Reconstruction of high-resolution paleomagnetic variations in the middle Gauss Chronozone including the upper and lower boundaries of the Mammoth reversed subchronozone

doi:10.21203/rs.3.rs-4861804/v1

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Reconstruction of high-resolution paleomagnetic variations in the middle Gauss Chronozone including the upper and lower boundaries of the Mammoth reversed subchronozone

https://doi.org/10.21203/rs.3.rs-4861804/v1

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In this study, we provide paleomagnetic records, including the upper and lower boundaries of the Mammoth reversed subchronozone, reconstructed from a continuous outcrop of the upper marine Pliocene in the southernmost part of the Boso Peninsula, which is more than 200 m thick. In this interval, the Mammoth reversed subchronozone was recognized in an interval of approximately 50 m thickness, including a polarity transition zone at its lower boundary of approximately 10 m thickness. Simultaneously, the upper boundary showed no polarity transition zone and a rapid polarity reversal, suggesting a probable lack of strata. Several significant variations were recognized in the paleomagnetic directions in the middle Gauss chronozone, some of which may qualify as new geomagnetic excursions. The relative paleointensity (RPI) around the lower Mammoth boundary was significantly attenuated compared to that of the average of the entire section. Additionally, the Virtual geomagnetic pole (VGP) path of the lower Mammoth boundary exhibits clustering around the North American Continent, moves to the southern hemisphere across the equatorial Pacific region, and clusters again in the eastern Antarctic continent immediately after migration.

The new paleomagnetic record obtained in this study is one of the highest temporal resolution continuous paleomagnetic records for the Late Pliocene. Therefore, the results of this study are expected to contribute to clarifying the geodynamic mechanisms and development of Pliocene magnetostratigraphy using RPI records.

paleomagnetism

rock-magnetism

geomagnetic reversals

magnetostratigraphy

Mammoth subchron

relative paleointensity

Boso Peninsula

The geomagnetic field fluctuates constantly and its form is not constant. The geomagnetic pole typically behaves as a random walk and moves around geographic poles; this is referred to as secular variation (Butler 1992).

Additionally, occasional excursions exist, in which the geomagnetic pole wanders beyond the normal range of secular variation, and reversals in which the geomagnetic pole switches to the opposite geographical pole. Among these variations, geomagnetic reversal, which occurs approximately once every several hundred thousand years, is the most significant event and provides essential information on geological timescales (e.g., Hagstrum et al. 2017; Tauxe and Opdyke 1982).

Yamazaki and Oda (2005) and Cannell et al. (2008) created RPI stack curves based on records from multiple deep-sea-bottom cores. Suganuma (2011) and Roberts et al. (2013) highlighted the potential of using the RPI to achieve precise stratigraphic correlations. Furthermore, geomagnetic secular variations are being applied to the dating of volcanic eruptions and the estimation of eruption processes, mainly after the Late Pleistocene event (e.g., Chiara et al. 2017; Anai et al. 2021; Hasegawa et al. 2023). However, the mechanisms of geomagnetic polarity reversals and excursions, let alone the nature of the geodynamo that drives the geomagnetic field, are not yet clarified and are under debate (e.g., Roberts 2008; Hoffman and Mochizuki 2012; Aubert 2013; Valet and Fournier 2016). The geomagnetic field is governed by fluid motion in the Earth's outer core, and understanding past variations in the Earth's magnetic field, including geomagnetic reversals and excursions, will contribute to exploring the state and behavior of the outer core (e.g., Loud et al. 1988; Hoffman and Singer 2004). It is also known that when a geomagnetic reversal or a geomagnetic excursion occurs, the intensity of the geomagnetic field is significantly attenuated, and the influx of galactic cosmic rays increases (e.g., McHargue et al. 1995; Beer et al. 2002), which is of interest to the paleoclimate and paleoenvironmental field in terms of its impact on climate and ecosystems (e.g., Christl et al. 2004; Burke et al. 2018). Thus, reconstructing long-term geomagnetic variations has important implications for solid Earth science.

Currently, many studies have used deep-sea bottom cores and revealed continuous and long-term paleomagnetic variations (e.g., Channell 2017; Ohno et al. 2012). However, deep-sea bottom cores provide a continuous paleomagnetic record with two problems. First, deep-sea bottom cores generally have low sedimentation rates, which offset the acquisition of magnetization (e.g., Roberts and Winklhofer 2004; Liu et al. 2008; Suganuma et al. 2011). This offset causes a time delay in the magnetization against the deposition. The delay time correlates with that of the sedimentation rate and is considered shorter at higher sedimentation rates (e.g., deMenocal et al. 1990; Bleil and von Dobeneck 1999). Therefore, targeting marine successions with much faster sedimentation rates is essential to elucidate the details of geomagnetic reversal (e.g., Okada et al. 2017; Konishi and Okada 2020).

