Haneda et al. Progress in Earth and Planetary Science
https://doi.org/10.1186/s40645-020-00354-y(2020) 7:44

Progress in Earth andPlanetary Science

RESEARCH ARTICLE

Open AccessA full sequence of the Matuyama–Brunhes
geomagnetic reversal in the Chiba
composite section, Central Japan
Yuki Haneda1,2,3* , Makoto Okada2, Yusuke Suganuma1,4 and Takahiro Kitamura2

Abstract
Geological records of the Matuyama–Brunhes (M–B) geomagnetic reversal facilitate the development of an age
model for sedimentary and volcanic sequences and help decipher the dynamics of the Earth’s magnetic field.
However, the structure of the geomagnetic field during the M–B geomagnetic reversal remains controversial due toits complex field behavior. In this study, we conducted paleo- and rock-magnetic analyses of samples from the
Chiba composite section (CbCS), a continuous and expanded marine succession in Central Japan, to reconstruct thefull sequence of the M–B geomagnetic reversal. We define an average stratigraphic position of the M–B boundaryand estimate its age based on three sections in the CbCS and a neighboring drill core, TB-2. The average
stratigraphic position of the M–B boundary in the CbCS is established at 1.1 ± 0.3 m above a widespread volcanicash bed (the Byk-E tephra). Assuming a chronological error associated with orbital tuning of 5 kyr and stratigraphic
uncertainty of 0.4 kyr, the M–B boundary in CbCS is at 772.9 ± 5.4 ka (1σ). The virtual geomagnetic pole, which is
calculated from the paleomagnetic directions, shows several short fluctuations between 783 and 763 ka, withconcomitant decreases in geomagnetic field intensity index. After termination of the field instabilities, the field
intensity recovered and became higher than before the M–B boundary, with a stable normal polarity direction. The
paleomagnetic records in the CbCS exhibit a field asymmetry between the axial dipole decay and field recovery,
providing a full sequence of the M–B reversal, suggesting that the non-axial dipole field dominated several times
during periods ca. 20 kyr long across the M–B boundary, due to depletion in the main axial dipole component. Our
results provide probably the most detailed sedimentary record of the M–B geomagnetic reversal and offer valuable
information to further understand the mechanism and dynamics of geomagnetic reversals.
Keywords: Paleomagnetism, Rock-magnetism, Matuyama–Brunhes geomagnetic reversal, Matuyama–Brunhes
boundary, Relative paleointensity, Chiba composite section, Chiba section, GSSP, Chibanian Stage/Age, Lower–
Middle Pleistocene boundary

1 Introduction
The Matuyama–Brunhes (M–B) geomagnetic reversal is
the last polarity reversal and one of the most studied
paleomagnetic events (e.g., Opdyke et al. 1966; Harrison1974). The mid-point of the polarity switch interval (the* Correspondence: yuuki.haneda.paleo@gmail.com
1
National Institute of Polar Research, 10-3 Midori-cho, Tachikawa, Tokyo
190-8518, Japan2
Department of Earth Sciences, Ibaraki University, 2-2-1 Bunkyo, Mito, Ibaraki
310-8512, Japan
Full list of author information is available at the end of the article

directional transition from a reversed polarity to normal
polarity) in the reversal (M–B boundary) is a criticalstratigraphic marker for calibrating the age of sedimentary and volcanic sequences (e.g., Channell et al. 2010;
Suganuma et al. 2015; Singer et al. 2019) and for defining the Lower–Middle Pleistocene boundary (Head et al.
2008; Head and Gibbard 2015; Suganuma et al. 2018).
The timing of the M–B boundary has been consistently
constrained to ca. 773 ka in recent studies based on theastronomical age of rapidly accumulated marine© The Author(s). 2020 Open Access This article is licensed under a Creative Commons Attribution 4.0 International License,
which permits use, sharing, adaptation, distribution and reproduction in any medium or format, as long as you give
appropriate credit to the original author(s) and the source, provide a link to the Creative Commons licence, and indicate ifchanges were made. The images or other third party material in this article are included in the article's Creative Commonslicence, unless indicated otherwise in a credit line to the material. If material is not included in the article's Creative Commonslicence and your intended use is not permitted by statutory regulation or exceeds the permitted use, you will need to obtain
permission directly from the copyright holder. To view a copy of this licence, visit http://creativecommons.org/licenses/by/4.0/.

(2020) 7:44

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sediments (e.g., Channell et al. 2010; Valet et al. 2014,2019; Suganuma et al. 2018) and radiometric dating of
volcanic ash and lava-flows (e.g., Suganuma et al. 2015;
Singer et al. 2019).
Geomagnetic field structure during the M–B geomagnetic reversal has also been intensively studied (e.g.,
Hartl and Tauxe 1996; Brown et al. 2004; Singer et al.
2005, 2019; Raisbeck et al. 2006; Sagnotti et al. 2014,2016; Balbas et al. 2018; Macri et al. 2018).
Paleomagnetic records from several marine sediments
and lavas, and numerical model simulations, show thatthe polarity switch spans several thousand years depending on site location (e.g., Channell and Guydo 2004;
Channell et al. 2010; Clement 2004; Singer et al. 2005;
Amit et al. 2010; Olson et al. 2011; Valet et al. 2016). On
the other hand, a remarkably short transition has been
identified in exposed lacustrine sediments in the Sulmona Basin of the Central Apennines, Italy (Sagnotti
et al. 2014, 2016) and marine sediments of the Valle diManche section from Calabria, Southern Italy (Macriet al. 2018). In these studies, the polarity switch at theM–B boundary, dated at ca. 786 ka, was inferred to havecompleted within a century (Sagnotti et al. 2014; Macriet al. 2018). Prior to the M–B boundary, a clear decay ingeomagnetic field intensity has been identified as a “precursor” event in deep-sea sediments (e.g., Hartl andTauxe 1996; Channell et al. 2009), lavas (e.g., Brown
et al. 2004; Singer et al. 2005), and ice (Raisbeck et al.
2006; Dreyfus et al. 2008). This precursor field decay is
generally dated to ca. 795 ka and appears to be of thesame magnitude as that at the M–B boundary (e.g.,
Hartl and Tauxe 1996; Gotton et al. 2007). Recent
paleomagnetic records from highly precisely 40Ar/39Ardated lavas have indicated the long-term processes ofthe M–B geomagnetic reversal (Balbas et al. 2018; Singer
et al. 2019). Based on these data, the field decay isthought to have begun ca. 30 kyrs before the polarity
switch at ca. 773 ka, and involved centennial-tomillennial-scale field instabilities and a depletion of theaxial dipole field (Balbas et al. 2018; Simon et al. 2018a;
Singer et al. 2019). However, volcanic sequences fail tocapture the entire M–B geomagnetic reversal record dueto the nature of sporadic volcanic eruptions, instead, recording near-instantaneous geomagnetic signals. Sedimentary sequences provide continuous records, but
geomagnetic recordings in slowly accumulated sediments are affected by vertical and temporal offsets between the sediment/water interface and the lock-in
depth of at least 10–15 cm, the zone where thepaleomagnetic record is fixed (e.g., Channell and Guydo
2004; Liu et al. 2008; Suganuma et al. 2010; Ménabréaz
et al. 2012; Simon et al. 2018b). This may generate asmoothing effect due to progressively changing
sedimentation-fabric processes (e.g., Suganuma et al.

