Hyodo et al. Progress in Earth and Planetary Science
https://doi.org/10.1186/s40645-020-00337-z(2020) 7:26

Progress in Earth andPlanetary Science

RESEARCH ARTICLE

Open AccessA centennial-resolution terrestrial
climatostratigraphy and Matuyama–
Brunhes transition record from a loesssequence in China
Masayuki Hyodo1,2* , Kenta Banjo2, Tianshui Yang3, Shigehiro Katoh4, Meinan Shi5, Yuki Yasuda1,6,
Jun-ichi Fukuda1,2, Masako Miki2 and Balázs Bradák7

Abstract
Terrestrial records of the last geomagnetic reversal often have few age constraints. Chronostratigraphy using
suborbital-scale paleoceanic events during marine isotope stage 19 may contribute to solving this problem. We
applied the method to an 8 m long, high-resolution paleomagnetic record from a loess sequence in China andrevealed millennial-to-sub-centennial scale features of the Matuyama–Brunhes (MB) transition. All samples were
subjected to progressive thermal demagnetization with 14–15 steps up to 650–680 °C. As a result, 96% of thesamples yielded a high-quality remanent magnetization. The MB transition terminated with a 75 cm thick zone withnine polarity flips. The polarity flip zone, dated at about 779–777 ka, began between the warm events “I” and “J”
and terminated at the end of the cooling event coincident with the lowest axial-dipole strength interval. Most
polarity flips occurred within 70 years. The virtual geomagnetic poles (VGPs) in the upper polarity flip zone clusteredin the SW Pacific region, where the MB transitional VGPs from lavas of the Hawaiian and Canary Islands andlacustrine deposits of Java also clustered. These sites were probably dominated by dipolar fields. The absence oftransitional fields across polarity flips implies a short time span for averaging fields due to a thin loessmagnetization lock-in zone. The reverse-to-normal polarity reversal dated at about 778 ka in Lingtai occurred at theend of the SW Pacific VGP zone, an important key bed for MB transition stratigraphy. The reversal is a goodcandidate for the main MB boundary. We found an excursion at about 766 ka spanning about 1 ka.
Keywords: Loess–paleosol, Geomagnetic reversal, Early–Middle Pleistocene transition, Loess magnetism, East
Asian monsoon

Introduction
Pleistocene marine and terrestrial sequences can provide
continuous records of geomagnetic field variations aswell as climate changes, on various time scales. Therefore, these records are useful for constructing “magnetoclimatostratigraphy,” a hybrid stratigraphy that enables
more precise stratigraphic correlations on global scales.
* Correspondence: mhyodo@kobe-u.ac.jp
1
Research Center for Inland Seas, Kobe University, Kobe 657-8501, Japan2
Department of Planetology, Kobe University, Kobe 657-8501, Japan
Full list of author information is available at the end of the article

For the marine isotope stage (MIS) 19 interglacial
period, decadal-to-centennial-resolution paleoceanic andpaleoclimate variation records have been obtained frommarine sediments (Kitaba et al. 2009, 2013, 2017; Ferretti
et al. 2015; Maegakiuch et al. 2016; Hyodo et al. 2017),
from which a number of centennial-to-millennial scale
climate events have been found (Hyodo et al. 2017). In
addition, many high-resolution Matuyama–Brunhes(MB) magnetic polarity transition records from various
regions have revealed submillennial-scale features ofgeomagnetic field changes (e.g., Wang et al. 2006; Hyodo© 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) 7:26

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et al. 2006, 2011; Channell et al. 2010; Okada et al.
2017). However, the detailed geomagnetic records havenever been analyzed together with the centennial-scale
climate changes. This may partly be a result of the lockin of magnetization of sediments being incompletely
understood (e.g., Hyodo 1984; deMenocal et al. 1990;
Suganuma et al. 2010).
Loess–paleosol sequences in the Chinese Loess Plateau(CLP) are excellent archives of geomagnetic field variations because loess–paleosol layers have stable remanent
magnetizations (e.g., Zhu et al. 1994; Yang et al. 2004,
2007, 2008). In addition, orbital-scale correlation of loessand paleosol units with marine oxygen isotope stages(MISs) has almost been established (Kukla 1987; Sunet al. 2006). Each paleosol layer is numbered and correlated, such as paleosol S0 (MIS 1), S1 (MIS 5), and S2(MIS 7). Furthermore, the Dansgaard–Oeschger andBond events recorded in loess–paleosol sequences(Porter and An 1995; Wang et al. 2005) may enable theconstruction of even millennial-scale climatostratigraphy.
The MB magnetic polarity transition, the last geomagnetic reversal, has been investigated often using loess–paleosol sequences (e.g., Spassov et al. 2001; Liu et al.
2008; Jin and Liu 2011; Yang et al. 2010). Many highresolution loess records of the transition are characterized by multiple, rapid polarity flips (Wang et al. 2006;
Yang et al. 2010; Spassov et al. 2001). However, thesequential polarity flips have seldom been correlated
between distant sites (Yang et al. 2010). The different
numbers of flips in each record may be one of the reasons. Therefore, the question of whether the rapid polarity flips are even real geomagnetic phenomena has been
discussed (Jin and Liu 2010). Nonetheless, the multiplepolarity flips during the MB transition are also evidenced
by terrestrial sequences of lava flows (Coe et al. 2004;
Singer et al. 2005) and marine and lacustrine sediments(e.g., Channell and Kleiven 2000; Channell et al. 2010;
Hyodo et al. 2006, 2011). Another unresolved problem isthat there are different climatostratigraphic interpretations
for the MB transition, which lies in the upper part of the S8
paleosol layer. The S8 paleosol layer is mainly correlated in
two ways: to MIS 21 (Ding et al. 2002; Sun et al. 2006; Haoet al. 2012) and to MIS 19 (Liu et al. 2008; Wang et al. 2006;
Ueno et al. 2019). The different correlations are strongly related to the lock-in depths of magnetization of the loess.
These problems of loess–paleosols may be resolved by combination with high-resolution climatostratigraphy.
In this study, a high-resolution and high-quality
paleomagnetic record was obtained from an 8 m thickloess–paleosol sequence in Lingtai, central CLP. We
evaluated the timing and millennial-to-centennial scale
behaviors of the MB transition in combination with the
published centennial-resolution paleoclimate record
from the same section (Ueno et al. 2019).

