Nearly syndepositional acquisitions of detrital and pedogenic 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 (2.5 cm) the polarity flip. The isolated DRM and PRM directions of samples deviate from each other by a few tens of degrees to more than 100°. In addition, such multi-component 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 within a short time after deposition, probably at decadal time scales. 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 in general paleomagnetic records such as that of Lingtai with a resolution of 2.5 cm.
Suborbital-scale climate-event stratigraphy-based age model
Loess and paleosol layers are correlated with glacial and interglacial periods, respectively. The χ variation of a loess sequence, an East Asian SM proxy, is generally used for stratigraphic correlation with marine oxygen isotope data (Kukla 1987). The sequential correlation with MISs from the top to paleosol layer S6 (MIS 17) is consistent among studies. However, the correlations of layers 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, and between the sea-level highstand MIS 19.3 and lowstand MIS 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 lock-in depths of magnetization, e.g., a few to several meters. Such large lock-in depths are supported by neither the laboratory 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 correlations in their orbital-scale changes; the SM strengthens and the WM weakens during interglacial periods, and vice versa during glacial periods (Kukla 1987; Maher 2016). In the high-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 mid-latitudes (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 based on 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 lowstand provides 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, and without the points of MIS 19.2 and even MIS 19.0 (based on 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 the age-control points are shown by the solid circle and the square 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 are plotted in Fig. 6. In addition to the magnetic polarity stratigraphy (Hyodo et al. 2006), diatom-based sea-level proxies (Maegakiuchi et al. 2016) and pollen-based paleoclimate data (Kitaba et al. 2009, 2013, 2017), both from 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 and water-temperature changes in the Northwest Pacific (Fig. 6h) (Hyodo et al. 2017). The planktonic oxygen isotope data of the core from IODP Site U1313 mainly reflect 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 the uppermost and lowermost control points with the mean accumulation rates of the nearest intervals.
In the present age model, we assume an absence of large 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 a loess sequence (Kang et al. 2018), indicates continuous dust deposition at a centennial scale. However, there is another case; the Laschamp geomagnetic excursion observed at many profiles in the CLP is absent in the profile near the Yellow River, which indicates either a hiatus 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 the early 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 is observed at all of the sites: the SM maximum “x3” and WM minimum at Lingtai and the warm events “G–H” in the Northwest Pacific and North Atlantic mid-latitudes. In the North Atlantic mid-latitudes, the alkenone-based sea surface temperature (SST) values are highest 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 (and Xifeng) 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). Around the 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 the warmest interval in the North Atlantic (Fig. 6i), and furthermore the much increased accumulation rate (see Fig. 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 no evergreen 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 and North Atlantic, within the warm event “I” (Fig. 6h, i). The MBpf interval may correlate with the interval from the base of short episode “b” to the top of episode “c” within event “I” in Osaka Bay (Fig. 6a, d). Episode “d” just after the warmest interval in Osaka Bay is absent in the Lingtai magnetic record, possibly due to resolution-related 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.
Post-reversal geomagnetic excursion
The geomagnetic excursion in paleosol S7, comprising only excursional fields with no full-reversal field (Figs. 3c and 4a), 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 the recent 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 much higher 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 each full 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 < 65 years, 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 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.
VGP clusters and few low-latitude VGPs
Our paleomagnetic data from Lingtai have only six excursional VGPs (lower than 45° in latitude), four of which 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 from lava 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 these sites were dominated by dipolar fields with a pole located in the SW Pacific.
There is no low-latitude VGP in Lingtai. Few low-latitude 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 thick sequence) (Hyodo et al. 2011). Sediment records a time-averaged 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 and Løvlie 2010; Zhao and Roberts 2010) and the χ values of modern 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 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 zone of a sediment.
MB transition stratigraphy from loess and lacustrine and lava 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 have never 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 (Coe et 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 over about 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 that cluster 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 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 are concentrated 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 is defined for the Lingtai sequence as in Fig. 8a, where the polarity 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 to the polarity change at 0 cm elevation (the base of the upper tuff (UT)) and the end to the polarity change at 336 cm elevation in Java. In this correlation model, the short 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 170 cm/ka for the Java sequence. The SW Pacific VGP interval is estimated to be about 120 years in Java, which is 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 normal-polarity intervals subsequently become longer than those of reverse-polarity intervals, as proposed in the Osaka Bay core (Hyodo et al. 2006). Consequently, radiometric dating of the MB transitionally magnetized lavas of Hawaii becomes important for the chronostratigraphy of the Early to Middle Pleistocene transition as well as the geomagnetic polarity timescale. Our age model estimate, ca. 777.8–777.9 ka, for the SW Pacific VGP interval is mostly consistent within errors for the reported Ar/Ar dates ranging from 770.8 ± 4.2 to 782.42 ± 6.4 ka (Coe et al. 2004; Singer et al. 2019) and the weighted mean age of 776 ± 2 ka (Singer et al. 2005). The short normal-polarity 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).