2011; Roberts et al. 2013; Valet and Fournier 2016).
Moreover, the sediment-recording smoothing process is
amplified during paleomagnetic measurements due tothe averaging effect of the magnetometer response window, when U-channels are used for the measurements(e.g., Oda and Xuan 2014; Philippe et al. 2018). These
complex behaviors and the lack of truly continuousinstantaneous recordings have resulted in an incomplete
picture of centennial-to-millennial-scale field behavior
during the M–B geomagnetic reversal.
A detailed paleomagnetic record of the M–B geomagnetic reversal with a centennial-scale resolution was reported from the 130 m thick Chiba composite section inCentral Japan (Fig. 1; Suganuma et al. 2015; Okada et al.
2017; Simon et al. 2019). On January 2020, the Chiba
section, the main portion of the Chiba composite section, was ratified by the Executive Committee of theInternational Union of Geological Sciences as a Global
Boundary Stratotype Section and Point (GSSP) for the
Middle Pleistocene Subseries (Chibanian Stage/Age)(Fig. 1c). The M–B boundary, the primary marker of the
GSSP, lies 1.1 m above the Byk-E tephra bed and isdated to 772.9 ka (Suganuma et al. 2015, 2018; Okada
et al. 2017; Simon et al. 2019). However, a complete sequence across the M–B geomagnetic reversal was notthen available because of a lack of data for the upper 36
m portion of the Chiba composite section.
We report new paleomagnetic and rock-magnetic
measurements for the upper part of the Chiba composite
section. We also combine the results with those of Suganuma et al. (2015), Okada et al. (2017), and Simon et al.
(2019), to reconstruct a full sequence of the M–B geomagnetic reversal in the interval between 794 and 748
ka, which covers the entire interval of depletion of thegeomagnetic field intensity, including the polarity switch
and preceding transitional phase. Additionally, new samples identify the M–B boundary from a sub-section of
the Chiba composite section, to confirm correlations between its sub-sections. For the first time using chronological and stratigraphic uncertainties between the Chiba
composite section (Suganuma et al. 2015; Okada et al.
2017; our results) and a neighboring sediment core
(Hyodo et al. 2016), we constrain the average age andstratigraphic position of the M–B boundary at the GSSP
for the Chibanian Stage/Age and Middle Pleistocene
Subseries/Subepoch.

Haneda et al. Progress in Earth and Planetary Science

1.1 Geological and stratigraphic settings

The Kazusa Group is the Pleistocene age marine succession exposed in the Boso Peninsula, Central Japan (Fig.
1a, b), which has been investigated in a number of studies (e.g., Kanehara et al. 1949; Kawai 1951; Shinada et al.
1951; Mitsunashi 1954; Mitsunashi et al. 1959; Nakajima1978). The group is characterized by excellent andHaneda et al. Progress in Earth and Planetary ScienceFig. 1 (See legend on next page.)(2020) 7:44

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Haneda et al. Progress in Earth and Planetary Science(2020) 7:44

Page 4 of 22(See figure on previous page.)
Fig. 1 Location map of the study sites and stratigraphy of the Kazusa Group. a Map of the Japanese archipelago. b Distribution of marine
successions on the Boso Peninsula and the location of the Chiba composite section. The black square indicates the area of the geological map in
Fig. 2. c Stratigraphy of the Kazusa Group after Kazaoka et al. (2015): (i) Niitsuma (1976), (ii) Okada and Niitsuma (1989), (iii) Pickering et al. (1999),and (iv) Tsuji et al. (2005)

continuous exposure and exceptionally high sedimentation rates (ca. 2 m/kyr) (Kazaoka et al. 2015). The presence of many volcanic ash beds in the Boso Peninsula
enables the reconstruction of a detailed stratigraphy(reviewed by Satoguchi and Nagahashi 2012; Fig. 1c). In
addition, well-preserved marine microfossils (calcareous
nannofossils, benthic and planktonic foraminifera, radiolaria, and diatoms) from the Kazusa Group have also
been used to reconstruct biostratigraphies and a foraminiferal oxygen isotope stratigraphy (Aoki 1968; Oda
1977; Sato et al. 1988; Okada and Niitsuma 1989; Pickering et al. 1999; Cherepanova et al. 2002; Suganuma et al.
2018; Fig. 1c). These features provide a continuous,
high-resolution chronostratigraphic framework.
The Chiba composite section in the middle part of the
Kokumoto Formation, Kazusa Group, comprises the
Yanagawa, Urajiro, Yoro River (including the Chiba section, the GSSP of the Chibanian Stage/Age), YoroTabuchi, and Kokusabata sub-sections (Figs. 2 and 3;
Suganuma et al. 2018). A summary of the stratigraphic
levels and locality of the Chiba composite section isshown in Table 1. The lithofacies of the middle part of
the Chiba composite section is dominated by bioturbated silty beds with minor sandy beds, showing no evidence of episodic deposition and hiatus (Nishida et al.
2016), although thicker sandy beds are observable in thelowermost and upper parts (Fig. 3). Numerous tephra
beds (Ku1, Ku2A, Ku2B, Tas, Tap, and Byk) occur in the
Chiba composite section and demonstrate good correlation between sub-sections (Fig. 3; Suganuma et al.
2018). The Byk-E tephra bed, originating from the Older
Ontake volcano in Central Japan (Takeshita et al. 2016),
was ratified as the base of the Chibanian Stage/Age.
1.2 Chronology of the Chiba composite section

The chronological framework for dating the Chiba composite section is provided by correlation between thebenthic oxygen isotope stratigraphy (Suganuma et al.
2018) and a sea-level proxy curve obtained from Ocean
Drilling Project (ODP) Site 1123, offshore New Zealand(Elderfield et al. 2012), assisted by a U-Pb radiometricage of the Byk-E tephra (Suganuma et al. 2015). The
resulting age model reveals that the Chiba composite
section corresponds to an age range from the late partof marine isotope stage (MIS) 20 (801.1 ka) to the earlypart of MIS 18 (747.6 ka) with the sedimentation rate
varying between 0.44 and 3.87 m/kyr (Suganuma et al.
2018; Haneda et al. 2020). In the Chiba section, the M–

B boundary is 1.1 m above the Byk-E tephra bed in aninterval with a sedimentation rate of 89 cm/kyr (Okada
et al. 2017; Suganuma et al. 2018). The Byk-E tephra has
a Th/U-corrected U-Pb zircon age of 772.7 ± 7.2 ka(Suganuma et al. 2015) and an astronomical age of
774.1 ka, which forms the basal age of the ChibanianStage (Suganuma et al. 2018).
1.3 Previous paleomagnetic studies in the Kokumoto
Formation

The M–B boundary was identified in the middle part of
the Kokumoto Formation of the Kazusa Group in thecentral Boso Peninsula by Nakagawa et al. (1969) andNiitsuma (1971, 1976) (Fig. 1c). These pioneering studies
identified the M–B boundary approximately 2 m below
the Byk-E tephra bed using alternating field (AF)demagnetization, which removes secondary remanencesby exposing a specimen to an alternating magnetic field(Niitsuma 1971; Okada and Niitsuma 1989; Tsunakawa
et al. 1995, 1999; Aida 1997).
Three paleomagnetic studies for the Chiba composite
section have provided detailed records of geomagnetic
field behavior across the M–B boundary (Suganuma et al.
2015; Okada et al. 2017; Simon et al. 2019). Suganuma
et al. (2015) performed thermal demagnetization, whichremoves secondary remanence by heating and coolingspecimens in a zero magnetic field. This was using specimens from the Chiba and Yanagawa sections between −6.55 and 5.75 m (indicated by blue horizontal bars in Fig.
3; Additional file 1: Table S1). The results of the thermal
demagnetization indicate the M–B boundary at 0.8 m
above the Byk-E tephra bed at the Yanagawa section—not
below the tephra bed as reported in earlier studies (Fig. 3).
This difference to previous studies was interpreted as due
to secondary remanence in the specimens, resulting fromhigh coercivity magnetic minerals, that were not removed
by AF demagnetization in the early studies. However, relative paleointensity could not be reconstructed, because
minerals in the specimens underwent heat-alteration during the thermal demagnetization process, producing new
magnetic phases (Suganuma et al. 2015).
Okada et al. (2017) conducted demagnetization using a
hybrid method of thermal demagnetization at 300 °C
followed by progressive AF demagnetization for
specimens from the Chiba and Yoro-Tabuchi sections between − 14.6 and 9.7 m (indicated by black horizontal bars
in Fig. 3; Additional file 2: Table S2). This hybrid
demagnetization method removes the secondary