Second, the target of paleomagnetic studies using deep-sea bottom cores dominated the Pleistocene because of the limited deep-sea bottom cores that reach back to the Pliocene (e.g., Valet and Meynadier 1993; Yamazaki and Yamamoto 2018; Hatfield et al. 2021). Thus, the continuous geomagnetic variations, including the four polarity reversal events in the Gauss chron that occurred in the Late Pliocene, need to be better understood.

Haneda and Okada (2019) reconstructed a paleomagnetic-oxygen isotopic composite stratigraphy from the Lower to Upper Pliocene, including the lower Mammoth boundary, from the Awa Group Anno Formation in the central Boso Peninsula. However, the record in the Mammoth subchron is yet to be discovered because of the hiatus. Hatfield et al. (2021) discovered characteristic variations in the relative paleointensity (RPI), including a Late Pliocene interval, in a deep-sea bottom core retrieved from the Caribbean Sea. However, they emphasized the necessity for records with higher temporal resolution.

In the southernmost part of the Boso Peninsula, the Chikura Group formed from the Late Pliocene to the Early Pleistocene and filled small basins scattered on the landward slope of the trench (e.g., Kawakami and Shishikura 2006; Yamamoto and Kawakami 2005). Few studies were conducted on the reconstruction of continuous paleomagnetic variation in the Chikura Group at high temporal resolution because of its complex geology with many faults and folds. Konishi and Okada (2020) reported detailed paleomagnetic variations, including the lower Olduvai boundary and the Feni subchronozone, indicating the high potential of the Chikura Group for paleomagnetic studies. Kameo et al. (2003) identified the last occurrence horizon of several calcareous nannofossils in the Mera Formation, Chikura Group. They showed that the strata exposed along the Mera Coast, including the Mammoth reversed subchron, were deposited during the Late Pliocene. Therefore, the Mera Coast is a suitable study area for reconstructing paleomagnetic variations during the Late Pliocene. The purpose of this study was to clarify the detailed magnetostratigraphy of the lower part of the Mera Formation, Chikura Group, and provide a continuous and high-resolution paleomagnetic record, including the upper and lower boundaries of the Mammoth subchronozone in the Gauss chronozone, which has not been well studied.

Miocene to Pleistocene marine successions are widely distributed on the Boso Peninsula in southeastern central Japan, facing Tokyo Bay and the Pacific Ocean (Fig. 1a, b). The Chikura Group, from the Upper Pliocene to the Lower Pleistocene, is distributed in the southernmost part of the Boso Peninsula (Kotake 1988; Kawakami and Shishikura 2006) (Fig. 1b, c). These sedimentary successions, which consist of siltstone-dominated siltstone and sandstone alternations intercalating numerous tephra beds, are thought to have filled deep-sea trenches and trench slope basins (Kawakami and Shishikura 2006). According to the rock facies, the Chikura Group is subdivided into the Shirahama, Shiramazu, Mera, Minamiasai, and Hata Formations (Kotake 1988). Kotake et al. (1995) constructed a chronostratigraphy of the Chikura Group based on a detailed geological survey and paleomagnetic study. Kawakami and Shishikura (2006) redefined the Chikura Group component, namely, that the Minamiasai Formation belongs to the Mera Formation. However, earlier studies of the chronostratigraphy of the Chikura Group (e.g., Kotake et al. 1995; Okada et al. 2012; Kameo and Okada 2016; Konishi and Okada 2020) follow the definition by Kotake (1988). This study follows the definition by Kotake (1988), which is consistent with that of earlier studies.

3.1 Geological observations and collection of paleomagnetic samples

The entire coastline extending south of the Tomisaki fishing port was treated as the "Mera Coast section" and a route map showing the distribution of outcrops was created based on field observations and aerial photographs taken by a drone (Additional file1: Figure S1). Additionally, geological observations were made in the lower to the middle part of the Mera Formation exposed on the Mera Coast, a geological column was prepared, and paleomagnetic samples as mini-cores with about 25-mm diameter were collected at every about 0.1–1.5 m intervals of the stratigraphic thickness from the siltstone layers (Fig. 2).

