Creative Commons Attribution 4.0 International License (http://creativecommons.org/licenses/by/4.0/), which permits unrestricted use, distribution, and reproduction in any medium, provided you give appropriate credit to the original author(s) and the source, provide a link to the Creative Commons license, and indicate if changes were made.

Authors’ Affiliations

(1)
Graduate School of Environment and Information Sciences, Yokohama National University, Yokohama, Japan
(2)
College of Science, Ibaraki University, Mito, Japan
(3)
Hiratsuka City Museum, Hiratsuka, Japan
(4)
Faculty of Environment and Information Sciences, Yokohama National University, Yokohama, Japan
(5)
Faculty of Science, Shizuoka University, Shizuoka, Japan

References

  1. Channell JET, Guyodo Y (2004) The Matuyama Chronozone at ODP Site 982 (Rockall Bank): evidence for decimeter-scale magnetization lock-in depths. In: Channell JET, Kent DV, Lowrie W, Meert JG (eds) Timescales of the paleomagnetic field, vol 145, Geophys Monogr Ser. AGU, Washington DC, pp 205–219. doi:10.1029/145GM15 View ArticleGoogle Scholar
  2. Channell JET, Mazaud A, Sullivan P, Turner S, Raymo ME (2002) Geomagnetic excursions and paleointensities in the Matuyama Chron at Ocean Drilling Program Sites 983 and 984 (Iceland Basin). J Geophys Res 107(B6). doi:10.1029/2001JB000491Google Scholar
  3. Channell JET, Labs J, Raymo ME (2003) The Réunion Subchronozone at ODP Site 981 (Feni Drift, North Atlantic). Earth Planet Sci Lett 215(1):1–12. doi:10.1016/S0012-821X(03)00435-7 View ArticleGoogle Scholar
  4. Channell JET, Curtis JH, Flower BP (2004) The Matuyama-Brunhes boundary interval (500-900 ka) in North Atlantic drift sediments. Geophys J Int 158(2):489–505. doi:10.1111/j.1365-246X.2004.02329.x View ArticleGoogle Scholar
  5. Channell JET, Xuan C, Hodell DA (2009) Stacking paleointensity and oxygen isotope data for the last 1.5 Myr (PISO-1500). Earth Planet Sci Lett 283(1):14–23. doi:10.1016/j.epsl.2009.03.012 View ArticleGoogle Scholar
  6. Channell JET, Mazaud A, Sullivan P, Turner S, Raymo ME (2013) Declination and inclination of ODP Site., pp 162–983. doi:10.1594/PANGAEA.808978 Google Scholar
  7. Clement BM (1991) Geographical distribution of transitional VGPs: evidence for non-zonal equatorial symmetry during the Matuyama-Brunhes geomagnetic reversal. Earth Planet Sci Lett 104(1):48–58. doi:10.1016/0012-821X(91)90236-B View ArticleGoogle Scholar
  8. Clement BM (2004) Dependence of the duration of geomagnetic polarity reversals on site latitude. Nature 428(6983):637–640. doi:10.1038/nature02459 View ArticleGoogle Scholar
  9. Constable C (2007) Non-dipole field. In: Gubbins D, Herrero-Bervera E (eds) Encyclopedia of geomagnetism and paleomagnetism. Springer, Netherlands, pp 701–704View ArticleGoogle Scholar
  10. Day R, Fuller M, Schmidt VA (1977) Hysteresis properties of titanomagnetites: grain size and composition dependence. Phys Earth Planet Int 13(4):260–267. doi:10.1016/0031-9201(77)90108-X View ArticleGoogle Scholar
  11. deMenocal PB, Ruddiman WF, Kent DV (1990) Depth of the post-depositional remanence acquisition in deep-sea sediments: a case study of the Brunches-Matuyama reversal and oxygen isotopic stage 19.1. Earth Planet Sci Lett 99:1–13. doi:10.1016/0012-821X(90)90066-7 View ArticleGoogle Scholar
