Paleomagnetic direction and paleointensity variations during the Matuyama–Brunhes polarity transition from a marine succession in the Chiba composite section of the Boso Peninsula, central Japan
© The Author(s) 2017
Received: 9 September 2016
Accepted: 9 March 2017
Published: 21 March 2017
KeywordsPaleomagnetism Magnetostratigraphy Matuyama–Brunhes boundary Paleointensity Marine isotope stage (MIS) 19 Chiba composite section
The Earth’s latest magnetic field reversal event, the Matuyama–Brunhes (M–B) boundary, is an important calibration point on the geological timescale, connecting sediments and volcanic rocks, and has therefore been the focus of a number of paleomagnetic studies. During the polarity transition of the M–B boundary as well as other reversals, the Earth’s geomagnetic field intensity dropped significantly (e.g., Valet et al. 2005; Valet and Fournier 2016), resulting in the increased production of cosmogenic radionuclides, including 10Be, in the upper atmosphere (Beer et al. 2002). Hence, the M–B boundary has also been recognized as a positive spike in the 10Be flux recorded in marine sediments (e.g., Suganuma et al. 2010; Valet et al. 2014) and in an Antarctic ice core (Raisbeck et al. 2006; Dreyfus et al. 2008). Therefore, changes in the geomagnetic field intensity and the 10Be production rate in the atmosphere also provide a timescale that has been widely used in the geochronology of marine sediments (e.g., Guyodo and Valet 1999; Yamazaki 1999; Channell and Kleiven 2000; Laj et al. 2000; Stoner et al. 2000; Kiefer et al. 2001; Christl et al. 2003, 2007; Horng et al. 2003; Valet et al. 2005, 2014; Yamazaki and Oda 2005; Yamamoto et al. 2007; Yamazaki and Kanamatsu 2007; Suganuma et al. 2008; Channell et al. 2008, 2009, 2010, 2014, 2016; Inoue and Yamazaki 2010: Macri et al. 2010; Mazaud et al. 2012, 2015). These geomagnetic field intensity data associated with the directional change also contain essential information about the Earth’s magnetic field reversal; however, the nature of the geomagnetic dynamo in the Earth’s outer core remains a controversial topic (e.g., Valet and Fournier 2016). One key issue is that the reversal process is relatively rapid compared to the geological timescale; therefore, adequate temporal resolution is required to describe the reversal process.
The M–B boundary has a frequently cited age of 780 ka, which is derived from astronomically tuned benthic and planktonic oxygen isotope records from the eastern equatorial Pacific (Shackleton et al. 1990). This marine, astronomically dated M–B boundary age, is supported by the 40Ar/39Ar ages of the Maui lavas at 775.6 ± 1.9 ka (Coe et al. 2004; Singer et al. 2005), amended to 781–783 ka by recent revisions to the reference age of the Fish Canyon Tuff sanidine standards for 40Ar/39Ar geochronology (Kuiper et al. 2008; Renne et al. 2011). However, an understanding of post-depositional remanent magnetization (PDRM) processes shows that lock-in of the geomagnetic signal occurs below the sediment–water interface in marine sediments (e.g., Roberts et al. 2013; Suganuma et al. 2011), which then tends to yield older ages for geomagnetic events than for depositions. Because this age offset is thought to be relative to the sedimentation rate (Suganuma et al. 2010), geomagnetic records with higher sedimentation rates should minimize the age offset that can occur due to the PDRM lock-in process. Indeed, younger astrochronological M–B boundary ages of 772–773 ka are given for high-sedimentation-rate records (Channell et al. 2010; Valet et al. 2014), particularly for one of the records with no PDRM lock-in delay detected by Valet et al. (2014). These M–B boundary ages are consistent with the records of cosmogenic nuclides in marine sediments (e.g., Suganuma et al. 2010) and an Antarctic ice core (Dreyfus et al. 2008). Recently, Suganuma et al. (2015) presented a new U–Pb zircon age of 772.7 ± 7.2 ka from a volcanic ash layer just below the M–B boundary in the Chiba composite section (a very rapidly deposited marine sediment) in Japan. This U–Pb zircon age, coupled with an astronomical age for the marine sediment, yields an M–B boundary age of 770.2 ± 7.3 ka. This is the first direct comparison of the astronomical age calibration, U–Pb dating, and geomagnetic reversal records for the M–B boundary; however, there has been no relative paleointensity record from the Chiba composite section, which is a key requirement for calibrating the geological timescale.