Methods/Experimental
Oriented block samples (each 15–20 cm high) were collected from a 7.8 m section covering the stratigraphy ofthe loess L9, paleosol S8, loess L8, and paleosol S7 layers(in ascending order). In the laboratory, each block was cut
into cubes, 2.2 cm on each side, using a knife and thendressed into 2 × 2 × 2 cm3 specimens using sandpaper.
Specimens were collected at 2.5 cm depth intervals
around the magnetic polarity transition interval and at 6
cm depth intervals on average throughout the section. In
total, 130 specimens were taken for paleomagnetic measurements, and 61 specimens were taken for rock magnetic experiments.
All specimens were subjected to progressive thermal
demagnetization (THD) at intervals of 30–50 °C up to680 °C. Natural remanent magnetization (NRM) of specimens was measured using a 2G cryogenic magnetometer.
Low-field magnetic susceptibility (χ) was measured using amagnetic susceptibility meter (SM-100, ZH Instruments,
Brno, Czech Republic). The anisotropy of low-field magnetic susceptibility (AMS) was measured using a KLY-3S
Kappabridge susceptometer (AGICO, Brno, Czech Republic). Anhysteretic remanent magnetization (ARM) was
imparted with a peak alternating field of 100 mT superimposed on a direct current biased field of 50 μT.
We performed thermomagnetic analyses in an air atmosphere on selected samples using an NMB-89 magneticbalance (Natsuhara Giken Corporation, Osaka, Japan). A
sample was heated from 50 to 700 °C and then cooled to
50 °C at a rate of 10 °C/min. Isothermal remanent
magnetization (IRM) acquisition, hysteresis experiments,and first-order reversal curve (FORC) measurements (Pikeet al. 1999; Roberts et al. 2000) were performed for selectedbulk loess and paleosol samples using a VSM MicroMag
3900 (Princeton Measurements Corporation, Westerville,
OH, USA). The IRM was measured at 100 levels from 0.1
to 1000 mT, and the data were analyzed using IRMUNMIX
2.2 software (Heslop et al. 2002). FORC distributions were
processed using FORCinel 1.18 software (Harrison andFeinberg 2008) with a smoothing factor of 3–5.

Hyodo et al. Progress in Earth and Planetary Science

Results
Thermomagnetic analysis, IRM spectra, and AMS

Representative thermomagnetic (J–T) curves for samples
from loess and paleosol layers (Fig. 1a, b) and the transition zone (Fig. 1c, d) exhibit the same features. All samples show magnetization reductions at approximately
580 and 680 °C, indicating the presence of magnetiteand hematite, respectively. In addition, the inflections in
the heating curves at around 300 °C and the absence ofcorresponding signals in the cooling curves imply thepresence of maghemite. The first derivatives of the heating curves (Fig. 1a–d) show that maghemite decomposes
mainly between 240 and 340 °C; it is then oxidized toHyodo et al. Progress in Earth and Planetary Science(2020) 7:26

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Fig. 1 Results of rock magnetic experiments and anisotropy of magnetic susceptibility measurements. a–d Temperature dependence ofmagnetization normalized by an initial value (J). The red/blue curves show heating/cooling. –dJ/dT indicates the first derivative of J–T heating
curves calculated at intervals of 10 °C. e The loss of total magnetization versus magnetic susceptibility (χ). f The magnetization loss by maghemite
decomposition versus χ, calculated as a sum of the first derivatives of the J–T curves between 240 and 340 °C subtracted by the backgroundmagnetization decrease of magnetite, a mean of the first derivatives between 140 and 240 °C. g, h Results of isothermal remanent magnetization
component analysis for loess and paleosol. i Anisotropy of magnetic susceptibility. The blue solid square/red open triangle/green open circle
symbols represent the maximum/intermediate/minimum axes of magnetic susceptibility, respectively

(2020) 7:26

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form hematite. The loess and paleosols contain coarse,
detrital magnetite particles with ultrathin maghemite
shells formed by low-temperature oxidation in thesurrounding air (van Velzen and Dekkers 1999; Liu et al.
2003, 2004). The inflection in the paleosol sample heating curve at around 300 °C is smaller than that in theloess sample heating curve, which is more clearly seen in
the first derivatives of the heating curves (Fig. 1a, b).
Thus, the relative content of maghemite in the paleosol
is much smaller than that in the loess.
The J–T curves show clear reductions in sample
magnetization after heating–cooling treatments, whichreflect the high-temperature oxidation of maghemite
and fine-grained magnetite to form weakly magnetic
hematite. The reductions range from 27 to 46% (Fig. 1).
The reduction levels depend on the χ values of the samples, as the smallest χ sample (nearly pristine loess) had
the largest reduction, and the largest χ sample (mature
paleosol) had the smallest reduction (Fig. 1e). Similar χdependence is observed in the reduction component ofmaghemite decomposition estimated using the first derivatives of the J–T heating curves (Fig. 1f). These results
imply that pedogenic fine-grained magnetites, whichmust be dominant in paleosols, contributed little to the
reduction. It is plausible that the reduction was attributable mainly to thermal oxidation of fine-grained detritalmagnetites and thermal decomposition of maghemites,
most of which were present as shells around coarse detrital magnetite grains. The present results indicate anear absence of pedogenic maghemites, consistent withprevious studies that showed that most of the maghemite in the paleosol samples is detrital (Spassov et al.
2003; Maher 2016; Hyodo et al. 2020). Therefore, thesuperfine (single-domain (SD)/superparamagnetic (SP))
ferrimagnets that must be contained in the mature
paleosols probably underwent neither maghemitization
nor thermal oxidation. This inference is consistent with
the TEM observation that pedogenic magnetites formed
authigenically within phyllosilicate grains, such that silicates protect the authigenic magnetites from oxidation(Hyodo et al. 2020).
The IRM was unmixed into three components withgradient maxima of approximately 30, 100, and > ~200mT of B1/2 for both the loess and paleosol bulk samples(Fig. 1 g, h). The only difference between the loess andpaleosol samples was their relative contributions. The
low B1/2 component (~ 30 mT) was predominant in the
paleosol samples, whereas the medium B1/2 component(~ 100 mT) was predominant in the loess samples. The
high B1/2 component (> ~ 200 mT) was moderately
represented in both the loess and paleosol samples. The
Bcr of SD–pseudo-single-domain (PSD)-sized intergrown magnetites ranges from approximately 20 to 50
mT (Heider et al. 1996). Considering the magnetic

carriers estimated from the thermomagnetic analyses,
the low B1/2 component must be dominated by pedogenic magnetites. The high B1/2 component comprises
detrital and pedogenic hematites, based on the TEM
observations that revealed the presence of interlayer
authigenic hematite particles within phyllosilicate grains(muscovite and chlorite), in addition to discrete detrital
hematite particles between silicate grains (Hyodo et al.
2020). The medium B1/2 component comprises coarse
detrital magnetites with maghemite rims. The higher
B1/2 values (~ 100 mT) for the coarse magnetites would
be caused by the maghemite rim, which increases thecoercivity of magnetite particles due to the enhanced
stress induced by the oxidation gradient between themaghematized rim and the magnetite core (van Velzen
and Zijderveld 1995). The correlation of the magnetic
minerals based on the J–T curves with the IRM components in this study is consistent with the correlation
based on the same experiments using many more
samples from different regions (Hyodo et al. 2020).
The AMS data show that the maximum axes represent
nearly horizontal directions and that the minimum axes
represent vertical ones (Fig. 1i). Therefore, the loess–paleosol sequence underwent little disturbance of
sediment structure. If they formed on silicate surfaces or
between silicate grains, the crystal growth directions of
pedogenic magnetic particles would be independent of
the depositional plane. The observed horizontal maximum axes (Fig. 1i) indicate that pedogenic magnetites
consist mainly of authigenic ones grown along layers ofphyllosilicates deposited horizontally.