3.2 Rock-magnetic experiments

The following rock magnetic experiments were conducted to evaluate the magnetic properties of the rock samples used in this study: Magnetic hysteresis and first-order reversal curves (FORCs) were measured to determine the magnetic domain states of the magnetic particles in the studied specimens. These measurements were performed on selected specimens using an alternating-gradient magnetometer (PMC MicroMag 3900 VSM; Lake Shore Cryogenics Inc., USA) at the Marine Core Research Institute, Kochi University (MACRI). We measured the magnetic hysteresis with a maximum magnetic field of 1.0 T. The magnetic domain state is categorized into the single-domain (SD), multidomain (MD), and magnetic vortex states, which are traditionally referred to as pseudo-single domain (PSD) (Roberts et al. 2017). The domain state depended on the size of the magnetic grains. The magnetic particles of the SD and/or vortex state, in which the grain size is much more impalpable than that of the MD, play an important role in reconstructing the geomagnetic field behavior because the magnetic stability is sufficiently high to maintain the primary component over geological timescales. Therefore, insight into the magnetic domains of magnetic grains within a studied sample is important for validating paleomagnetic data. The Day diagrams (Day et al. 1977) were used to evaluate the broad distribution of the magnetic domain states of the ferrimagnetic minerals based on the ratios of saturation remanence to saturation magnetization (Mrs/Ms) and remanence coercivity to coercive force (Hcr/Hc), which were observed from the magnetic hysteresis curves (Fig. 3b).

We measured 300 FORCs for each sample with a field increment of 2 mT and an average duration of 150 ms in a 1 T saturating field. The FORC diagram is a contour plot of the coercivity distribution of magnetic grains within a sample and their interaction field intensities (e.g., Pike et al. 1999; Roberts et al. 2000; Muxworthy and Roberts 2007). FORC distributions were calculated using the free FORCinel software (Harrison and Feinberg 2008). We used a smoothing factor of 7 (Fig. 3d–f).

King et al. (1982) reported that the magnetite grain size correlates with the ratio of the anhysteretic remanent magnetization (ARM) susceptibility to the magnetic susceptibility. The ARM was applied to the sample using a DC field of 0.03 mT and an AC field of 80 mT..

Two experiments were performed on the selected samples to reveal the magnetic minerals and their characteristics: a thermomagnetic experiment and a low-temperature magnetic experiment. The thermomagnetic experiments were performed using a magnetic balance (NMB-89; Natsuhara-Giken, Japan) located at the MACRI. The samples were heated to 700 ℃ and then cooled back to 50 ℃ in the air and a vacuum under 0.3 T of the applied field. We inferred the magnetic minerals present in the samples based on their blocking/unblocking temperatures and the temperatures of thermal alteration, which were obtained from the J-T curves (Fig. 4a-f). Low-temperature magnetic experiments were performed using a magnetic property measurement system consisting of SQUID sensors (MPMS XL-5; Quantum Design, USA) located at the MACRI. The samples were magnetized in a 1 T field warmed to 300 K after cooling to 10 K in a zero-T field. Such experiments suggested that the magnetic phase transition temperatures of the magnetic minerals included in the samples were not altered by heating and/or oxidation (Fig. 4g–i).

3.3 Paleomagnetic measurements

The natural remanent magnetization (NRM) of each specimen was measured using a pass-through cryogenic magnetometer (SRM-760R; 2G Enterprises, USA) at the MACRI, or a spinner magnetometer (JR-6; AGICO, Czech Republic) at Ibaraki University. We attempted to remove the secondary remanence from the NRM using three methods: progressive alternating field demagnetization (pAFD), progressive thermal demagnetization (pThD), and a combination of thermal demagnetization and pAFD, referred to as the hybrid method (Okada et al 2017; Konishi and Okada 2020). For the hybrid method, the heating condition was set to 250 ℃ to minimize the thermal alteration of the samples. This temperature is the lower limit at which secondary magnetization is removed by thermal demagnetization.

The pAFD was performed at 5–10 mT intervals, ranging from 5 mT to 80 mT. This procedure utilized either an AF demagnetizer equipped with a set of electrostatic three-axis AF demagnetization coils installed on the pass-through magnetometer at MACRI, or an AF demagnetizer DEM-8601 (Natsuhara-Giken, Japan) located at Ibaraki University. The pThD was performed at 50 ℃ intervals, ranging from 100 ℃ to a maximum of 600 ℃, using a thermal demagnetizer TD-48 (ASC Scientific, USA) at Ibaraki University and TDS-1 (Natsuhara-Giken, Japan) at MACRI.

The pAFD treatment was performed on 19 specimens from 19 horizons at MACRI. pThD was performed on 19 specimens from 19 horizons at Ibaraki University and six specimens from six horizons at MACRI. The hybrid treatment was performed on 579 specimens from 193 horizons at MACRI and 219 specimens from 73 horizons at Ibaraki University. These were performed to remove the secondary components with coercivities and unblocking temperatures lower than those of the magnetic grains responsible for the primary component.