  12. Dunlop DJ (2002) Theory and application of the Day plot (Mrs/Ms versus Hcr/Hc), 1. Theoretical curves and tests using titanomagnetite data. J. Geophys Res 107(B3). doi:10.1029/2001JB000486Google Scholar
  13. Eto T (1986) Stratigraphy of the Hayama Group in the Miura Peninsula, Japan. Sci Rep Yokohama National Univ (Sec II) 33:67–106Google Scholar
  14. Hoffman KA (1991) Long-lived transitional states of the geomagnetic field and the two dynamo families. Nature 354(6351):273–277View ArticleGoogle Scholar
  15. Hoffman KA (1992) Dipolar reversal states of the geomagnetic field and core-mantle dynamics. Nature 359(6398):789–794View ArticleGoogle Scholar
  16. Hoffman KA, Singer BS, Camps P, Hansen LN, Johnson KA, Clipperton S, Carvallo C (2008) Stability of mantle control over dynamo flux since the mid-Cenozoic. Phys Earth Planet Int 169(1):20–27. doi:10.1016/j.pepi.2008.07.012 View ArticleGoogle Scholar
  17. Hyodo M (1984) Possibility of reconstruction of the past geomagnetic field from homogeneous sediments. J Geomag Geoelectr 36:45–62. doi:10.5636/jgg.36.45 View ArticleGoogle Scholar
  18. Irving E, Major A (1964) Post-depositional detrital remanent magnetization in a synthetic sediment. Sedimentology 3(2):135–143. doi:10.1111/j.1365-3091.1964.tb00638.x View ArticleGoogle Scholar
  19. Kent DV (1973) Post-depositional remanent magnetisation in deep-sea sediment. Nature 246(5427):32–34View ArticleGoogle Scholar
  20. King JW, Banerjee SK, Marvin J (1983) A new rock-magnetic approach to selecting sediments for geomagnetic paleointensity studies: application to paleointensity for the last 4000 years. J Geophys Res 88:5911–5921View ArticleGoogle Scholar
  21. Kirschvink JL (1980) The least-squares line and plane and the analysis of palaeomagnetic data. Geophys J Int 62:699–718. doi:10.1111/j.1365-246X.1980.tb02601.x View ArticleGoogle Scholar
  22. Kitazaki T, Majima R (2003) A slope to outer-shelf cold-seep assemblage in the Plio-Pleistocene Kazusa Group, Pacific side of central Japan. Paleontol Res 7:279–296View ArticleGoogle Scholar
  23. Kusu C, Nozaki A, Okada M, Wada H, Majima R (2014) Lithology and upper boundary of the Olduvai Subchronozone in a core recovered from the middle Kazusa Group (Lower Pleistocene) on the Miura Peninsula, Pacific side of central Japan. J Geol Soc Jpn 120:53–70. doi:10.5575/geosoc.2014.0002 View ArticleGoogle Scholar
  24. Laj C, Mazaud A, Weeks R, Fuller M, Herrero-Bervera E (1991) Geomagnetic reversal paths. Nature 351:477. doi:10.1038/351447a0 View ArticleGoogle Scholar
  25. Lisiecki LE, Raymo ME (2005) A Pliocene-Pleistocene stack of 57 globally distributed benthic δ18O records. Paleoceanography 20:PA1003. doi:10.1029/2004PA001071 Google Scholar
  26. Mazaud A, Channell JET (1999) The top Olduvai polarity transition at ODP Site 983 (Iceland Basin). Earth Planet Sci Lett 166:1–13. doi:10.1029/2001JB000491 View ArticleGoogle Scholar
  27. Mazaud A, Channell JET, Xuan C, Stoner JS (2009) Upper and lower Jaramillo polarity transitions recorded in IODP Expedition 303 North Atlantic sediments: implications for transitional field geometry. Phys Earth Planet Int 172(3):131–140. doi:10.1016/j.pepi.2008.08.012 View ArticleGoogle Scholar