Hyodo et al. Progress in Earth and Planetary Science

Hysteresis loops and FORC diagrams

The hysteresis loops of samples with different magnetic
susceptibilities show slight differences in magnetic coercivity (Hc) and loop shape (Fig. 2) between loess andpaleosol samples. The higher Hc of the loess samples,
dominated by detrital magnetic particles (Fig. 2a), would
be caused mainly by detrital magnetites with a maghemite shell. The slightly “wasp-waisted” shape of the loop
for the paleosol samples, dominated by pedogenic magnetic particles (Fig. 2b, c), may be caused by a mixtureof two magnetic components with strongly contrasting
coercivities (Roberts et al. 1995). The two components
are probably pedogenic magnetites and hematites. The
FORC diagrams of both the loess and paleosol samples
show a peak at < 20 mT, which indicates a dominanceof SD particles, and also show modestly divergent distributions at low coercivities, which indicates the presenceof vortex-state particles (Fig. 2). The slightly higher coercivity and broader vertical spread for the loess sample isprobably the effect of coarse detrital magnetites with amaghemite shell. In contrast, the lower coercivity andsimilar FORCs (Fig. 2b, c) of the paleosol reflect theHyodo et al. Progress in Earth and Planetary Science(2020) 7:26

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Fig. 2 Results of hysteresis and FORC measurements. The hysteresis curves and FORC diagrams are for loess (a) and paleosol (b, c) samples

dominance of pedogenic magnetites. The asymmetrical
diverging contours lying close to the origin in all FORCs
are likely due to SP particles (Pike et al. 2001; Roberts
et al. 2000).
Progressive thermal demagnetization

As shown by the typical results of progressive THD (Fig. 3a,b), most of the samples have NRMs that show quite linear
vectorial decay toward the origin with temperatures ranging
from 200–300 to 620–680 °C. The range of temperature implies that the NRM is carried by magnetite and hematite,and each mineral component has the same direction. Such
NRM data were obtained from about 70% of the samples,
all collected from the full-normal or full-reverse polarity
zones, independent of the χ or the degree of pedogenesis.
The remaining samples mostly had a single NRM component with a temperature range from 200–300 to 530–590 °C
carried by magnetite (Fig. 3c, d). These samples had no
isolatable higher-temperature (> 590 °C) component carriedby hematite. Several samples had multi-component NRMs
(Fig. 3e–i) with weak intensity. These samples were all froma magnetic polarity flip zone and just above it. Five samples
had a quite low-intensity NRM that showed no linear decayof the magnetic vector. These samples were also from themagnetic polarity flip zone, or just below it.
Paleomagnetic directions

Characteristic remanent magnetization (ChRM) directions
were determined by principle component analysis using the
THD data of 125 samples, 2 of which were from the same
depth (197 cm). Some of the results are listed in Table 1.
The average number of demagnetization steps of the data
used for the analysis was 10.4, and the mean maximum angular deviation (MAD) (Kirschvink 1980) was 3.6. Most of
the ChRMs were carried by magnetite (Fig. 3a–d). Of the
125 samples, 4 had another component carried mainly byhematite (Fig. 3e–h). The two components had relatively
close directions (Fig. 3f–h), except for the sample at 337
cm, which had magnetite-carrying normal-polarity andhematite-carrying reverse-polarity components (Fig. 3e).
The sample was located at a depth underlain by a shortreverse-polarity episode and overlain by a normal-polarity
succession (Table 1). From this situation, the reversepolarity hematite-carrying component was probably acquired first, followed by acquisition of the normal-polarity
magnetite-carrying component. This delay in acquisition

Hyodo et al. Progress in Earth and Planetary Science(2020) 7:26

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Fig. 3 Vector end-point diagrams. The solid/open circles show projections on the horizontal/vertical planes, respectively. Red lines in a, b show
single normal/reverse components obtained from most of samples. In c to h, blue/orange lines show magnetite-/hematite-carried components,
respectively. i Pink/blue/orange arrows show low/middle/high-temperature components, respectively

would reflect the time to the adjacent sample (5 cm
distance in this case). The former should be the resultof detrital remanent magnetization (DRM), and thelatter should be the result of pedogenic remanent
magnetization (PRM). The sample at 357.5 cm comprised three components with temperature ranges of
250–450 °C (low), 450–560 °C (middle), and 560–650 °C(high) (Fig. 3i). The low/middle- and high-temperaturecomponents would be carried mainly by magnetite andhematite, respectively. The low-temperature magnetitecarrying component and the hematite-carrying component
had similar low-inclination reverse-polarity directions,which probably indicate DRM. On the other hand, themiddle-temperature magnetite-carrying component had alow-inclination normal-polarity direction, which probably
indicates PRM (Fig. 3i). The depth of the three-component
sample was 2.5 cm beneath/above normal-polarity horizons(Table 1). The reverse-polarity field carried by the
DRM should reflect a brief episode that occurred
within quite a short time (< 5 cm thickness). For thisdepth, we calculated a composite ChRM of low- andhigh-temperature components.
The ChRM directions and virtual geomagnetic pole
(VGP) latitudes clearly define polarity boundaries(Fig. 4a–c). The reverse-polarity zone below 415 cm
is overlain by a mixed-polarity zone consisting of 28

(2020) 7:26

Hyodo et al. Progress in Earth and Planetary Science

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Table 1 Results of paleomagnetic analysis
Depth

Age

T1

T2

(cm)

(ka)

(°C)

(°C)

No of steps

154.0

764.37

200

680

11

166.5

764.69

250

650

176.0

764.93

200

400

MAD

Dec

Inc

VGP lat.

VGP lon.

(°)

(°)

(°)

(°)