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Fig. 2 Location of the Chiba composite section (CbCS). a Geological map of the central Boso Peninsula after Suganuma et al. (2018). Red stars
indicate locations of the sub-sections of the CbCS. b–d Detailed maps of the sub-sections of the CbCS, as indicated by the black squares in aremanence, which cannot be removed by AF
demagnetization alone, importantly with minimum
heat-alteration, therefore allowing reconstruction ofrelative paleointensity (such a procedure is widelyused for similar reasons in many other sedimentbased studies). Based on this hybrid approach, theHaneda et al. Progress in Earth and Planetary Science(2020) 7:44

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Fig. 3 Lithological columns of the sub-sections of the Chiba composite section (and TB-2 core) after Suganuma et al. (2018) and Hyodo et al.
(2016, 2017). Blue, black, green, and red horizontal bars indicate sampling levels in Suganuma et al. (2015), Okada et al. (2017), Simon et al. (2019),
and this study, respectively. Green squares indicate sampling levels for 10Be measurements of Simon et al. (2019). The stratigraphic correlation
between the Chiba composite section and the TB-2 core is indicated by red solid lines. TBxx, TB2-xx, TB3-xx, TB4-xx, TBHxx, and YNGxx are sample
codes. Byk-A, Byk-E, Ku1, Ku2A, Ku2B, Koss-1B, Tap-A, Tap-B, Tas-A, Tas-B, and Tas-C are local tephra marker beds

M–B boundary was observed, with a depleted geomagnetic field intensity at 1.1 m above the Byk-E
tephra bed in the Chiba section (Fig. 3). The redefined stratigraphic position of the M–B boundary is

supported by paleomagnetic data from the TB-2 sediment core drilled 190 m northeast of the Chiba section (Hyodo et al. 2016; Figs. 3 and 4). However, therecovery phase of the geomagnetic field (after theHaneda et al. Progress in Earth and Planetary Science(2020) 7:44

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Top

Bottom Top

Bottom

b

Chiba

2.20

-15.5035°17’39.6”N /140°08’47.6”E

35°17'36.9"N /140°08'47.2"Eare almost simultaneous with deposition. However, paleoand rock-magnetic records from the strata above 37.5 m
have not yet been obtained. In addition, although the M–
B boundary is located 1.1 m above the Byk-E tephra bedin the Chiba section, this boundary has not been examined in the Yoro-Tabuchi section, 70 m northeast of the
Chiba section (Figs. 2c and 4).

Yoro Riverc

95.50-15.5035°17’59.6”N /140°08’34.8”E

35°17'36.9"N /140°08'47.2"E

2 Methods

YoroTabuchi

37.50

0.05

35°17’48.1”N /140°09'02.1"E

35°17'41.1"N /140°08'49.7"E

Kokusabatad 57.07

7.87

35°18'43.2"N /140°11'45.7"E

35°18'32.7"N /140°11'53.6"E

Yanagawad

5.60-15.9335°17'08.9"N /140°07'52.9"E

35°17'05.8"N /140°07'51.0"E

Urajirod-19.67 -34.21

35°16'52.4"N /140°07'28.2"E

35°16'49.4"N /140°07'28.3"E

Table 1 Summary of the stratigraphic interval based on theoxygen isotope sampling levels and locality of the Chiba
composite section
SectionLevel (m)aLatitude / LongitudeaStratigraphic levels are indicated as stratigraphic distance from the Byk-E
tephra bedb
The stratigraphic interval of the Chiba section is based on the full range ofthe outcropc
The Yoro River section contains the Chiba sectiond
The stratigraphic levels are converted into the scale of the Yoro River andYoro-Tabuchi sections

polarity switch) was not observed by Okada et al.
(2017) due to a lack of samples collected from the
Byk-E tephra bed above the 9.7 m level.
Simon et al. (2019) undertook paleomagnetic measurements of samples from the Yoro-Tabuchi section between
9.8 and 37.5 m (indicated by green horizontal bars in Fig.
3; Additional file 2: Table S2) and beryllium isotope measurements of samples from the Chiba and Yoro-Tabuchi
sections up to 37.5 m (indicated by pink squares in Fig. 3).
Based on their measurements, the paleomagnetic direction
and relative paleointensity records were obtained, alongwith a beryllium isotope record (10Be, 9Be, and 10Be/9Be
ratio), as independent proxies of the geomagnetic field intensity. These authors suggested that the M–B boundary
involved geomagnetic dipole moment collapse and observed a sequence of rapid directional oscillations between
the M–B boundary and the later field recovery phase, inwhich a stable dipole is reestablished on the opposite direction. The directional oscillations were identified as instability clusters (IC1 = 771.9–768.5 ka, IC2 = 768.5–
765.2 ka, and IC3 = 763.5–762.5 ka) by Simon et al. (2019).
These previous paleomagnetic studies of the Chiba
composite section have reconstructed a centennialresolution paleomagnetic record from the stratigraphic
interval between – 14.6 and 37.5 m from the Byk-E tephra
bed. Exceptionally high sedimentation rates of the Chiba
composite section are thought to limit the lock-in depth
and smoothing effect of the geomagnetic field recording
process (e.g., Suganuma et al. 2011; Valet and Fournier2016). Therefore, the sedimentary succession documents
short-duration instabilities of the geomagnetic field that

2.1 Paleomagnetic core samples

We collected 25 mm-diameter drill core samples from theupper part of the Yoro River section and the lowermost
and middle parts of the Yoro-Tabuchi section using a
portable drill (Figs. 3, 4, and 5). For the Yoro River section(above 37.5 m), the drill cores were collected from 64
levels with a ca. 50 cm spacing (the same interval as used
for deriving the foraminiferal oxygen isotope stratigraphyof Suganuma et al. 2018 and Haneda et al. 2020). The locations of the collected samples are indicated by red horizontal bars in Fig. 3 and red letters in Fig. 5. In thelowermost (0 m to 3.8 m) and middle (11.6 m to 13.4 m)
part of the Yoro-Tabuchi section, drill cores were collected from 34 levels and 9 levels with a ca. 10 to 20 cm
spacing, as indicated by red horizontal bars in Fig. 3 andred letters in Fig. 4. These samples were oriented with amagnetic compass. Each core was cut into 2 cm-longspecimens for subsequent paleomagnetic and rockmagnetic measurements at Ibaraki University and National Institute of Polar Research (NIPR).
2.2 Paleomagnetic measurements