In this paper, we performed a high-resolution paleomagnetic analysis for the Yoro-Tabuchi and Yoro-River sections of the main part of the Chiba composite section in the Kokumoto Formation of the Kazusa Group, Japan, to provide very detailed records of the virtual geomagnetic poles (VGP) and the relative paleointensity changes through the M–B boundary. This record provides one of the most detailed descriptions of the M–B polarity transition obtained thus far from marine sediments and will therefore be key for understanding the dynamics of the geomagnetic dynamo and for calibrating the geological timescale.
Geology of the studied section and samples
The Kokumoto Formation represents an expanded and well-exposed sedimentary succession across the Lower–Middle Pleistocene boundary, particularly in the Chiba composite section (Yoro-Tabuchi, Yoro-River, Yanagawa, and Kogusabata sections) (Fig. 1b). The predominant silty beds of the Chiba composite section are intensely bioturbated, and there is a lack of evidence for episodic deposition such as slumps or muddy turbidites, although minor sandy beds are intercalated within the silty section, particularly in the lower part of the section (Nishida et al. 2016). Marine oxygen isotope records reveal continuous deposition from MIS 21 to MIS 18 with glacial and interglacial cycles corresponding to sandstone- and siltstone-dominated units, respectively (Okada and Niitsuma 1989; Pickering et al. 1999; Suganuma et al. 2015). The Byk-E tephra is widely distributed in the area and provides an excellent stratigraphic marker for the Lower–Middle Pleistocene boundary in the Chiba composite section (Suganuma et al. 2015). The Byk-E tephra bed consists of white, glassy, fine-grained ash and has 1–3 cm in thickness (Kazaoka et al. 2015).
Mini-cores with 1 in. diameters were collected at 213 horizons with a 10-cm stratigraphic interval using a portable engine drill and covering a 29-m succession across the Byk-E tephra bed (Fig. 2). All mini-cores were oriented with a magnetic compass before being removed from the outcrop. Each mini-core was cut into approximately 2-cm-long specimens. The samples for the oxygen isotopic analysis were collected at 72 horizons with a 20-cm stratigraphic interval covering an 18-m succession (Fig. 2). Sandstones were avoided for the sampling.
Paleomagnetic and rock-magnetic measurements
To determine the grain size, the composition of the magnetic materials, and the stability for thermal demagnetization (ThD) analysis, the following rock-magnetic and paleomagnetic measurements were taken in addition to the previous paleomagnetic studies conducted by Suganuma et al. (2015).
Before any other measurements were taken, low-field magnetic susceptibility (volumetric) measurements were taken on all specimens using a Kappabridge susceptibility meter (KLY-3; AGICO, Brno, Czech Republic) at Ibaraki University. The natural remanent magnetization (NRM) was measured using a three-axis cryogenic magnetometer (SRM-760R; 2G Enterprises, USA) installed in a magnetically shielded room at the National Institute of Polar Research (NIPR). Stepwise alternating field demagnetization (AFD) was performed in 2.5- to 10-mT increments up to 80 mT using an AF demagnetizer with a set of static 3-axis AF coils installed on the magnetometer. Stepwise thermal demagnetization (ThD) was performed in 20–50 °C increments up to 700 °C using thermal demagnetizers (TDS-1; Natsuhara-Giken, Japan) at the NIPR.