1.6

7.7

48.0

81.1

− 122.1

11

4.1

− 7.5

55.6

83.8

29.8

5

13.3

− 4.4

− 25.441.4− 67.0197.0765.55

200

560

9

7.1

51.1

− 64.3− 4.7− 109.0

197.0

765.55

250

530

7

5.9

56.7

− 56.80.5− 114.2

204.5

765.77

200

650

12

6.8

144.1

2.5

− 40.6

157.9

209.5

765.91

200

590

10

3.1

2.8

55.3

87.6

177.4

212.0

765.98

200

590

10

5.6

− 4.1

46.9

82.2

− 44.7230.0

766.50

250

620

11

2.6

− 4.4

50.9

85.0

− 24.5247.0

767.99

200

590

10

3.9

0.5

54.5

89.6

− 160.4

258.0

768.95

200

560

7

1.9

− 5.2

48.0

82.6

− 34.9271.0

770.09

350

620

8

2.9

− 5.7

49.7

83.4

− 24.3281.0

771.32

200

620

11

2.7

1.1

55.0

89.0

168.5

290.0

772.48

450

680

8

6.4

18.5

32.5

66.1

− 120.8

300.0

773.76

200

590

10

5.7

− 16.1

46.8

74.6

− 5.3

304.0

774.28

300

620

9

3.1

7.2

43.1

78.3

− 106.5

309.0

774.84

200

530

8

5.4

− 7.0

50.2

82.9

− 15.4314.0

775.32

200

560

9

5.4

− 10.1

50.7

80.8

− 3.0

321.5

776.04

300

560

7

11.3

35.2

68.5

59.9

152.7

326.5

776.52

300

530

6

12.1

9.6

41.6

76.1

− 112.1

331.5

776.82

350

590

7

9.3

23.4

58.8

70.9

176.8

332.0

776.84

250

560

8

8.9

− 2.0

40.4

77.9

− 63.8337.0

776.96

450

590

5

9.9

16.7

21.7

61.8

− 109.2

342.0

777.09

200

560

9

7.2

184.0

− 61.2− 82.1− 50.5347.0

777.21

250

590

9

10.0

126.2

− 34.4− 40.1− 166.8

350.0

777.29

350

590

7

11.8

25.0

34.8

62.8

− 133.4

355.0

777.41

250

530

7

6.0

− 7.4

43.3

78.3

− 37.6357.5

777.47

250

650

8

6.9

160.1

− 1.6

− 51.1

140.2

360.0

777.54

250

590

9

8.4

7.0

50.9

83.2

− 134.0

362.5

777.60

250

530

7

6.8

16.6

56.5

76.4

− 176.1

363.0

777.61

250

560

8

5.8

− 6.8

44.1

79.1

− 38.3368.0

777.74

250

590

9

10.2

9.2

45.9

79.0

− 120.5370.5

777.80

250

620

11

6.4

161.3

− 21.1− 60.5

147.1

373.0

777.86

200

590

10

5.6

160.1

− 28.1− 63.1

154.0

375.5

777.93

200

620

11

1.4

176.8

− 46.4− 82.2

128.7

379.5

778.03

200

590

10

5.7

0.2

45.0

81.5

− 73.8382.0

778.09

300

620

9

11.6

183.8

− 56.9− 86.1− 22.3384.5

778.15

200

680

13

2.6

167.0

− 57.6− 79.1− 141.8

387.0

778.21

200

680

13

2.5

164.5

− 40.6− 72.2

160.3

389.5

778.27

300

530

6

7.7

164.1

− 37.2− 71.8

149.0

392.0

778.34

200

680

13

1.3

180.3

− 43.5− 80.3

105.8

394.5

778.40

200

680

13

1.3

178.3

− 49.2− 84.8

124.0

395.0

778.42

250

680

12

6.9

179.1

− 57.3− 87.0− 86.5(2020) 7:26

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Table 1 Results of paleomagnetic analysis (Continued)
Depth

Age

T1

T2

(cm)

(ka)

(°C)

(°C)

397.5

778.48200

680

No of steps
13

MAD
2.2

Dec

Inc

VGP lat.

VGP lon.

(°)

(°)

(°)

(°)

170.1

− 45.1− 78.1

155.4

400.0

778.55

200

680

12

2.5

167.6

− 48.5− 78.1

172.9

402.5

778.61

200

680

12

1.3

175.1

− 43.7− 79.6

132.7

405.0

778.68

200

680

13

1.7

169.2

− 52.2− 80.8− 173.5

406.0

778.71

250

680

12

1.7

176.5

− 52.1− 86.3

159.7

408.5

778.77

200

680

13

2.4

163.4

− 47.3− 74.4

176.9

411.0

778.84

300

530

6

9.8

− 88.3

71.3

29.3

67.4

413.5

778.90

300

530

6

9.8

− 47.0

53.3

51.6

29.6

416.0

778.97

300

560

7

4.1

198.4

− 45.7− 72.4

39.0

418.5

779.04

300

560

7

3.5

160.2

− 45.0− 71.0176.3

T1, T2 Temperature ranges of thermal demagnetization data used for principal component analysis

Fig. 4 Paleomagnetic results. a, b Declination and inclination of a ChRM calculated by the principal component analysis. c Latitude of a VGP
calculated from the ChRM direction. d MAD (Kirschvink 1980) of the ChRM. e Intensity of NRM after demagnetization at 300 °C (NRM300 °C). f
Intensity of NRM300 °C normalized by magnetic susceptibility χ. g Intensity of NRM300 °C normalized by ARM. h NRM components of temperature
ranges of 200–350 °C, 350–590 °C, and > 590 °C. i Contribution to the total NRM of the components of temperature ranges of 200–350 °C, 350–
590 °C, and > 590 °C. The blue band shows the MB polarity flip zone (see text)

(2020) 7:26

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horizon data, which is further overlain by the normal-polarity zone above 341 cm. The normal-polarity
zone includes an excursion at 206–176 cm that consists offour excursional fields with VGP latitudes shallower than
45°. Throughout the section, the paleomagnetic directions
are dominated by the full-normal/reverse-polarity directions. Except for those in the excursion, there are only twotransitional directions, both within the polarity flip zone.
The reverse- and normal-polarity zones correlate with the
Matuyama and Brunhes chrons, respectively. The mixedpolarity zone includes nine polarity flips. Each polarity flip
occurred within 2.5–5.0 cm of depth. Similar high numbers of rapid polarity changes have often been observed inhigh-resolution MB transition loess records (e.g., Yanget al. 2010; Jin and Liu 2011). We named the zone the“Matuyama–Brunhes polarity flip (MBpf) zone” (Uenoet al. 2019), which is the part of the MB transition that
contains the low field intensity interval.
Six samples, from within the MBpf zone and justabove it and within the excursion (Fig. 4d), had ChRMs
with MAD values > 10. They had quite low intensities of
the NRMs demagnetized at 300 °C (NRM300 °C) and thenormalized NRMs (NRM300 °C/χ and NRM300 °C /ARM)(Fig. 4d–g). Thus, the MAD values may reflect mainly
the variability of field direction and partly the field intensity. The normalized NRM intensity curves may showthat the upper ends of the curves do not reach the broad
paleointensity peak spanning about 760–715 ka (Guyodo
and Valet 1999).
High-resolution MB transition records from the CLP,
including the present one, obtained by progressive THDs
of samples collected at 2–3 cm depth intervals, havecommon features (Spassov et al. 2001; Wang et al. 2006;
Jin and Liu 2010; Yang et al. 2010). All have a multiplepolarity flip zone and comprise paleomagnetic fields
dominated by full-normal/reverse-polarity directions. On
the other hand, there are differences between our record
and those of others. The latter include more directions
with VGP latitudes between 45° and − 45° and havelarger MADs, which may be the result of errors due tosmaller demagnetization steps used in ChRM calculations. Using data of only five steps (Spassov et al. 2001)
or five steps/more than five steps (Yang et al. 2010)
results in larger mean MAD values > 10 for the MBpf
zone, values that are much larger than the value of 5.6
in this study. The polarity flip zone of each site differs inthickness (about 50–100 cm), with different numbers offlips (7 to 13); we do not include the top short episodethat lies far (25–50 cm) above the MBpf zone in the
Baoji and Xifeng sections (Yang et al. 2010). These differences would reflect dust accumulation rates, samplingintervals, and physical/chemical properties related tomagnetization acquisition. As far as we can compare thepresent and previous high-resolution MB transition

records, detailed stratigraphic correlations of short episodes between sites seem to be difficult. This may indicate centennial-scale variability in dust accumulation
rates.