The natural remanent magnetization (NRM) for each specimen was measured using a three-axis cryogenic magnetometer SRM-760R (2G Enterprises, USA) in a magneticallyshielded room at the NIPR. To extract the primary magnetic component, we removed the secondary remanencefrom the NRMs using a hybrid of thermal demagnetization
to 300 °C and AF demagnetization (Okada et al. 2017). The
thermal demagnetization was conducted using a TDS1 (Natsuhara Giken, Japan) at the NIPR. The progressive AF demagnetization was performed in 5 to 10
mT increments between 5 and 20 mT before thethermal demagnetization and between 20 and 80 mT
after the thermal demagnetization, using static threeaxis AF demagnetization coils in an SRM-760R magnetometer at the NIPR.
We
identified
thecharacteristic
remanent
magnetization (ChRM) at peak AF fields between 30
to 50 mT after the thermal demagnetization at 300 °C(Okada et al. 2017; Simon et al. 2019) and calculated
paleomagnetic directions by principal component analysis (Kirschvink 1980) using these step points withanchoring to the origin. The virtual geomagnetic pole
(VGP) was calculated using the ChRM directions with

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Fig. 4 Sampling location in the lower part of the Yoro River section (Chiba section) and the Yoro-Tabuchi section. Circles and squares indicate
sampling sites for paleomagnetic (green) and oxygen isotope (blue) analyses. Sample codes for samples collected by previous studies (Okada
et al. 2017; Suganuma et al. 2018; Simon et al. 2019; Haneda et al. 2020) and this study are shown by black and red letters and numbers,
respectively. The drilling site for the TB-2 sediment core is indicated by a red star (Hyodo et al. 2016)

maximum angular deviation (MAD) of ≤ 15°. We
identified magnetic polarities based on VGP latitudes,
where VGP latitudes > 45° were defined as normal
polarity, latitudes < − 45° were defined as reversed
polarity, and latitudes of − 45 to + 45° were defined

as intermediate polarity. In this study we define theM–B boundary as the mid-point of the polarity
switch interval between the older reversed polarity(VGP latitudes < − 45°) to normal polarity (VGP latitudes > 45°).

Haneda et al. Progress in Earth and Planetary Science(2020) 7:44

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Fig. 5 Sampling location in the upper part of the Yoro River section. Green and blue circles indicate sampling sites for paleomagnetic andoxygen isotope analyses, respectively. Sample codes of samples collected by previous studies (Suganuma et al. 2018; Haneda et al. 2020) and thisstudy are shown by black and red letters and numbers, respectively

2.3 Rock-magnetic measurements

To determine rock-magnetic properties and calculate
the relative paleointensity indices, we measured low-fieldmagnetic susceptibility (kLF), laboratory-induced magnetizations of anhysteretic remanent magnetization (ARM),
and isothermal remanent magnetization (IRM).
Measurements of the kLF were performed on specimens using a KLY-3S Kappabridge susceptibility meter(AGICO, Czech Republic) at Ibaraki University, before
any paleomagnetic or rock-magnetic measurements were
made. For specimens whose NRM was demagnetized by
the hybrid method, ARM was acquired in a 30 μT DC
field with an 80 mT alternating field using the SRM760R magnetometer at the NIPR. In each case, after
ARM
acquisition,
werepeatedthehybriddemagnetization (thermal demagnetization at 300 °C and(2020) 7:44

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AF demagnetization) at 10 mT increments between 20
and 80 mT. ARM susceptibility (kARM) was calculated
from the initial ARM intensity.
kARM is mainly preferentially acquired by small magnetic grains with weak magnetostatic interaction. On theother hand, kLF mainly reflects the total amount of magnetic minerals in a specimen, irrespective of the magnetic grain size. Therefore, the kARM/kLF ratio is aconventional proxy for the grain size of magnetic particles in a specimen; an increase in the kARM/kLF ratio indicates the relative contribution of finer magnetic
particles and vice versa (e.g., Yamazaki 2008; Yamazaki
and Ikehara 2012; Shin et al. 2019).
After the ARM experiments, we exposed selected
specimens to saturated IRM (SIRM) in a 1.5 T fieldusing a MMPM-9 pulse magnetizer (Magnetic Measurements Ltd., UK). The SIRM was measured using a JR-6A
magnetometer (AGICO, Czech Republic) and demagnetized using a TD-48 thermal demagnetizer (ASC Scientific, USA) and AF demagnetizer DEM-8601C(Natsuhara-Giken, Japan) at Ibaraki University. Resultsfrom the hybrid demagnetization method used on the
ARM, along with the SIRM, were used to calculate relative paleointensity indices as outlined below.
To determine S-ratios, IRM acquisition was performedon unprocessed specimens. S-ratios indicate the relative
abundance of antiferromagnetic minerals (such magnetically hard hematite and goethite) in a mixture withferrimagnetic minerals (such as magnetically soft magnetite and maghemite); values close to unity indicatethat the remanence is dominated by magnetically soft
ferrimagnets and vice versa (Bloemendal et al. 1992; Evans and Heller 2003). We imparted SIRM on selected
specimens from the Yoro River and Yoro-Tabuchi sections above 9.7 m using a 1.5 T field using the MMPM-9
pulse magnetizer (Magnetic Measurements Ltd., UK).
The IRM was then measured using the JR-6A magnetometer (AGICO, Czech Republic) at Ibaraki University.
IRMs at 0.1 T and 0.3 T fields were acquired in the opposite direction of the SIRM. The S-ratio-0.1T and Sratio-0.3T were calculated as

calculated the relative paleointensity by normalizing thevector difference of the NRM between 30 mT and 50
mT (after thermal demagnetization at 300 °C) with theARMintensity
removedbetweenthesamedemagnetization steps after thermal demagnetization at
300 °C (Okada et al. 2017; Simon et al. 2019). However,the pseudo-Thellier approach of Tauxe et al. (1995) orother related method has been widely employed to deduce the paleointensity. In the approach of Tauxe et al.
(1995), the slope of the NRM intensity left at a certain
demagnetization step normalized against ARM or IRM
intensity gain at the same imparted field step was usedas the paleointensity proxy. In practice, the slope of the
NRM intensity left at selected demagnetization steps
against the ARM or IRM left at the samedemagnetization steps has been more widely used (e.g.,
Channell et al. 2002; Valet et al. 2014). Moreover,normalization by ARM may over-compensate magnetic
concentration changes because ARM can be sensitive tomagnetostatic interaction among grains (Yamazaki and
Kanamatsu 2007). As a result, the efficiency of ARM acquisition in specimens can decrease with an increase inmagnetic concentration.
Therefore, we normalized the NRM against the ARM(or IRM) as NRM30–50/ARM30–50 (or NRM30–50/IRM30–50) from NRM lost versus ARM (or IRM) lost between
the 30 and 50 mT demagnetization steps (after thermal
demagnetization at 300 °C). We employed these ratios as
relative paleointensity indices. kLF was not used as anormalizer in this study, because it may be less effective
due to possible influence from diamagnetic and paramagnetic minerals and large ferrimagnetic grains that do
not carry NRM (Shin et al. 2019). The Pearson linear
correlation coefficient (R-valueARM and R-valueIRM), associated with the slope of the NRM intensity left at certain demagnetization steps against the ARM or IRM left
at the same demagnetization steps, was calculated tomonitor the goodness of the slope fit (Channell et al.
2002; Shin et al. 2019). A comparison of the proxies
used for relative paleointensity between this study andthe previous studies (Okada et al. 2017; Simon et al.
2019) is in Figure S1.