To evaluate the magnetic grain concentrations in a specimen, remanence after the acquisition of anhysteretic remanent magnetization (ARM) was measured. The ARM acquisition was performed in a 0.03-mT DC field with an 80-mT AF using the SRM-760R magnetometer at the NIPR. The isothermal remanent magnetization (IRM), regarded as saturation IRM (SIRM), was imparted at 1.5 T using a pulse magnetizer (MMPM-9; Magnetic Measurements, UK) at the NIPR. Then, IRM of 0.1 and 0.3 T was acquired in the opposite direction of the initial IRM, and the S-ratio0.1T and S-ratio0.3T were calculated following the definition of Bloemendal et al. (1992).
Magnetic hysteresis was measured with a maximum magnetic field of 0.5 T for selected specimens using an alternating gradient magnetometer (PMC MicroMag 2900 AGM; Lake Shore cryogenics Inc., USA) at the NIPR. The ratio of saturation magnetization to saturation remanence (Mrs/Ms) is commonly used as a proxy for the magnetic grain size of ferrimagnetic particles (Day et al. 1977). Magnetic hysteresis properties were also obtained using first-order reversal curve (FORC) diagrams, which provide enhanced mineral and domain state discrimination (Pike et al. 1999; Roberts et al. 2000; Muxworthy and Roberts 2007).
Thermomagnetic experiments were performed on selected specimens using a thermomagnetic balance (NMB-89; Natsuhara-Giken, Japan) at the Center for Advanced Marine Core Research, Kochi University. The specimens were heated in air and in a vacuum from room temperature up to 700 °C in a field of 300 mT.
Oxygen isotope analysis
The rock samples for the oxygen isotopic analysis were disaggregated primarily using Na2SO4 and partly using a high-voltage pulse power fragmentation system (SELFRAG Lab; SELFRAG AG, Switzerland) installed at the NIPR. The non-magnetic fraction, including foraminiferal tests, was concentrated using an isodynamic separator at Ibaraki University. We manually picked benthic foraminifera from the non-magnetic fraction for each sample. Oxygen isotopic measurements were taken with an MAT 253 mass spectrometer with a Kiel IV carbonate device installed at the Department of Geology and Paleontology, National Museum of Nature and Science. Jcp-1 and NBS-19 were used as standards to calibrate the measured isotopic values to the Vienna Pee Dee Belemnite (VPDB). The standard deviation of the oxygen isotopic measurements was calculated as 0.038‰ from 119 measurements of NBS-19 working standard samples. We used Bolivinita quadrilatera and Cibicides spp., which were the dominant species yielded from this succession, for the isotopic measurements. Okada et al. (2012) reported that B. quadrilatera has δ18O values identical to the genus Uvigerina, which is thought to have an equilibrium δ18O value with the bottom water. Shackleton and Hall (1984) reported 0.64‰ as the average δ18O difference of Uvigerina spp. minus Cibicidoides wuellerstorfi. In the Chiba composite section, Suganuma et al. (2015) used this value to adjust the δ18O measurements of the Cibicides spp., which is a genus closely related to Cibicidoides, to those of the Uvigerina spp. because the average δ18O difference between them (0.74 ± 0.18‰; 95% confidence limit of the average) reasonably matched the value reported by Shackleton and Hall (1984). We therefore corrected the δ18O values of the Cibicides spp. to those of B. quadrilatera by adding 0.64‰, in accordance with Suganuma et al. (2015).
These data indicate that a remarkably deep PDRM lock-in reported by Okada and Niitsuma (1989) most likely originated from an overprint of magnetization due to the formation of secondary magnetic mineral under the sediment surface. Thus, the ThD analyses at 300 °C are effective for removing the secondary component, indicating that the deeper M–B boundary horizons reported by previous studies (Niitsuma 1971; Okada and Niitsuma 1989; Aida 1997) should be revised by the data represented in this study. Accordingly, a detailed VGP path for the Yoro-Tabuchi and Yoro-River sections was established at a 10-cm resolution and clearly identifies the geomagnetic polarity reversal (Fig. 9a).