Hyodo et al. Progress in Earth and Planetary Science

Temperature components of NRM

A viscous remanent magnetization (VRM) component
was almost completely demagnetized at 200 °C. We divided the NRM above 200 °C into components between
200 and 350 °C (NRM200–350 °C), between 350 and 590 °C(NRM350–590 °C), and above 590 °C (NRM> 590 °C), which
would be carried by magnetite + maghemite, magnetite,and hematite, respectively. For the S8 and S7 paleosol
layers, all three components increase as χ increases.
Thus, each component includes pedogenic components.
However, except for the mature paleosol intervals, not allcomponents match in variations; NRM200–350 °C changes
consistently with NRM> 590 °C throughout and inconsistently with NRM350–590 °C in part (Fig. 4 h, i). Therefore,
again, the low-temperature magnetite-carrying component(NRM200–350 °C) and the hematite-carrying component(NRM> 590 °C) are probably dominated by DRM, whereas
the middle-temperature magnetite-carrying component(NRM350–590 °C) mainly reflects PRM.
The vertical change in ARM shows a rapid increase
above ca. 600–580 cm, with a broad peak from 560–500
cm, followed by a gradual decrease until about 270 cm(Fig. 5a). Above 200 cm, it repeats another increase/decrease that is relatively gradual. The NRM component <200 °C (NRM< 200 °C) is probably dominated by a VRM
with the present field direction, carried by multi-domain(MD) magnetites. The NRM<200 °C, magnetic grain size
proxy (ARM/χ), and χ show a rapid increase at 600–580
cm, like the ARM (Fig. 5a–c, e). Above 580 cm up to
440 cm, the curves show complex changes. The NRM<
200 °C has double peaks, correlated with the double peaksof the summer monsoon (SM) proxy χFD (Fig. 5b, d). At
these peaks, the magnetic grain size (ARM/χ) has
depressions that indicate an increase in coarse magnetite
particles under strong pedogenesis. The TEM observation of a 1–10 μm grain-size host silicate fraction
revealed that Lingtai paleosol S8 samples included
several-hundred-nanometer-long magnetite and 1–1.5
μm long hematite particles between layers of phyllosilicates (Hyodo et al. 2020). This TEM observation implies
that authigenic magnetites could coarsen to MD-sizeand increase VRM at SM peaks. The median size, a winter monsoon (WM) proxy, and the ARM/χ represent
anti-phase changes with χFD on the whole. However, theARM/χ and the median size represent different changes
in the depth interval from about 440 to 330 cm (Fig. 5c,f), where the grain size of silicates has been affected by
WM strengthening related to an intensive increase ofgalactic cosmic rays (Ueno et al. 2019). For this interval,

Hyodo et al. Progress in Earth and Planetary Science(2020) 7:26

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Fig. 5 Magnetic grain size and monsoon strength parameters. a ARM. b NRM component < 200 °C (dominated by a VRM). c A magnetic grain
size proxy (ARM/χ). d–f Frequency dependence of magnetic susceptibility (χFD), magnetic susceptibility (χ), and median grain size after Ueno et al.
(2019). The arrows show the features of the rapid increase at ca. 600–580 cm. The arrow heads show peaks and depressions between 580 and
440 cm. The blue band shows the MB polarity flip zone (see text)

instead of the median grain size, the ARM/χ shows antiphase variations with the SM proxy χFD (Fig. 5c, d).

Discussion
Nearly syndepositional acquisitions of detrital andpedogenic remanences

The detailed THDs in this study revealed that the DRM
and PRM components have the same directions throughout the sequence, except at the MB transition zone. Thus,
the DRM and PRM directions are indistinguishable as far
as they are acquired in the geomagnetic secular variation.
The multi-component NRMs are present only in the MB
transition zone, especially just above or just below (2.5cm) the polarity flip. The isolated DRM and PRM directions of samples deviate from each other by a few tens ofdegrees to more than 100°. In addition, such multicomponent samples do not appear over two successive
horizons. These results imply that the PRM was probably
acquired mostly within 2.5 cm below the depth of DRM
acquisition and that the polarity flip occurred mostly
within a 2.5 cm thickness of loess. The isolated single

reverse-polarity short episode at 357.5 cm (Fig. 3i) implies
the occurrence of a normal–reverse–normal polarity
change within 2.5–5 cm.
Laboratory experiments imply syndepositional acquisition of DRM in loess (Wang and Løvlie 2010; Zhao
and Roberts 2010). In addition, the χ profiles of modern
loess sections (Yang et al. 2015; Kang et al. 2018) demonstrate that the majority of pedogenic ferrimagnets
are formed in the subsurface layer, i.e., the χ shows
mature-paleosol values (> 100 × 10−8 m3 kg−1) at 0-cm
depth. The χ of modern soils represents a linear relationship with modern precipitation, providing a basis
for a modern analog method to estimate past summer
precipitation (Maher and Thompson 1995; Balsam
et al. 2011). Therefore, PRM would be acquired withina short time after deposition, probably at decadal timescales. Thus, small lock-in depths are likely for PRM as
well as DRM. Therefore, the acquisition of DRM and
PRM by loess–paleosols is nearly syndepositional ingeneral paleomagnetic records such as that of Lingtai
with a resolution of 2.5 cm.

Hyodo et al. Progress in Earth and Planetary Science(2020) 7:26

Suborbital-scale climate-event stratigraphy-based age
model

Loess and paleosol layers are correlated with glacial andinterglacial periods, respectively. The χ variation of aloess sequence, an East Asian SM proxy, is generally
used for stratigraphic correlation with marine oxygenisotope data (Kukla 1987). The sequential correlation
with MISs from the top to paleosol layer S6 (MIS 17) is
consistent among studies. However, the correlations oflayers from paleosol S7 to paleosol S8, within which lies
the MB transition, differ. The MIS 19 interglacial period
has generally been correlated with S7 (Ding et al. 2002;
Sun et al. 2006; Hao et al. 2012) or S8 (Wang et al. 2006;
Liu et al. 2008; Ueno et al. 2019). In consideration of the
MB boundary lying in S8 or the S8–L8 transition, andbetween the sea-level highstand MIS 19.3 and lowstandMIS 19.2 (Bassinot et al. 1994; Ferretti et al. 2015;
Hyodo and Kitaba 2015), the latter correlation seems
more likely. In addition, the correlation of S7 with MIS
19 requires the assumption of extraordinarily large lockin depths of magnetization, e.g., a few to several meters.
Such large lock-in depths are supported by neither thelaboratory experiments (Wang and Løvlie 2010; Zhao
and Roberts 2010) nor the loess records of the geomagnetic excursion and secular variation without significant
downward shifts (Zhu et al. 1994; Heslop et al. 1999).
The East Asian SM and WM show inverse correlationsin their orbital-scale changes; the SM strengthens and theWM weakens during interglacial periods, and vice versa
during glacial periods (Kukla 1987; Maher 2016). In thehigh-resolution monsoon records of S7–L9 from Lingtai
and Xifeng, even millennial-scale climate events exhibit
anti-phase changes that are well correlated with MIS 19
Northern Hemisphere mid-latitude climate episodes(Ueno et al. 2019). In this study, we adopted an age model
based on sequential correlation of more than 10 climate
events, using the brief sea-level drop/cooling events observed in the Northwest Pacific and North Atlantic midlatitudes (Ueno et al. 2019). At present, this is the only age
model with multiple suborbital-scale age constraints. The
chronology for the original climatostratigraphy is basedon orbital tuning with at least three age controls for the
MIS 19 interval, related to MIS 20.0, MIS 19.2, and MIS
19.0 (Hyodo et al. 2017). The MIS 19.2 sea-level lowstandprovides quite an important age constraint for the MB
transition. The deep-sea benthic δ18O data are usually orbitally tuned with only the control point of MIS 20.0, andwithout the points of MIS 19.2 and even MIS 19.0 (basedon the assumption of a constant accumulation rate),
whose published astronomical ages can be revised to several thousand years older assuming the point of MIS 19.2
(Hyodo and Kitaba 2015).
The time variations of the Lingtai magnetic polarity,
SM, and WM proxies are plotted in Fig. 6a–c, where thePage 11 of 18