Haneda et al. Progress in Earth and Planetary ScienceS − ratio − 0:3Tð − 0:1TÞ ¼

− IRM − 0:3Tð − 0:1TÞ =SIRM þ 1 =2

3 Results and discussion
according to Bloemendal et al. (1992).
2.4 Relative paleointensity

Proxies of magnetic grain concentration, kLF, ARM, or
IRM, are usually employed to normalize NRM when calculating the relative paleointensity index (e.g., Kent andOpdyke 1977; Tauxe 1993; Tauxe and Shackleton 1994;
Guyodo and Valet 1999; Yamazaki 1999; Valet 2003;
Yamazaki and Kanamatsu 2007; Suganuma et al. 2008).
Previous studies of the Chiba composite section

3.1 Remanent magnetization

Representative orthogonal vector diagrams (Zijderveld1967) of the NRMs are shown in Fig. 6. Some of the datafor intermediate and reversed polarities before and after
the M–B boundary are taken from Okada et al. (2017) andSimon et al. (2019) (Fig. 6c, d, h, and i). The magnetizationfrom the Chiba composite section consists of one or morecomponents (Fig. 6). The component before thermal
demagnetization at 300 °C generally is the secondary component. After thermal demagnetization at 300 °C, several

Haneda et al. Progress in Earth and Planetary Science(2020) 7:44

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Fig. 6 Orthogonal vector diagrams for representative results of the hybrid demagnetization method. a, b Results indicating normal polarity fromthe upper part of the Yoro River section (this study). c, d Results indicating intermediate polarity in the interval of instability clusters (IC1–IC3)
from the Yoro-Tabuchi section (c Simon et al. 2019; d Okada et al. 2017). e–g Specimen data across the M–B boundary in the lowermost part of
the Yoro-Tabuchi section (this study). h, i Results from the Matuyama reversed polarity interval from the Chiba section (Okada et al. 2017). Closed
and open circles indicate horizontal and vertical components, respectively. Closed and open black data points were used to calculatecharacteristic remanent magnetization in the principal component analysis

specimens showed a ChRM between 30 and 50 mT, which
decays toward the origin (Fig. 6b–i).
Okada et al. (2017) revealed that some samples fromthe Chiba composite section became unstable to thermal
demagnetization at temperatures above 300 °C, and thefollowing AF demagnetization between 30 and 50 mT

appeared to reflect-well the ChRM decay towards theorigin. Because the samples have a single Curie/Néeltemperature at 580 °C (Suganuma et al. 2015; Okada
et al. 2017), the component between 30 and 50 mT isthought to be carried by magnetite. Based on the resultof Suganuma et al. (2015) and Okada et al. (2017), also(2020) 7:44

Page 12 of 22

the component between 30 and 50 mT, probably carriedby magnetite, is identified as the primary remanence. Repeated thermomagnetic experiments performed in the
air show no distinct heat-alteration of magnetic minerals
until 400 °C (Okada et al. 2017). Therefore, we calculated
paleomagnetic directions of the ChRM between the 30
and 50 mT steps, following the thermal demagnetization
at 300 °C.

decrease to 0.86 and 0.46, respectively (Fig. 7m). These
decreases are accompanied by decreases in the relative
paleointensity indices, linear correlation coefficients (RvalueARM and R-valueIRM), NRM intensity, and proxiesof magnetic concentration (Fig. 7e–h, j–k, and m). No
distinct lithological change was observed in this stratigraphic interval in the field. As the interval above 67 mcorresponds to just after the full glacial interval of MIS
18, in the oxygen isotope stratigraphy (Fig. 7b; Haneda
et al. 2020), the change in the magnetic properties may
have originated from local environmental change (e.g.,
supplies of terrestrial and/or organic matter to the basin,hydrographic change in the bottom water, sea levelchange, oceanic circulation changes, or oxygenation).
The indices above the 67 m level may not represent realistic variation in geomagnetic relative intensity. Therefore, we used an S-ratio-0.3T of less than 0.9 as a cut-off
limit for useful relative paleointensity indices, so the indices above the 67 m level are not used in the following
discussions.
The VGP latitude profile above the 38 m level (indicated by red open circles in Fig. 7d) exhibits stable normal polarity with three intermediate polarities at 42.4,
54.0, and 69.2 m (indicated by gray arrows in Fig. 7d).
The NRM30-50/ARM30-50 ratio above the 38 m level isgenerally higher than the overall average of 0.08 andwidely fluctuates between 0.04 and 0.29 in this interval(Fig. 7e). In comparison, the VGP latitude and rockmagnetic properties remain relatively stable.
The NRM30-50/IRM30-50 profile was also obtained fromselected specimens obtained from the Chiba composite
section (Fig. 7g). This ratio shows similar variations to
the NRM30-50/ARM30-50 ratio with the minimum value(at 1.9 m) co-occurring with the M–B boundary, a slightincrease between 6 and 13 m, and a recovery phase to
full intensity at 15–25 m. The high linear correlation coefficient between NRM30-50/ARM30-50 and NRM30-50/
IRM30-50 of 0.94 indicates no systematic differences between the relative paleointensity indices obtained fromthe Chiba composite section.
We calculated the correlations between the relative
paleointensity indices and the normalizers to test for independence. The linear correlation coefficients for ARM
and NRM30-50/ARM30-50, and SIRM and NRM30-50/
IRM30-50 are 0.44 and 0.52, respectively (Additional file
4: Figure S2a–b). These weak positive correlations imply
that the relative paleointensity indices could be partiallycontaminated by the normalizers. On the other hand,
the linear correlation coefficients for authigenic 10Be/9Be
versus NRM30-50/ARM30-50 and NRM30-50/IRM30-50 are− 0.72 and − 0.70, respectively (Additional file 4: Figure
S2c–d). Because the cosmogenic nuclide 10Be production is mainly modulated by variations in magnetospheric shielding on a millennial time scale, 10Be

Haneda et al. Progress in Earth and Planetary Science

3.2 Rock-magnetic properties and paleomagnetic
directions in the Yoro River section

The rock magnetic properties in the upper part of the
Yoro River section (indicated by red and orange circles
in Fig. 7; Additional file 2: Table S2) are relatively
homogenous as are those from lower in the Chiba composite section. Proxies of magnetic grain concentration(kLF, kARM) show simultaneous and minor fluctuations
by a factor of < 10 throughout the Chiba composite section (Fig. 7k). The kARM/kLF ratio—a proxy of magnetic
grain size—varies between 1.5 and 4.4 throughout thesection (Fig. 7l). The S-ratio-0.3T, which represents therelative abundance of antiferromagnetic minerals in a
mixture with ferrimagnetic minerals, fluctuates moderately throughout the Chiba composite section except in
the Yoro River section above 67 m (Fig. 7 m). The average S-ratio-0.3T of 0.96 indicates that the magnetic grains
throughout the Chiba composite section mostly consistof ferrimagnetic minerals (Frank and Nowaczyk 2008).
While the S-ratio-0.1T fluctuates between 0.66 and 0.90
in the Chiba and lower Yoro-Tabuchi sections (Okada
et al. 2017), the fluctuation amplitude is suppressed in
the upper parts of these sections (Fig. 7m). The periodic
decrease of the S-ratio-0.1T in the Yoro-Tabuchi and
Chiba sections is thought to be not attributed to the increase in the proportion of hematite or goethite, butmainly to the increase (decrease) in the amount of fine(coarse) size magnetic grain, because of its cooccurrence with changes in kARM/kLF (Yamazaki et al.
2003; Suganuma et al. 2008; Okada et al. 2017). This interpretation is supported by the high average value ofthe S-ratio-0.3T (0.96), indicating the dominance of ferrimagnetic minerals throughout the Chiba composite section (Frank and Nowaczyk 2008; Fig. 7m). Additionally,
the amplitude of the kARM/kLF changes is limited to afactor of < 3. These data indicate that the magnetic
properties throughout the Chiba composite section are
generally uniform, although minor grain size changes
occur in the lower interval. Therefore, the rockmagnetic data of the Chiba composite section fulfill theestablished criteria for the construction of a robust relative paleointensity record with respect to magnetic grain
concentration, magnetic grain size, and magnetic mineralogy (Tauxe 1993). However, the S-ratio-0.3T and Sratio-0.1T in the Yoro River section above the 67 m level