  28. Merrill RT, McFadden PL (1999) Geomagnetic polarity transitions. Rev Geophys 37(2):201–226View ArticleGoogle Scholar
  29. Mitsunashi T, Suda Y (1980) Geological map of Japan 1: 200 000, Otaki. Geol Surv Japan, TsukubaGoogle Scholar
  30. Mitsunashi T, Ono K, Suda Y (1980) Geological map of Japan 1:200 000, Yokosuka. Geol Surv Japan, TsukubaGoogle Scholar
  31. Nozaki A, Majima R, Kameo K, Sakai S, Kouda A, Kawagata S, Wada H, Kitazato H (2014) Geology and age model of the Lower Pleistocene Nojima, Ofuna, and Koshiba Formations of the middle Kazusa Group, a forearc basin-fill sequence on the Miura Peninsula, the Pacific side of central Japan. Island Arc 23:157–179. doi:10.1111/iar.12066 View ArticleGoogle Scholar
  32. Oba T, Irino T, Yamamoto M, Murayama M, Takamura A, Aoki K (2006) Paleoceanographic change off central Japan since the last 144 000 years based on high-resolution oxygen and carbon isotope records. Glob Planet Change 53:5–20View ArticleGoogle Scholar
  33. Ohno M, Murakami F, Komatsu F, Guyodo Y, Acton G, Kanamatsu T, Evans HF, Nanayama F (2008) Paleomagnetic directions of the Gauss-Matuyama polarity transition recorded in drift sediments (IODP Site U1314) in the North Atlantic. Earth Planets Space 60:e13–e16View ArticleGoogle Scholar
  34. Okada M, Niitsuma N (1989) Detailed paleomagnetic records during the Brunhes-Matuyama geomagnetic reversal, and a direct determination of depth lag for magnetization in marine sediments. Phys Earth Planet Int 56:133–150View ArticleGoogle Scholar
  35. Passier HF, De Lange GJ, Dekkers MJ (2001) Magnetic properties and geochemistry of the active oxidation front and the youngest sapropel in the eastern Mediterranean Sea. Geophys J Int 145(3):604–614View ArticleGoogle Scholar
  36. Sakamoto T, Sakai A, Hata M, Unozawa A, Oka S, Hiroshima T, Komazawa M, Murata Y (1987) Geological Map of Japan 1: 200 000, Tokyo. Geol Surv Japan, TsukubaGoogle Scholar
  37. Seno T, Takano T (1989) Seismotectonics at the trench-trench-trench triple junction off central Honshu. Pure Appl Geophys 129:27–40View ArticleGoogle Scholar
  38. Shackleton NJ, Berger A, Peltier WR (1990) An alternative astronomical calibration of the lower Pleistocene timescale based on ODP Site 677. Trans R Soc Edinburgh Earth Sci 81:251–261View ArticleGoogle Scholar
  39. Suganuma Y, Yokoyama Y, Yamazaki T, Kawamura K, Horng CS, Matsuzaki H (2010) 10Be evidence for delayed acquisition of remanent magnetization in marine sediments: implication for a new age for the Matuyama-Brunhes boundary. Earth Planet Sci Lett 296(3):443–450. doi:10.1016/j.epsl.2010.05.031 View ArticleGoogle Scholar
  40. Suganuma Y, Okuno J, Heslop D, Roberts AP, Yamazaki T, Yokoyama Y (2011) Post-depositional remanent magnetization lock-in for marine sediments deduced from 10Be and paleomagnetic records through the Matuyama-Brunhes boundary. Earth Planet Sci Lett 311(1):39–52. doi:10.1016/j.epsl.2011.08.038 View ArticleGoogle Scholar