age-control points are shown by the solid circle and thesquare symbols denoted “1 to 10b” defined in the core
TB2 of the Chiba Section, central Japan, and the IODP
Site U1313 core in the North Atlantic (Hyodo et al.
2017; Ueno et al. 2019). “x1 to x4” represent SM features
defined in the Xifeng data (Ueno et al. 2019). For
comparison, the high-resolution magnetic polarity
stratigraphy and the paleoceanic/paleoclimate data areplotted in Fig. 6. In addition to the magnetic polarity
stratigraphy (Hyodo et al. 2006), diatom-based sea-levelproxies (Maegakiuchi et al. 2016) and pollen-based
paleoclimate data (Kitaba et al. 2009, 2013, 2017), bothfrom Osaka Bay, were obtained (Fig. 6d–g). The 10-year
resolution biogenic productivity proxy record (Ca/Ti)
from the TB2 core dominantly reflects the sea-level andwater-temperature changes in the Northwest Pacific(Fig. 6h) (Hyodo et al. 2017). The planktonic oxygenisotope data of the core from IODP Site U1313 mainlyreflect sea level, water temperature, and salinity. The age
models for these datasets are based on orbital tuning
using the ice-volume model (Fig. 6j) (Hyodo et al. 2017).
The age controls used for the tuning include the MIS
19.2 sea-level lowstand, an important key point for the
MB transition (Hyodo and Kitaba 2015). The depths of
the Lingtai sequence were dated by linear interpolation
between the control points and extrapolation beyond theuppermost and lowermost control points with the mean
accumulation rates of the nearest intervals.
In the present age model, we assume an absence oflarge hiatuses in the loess sequence. The sequential correlation of millennial-scale climate events (Fig. 6) may
support the assumption, although the correlation does
not certify the absence of submillennial-scale hiatuses. A
series of centennial-scale climate variations over the past
1000 years, including the Little Ice Age recorded in aloess sequence (Kang et al. 2018), indicates continuous
dust deposition at a centennial scale. However, there isanother case; the Laschamp geomagnetic excursion
observed at many profiles in the CLP is absent in the
profile near the Yellow River, which indicates either ahiatus or a large decrease in the dust accumulation rate
(Zhu et al. 2007). Thus, we need to be careful when discussing centennial-scale events.
Background climate for the MB transition

Based on the quantitative climate estimate from pollen
assemblage data (Kitaba et al. 2013), the climate of theearly MIS 19 interglacial is characterized by a cooling
event (783–777 ka) intercalated by the warmest (777–775 ka) and a warm (786–783 ka) interval (Fig. 6). The
climate signal corresponding to the warmest interval isobserved at all of the sites: the SM maximum “x3” andWM minimum at Lingtai and the warm events “G–H”
in the Northwest Pacific and North Atlantic mid-

Hyodo et al. Progress in Earth and Planetary Science(2020) 7:26

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Fig. 6 Comparison of magnetic, climatic and paleoceanic variations. a–c The data of Lingtai. a The magnetic polarity stratigraphy. b The frequency
dependence of magnetic susceptibility χFD (an SM proxy). The dots and squares represent age controls of the cool/low sea-level events and “x1–x5”
show the SM events defined in the Xifeng loess sequence (Ueno et al. 2019). c The coarse grain size (a WM proxy) (Ueno et al. 2019). d–g The datafrom Osaka Bay. d The magnetic polarity stratigraphy (Hyodo et al. 2006). e Diatom-based salinity (sea-level proxy) (Maegakiuchi et al. 2016). f Mean
temperatures of the coldest month (MTCO) and warmest month (MTWA) (Kitaba et al. 2013). g Proportion of deciduous trees (warm proxy) (Hyodo
and Kitaba 2015). h The Ca/Ti ratio reflects biogenic productivity in Northwest Pacific based on the Chiba Section core TB2 (Hyodo et al. 2017). The
dots designated “1–10b” show cool/low sea-level events, and “A–K” show warm/high sea-level events. i Planktic δ18O from the mid-latitude North
Atlantic, with iceberg discharge events and the highest sea surface temperature interval (after Ferretti et al. 2015). The age model is after Hyodo et al.
(2017). j Ice-volume model (after Hyodo et al. 2017). k Virtual dipole moment (after Guyodo and Valet 1999). Black/white/hatched areas in themagnetic polarity stratigraphy: normal-polarity/reverse-polarity/excursion. M: Matuyama. B: Brunhes. MBpf: Matuyama-Brunhes polarity flip (see text).
“MB” in h and i represents the Matuyama–Brunhes boundary. The arrow heads in e, h, and i show the MIS 19.3 sea-level highstandlatitudes. In the North Atlantic mid-latitudes, thealkenone-based sea surface temperature (SST) values arehighest within and just after the warmest interval(Ferretti et al. 2015). The cooling event that occurred
during the low geomagnetic field interval, which has
been interpreted as having been caused by increased
galactic cosmic ray flux and low cloud cover (Kitaba
et al. 2012, 2013, 2017), is observed in Lingtai (andXifeng) as a WM intensification event due to the more
highly cooled continent that resulted from an umbrella
effect of low cloud (Fig. 6g) (Ueno et al. 2019). Aroundthe sea-level highstand MIS 19.3 in the North Atlantic,
the cooling event may have affected the flattened planktic δ18O curve and the 1–2 °C lower SST than in thewarmest interval in the North Atlantic (Fig. 6i), andfurthermore the much increased accumulation rate (seeFig. 5 of Hyodo and Kitaba 2015, and Fig. S10 of Hyodo
et al. 2017). The relatively low-resolution pollen assemblage data after 770 ka show a cool climate with noevergreen tree pollen, but the sum of deciduous broadleaved tree pollen (Fagus and Quercus (deciduous))
shows two relatively warm intervals at about 767–764 ka
and 762–757 ka, respectively. The warm/high sea-level
events “A” to “C” occurred during the former warm
interval, and the latter warm-interval climate seems to
be reflected in the paleoceanic data of the Northwest
Pacific and North Atlantic mid-latitudes (Fig. 6h, i).
The MBpf zone, dated at about 779 to 777 ka, lies in
the latest stage of the cooling event. It occurred during
the weakest dipole field, which was < 1/4 the present
field (Fig. 6k). Therefore, the frequent magnetic polarity
reversals are likely strongly related to the low geomagnetic field. The MBpf interval includes the main MB
boundary reported from the Northwest Pacific andNorth Atlantic, within the warm event “I” (Fig. 6h, i).
The MBpf interval may correlate with the interval fromthe base of short episode “b” to the top of episode “c”
within event “I” in Osaka Bay (Fig. 6a, d). Episode “d”

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just after the warmest interval in Osaka Bay is absent in
the Lingtai magnetic record, possibly due to resolutionrelated problems. The short episode observed just above
the sharp peak of χ (“x3”) in Xifeng (Yang et al. 2010),
dated at about 775 ka by the age model (Ueno et al.
2019), correlates well with episode “d” in Osaka Bay.
The Lingtai sequence did not record an event comparable to episode “a” in the warm interval in Osaka Bay(Fig. 6d), which correlates well with the precursor to the
MB reversal in the Luochuan loess sequence (Jin et al.
2012). The precursor in Luochuan is recorded in the
flattened broad χ peak that is correlated to the SM peak“x5” in Lingtai. The absence of the precursor in the Lingtai
loess sequence may be due to a hiatus or a slower accumulation rate, relative to that of the Luochuan sequence.

outer core may reverse on time scales < 500 years(Gubbins 1999). Polarity changes with few transitional
directions have been observed often in high-accumulation
rate and fine-grained sediments (e.g., Hyodo et al. 2006,
2011; Sagnotti et al. 2014). This will be discussed further
later.