Haneda et al. Progress in Earth and Planetary Science(2020) 7:44

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Fig. 7 Stratigraphic profiles of paleomagnetic directions and rock-magnetic properties. a Latest age model for the CbCS (Suganuma et al. 2018).
b Benthic oxygen isotope stratigraphy (Haneda et al. 2020). c Maximum angular deviation (MAD) with the 15° cut-off limit indicated, for the
principal component analysis. d VGP latitude calculated from the results of the hybrid demagnetization method. e–h NRM30-50/ARM30-50 and
NRM30-50/IRM30-50 as relative paleointensity indices calculated following the pseudo-Thellier approach of Tauxe et al. (1995), and linear correlation
coefficients (R-valueARM and R-value IRM) of the indices. i Authigenic 10Be/9Be (Simon et al. 2019) as an independent proxy of the field intensity. j
Natural remanent magnetization intensity. k Low-field (kLF) and ARM magnetic susceptibility (kARM) as proxies of magnetic grain concentration. lkARM/kLF ratio as a proxy of magnetic grain size. m S-ratio at − 0.3 T and − 0.1 T as a proxy of the relative abundance of antiferromagnetic toferrimagnetic minerals. New data points obtained from the Yoro-Tabuchi section are shown as blue and light blue open circles and lines, andnew data points from the Yoro River section are shown as red and orange open circles and lines. Data points from previous studies (Okada et al.
2017; Simon et al. 2019) are shown as black and gray closed circles and lines. The assignment of marine isotope stages is after Haneda et al.
(2020). Instability clusters (IC1–IC3) are after Simon et al. (2019)

concentration records from marine sediments constitutean independent proxy of the geomagnetic dipole moment,
providing that secondary depositional processes have been
removed through a normalization procedure (e.g., Simon
et al. 2016). The similarity of NRM30-50/ARM30-50 and
NRM30-50/IRM30-50 ratios with authigenic 10Be/9Be demonstrates that the relative paleointensity from the Chiba
composite section is reliable proxies for the geomagnetic
field intensity variation. Therefore, we used the highresolution NRM30-50/ARM30-50 ratio as the best relative
paleointensity index.
3.3 Paleomagnetic directions of the lowermost part of the
Yoro-Tabuchi section

Paleomagnetic directions of the lowermost part of the
Yoro-Tabuchi section were obtained from 16 levels from

specimens in which the ChRM component MAD was≤ 15° (Additional file 5: Table S3). The polarity
switch from a reversed polarity (VGP latitude < −45°) to a normal polarity (VGP latitude > 45°) occurs
in the stratigraphic interval between 1.1 and 2.1 m.
The M–B boundary, defined as a middle of the polarity switch interval, is located 1.6 m above the Byk-E
tephra bed in the Yoro-Tabuchi section (Fig. 8; Table
1). After the polarity switch is completed, the VGP
latitude largely rebounds across the Equator. The
stratigraphic position of the M–B boundary and subsequent directional instability are consistent with previous paleomagnetic data from the Yanagawa section(Suganuma et al. 2015), the Chiba section (Okada
et al. 2017), and the TB-2 sediment core (Hyodo
et al. 2016).

Haneda et al. Progress in Earth and Planetary Science(2020) 7:44

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Fig. 8 Comparison of the M–B boundary observed from the Chiba composite section and the TB-2 sediment core. VGP latitudes based onthermal demagnetization in Suganuma et al. (2015) are shown as blue cross marks. VGP latitudes based on the hybrid demagnetization method
are shown as black closed circles and lines (Okada et al. 2017) and red open circles and lines (this study). The TB-2 sediment core was drilled 190
m northeast of the Chiba section, and VGPs obtained from the core as reported by Hyodo et al. (2016) are shown as blue closed circles and lines(AFD) and red closed circles and lines (ThD). Lithology of the TB-2 sediment core was from Hyodo et al. (2016, 2017). Blue, black, and red
horizontal bars indicate the sampling level of Suganuma et al. (2015), Okada et al. (2017), and this study, respectively. Black and red dashed lines
show the mid-thickness of polarity switch indicating the M–B boundary and the correlation of the volcanic ash

The NRM30-50/ARM30-50 ratio as a relative paleointensity index was obtained from the 16 specimens in which
the ChRM component MAD was ≤ 15° (Fig. 7e–f; Additional file 5: Table S3). The depletion of the relative
paleointensity occurs 0.2 m above the Byk-E tephra bed,
although the values are already lower than the averagefor the whole Chiba composite section (Fig. 7e). Theinterval of minimum relative paleointensity is observed
from 3.0 m above the Byk-E tephra bed and is followedby subsequent directional instabilities, a phenomenon
also observed in the Chiba section (Okada et al. 2017).
3.4 Comparison of the M–B boundary in the Chiba
composite section

In this study, the M–B boundary is defined as a midpoint of the polarity switching interval from a reversed
polarity (VGP latitude < − 45°) to a normal polarity(VGP latitude > 45°). An alternative definition is thestratigraphic position at which the VGP latitude = 0.
However, using this definition, the M-B boundary is
identified using only two intermediate polarity data
points, in the case of the Chiba composite section.
Therefore, we adopt the former definition of the M–B
boundary, which is more robust.
The M–B boundary has been identified in the Yanagawa(Suganuma et al. 2015), Chiba (Okada et al. 2017), andYoro-Tabuchi (this study) sections of the Chiba composite
section, and in the TB-2 sediment core (Hyodo et al.
2016) (Fig. 8; Table 1). The depth scale of the TB-2 sediment core was converted into the equivalent thickness
scale of the Chiba composite section, using 7 tie points
based on lithology and tephrostratigraphy (indicated by
red solid lines in Fig. 3). The average polarity switch occurs in an interval between 0.6 ± 0.3 and 1.6 ± 0.4 m (1σ)
above the Byk-E tephra bed (Table 2). The average midpoint of the polarity switch in the three sections and sediment core is 1.1 ± 0.3 m (1σ) above the Byk-E tephra bed(Table 2). Based on the age model of Suganuma et al.
(2018), the M–B boundary is dated at 772.9 ± 0.4 ka (1σ)
and the polarity switch is completed within 1.1 ± 0.4 kyrs(1σ), although the field intensity minimum and directional
instability follow it (Fig. 8; Simon et al. 2019). Assuming a
potentially chronological uncertainty of 5 ka in orbital tuning methods (Simon et al. 2019; Valet et al. 2019), theaverage age of the M–B boundary in the Chiba composite
section is estimated as 772.9 ± 5.4 ka (1σ). This is consistent with the inferred ages obtained from sedimentary archives (e.g., Channell et al. 2010; Valet et al. 2019).
3.5 Millennial-scale VGP deviation and its distribution

The time evolution of the VGP latitude and the VGP locations in the Chiba composite section are shown inHaneda et al. Progress in Earth and Planetary Science(2020) 7:44

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Table 2 Summary of the M–B boundary in the Chiba composite section
Stratigraphic distance from the Byk-E (m)Age (ka)1

Start

End

Mid-point

Start

End

Mid-point

Duration

Yoro-Tabuchi/hybrid

1.07

2.09

1.58

772.96771.79772.38

1.17This study

Yanagawa (and Chiba)/ThD

0.35

1.25

0.80

773.75772.73773.24

1.02

Suganuma et al. (2015)

Chiba/hybrid

0.35

1.85

1.10

773.75772.06772.91

1.69

Okada et al. (2017)
Hyodo et al. (2016)