The concentration-dependent parameters, including k and k ARM, all vary by a factor of less than 4; the magnetic grain-size-dependent parameter k/k ARM varies by a factor of less than 2. The relatively constant concentration and grain size of the magnetic grains satisfy the criteria suggested for the construction of relative paleointensity proxies (e.g., Tauxe 1993). The magnetic-concentration-sensitive parameters, ARM and k, are often used to normalize the NRM for constructing paleointensity proxies (e.g., Valet 2003; Suganuma et al. 2008, 2009). However, the low-temperature component is thought to be a secondary acquired “noise” with respect to the primary magnetic signal, as shown by rock-magnetic experiments. In this study, we use the ARM after ThD at 300 °C as a normalizer for the concentration of magnetic grains that carries a primary signal to avoid the secondary low-temperature component. The ratio of NRM300/ARM300 (both proxies are coercivity fractions between 30 and 50 mT for the NRM and ARM vectors after thermal demagnetization at 300 °C) is used as a paleomagnetic paleointensity proxy. Although a similar method was used by Wu et al. (2015), they used magnetic susceptibility as a normalizer. The uppermost diagram of Fig. 6 displays a prominent low in the relative paleointensity proxy at the directional change zone of the M–B boundary (indicated by a shaded bar).
Oxygen isotope curve and age model
Thickness from Byk-E (m)
Thickness without sand (m)
Sedimentation rate (m/kyr)
The M–B boundary is detected in the interval showing a sedimentation rate of 61 cm/kyr, which is a relatively low rate compared to other parts of the Chiba composite section, but it might still be high enough to minimize the lock-in effect (Suganuma et al. 2010) and to reconstruct a high-resolution paleomagnetic record during the M–B polarity transition.
Matuyama–Brunhes polarity at the Yoro-River and Yoro-Tabuchi sections and the VGP path
The M–B boundary for the Kokumoto Formation and detailed geomagnetic behavior during the polarity transition have been reported by Niitsuma (1971) and Okada and Niitsuma (1989). In these reports, the M–B boundary was considered to be 1–2 m below the Byk-E tephra. In addition, Tsunakawa et al. (1999) reconstructed the geomagnetic field variability during the M–B polarity transition by applying a deconvolution technique using a continuous paleomagnetic record from the Kokumoto Formation. These paleomagnetic data were obtained using only AF demagnetization techniques; however, Suganuma et al. (2015) indicated that these previous studies were not successful in removing the overprint on the magnetic signals based on thermal demagnetization for the specimens from same section and concluded that their VGP records needed revision.
The VGP path for the M–B boundary from the Yoro-River and Yoro-Tabuchi sections appears to be one of the most detailed illustrations of the geomagnetic polarity reversal obtained from marine sediments. This path shows that the VGP swings back and forth several times after the “polarity switch.” These swings are thought to be “rebounds” after the “polarity switch” (e.g., Valet et al. 2014) corresponding to the M–B boundary. These rebound-like VGP swings contain clustering features in South Asia or the equatorial western Pacific and in North America. The VGP clustering in North America is likely to be consistent with that observed in marine sediment cores from the North Atlantic Ocean 983B (Channell and Kleiven 2000) and the Indian Ocean V16-58 (Clement and Kent 1991) as well as in the Tahitian lavas (Hoffman and Mochizuki 2012).
Relative paleointensity and comparison with other records
Figure 10 also compares our relative paleointensity record with the MD90-0961 record from the Indian Ocean (Valet et al. 2014), the ODP Site 983 and the IODP Site U1308 from the Iceland Basin in the North Atlantic (Channell et al. 1998, 2010; Channell and Kleiven 2000), and the paleointensity stack curve of PISO-1500 (Channell et al. 2009). The 10Be data from MD97-2143 (Suganuma et al. 2010), MD90-0961 (Valet et al. 2014), and the EPICA Dome C (Raisbeck et al. 2006; Dreyfus et al. 2008) are also shown for comparison. Although small discrepancies exist in the ages and sharpness of the peaks and valleys, these paleointensity records generally show similar long-term variations and patterns. A distinctive common feature of these records is a paleointensity low at approximately 770–780 ka, which corresponds to the M–B boundary. These consistencies, including the horizons of the M–B boundary and its reversal interval in these records, support the relevance of the oxygen isotope age model for the Chiba composite section.