  41. Takahashi N, Mitsuoka T, Katoh A, Yokoyama K (2005) Correlation of the key tephra bed ‘Kd38’ occurring near the boundary between the Tertiary and Quaternary in the southern part of the Kanto Province, central Japan: correlation in the Kazusa Group and the Chikura Group, Boso Peninsula. J Geol Soc Jpn 111:371–388View ArticleGoogle Scholar
  42. Tate Y, Majima R (1998) A chemosynthetic fossil community related to cold seeps in the outer shelf environment: a case study in the Lower Pleistocene Koshiba Formation, Kazusa Group, central Japan. J Geol Soc Jpn 104:24–41View ArticleGoogle Scholar
  43. Tauxe L (1993) Sedimentary records of relative paleointensity of the geomagnetic field: theory and practice. Rev Geophys 31:319–354View ArticleGoogle Scholar
  44. Tauxe L, Yamazaki T (2007) Paleointensities. In: Schubert G (ed) Treatise on geophysics, vol 5, 1st edn. Elsevier, Amsterdam, pp 509–563View ArticleGoogle Scholar
  45. Tauxe L, Herbert T, Shackleton NJ, Kok YS (1996) Astronomical calibration of the Matuyama–Brunhes boundary: consequences for magnetic remanence acquisition in marine carbonates and the Asian loess sequences. Earth Planet Sci Lett 140:133–146. doi:10.1016/0012-821X(96)00030-1 View ArticleGoogle Scholar
  46. Tauxe L, Steindor JL, Harris A (2006) Depositional remanent magnetization: toward an improved theoretical and experimental foundation. Earth Planet Sci Lett 244:515–529. doi:10.1016/j.epsl.2006.02.003 View ArticleGoogle Scholar
  47. Tric E, Laj C, Jéhanno C, Valet JP, Kissel C, Mazaud A, Iaccarino S (1991) High-resolution record of the Upper Olduvai transition from Po Valley (Italy) sediments: support for dipolar transition geometry? Phys Earth Planet Int 65(3):319–336View ArticleGoogle Scholar
  48. Unozawa A, Oka S, Sakamoto T, Komazawa M (1983) Geological map of Japan 1:200 000, Chiba. Geol Surv Japan, TsukubaGoogle Scholar
  49. Utsunomiya M, Majima R (2012) Paleobathymetries of the Plio-Pleistocene Urago and Nojima Formations, Kazusa Group, Miura Peninsula, central Japan: revision on the basis of molluscan fossils from new localities. Fossils (Palaeontological Soci Jpn) 91:5–14Google Scholar
  50. Valet JP, Meynadier L, Guyodo Y (2005) Geomagnetic dipole strength and reversal rate over the past two million years. Nature 435(7043):802–805View ArticleGoogle Scholar
  51. Valet JP, Fournier A, Courtillot V, Herrero-Bervera E (2012) Dynamical similarity of geomagnetic field reversals. Nature 490(11491):89–93. doi:10.1038/nature11491 View ArticleGoogle Scholar
  52. Verosub KL (1977) Depositional and postdepositional processes in the magnetization of sediments. Rev Geophys 15(2):129–143View ArticleGoogle Scholar
  53. Wada H, Fujii N, Niitsuma N (1984) Analytical method of stable isotope for ultra-small amounts of carbon dioxide with MAT250 mass-spectrometer. Geosci Repts Shizuoka Univ 10:103–112Google Scholar
  54. Yamazaki T, Oda H (2005) A geomagnetic paleointensity stack between 0.8 and 3.0 Ma from equatorial Pacific sediment cores. Geochem Geophys Geosys 6(11):Q11H20. doi:10.1029/2005GC001001 View ArticleGoogle Scholar
  55. Zijderveld JDA (1967) A. C. demagnetization of rocks: analysis of results. In: Collinson DW, Creer KM, Runcorn SK (eds) Methods in palaeomagnetism. Elsevier, Amsterdam, pp 254–286Google Scholar

Copyright

© The Author(s). 2016

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