Hyodo et al. Progress in Earth and Planetary Science

Post-reversal geomagnetic excursion

The geomagnetic excursion in paleosol S7, comprising only
excursional fields with no full-reversal field (Figs. 3c and4a), lies around warm event B and is dated at ca. 766 ka,
with an interval of about 1 ka. This excursion can be correlated with neither episode “e” at about 762 ka (Fig. 6d) nor
the Stage 17 excursion ranging in depth from 356 to 352 m
in the Osaka Bay 1700 m core (Biswas et al. 1999). The
latter is dominated by full-reversal fields. According to therecent linear age model for the MIS 17 marine clay layer(Kitaba et al. 2013), the Stage 17 excursion ranges in age
from 710 to 703 ka. The excursion in Lingtai is stratigraphically well correlated with the Baoji A excursion in the S7
paleosol layer (Yang et al. 2007), although the detailed data
represent a complex field behavior that is frequently interrupted by normal-polarity intervals (Yang et al. 2010). The
complex field behavior may be the result of the muchhigher resolution; the excursion comprises 23 data points
at 2.2 cm intervals on average.
Rapid polarity changes during the MB polarity flip zone

The polarity reversals are numbered 1–9, as in the VGP
path (Fig. 7). The age model shows that each occurred
within 65 years (2.5 cm sample spacing), except for no.
9, which occurred within ca. 130 years (5 cm sample
spacing). The span of each polarity zone ranges from 65
years for nos. 3–4/6–7 to 800 years for nos. 2–3. The
minimum polarity intervals (65 years) consist of a single
NRM record that comprises two components with almost antipodal directions for nos. 3–4 (Fig. 3i) and one
component for nos. 6–7. These results imply that eachfull polarity reversal, and even a full round of polarity
reversals, occurred within 65 years and that the DRM
and PRM were acquired within a time difference < 65years, the limit for our 65-year resolution paleomagnetic
record.
Such rapid reversal may have been caused by the fields
in the liquid outer core of the Earth. The field in the

VGP clusters and few low-latitude VGPs

Our paleomagnetic data from Lingtai have only six
excursional VGPs (lower than 45° in latitude), four ofwhich are in the excursion zone and two of which are
adjacent to the polarity flip in the MBpf zone (Fig. 4). It
is noted that the relatively low-latitude (40–65°S) VGPs
concentrate in the upper MBpf zone, all just before/after
the polarity flip. In addition, they cluster in the SW
Pacific region (Fig. 7), where MB transitional VGPs fromlava flows of the Hawaiian and Canary Islands (Coe et al.
2004; Singer et al. 2005) and lacustrine deposits of Java
(Hyodo et al. 2011) cluster. The results imply that thesesites were dominated by dipolar fields with a polelocated in the SW Pacific.
There is no low-latitude VGP in Lingtai. Few lowlatitude VGPs may be a feature of high-resolution polarity
transition records of high-accumulation-rate fine-grained
sediments; marine clays in Osaka Bay (accumulation rate;
60 cm/ka) (Hyodo et al. 2006) and lacustrine clays/silty
clays in Java (the MB transition spans about 700 cm thicksequence) (Hyodo et al. 2011). Sediment records a timeaveraged geomagnetic field. The time represented by the
sample thickness is a minimum duration for averaging,and it lengthens in DRM due to gradual post-depositional
compaction depending on the sediment materials (Hyodo
1984; Hyodo et al. 1993). The time for a 2 cm thick sample in the present study is estimated to be about 50 years
based on our age model. An averaging time of fields significantly longer than 50 years is implausible in this study
because the data resolved rapid polarity reversals within
65 years and the multi-component NRM adjacent to a polarity reversal did not span two horizons (5 cm/130 years).
As mentioned above, small lock-in depths/time for loess
are also supported by laboratory experiments (Wang andLøvlie 2010; Zhao and Roberts 2010) and the χ values ofmodern soils (Maher and Thompson 1995; Balsam et al.
2011; Yang et al. 2015; Kang et al. 2018).
There may be transitional fields with low-latitude
VGPs, as shown by the data from the Canary Islands(Fig. 7b). However, they were probably short in persistence time and/or weak in strength and so easily averaged
out in sediment magnetization. Under the dominant dipolar fields (Fig. 7b), slow accumulation rates and/or
thick lock-in zone sediments can have magnetizations of
VGPs with various latitudes, including low latitudes,
forged by the vector sum of two wide-angle direction
components with various ratios. Thus, the presence orHyodo et al. Progress in Earth and Planetary Science(2020) 7:26

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Fig. 7 The Matuyama–Brunhes (MB) transition VGPs. a The VGP pass during the MB polarity flip zone. The polarity changes are numbered 1–9.
The pass after no. 5 polarity change is shown by broken lines. b The MB transitional VGPs clustered in the southwest Pacific from the four
terrestrial sites of Lingtai (this study), Java (Hyodo et al. 2011), and the Hawaii and Canary Islands (Coe et al. 2004; Singer et al. 2005)

absence of low-latitude VGPs can be used as a criterion
to assess the fidelity and resolution of transition records
determined by the accumulation rate and lock-in zoneof a sediment.
MB transition stratigraphy from loess and lacustrine andlava sequences

The age of the MB reversal has been discussed by combining the radiometric dates of lavas and the astronomical ages of deep-sea sediments (e.g., Channell et al.
2010; Singer et al. 2019). However, researchers havenever discussed the stratigraphic position of the radiometrically dated lavas’ MB transition fields in the sediment record of the MB transition, which spans several

thousand years. Here, we discuss the stratigraphic position of the MB transitional fields of Hawaiian lavas (Coeet al. 2004; Singer et al. 2005) in comparison with the
MB transition stratigraphy records.
The horizon mean paleomagnetic directions from lacustrine clays of Java reveal the MB transition overabout a 700 cm thick sequence characterized by few
low-latitude VGPs (Hyodo et al. 2011) (Fig. 8b). There
are only four reverse-polarity low-latitude VGPs thatcluster in the South Pacific. Thermoremanent magnetizations of lava flows in the Haleakala section of Maui
Island, Hawaii, yield snapshot fields of the MB transition(Coe et al. 2004). A vertical plot of the VGP latitude
against flow unit numbers shows a succession of nine

Hyodo et al. Progress in Earth and Planetary Science(2020) 7:26

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Fig. 8 Comparison of terrestrial sequence records of the MB transition VGPs. a The age plot of the VGP latitudes around the MB polarity flip
(MBpf) zone from Lingtai. The short polarity zones in the MBpf are numbered “n1–n4” for normal polarity and “r1–r4” for reverse polarity inascending order. Polarity flip no. 5 is after Fig. 7a. b The elevation plot of the VGP latitudes of a horizon mean paleomagnetic direction with anerror bar of α95 from Sangiran, Java (After Hyodo et al. 2011). UT: Upper Tuff. c The VGP latitudes from lavas of the Haleakala section on Maui
Island, Hawaii (after Coe et al. 2004), plotted to the flow unit number. The Ar/Ar ages with an asterisk are after Singer et al. (2019). The thick andthin red lines show the correlations for the upper and lower boundaries of the SW Pacific VGP zone, respectively. MBB: the main MB boundary

lava flows with almost the same VGP located in the SW
Pacific just above the short normal-polarity interval(Fig. 8c). At Lingtai, the no. 5 polarity reversal is underlain by two successive horizons of the SW Pacific VGP(over 5 cm/130 years), above which two isolated horizons also show the SW Pacific VGP (Fig. 8a). In Java,the SW Pacific VGPs are concentrated in three successive horizons from 133 to 153 cm in elevation. In
Hawaii, as mentioned above, the SW Pacific VGPs areconcentrated in nine successive lava flow units. Thus, it
is reasonable to correlate these successive SW Pacific
VGP intervals among Lingtai, Java, and Hawaii (Fig. 8).