We estimated the characteristic remanent magnetization (ChRM) using the principal component analysis of Kirschvink (1980). For pAFD and pThD, we used at least five linearly aligned steps toward the origin for each specimen. For the hybrid method, we used seven steps from 20 to 50 mT for all specimens in the analysis. The maximum angular deviation (MAD) related to each ChRM calculation was generally below 5°. In the Results and Discussion section, we avoid using ChRMs with MADs greater than 15°. Additionally, the Fisher statistics (Fisher, 1953) were performed at each horizon, and criteria of α95 ≤ 20° for stable polarity intervals were established to avoid using data that did not meet the criteria.

We based the paleomagnetic polarity determination on the virtual geomagnetic pole (VGP) latitude as follows: a VGP latitude greater than + 45° was considered normal polarity, less than 45° was reversed polarity, and an absolute value less than 45° was intermediate polarity.

When calculating the RPI, ARM or isothermal remanent magnetization (IRM) is often used as a proxy for the magnetic grain concentration to normalize NRM (e.g., Channell et al. 2016; Konishi and Okada 2020). In this study, based on the judgment that the rock magnetic properties were not problematic, the NRM/ARM slope in the 20–50 mT interval of the pAFD using the hybrid method was used as the RPI proxy.

4.1 Rock-magnetic properties

The results of representative hysteresis measurements of the specimens are shown in Fig. 3a. The hysteresis loop does not exhibit the wasp-waisted structure reported by Roberts et al. (1995); therefore, the contribution of the high coercivity or superparamagnetic grain component is considered sufficiently small. Additionally, the hysteresis loops show saturation below 300 mT. Data points on the Day plot (Day et al. 1977) shown in Fig. 3b fall within a limited range of magnetic grain sizes that correspond to the “pseudo-single domain (PSD)” size based on the definition of Dunlop (2002) is likely to retain sufficient remanent magnetization. The magnetic mineral particles were sufficiently small to support the Day-plot results, as shown in Fig. 3c. Furthermore, when the coercivity peak shown in the FORC diagrams was higher, the position on the Day plot of the corresponding specimen was closer to the SD area (Fig. 3d-f). The FORC diagram exhibits a limited spread distribution along the Bu axis. The coercivity reaches a maximum of 100 mT, with a peak ranging from 20 to 40 mT. These results support the aforementioned findings and are indicative of the 'PSD (vortex state)' behavior as proposed by Roberts et al. (2017). These parameters suggest that the sample comprises magnetic particles, including SD grains, that can retain magnetization throughout the geological timescale.

The results of the thermomagnetic and low-temperature magnetic experiments on the three selected specimens are shown in Fig. 4a–i. The thermomagnetic experiments performed in both air and a vacuum demonstrate that the specimens have a single Curie/Néel temperature at approximately 560 ℃ (Fig. 4a–f). According to these results, Ti-poor magnetite is probably the dominant ferromagnetic mineral responsible for magnetization (e.g., Hunt et al. 1995). In the thermomagnetic results in air, the normalized induced magnetization (J/J0) through cooling was smaller than that through heating (Fig. 4a–c). Contrastingly, in the cooling process under vacuum conditions, the normalized induced magnetization (J/J0) through cooling was more significant than that through heating (Fig. 4d-f). This also suggests that heating in air may have produced fewer magnetic minerals, such as hematite, by high-temperature oxidation, whereas heating in vacuum may have created new ferromagnetic minerals.

Sample LMC93 has a small inflection point of 300–350°C when heating the air. However, it is unclear when heated in vacuum, which may indicate the formation of ferromagnetic minerals by oxidation (Fig. 4b, e).

Low-temperature magnetic experiments showed an apparent Verwey transition temperature of approximately 100–110 K (Fig. 4c, f, i). The temperature moved slightly toward the lower side of the range, possibly because of surface oxidation or the effect of titanium in magnetic minerals (e.g., Özdemir et al. 1993; Jackson and Moskowitz 2021). These results indicate that the major magnetic carriers were nearly homogeneous, regardless of the sample.

Figures 5e and 5f show the magnetic susceptibility (κ) and ARM susceptibility (κARM) and the ratios of both parameters to evaluate the relationship of rock-magnetic characteristics to stratigraphy. In the κARM/κ ratios, almost no significant spikes are observed, indicating that the rock magnetic characteristics are uniform (Fig. 5f). Some significant spikes in the κ and κARM/κ variations, which may correspond to scattered tephra or sandy layers, are observed. However, apart from these, the variations were within an order of magnitude. This suggests that the samples from the Mera Coast meet the criteria of Tauxe (1993) and are suitable for estimating RPIs.