Unfortunately, a recovery of the paleointensity to the normal level seen before the reversal is not observed in our record due to a shortage of sampling horizons in the upper section. However, the generally observed common variations seen in all relative paleointensity records from widely separated areas with different sedimentary responses to climate changes should represent the true geomagnetic field behavior.
The M–B boundary age
Based on our age model, an age of 771.7 ka is assigned to the M–B boundary in the Yoro-River and Yoro-Tabuchi sections of the Chiba composite section (Fig. 8). The duration of the M–B directional transition zone observed between 0.25 and 1.95 m in our record is estimated to be 2.8 kyr, which is consistent with those of high-resolution marine sedimentary records from the North Atlantic Ocean (2.9–6.2 kyr) (Channell et al. 2010). The M–B boundary “precursor” (Hartl and Tauxe 1996), which predates the M–B boundary by ~18 kyr (e.g., Valet et al. 2014), is not observed in our record.
This M–B boundary age of 771.7 ka is apparently younger than the frequently cited astrochronological age of 777.8–780.1 ka for the M–B boundary based on marine records of low depositional rate (e.g., Shackleton et al. 1990; Lisiecki and Raymo 2005; Pilans and Gibbard 2012). In contrast, Suganuma et al. (2015) recently presented a new U–Pb zircon age of 772.7 ± 7.2 ka from the Byk-E tephra and gave an age of 770.2 ± 7.3 ka for the M–B boundary based on the depositional time interval between the tephra and the M–B boundary. According to our new age model, the sedimentation rate of the section including the M–B boundary is deduced as 61 cm/kyr, which provides an age for the VGP midpoint horizon of the directional transition zone that is 1.8 kyr younger than the depositional age of the Byk-E tephra. Based on this age model, we recalculate the M–B boundary age using the U–Pb zircon age of Byk-E as 770.9 ± 7.3 ka (error includes uncertainty in orbital tuning), showing remarkable consistency with the M–B boundary age of 771.7 ka derived by the correlation of the oxygen isotope records.
The age of 771.7 ka is also consistent with the astrochronological ages obtained from the high-sedimentation-rate records in the North Atlantic (773.1 ka; Channell et al. 2010) and in the equatorial Indian Ocean (772 ka; Valet et al. 2014). Recent reports of the ages of the 10Be flux anomaly from marine sediments in the equatorial Indian (772 ka; Valet et al. 2014) and Pacific (770 ka; Suganuma et al. 2010) Oceans are also consistent with our M–B boundary age estimate. Because the paleomagnetic records from sections with higher sedimentation rates are thought to be less affected by the PDRM lock-in (e.g., Suganuma et al. 2010, 2011), more reliable records with higher sedimentation rates may provide younger M–B boundary ages.
A rapid polarity transition and older M–B boundary ages have recently been reported (Sagnotti et al. 2014, 2016) for the last reversal from a paleolacustrine sequence from the central Apennines, Italy. However, since continental sediments likely behave differently than marine environments, further studies would be needed to compare the reversal timing between the continental and marine records. Although M–B boundary age and reversal duration may depend on the site location or on the local non-dipole field configuration, further investigation of suitable stratigraphic sequences, particularly from marine sediments, is still needed to understand the exact timing and nature of the geomagnetic field reversal.