Centennial-scale magnetic polarity stratigraphy isdefined for the Lingtai sequence as in Fig. 8a, where thepolarity zones are numbered “n1–n4” for normal polarity and “r1–r4” for reverse polarity in ascending order.
We propose correlating the beginning of the MBpf tothe polarity change at 0 cm elevation (the base of theupper tuff (UT)) and the end to the polarity change at
336 cm elevation in Java. In this correlation model, theshort polarity zone r3 is absent in the Java sequence,possibly due to the large sample intervals (10 cm) or
filtering out by a magnetization lock-in zone. The correlation provides an average accumulation rate of about(2020) 7:26

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170 cm/ka for the Java sequence. The SW Pacific VGP
interval is estimated to be about 120 years in Java, whichis consistent with the estimate of about 130 years for
Lingtai. Based on these estimates, the nine MB transitional
Hawaiian lavas would have erupted within a period of
120–130 years, in the late stage of warm event “I”.
We also propose defining the main MB boundary at
the no. 5 reverse-to-normal polarity change (the r2/n3
boundary) (Fig. 8a), because the time spans of normalpolarity intervals subsequently become longer than thoseof reverse-polarity intervals, as proposed in the Osaka
Bay core (Hyodo et al. 2006). Consequently, radiometricdating of the MB transitionally magnetized lavas of
Hawaii becomes important for the chronostratigraphy of
the Early to Middle Pleistocene transition as well as thegeomagnetic polarity timescale. Our age model estimate,ca. 777.8–777.9 ka, for the SW Pacific VGP interval ismostly consistent within errors for the reported Ar/Ar
dates ranging from 770.8 ± 4.2 to 782.42 ± 6.4 ka (Coeet al. 2004; Singer et al. 2019) and the weighted mean
age of 776 ± 2 ka (Singer et al. 2005). The short normalpolarity zones n1/n2 and the short reverse-polarity zones
r3/r4 can be regarded as parts of the precursor and rebound episodes of the main MB reversal, respectively(Valet et al. 2012).

in Lingtai and Java and is a useful key bed for MB transition stratigraphy. The reverse-to-normal polarity reversal
dated at 778 ka at Lingtai occurs at the upper end of the
SW Pacific VGP zone in Java and Hawaii, as well as in
Lingtai. The reversal is a good candidate for the main MB
boundary. (7) We found an excursion that occurred
during warm event B, which is dated at about 766 ka andspanned about 1 ka.

Hyodo et al. Progress in Earth and Planetary Science

Conclusion
We performed detailed magnetic analyses of the loess–paleosol sequence at Lingtai in the CLP with centennialresolution climatostratigraphy correlated with paleoceanic
events in the Northwest Pacific and North Atlantic midlatitudes. Our chronology is based on orbital tuning with
control points, including the sea-level lowstand MIS 19.2,
an important anchor point for the MB transition chronology. Based on the results, we drew the following conclusions. (1) The pedogenic magnetite-carrying remanent
magnetization was mostly acquired within 2.5 cm below
the acquisition depth of DRM. (2) The thickness of thelock-in zone of both DRM and PRM in the loess does notexceed 5 cm and is most probably < 2.5 cm. (3) The MB
transition terminated with the multiple-polarity flip zone
coinciding with the lowest dipole strength. The polarity
flip zone, dated at 779–777 ka and tentatively named the
MBpf zone, began between the warm events “I” and “J”
and terminated with the end of the cooling event. (4) Mostof the polarity flips during the MBpf zone occurred within
2.5 cm (65 years), i.e., they did not occur over two successive horizons (5.0 cm/130 years). (5) Four VGPs adjacent
to a polarity flip in the upper part of the MBpf zone cluster in the SW Pacific, where MB transitional VGPs fromlavas of the Hawaiian and Canary Islands and lacustrine
deposits of Java also cluster. These sites would have been
governed by dipolar fields in the late MBpf interval. (6)The SW Pacific VGP zone probably spans 120–130 years

Abbreviations
MB: Matuyama–Brunhes; VGP: Virtual geomagnetic pole; MIS: Marine isotope
stage; CLP: Chinese Loess Plateau; NRM: Natural remanent magnetization;
χ: Low field magnetic susceptibility; AMS: Anisotropy of low-field magnetic
susceptibility; ARM: Anhysteretic remanent magnetization; AF: Alternating
field; DC: Direct current; IRM: Isothermal remanent magnetization;
FORC: First-order reversal curve; Bcr: Remanent coercive force; SD: Single
domain; SP: Superparamagnetic; B1/2: The field at which half of the saturation
IRM is reached; PSD: Pseudo-single domain; Hc: Magnetic coercivity;
ChRM: Characteristic remanent magnetization; THD: Thermal
demagnetization; MAD: Maximum angular deviation; DRM: Detrital remanent
magnetization; PRM: Remanent magnetization carried by pedogenic
magnetic particles; MBpf: Matuyama–Brunhes polarity flip; VRM: Viscous
remanent magnetization; MD: Multi-domain; χFD: Frequency dependence ofmagnetic susceptibility; SM: Summer monsoon; WM: Winter monsoon;
IODP: Integrated Ocean Drilling Program; MTWA: Mean temperature of thewarmest month; MTCO: Mean temperature of the coldest month; SST: Sea
surface temperature
Acknowledgements
Part of this study was performed in collaboration with the Center for
Advanced Marine Core Research, Kochi University (reference numbers:
12A002, 13A011, 14A023, and 15A001) and with the support of JAMSTEC. We
thank M. Tobi and K. Takasaki for fieldwork. Constructive comments fromHounslow and an anonymous reviewer helped to improve the paper.
Authors’ contributions
MH designed the study and wrote the paper. KB, MH, YY, and MS conducted
the experiments, and MH, TY, and SK made the field works. JF, MM, and BB
contributed to the discussion, and all contributed to editing the paper. All
authors read and approved the final manuscript.
Funding
This study was supported by grant numbers 19340151, 20654043, 26610162,
and 18H03609 from the Japanese Society for the Promotion of Science.
Availability of data and materials
Requests for materials should be addressed to Masayuki Hyodo (e-mail:mhyodo@kobe-u.ac.jp).
Competing interests
The authors declare that they have no competing interests.
Author details
1
Research Center for Inland Seas, Kobe University, Kobe 657-8501, Japan.
2
Department of Planetology, Kobe University, Kobe 657-8501, Japan. 3Schoolof Earth Sciences and Resources, China University of Geosciences, Beijing
100083, China. 4Museum of Nature and Human Activities, Hyogo, Sanda
669-1546, Japan. 5School of Ocean Sciences, China University of Geosciences,
Beijing 100083, China. 6Institute of Seismology and Volcanology, Hokkaido
University, Sapporo 060-0810, Japan. 7Department of Physics, EPS, Universityof Burgos, Burgos 09006, Spain.
Received: 31 October 2019 Accepted: 19 May 2020

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