4.2 Remanent magnetization

Typical examples of the pAFD, pThD, and hybrid method results are shown as orthogonal vector diagrams in Fig. 6. The demagnetization paths for the pAFD (Fig. 6a, d) indicate that the NRMs consist of two magnetic components: low-coercivity (LC) components, which are removed at a peak field of 15–20 mT, and the remaining high-coercivity (HC) components. The HC components exhibit an almost linearly demagnetized behavior but slightly curvilinearly toward the origin (gray arrows in Fig. 6a, d). The demagnetization paths for the pThD (Fig. 6b, e) also exhibit two components like in the pAFD: the low-temperature (LT) components which are demagnetized at 250°C, and the remaining high-temperature (HT) components. Similar to the HC components in the pAFD, the HT components exhibited slight curvilinear demagnetization toward the origin (red arrows in Fig. 6b, e). In both methods, the curvilinear early declining paths toward the origin indicated that the HC and HT components were a mixture of two different components. Contrastingly, the demagnetization paths for the hybrid method (Fig. 6c, f) exhibit that the components remained after 15–20 mT of AFD and 250°C of ThD almost demagnetized straight toward the origin with no curvilinear behaviors (purple arrows in Fig. 6c, f). This suggests that the secondary components were effectively removed by the combination of pAFD and ThD. In addition to the earlier study conducted by the Chikura Group (Konishi and Okada 2020), we assumed that the HC component in pAFD consists of low- and high-temperature components, expressed as _HC_LT and _HC_HT, respectively. As mentioned earlier, the HT component in pThD consists of low- and high-coercivity components, expressed as _HT_LC and _HT_HC, respectively. In the LMC93 specimens, the demagnetization paths in pAFD and pThD were similar, although in the hybrid method (Fig. 6a–c). This may be because the secondary component consisting of _HC_LT was not removed in pAFD, whereas the secondary magnetization of _HT_LC was not removed in pThD. The hybrid method effectively demagnetized the secondary magnetizations and separated the presumed primary components. In the LMC73 specimens, the demagnetization path in the pAFD linearly declined toward the origin (Fig. 6d–f). However, that in pThD shows curvilinearly until 400°C, which indicates that the secondary component consisting of a part of the LC component is not removed in pThD until 400°C (Fig. 6e). However, in this case, the demagnetization path in the hybrid method decreased linearly toward the origin after at least 20 mT, indicating that the hybrid method was effective (Fig. 6f).

As a result, we employed seven demagnetization steps between 20 and 50 mT using the hybrid method to calculate the ChRMs for paleomagnetic discussion.

In this study, we applied a reversal test to confirm that ChRMs were successfully extracted without secondary components. In the reversal test, if the ChRMs from reversed polarity were distributed symmetrically with those from normal polarity, they could be treated as a primary record without any secondary components. For the reversal test, we used all ChRMs except for the stratigraphic intervals showing extreme variations, such as reversal boundaries and geomagnetic excursions. The distributions of the ChRM directions shown in the stereonet plot (Fig. 7) were separated into the following three groups: the lower normal polarity zone (0–110 m), reversed polarity zone (110–153 m), and upper normal polarity zone (153–200 m). The α95 circle of ChRMs in each polarity zone was completely separated from each other, indicating that our data did not pass the reversal test (Fig. 7). The overall mean direction of the lower normal polarity zone is very close to the geomagnetic axial dipole (GAD) direction (D = 0, I = ± 54.4) assumed for the study area. Contrastingly, the mean directions of the reversed and upper normal polarity zones were far from the GAD direction. In the reversed-polarity zone, the declinations differed counterclockwise from the GAD, and the inclinations were shallower. In the upper normal polarity zone, the decline was clockwise relative to the GAD. These trends are consistent with that of the influence of normal polarity secondary magnetization acquired after the tilting of the formation because the strata in the study area are tilted south. However, the reversed polarity zone exhibits an unusually shallow inclination that was shallower than the expected inclination (± 54.4°) in this study area for almost all horizons (Fig. 5b). The mean inclination for the reversed polarity zone is -41.6°; hence, the difference is 12.8° (Fig. 7). This may indicate an inclination anomaly observed in long-term (for the past several million years) deep-sea core records, not due to overprinting of secondary magnetizations, as reported and discussed by Yamazaki and Yamamoto (2018) and Hatfield et al. (2021). Although the inclination anomaly observed in the reversed polarity zone in our record was not due to secondary magnetization, declinational deviations in the ChRM distributions remained. Thus, it can be concluded that our data did not pass the reversal test, and the effect of the secondary magnetization remained. However, because the components derived from the secondary magnetizations are small, the directional effects of the secondary magnetizations in the ChRMs derived using the hybrid method probably do not primarily affect the geomagnetic reversals and excursions. Therefore, we used the ChRMs from the Hybrid method for later discussion.