Section/Demag. method

2

Reference

TB2 sediment core /ThD

0.68

1.23

0.96

773.37772.76773.07

0.61

Average

0.6

1.6

1.1

773.5

772.3

772.9

1.1

SD-1σ

0.3

0.4

0.3

0.4

0.5

0.4

0.4

1

Age model of Suganuma et al. (2018) is used to calculate the ages of the M–B boundary and the durations
2
Depth scale of the TB-2 sediment core is converted into thickness scale in the Chiba composite section based on the lithology and tephrostratigraphy

Figs. 9, 10, and 11 alongside the relative paleointensity
index (NRM30-50/ARM30-50) and Be ratio-derived dipole
moment (BeDM) profiles. Each temporal profile was calculated based on the latest age model of the Chiba composite section by Suganuma et al. (2018). The longest
section record shown in Fig. 9 is integrated with datafrom the Chiba composite section and new results fromthe Yoro River section above 37.5 m, which were obtained using the hybrid demagnetization method. The
new record from the lowermost part of the YoroTabuchi section spans between 774.1 and 769.9 ka,
which corresponds to the period from the polarity
switch to the early phase of the directional instability(Fig. 10). The results of the progressive thermal
demagnetization from the Chiba and Yanagawa sections
show the VGP path and the VGP latitude profile between 787.1 and 767.7 ka (Fig. 11; Suganuma et al.
2015). The VGP rebounds, which are large deviations
from the north pole, after the M–B boundary are directional instability clusters (IC1, IC2, and IC3 in Figs. 9,
10, and 11; Simon et al. 2019).
The VGP latitude profiles show millennial-scale
fluctuations with intermittent depletions in relative
paleointensity during periods where the BeDM waslower than its modern level (Fig. 9; Simon et al.
2019). The VGPs in the Chiba composite section
prior to the M–B boundary intermittently deviate at
ca. 783, 780, and 775 ka and mainly plot in the southeastern Pacific (Figs. 9 and 11). The VGPs thenmainly shift from the Southern Hemisphere to Northern Hemisphere at ca. 773 ka without the longitudinal
preference of the VGP path suggested by Clement(1991) and Laj et al. (1991). The VGPs of the directional instability clusters are mainly located in North
America and the northeastern Pacific and are scattered across Eurasia and the western Pacific (Fig. 9,10, and 11). After the termination of IC3 at ca. 763
ka, the VGPs is mostly stable north of 70°N.
Millennial-scale geomagnetic field instability has been
identified as a characteristic feature of the geomagnetic
reversal process (e.g., Valet et al. 2012; Balbas et al.

2018; Singer et al. 2019). Valet et al. (2012) proposed thepresence of three successive phases in both reversed–
normal and normal–reversed transitions during the last
180 Ma. Based on ten volcanic records, they suggested
phasing of (i) a precursor event, (ii) a 180° polarity
switch, and (iii) a rebound. Based on accurate 40Ar/39Ar
dating of lava flows, Singer et al. (2019) have argued that
the M–B geomagnetic reversal began in association witha large VGP deviation and field depletion at ca. 795 ka, a
process recording the dominance of a non-axial dipole
field, due to the collapse of the axial dipole. The intervalof VGP fluctuation and depletion of the relative paleointensity at 783 ka in the Chiba composite section are consistent with the geomagnetic field instability recorded inlava flows in Chile and Guadeloupe at ca. 784 ka (Fig.
12; Singer et al. 2019). The VGP at ODP sites 983 and984 in the North Atlantic shows repeatedly large fluctuation, which begins from ca. 780 and 783 ka, respectively(Fig. 12b; Channell and Guydo 2004). These VGP fluctuations observed in North Atlantic cores may correspondto the VGP fluctuations in the Chiba composite section
seen from 783 ka (Fig. 12).
Simon et al. (2019) reported the succession of large
VGP deviations and low BeDM between 773.9 and
768.5 ka in the Chiba composite section corresponding
to the interval between the polarity switch and IC1. The
large VGP deviations just after the M–B boundary agreewith the polarity switch and subsequent rebound phasesof the successive reversal as modeled by Valet et al.
(2012) (Simon et al. 2019). The directional instability associated with the geomagnetic reversal terminated ca. 3
kyr prior to the field recovery at ca. 760 ka. After thefield recovery, following the termination of IC3, theVGP profile is generally stable, showing only occasional
deviations. Because these deviations consist of single
data points and are not accompanied by large decreases
in the relative paleointensity, it is reasonable to consider
them as noise. The relative paleointensity after the fieldrecovery is higher than that before the field decay, although the rock magnetic properties are relatively stable(Figs. 7 and 12). This field asymmetry across the M-B

Haneda et al. Progress in Earth and Planetary Science(2020) 7:44

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Fig. 9 VGP paths (top plot), paleomagnetic intensity indices (middle plot), and time evolution of the VGP latitude (bottom plot) of the CbCS. The
integrated results of the hybrid demagnetization method from the Yoro River and Yoro-Tabuchi sections (Okada et al. 2017; Simon et al. 2019;
this study). The relative paleointensity is based on the newly calculated NRM30-50/ARM30-50 reported in this study. The BeDM profile, shown by ablue band, is from Simon et al. (2019). Blue vertical shades and gray vertical dashed line indicate millennial-scale VGP fluctuations and the M-B
boundary, respectively. Ages of the Byk-A and Byk-E tephra beds are indicated by red arrows. IC, instability cluster; PS, polarity switch

boundary is consistent with the virtual axial dipole moment series from the relative paleointensity global stacks(Fig. 12c; Valet et al. 2005; Channell et al. 2009). The
paleomagnetic record of the Chiba composite section
shows the full sequence of the M–B geomagnetic reversal between 783 and 763 ka, indicating the dominant
contribution of the non-axial dipole field, during the depletion of the main axial dipole, at ca. 20 kyrs interval
across the M–B boundary at ca. 773 ka.

4 Conclusions
Paleomagnetic and rock magnetic study of the Chiba
composite section located in the Boso Peninsula, Central
Japan, allows reconstruction of the full sequence of theM–B geomagnetic reversal in the interval between 794
and 748 ka. This covers the entire interval of depletionof the geomagnetic field intensity, including the polarity
switch and preceding transitional phase. We have defined an average stratigraphic position and age of theHaneda et al. Progress in Earth and Planetary Science(2020) 7:44

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Fig. 10 VGP paths (top plot), paleomagnetic intensity indices (middle plot), and time evolution of the VGPlatitude (bottom plot) fromthe lowermost part of the Yoro-Tabuchi section. The relative paleointensity is based on the NRM30-50/ARM30-50. The BeDM profile, shown by ablue band, is from Simon et al. (2019). Blue vertical shades and gray vertical dashed line indicate millennial-scale VGP fluctuations and the M-B
boundary, respectively. Ages of the Byk-A and Byk-E tephra beds are indicated by red arrows. IC, instability cluster; PS, polarity switch

M–B boundary between three (sub) sections in the
Chiba composite section and a neighboring drilled core
based on newly obtained and previously reported data.
The VGP latitude and relative paleointensity results
show a clearly defined polarity switch, with a geomagnetic field minimum, 1.6 m above the Byk-E tephra bedin the Yoro-Tabuchi section. This is consistent with thepreviously reported paleomagnetic data from the Chiba
composite section. The average stratigraphic position of

the M–B boundary between the Chiba, Yoro-Tabuchi,
and Yanagawa sections, and the TB-2 sediment core, lies
1.1 ± 0.3 m (1σ) above the Byk-E tephra bed. Accounting
for uncertainty due to a 5 kyr chronological error withthe orbital tuning and a stratigraphic deviation of 0.4
kyr, the age of the M–B boundary in the Chiba composite section is estimated at 772.9 ± 5.4 ka (1σ).
The results from the upper part of the Yoro River section exhibit homogeneous rock-magnetic properties andHaneda et al. Progress in Earth and Planetary Science(2020) 7:44