Despite a long history of paleomagnetic studies, no consensus has been reached on the nature of geomagnetic field reversals. In addition, refining the chronology for geomagnetic polarity reversals, such as the M–B boundary, is very important for precise correlations among sediments, ice cores, and lavas. Because the geomagnetic polarity reversal is a relatively rapid process in terms of the geological timescale, polarity transition records with a higher time resolution are essential to address these topics. In this article, we report a high-resolution paleomagnetic and oxygen isotope record for the M–B polarity transition from a continuous marine succession in the Yoro-River and Yoro-Tabuchi sections of the main part of the Chiba composite section in the Kokumoto Formation of the Kazusa Group in Japan. Rock-magnetic experiments indicate that magnetic carriers contained in the samples are mainly composed of PSD-sized magnetite. The thermomagnetic experiments show that magnetizations of the samples are mostly stable up to 400 °C. The variations in rock-magnetic properties are relatively homogeneous throughout the Yoro-River and Yoro-Tabuchi sections. Progressive alternating field demagnetization after thermal demagnetization at 300 °C reveals a ChRM direction change that indicates a clear “polarity switch” corresponding to the M–B boundary in the section. The relative paleointensity records also show a significant paleointensity minimum near the M–B boundary. A newly obtained high-resolution oxygen isotope chronology indicates that the M–B boundary is located in the middle of MIS 19 and yields an age of 771.7 ka for the boundary. This new M–B boundary age is consistent with the findings based on the latest astronomically tuned high-resolution marine sedimentary records, Antarctic ice cores, and the recalculated age of 770.9 ± 7.3 ka deduced from the U–Pb zircon age of the Byk-E tephra using the new age model based on oxygen isotopes. This record shows one of the most detailed behaviors of the M–B polarity transition that has been obtained thus far from marine sediments and will therefore be key for understanding the dynamics of the geomagnetic dynamo. In addition, the Chiba composite section is one of the candidate sites for the Lower–Middle Pleistocene Boundary GSSP; therefore, this record has certain merit for calibrating the geological timescale, including the use of methods such as astronomical tuning, U–Pb dating, and magnetostratigraphy for the M–B boundary.
Global Boundary Stratotype Section and Point
National Institute of Polar Research
- M–B boundary:
characteristic remanent magnetization
virtual geomagnetic pole
natural remanent magnetization
alternating field demagnetization
anhysteretic remanent magnetization
isothermal remanent magnetization
saturation isothermal remanent magnetization
maximum angular deviation
marine isotope stage
MO proposed the topic and conceived and designed the study. YS assisted in the fieldwork and rock-magnetic experiments, in the discussion of the data, and in the drafting of the manuscript. YH assisted in the fieldwork, carried out the experimental study, and developed the figures. OK assisted in the fieldwork. All authors read and approved the final manuscript.
Makoto Okada is a professor at Ibaraki University, Japan. He received his Ph.D. from the University of Tokyo. His research interests include magnetostratigraphy and oxygen isotope stratigraphy and their applications in paleoceanography and tectonic events during the Cenozoic time in Japan and the surrounding areas. Yusuke Suganuma is an associate professor at the National Institution for Polar Research, Japan. He received his Ph.D. from the University of Tokyo. His research interests focus on cosmogenic radionuclides and their application in earth sciences, including precise dating of geomagnetic reversals and the growth and decay timing of the Antarctic ice sheets. Yuki Haneda is a Ph.D. student at Ibaraki University. The main themes of his research include magnetostratigraphy and oxygen isotopic stratigraphy of the Upper Pliocene to the Lower Pleistocene distributed around central Japan and their utilization for paleoceanographic environment reconstruction in the northwestern marginal region of the Pacific Ocean during and after Northern Hemisphere glaciation. Osamu Kazaoka is a chief researcher at the Research Institute of Environmental Geology, Chiba Prefecture, Japan. He received his Ph.D. from Osaka City University. He is a specialist in sedimentary geology, particularly in Pleistocene and Holocene marine sequences, and environmental geology focusing on the liquefaction of man-made sedimentary formations.
We thank Martin J. Head and Masayuki Hyodo for helpful discussions on this project. We also thank Nozomi Suzuki and Yoshimi Kubota for help with the oxygen isotope measurements at the National Museum of Nature and Science and Yuhji Yamamoto who supported the rock-magnetic measurements at the Kochi Core Center through a joint use system (Grants 15A046, 15B041, 14A017, and 14A018). We are also grateful for the constructive comments from anonymous referees. This work was partially supported by JSPS KAKENHI, Grant Numbers 16H04068 and 15K13581, by a donation from Hisashi Nirei, and by special funding from the Director of the National Institute of Polar Research, Japan.
The authors declare that they have no competing interests.
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