4.3 Magnetostratigraphy

The VGP latitudes plotted in Figs. 5c and 8a show polarity reversals between 103.66 m and 103.76 m and between 152.3 m and 153.25 m. In the horizon of the lower normal polarity zone (1.51–103.76 m), the last occurrence (LO) of Reticulofenestra minutula var. A was identified by Kameo et al. (2003) (Fig. 5). Additionally, the LO of Reticulofenestra minutula var. A was dated to 3.31–3.44 Ma by Kameo and Takayama (1999). This indicates that the lower normal polarity zone corresponds to C2An.3n, and the upper normal polarity zone from 153.25 m to 199.77 m corresponds to C2An.2n accordingly. Therefore, the two polarity switches associated with the reversed polarity zone correspond to the lower and upper boundaries of the Mammoth subchronozone. Ogg (2020) assigned ages of 3.33 Ma and 3.207 Ma to the lower and upper Mammoth boundaries, respectively.

Assuming that the polarity reversal boundaries are at 103.71 m and 152.775 m, which are the stratigraphic midpoints at the boundary zones, the sedimentation rate for the Mammoth subchronozone is approximately 38.9 cm/kyr. This sedimentation rate is higher than that of any deep-sea bottom core and provides essential information for understanding geomagnetic reversal events and excursions.

Several significant variations were found in the interval from C2An.3n to C2An.2n, resulting in VGP latitudes <|±45| or comparable. Oda (2005) stated that one of the definitions of an excursion is when the VGP varies beyond 45°. Therefore, as these intervals may correspond to geomagnetic excursions, we tentatively named them as Mera 1 to 7 from the lower level. Regarding Mera 3, although the mean VGP latitudes calculated from the mean direction do not satisfy <|±45|, some specimens satisfy this criterion (see Additional file 2 Table S2). Mera 4, which is located near the top of the Mammoth subchronozone, also did not meet this criterion. However, its declination and inclination behaviors are similar to those observed just before the polarity reversal, as reported by Berbera and Jicha (2022). Therefore, these variations likely capture the behavior during the transition of an excursion; therefore, we included them as the named intervals. However, the paleomagnetic record near the upper Mammoth boundary obtained from the Mera Formation is characterized by a rapid reversal without a polarity transition period, suggesting a lack of stratigraphy. The lack of partial stratigraphy makes it difficult to accurately compare it with that of other upper Mammoth boundary studies (e.g., Linssen 1999; Bervera and Jicha 2022). Utsunomiya et al. (2023) highlighted that the stratigraphic deficit at 4.5–3.2 Ma in marine successions distributed in the Miura Peninsula and eastern Boso Peninsula may be caused by submarine landslides. Although a direct comparison cannot be made because the sedimentary basins are different from those in this study area, it is possible that a similar tectonic setting caused a scouring event leading to a stratigraphic lack, which is a potential implication worth exploring in future work.

The highest temporal resolution paleomagnetic stratigraphy covering the Late Pliocene from the northwest Pacific Ocean region is probably that of Haneda and Okada (2019, 2022). However, the excursion-like significant variations found in this study were not detected in their records. This may be because the Mera Formation used in this study has a more significant average sedimentation rate and provides a higher temporal resolution record than the Anno Formation used by Haneda and Okada (2019; 2022). However, because it includes horizons with coincident susceptibility spikes and peaks, such as Mera 5, its reliability must be examined carefully (Fig. 5e).

4.4 VGP and RPI variations in the lower Mammoth transition zone

Here, we describe the variations in VGPs and RPIs in the lower Mammoth boundary zone in detail. We define the stratigraphic interval from where the NRM/ARM slope falls below 0.05 to where it exceeds 0.05 as the "RPI minimum zone.” Additionally, we call the interval where the VGP latitude moves from the high northern to high southern hemisphere latitudes the "polarity switch" interval.

The declinations, inclinations, VGP latitudes, and RPIs corresponding to the lower Mammoth boundary in the Mera Coast section are shown in Fig. 9a, and a comparison with the lower Mammoth boundary records from other regions is shown in Fig. 9b. The records obtained in this study include a pronounced RPI minimum near the lower Mammoth boundary, indicating that the geomagnetic field was extremely weak during the polarity transition, similar to the findings of Haneda and Okada (2022). Focusing on the inclinations, a shallowing trend was observed as the RPI decreased. The behavior of the VGP variations shows several characteristic features, which we named A to D (Fig. 8b). The VGP gradually moves southward as its inclination becomes shallower, forming a cluster of VGPs near North America (Area A). Subsequently, it transitioned largely into the southern hemisphere, traversing the Indian Ocean, African continent, and passed thorough equatorial Pacific (Area B) before returning to the northern hemisphere. The VGP then crosses the equatorial Pacific and re-enters the southern hemisphere. This VGP path closely resembles the first rebound observed by Haneda and Okada (2022), indicating the high likelihood that the two correspond.