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Fig. 11 VGP paths (top plot) and time evolution of the VGP latitude (bottom plot) from the Chiba and Yanagawa sections. Data are based on the
progressive thermal demagnetization of Suganuma et al. (2015). The temporal profile was calculated from the most recent age model of
Suganuma et al. (2018). Blue vertical shades and gray vertical dashed line indicate millennial-scale VGP fluctuations and the M-B boundary,respectively. Ages of the Byk-A and Byk-E tephra beds are indicated by red arrows. IC, instability cluster; PS, polarity switcha higher field intensity than that before the M–B boundary, with stable normal polarity directions. Integrating
the previous paleomagnetic results from the Chiba composite section shows the geomagnetic direction and relative paleointensity repeatedly deviate from the norm
between ca. 783 and 763 ka. This suggests that the contribution of the non-axial dipole field was dominant during the depletion of the main axial dipole and spannedat least 20 kyr across the polarity switch at ca. 773 ka.
Our results provide a detailed and expanded sedimentary record of the M–B geomagnetic reversal and offer
valuable new information to further understand themechanisms and dynamics of geomagnetic reversals.

5 Supplementary information
Supplementary information accompanies this paper at https://doi.org/10.
1186/s40645-020-00354-y.

Additional file 1: Table S1. Results of the thermal demagnetizationfrom the Chiba and Yanagawa sections reported by Suganuma et al.
(2015). Age is recalculated from the latest age model of the CbCS by
Suganuma et al. (2018).
Additional file 2: Table S2. Paleomagnetic and rock magnetic results
for the Yoro River and Yoro-Tabuchi sections by Okada et al. (2017), Simon et al. (2019), and this study.
Additional file 3: Figure S1. Comparison between new and previous
relative paleointensity records of the CbCS.
Additional file 4: Figure S2. Comparison of the relative paleointensity
indices, normalizers, and authigenic 10Be/9Be ratio: (a) Anhysteretic
remanent magnetization (ARM) intensity against NRM30-50/ARM30-50; (b)Saturation isothermal remanent magnetization (SIRM) intensity against
NRM30-50/ARM30-50; (c) Authigenic 10Be/9Be against NRM30-50/ARM30-50; (d)Authigenic 10Be/9Be against NRM30-50/IRM30-50. The result of the IRM
experiment of TB40 was not used because of its large deviations.
Additional file 5: Table S3. Paleomagnetic and rock magnetic results
for the lowermost part of the Yoro-Tabuchi section obtained by thisstudy.

Haneda et al. Progress in Earth and Planetary Science(2020) 7:44

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Fig. 12 Comparisons between paleomagnetic records. a VGP latitude from the Chiba composite section. The black line and cross marker are the
VGP profile, and the duration and mid-point of the polarity switch from the Chiba, Yoro-Tabuchi, and Yoro River sections (Okada et al. 2017;
Simon et al. 2019; this study). The red line and cross marker indicate the VGP profile and duration and mid-point of the polarity switch from the
Yoro-Tabuchi section, respectively (this study). The light blue line and cross marker indicate the VGP profile and duration and mid-point of thepolarity switch from the Chiba and Yanagawa sections, respectively (Suganuma et al. 2015). b VGP latitude based on marine sediment cores in
the North Atlantic (green: ODP 983, orange: ODP 984 in Channell and Guydo 2004) and lavas and volcanic ash beds (Singer et al. 2019).
Horizontal bars indicate chronological errors (2σ) of 40Ar/39Ar dating. c Proxies of field intensity. Black and red lines indicate profiles of the
NRM30-50/ARM30-50 as the relative paleointensity index from the Chiba, Yoro-Tabuchi, and Yoro River sections (Okada et al. 2017; Simon et al. 2019;
this study) and the Yoro-Tabuchi section (this study), respectively. The profiles of virtual axial dipole moment (VADM) from the global relative
paleointensity stack records are indicated by orange (Channell et al. 2009) and green dashed lines (Valet et al. 2005)

Abbreviations
AF demagnetization: Alternating field demagnetization; ARM: Anhysteretic
remanent magnetization; BeDM: Be ratio-derived dipole moment;
CbCS: Chiba composite section; ChRM: Characteristic remanent
magnetization; DRM: Detrital remanent magnetization; IC: Instability cluster;
IRM: Isothermal remanent magnetization; MAD: Maximum angular deviation;
M–B boundary: Matuyama–Brunhes boundary; M–B geomagnetic
reversal: Matuyama–Brunhes geomagnetic reversal; MIS: Marine isotope
stage; NRM: Natural remanent magnetization; ODP: Ocean Drilling Program;

SIRM: Saturation isothermal remanent magnetization; VGP: Virtual
geomagnetic pole; kARM: ARM susceptibility; kLF: Low-field magnetic
susceptibility
Acknowledgements
We would like thank Yuka Shimota, Keisuke Yamamoto, and Nanako
Yokokawa for collecting samples and Masayuki Hyodo for sharing the data of
the TB-2 sediment core. Valuable comments made by Quentin Simon greatly
improved the paper. We also greatly appreciate the kind support given by(2020) 7:44

Page 20 of 22

the Tabuchi district of Ichihara City, Chiba Prefecture, the Agency for Cultural
Affairs, and the Ministry of Education, Culture, Sports, Science and Technology. We would like thank Editage (www.editage.jp) for the English language
editing. This paper greatly improved by constructive comments by reviewers
Mark W. Hounslow and Chuang Xuan, and editor Ryuji Tada. The sample collection within the Natural Monument site was performed with the permission and under the witness of Ichihara City, the administrator of the site.

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Haneda et al. Progress in Earth and Planetary Science

Authors’ contributions
YH carried out the experimental study and data analyses and prepared themanuscript. OM and YS proposed the topic, designed the study, assisted in
the data analyses, and prepared the manuscript. TK performed the rockmagnetic analysis. All authors read and approved the final manuscript.
Authors’ information
YH is a post-doctoral fellow at the Geological Survey of Japan, AIST, Japan, inpresent, and has belonged to National Institute of Polar Research (NIPR),
Japan, during conducting the experimental study; MO is a professor at Ibaraki University, Japan; YS is an associate professor at NIPR and at the Graduate
University for Advanced Studies (SOKENDAI), Japan; and TK is a master’scourse student at the Graduate School of Science and Engineering, Ibaraki
University, Japan.
Funding
This study was supported by a JSPS KAKENHI grant (16H04068, 17H06321,
and 19H00710) awarded to MO and YS, a NIPR Advanced Project (KP-7 and
KP301) awarded to YS, and a Geoscience Research Grant from the Tokyo
Geographical Society awarded to YH.
Availability of data and materials
Please contact author for data requests.
Competing interests
The authors declare that they have no competing interest.
Author details
1
National Institute of Polar Research, 10-3 Midori-cho, Tachikawa, Tokyo
190-8518, Japan. 2Department of Earth Sciences, Ibaraki University, 2-2-1
Bunkyo, Mito, Ibaraki 310-8512, Japan. 3Present Address: Geological Survey ofJapan, AIST, 1-1-1 Central 7, Higashi, Tsukuba, Ibaraki 305-8567, Japan.
4
Department of Polar Science, School of Multidisciplinary Sciences, The
Graduate University for Advanced Studies, SOKENDAI, Midori-cho 10-3,
Tachikawa, Tokyo 190-8518, Japan.
Received: 29 September 2019 Accepted: 19 July 2020

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