The RPI minimum zone at the lower Mammoth boundary defined in the Mera Coast section, which is from 99.2 m to 103.7 m, reaches about 4.5 m in thickness. It is similar to that of the Shikoma River section (Haneda and Okada 2022), approximately 3.6 m. However, a stratigraphic lack exists on the outcrop between MC38 (102.22 m) and MC39 (103.36 m) just before the polarity switch, which possibly reduces the temporal resolution. Therefore, the second rebound phenomenon reported by Haneda and Okada (2022) may not be clear for the Mera Coast section. In the second rebound reported by Haneda and Okada (2022), the VGP passed from the western Atlantic to the eastern equatorial Pacific (near Area B) immediately after the polarity reversal, and the variability converged in East Antarctica formed a cluster (Area C). In the Mera Coast section, the VGP passes through the South Atlantic immediately after the polarity reversal, forming a cluster from eastern Antarctica to the South Pacific (Area D). The VGP variation converged to Area C. The nature of VGP clustering observed in this study around Area C was reported from the Punta Piccola section in Italy (Linssen 1991), and that around Area D was also reported from lava in Hawaii (Berbera and Coe 1999). These results indicate that the Mera Coast records detailed behavior immediately after polarity reversal, and combined with those of earlier studies (Linssen 1991; Bervera and Coe 1999; Haneda and Okada 2022), they reveal a complete picture of a lower Mammoth polarity reversal event.

Constable (2007) noted the dominance of non-axial dipole fields in modern times in North America, the South Atlantic, Siberia, and Western Australia. Hoffman et al. (2020) also noted that the magnetic flux features observed at the core-mantle boundary may have been maintained for an extended period of geological time in the Southern Hemisphere, particularly because the magnetic flux belt off the west coast of Australia may have been almost stagnant for the past 3 million years.

By integrating the studies at the lower Mammoth reversal boundary, the VGP paths pass through regions where common nonaxial dipole magnetic fields dominate when the geomagnetic field strength is minimal. This strongly suggests that the VGP path is affected by non-axial dipole magnetic fields during geomagnetic polarity reversal.

4.5 Characteristics of RPI variation in mid-Gauss Chron

The RPI records, a significant outcome of this study, are presented along with IODP U1396 (Hatfield et al. 2021), ODP Leg 138 (Valet and Meynadier 1993), and IODP U1335 (Yamazaki and Yamamoto 2018) (Fig. 9). The virtual axial dipole moment (VADM) record of U1335 (Yamazaki and Yamamoto 2018) was derived from the RPI value using the absolute paleointensity (API) record of the volcanic rocks. Furthermore, Tanaka and Yamamoto (2016) computed VADM averages for 3.5–2.5 Ma based on the API obtained from volcanic rocks.

The RPI variations of U1396, ODP Leg 138, and U1335 were compared by Hatfield et al. (2021), who observed a decaying trend in the peaks and troughs in the C2An.3n period. The results of our study align well with records from deep-sea cores and indicate the characteristics of paleointensity variations during this period. Additionally, the RPI variation of C2An.2n obtained in our study was distinct and closely matched with that of U1396 and ODP Leg 138. These findings reinforce the potential use of paleomagnetic intensity variations as a dating tool, as discussed by Roberts et al. (2013) and Suganuma (2011). Given that the Late Pliocene is a comparable period of global warming, the unique RPI variations identified in our study could significantly enhance the age constraints and advance climate models.

The RPI values in this study were significantly lower during C2An.2n than during Chrons C2An.2r and C2An.3n. However, the ODP Leg 138 record showed a lower RPI value that continued until C2An.1r and then recovered to its original level after polarity reversal at the top boundary of the Kaena subchron. In the U1396 record, the RPIs in the early parts of C2An.2r and C2An.1r were weak; however, those in C2An.2n were not remarkable. Local non-axial dipole magnetic fields may cause such differences in relative values over a long period of observation. Alternatively, the ferromagnetic minerals responsible for secondary magnetization may have influenced the calculation of the RPI because the rock magnetic properties of the samples varied from region to region. In the future, extracting the characteristics of the rock magnetic properties in each region may recover continuous RPI from samples from various regions, including deep-sea cores and terrestrial and marine strata. It would also be helpful to continuously reconstruct absolute paleomagnetic intensities from volcanic rocks and to estimate the influx of galactic cosmic rays through ¹⁰Be measurements.

This study reconstructed paleomagnetic records of the mid-Gauss chronozone, including the top and bottom boundaries of the Mammoth subchronozone, from a continuous outcrop over 200 m in thickness exposed along the Mera Coast in the southernmost part of the Boso Peninsula. This section identified the lower Mammoth boundary between 103.66 m and 103.76 m and the upper boundary between 152.3 m and 153.25 m. Additionally, significant paleomagnetic variations were identified at 1.02 m, 55.93 m, 88.21 m, 171.04–173.91 m, 183.39 m, and 192.05 m. These results suggest that previously unreported geomagnetic excursions may have occurred during the Late Pliocene. They will contribute to the reconstruction of a more detailed chronostratigraphy during the Pliocene.

The average sedimentation rate within the Mammoth reversed subchronozone in the Mera Coast section accounts for approximately 38.9 cm/kyr, surpassing any deep-sea bottom core covering the Pliocene. This underscores the significance of our study as one of the highest temporal resolution paleomagnetic variation records, including the Mammoth reversed subchron.

The VGP path at the lower Mammoth boundary in the Mera Coast section formed a cluster around the North American continent, moved to the southern hemisphere through the equatorial Pacific region, and formed a cluster in the western Antarctic continent immediately after migration. This migration pattern suggests that polarity reversal proceeds under the influence of nonaxial dipole magnetic fields. This finding has significant implications for our understanding of geomagnetic reversal mechanisms.

The shapes of the RPI variations obtained in this study agree well with those of deep-sea bottom core records obtained from distant areas, indicating the possibility of using the RPI to precisely contrast stratigraphic levels. However, the RPI during the C2An.2n period in this study was significantly lower than that from deep-sea bottom cores, indicating that issues to address still exist.

The data obtained in this study were from the Late Pliocene, a period with few reported cases. They may contribute to the development of paleointensity stratigraphy and clarify the geomagnetic reversal mechanism.

ChRM, characteristic remanent magnetization; VGP, virtual geomagnetic pole; NRM, natural remanent magnetization; AFD, alternating field demagnetization; ThD, thermal demagnetization; ARM, anhysteretic remanent magnetization; IRM, isothermal remanent magnetization; MAD, maximum angular deviation; RPI, relative paleointensity; API, absolute paleointensity

Ethics approval and consent to participate

Not applicable

Consent for publication

Not applicable

Availability of data and materials

The datasets used and/or analyzed in this study are available from the corresponding author upon request.

Competing interests

The authors declare that they have no competing interests.

Funding

This study was supported by the Japan Society for the Promotion of Science KAKENHI, grant numbers 19H00710, 21H04523 and 24H00016.

Authors' contributions

AT designed the study, conducted fieldwork and rock magnetic experiments, discussed the data, and prepared the manuscript. MO conceived and designed the study and prepared the manuscript. RH assisted with fieldwork and rock magnetic experiments in the discussion of the data.

Acknowledgments

We thank R. Takizawa, K. Ochi, Y. Ebe, and J. Songeons for collecting the paleomagnetic samples. This study was performed as part of a joint program at the Marine Core Research Institute, Kochi University (21B048, 22A026/22B024, 23A009/23B007).

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RoutemapoftheMeraCoastSection.pdf
Additional file 1 .pdf Route map of the Mera Coast Section Route map showing the distribution of outcrops along the Mera Coast. Blue circles indicate the sampling sites.
2024ATPaleomagneticdata.xlsx
Additional file 2 .xls 2024_AT_Paleomagnetic data The results of rock magnetic and paleomagnetic measurements from the Mera Coast Section are shown in Figures 5 and 8. Table S1. Data measured at Ibaraki University. Table S2. Data measured at the Marine Core Research Institute, Kochi University.
GraphicalAbstract3.jpg

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Reconstruction of high-resolution paleomagnetic variations in the middle Gauss Chronozone including the upper and lower boundaries of the Mammoth reversed subchronozone

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Abstract

Figures

1 Introduction

2 Geological setting

3 Methods

3.1 Geological observations and collection of paleomagnetic samples

3.2 Rock-magnetic experiments

3.3 Paleomagnetic measurements

4 Results and discussion

4.1 Rock-magnetic properties

4.2 Remanent magnetization

4.3 Magnetostratigraphy

4.4 VGP and RPI variations in the lower Mammoth transition zone

4.5 Characteristics of RPI variation in mid-Gauss Chron

5 Conclusions

Abbreviations

Declarations

References

Supplementary Files

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