M5) planktonic foraminifera zone boundary: constraints from the zircon geochronology and magnetostratigraphy of early Miocene sediments in Ichishi, Japan

Chemical Geology 530 (2019) 119333

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Age of the N7/N8 (M4/M5) planktonic foraminifera zone boundary: constraints from the zircon geochronology and magnetostratigraphy of early Miocene sediments in Ichishi, Japan

T

Hiroyuki Hoshia,*, Hideki Iwanob, Tohru Danharab, Hikoma Oshidac, Hiroki Hayashid, Yukito Kuriharae, Yukio Yanagisawaf a

Department of Earth Sciences, Aichi University of Education, Kariya, Aichi, Japan Kyoto Fission-Track Co., Kyoto, Japan Aichi Prefectural Kaisho High School, Yatomi, Aichi, Japan d Institute of Environmental Systems Science, Shimane University, Matsue, Japan e Faculty of Education, Mie University, Tsu, Japan f Geological Survey of Japan, AIST, Tsukuba, Ibaraki, Japan b c

ARTICLE INFO

ABSTRACT

Editor: Balz Kamber

The age assigned to the Miocene N7/N8 (M4/M5) planktonic foraminifera zone boundary is markedly different from the astronomically tuned Neogene time scales published in 2004 (ATNTS2004) and 2012 (ATNTS2012). This difference is primarily because of the lack of reliable assignment of the zone boundary to magnetostratigraphy, which needs to be resolved for the better refinement of the geologic time scale. Our previous biostratigraphic results from an onland sedimentary sequence of the upper Ichishi Group in central Japan allow location of the stratigraphic position of the lowest occurrence datum (LOD) of Praeorbulina sicana, a key planktonic foraminifera species for defining the N7/N8 (M4/M5) zone boundary. In this study, we performed laser ablation-multiple collector-inductively coupled plasma mass spectrometry U–Pb dating on zircons from two felsic tuff beds in the Ichishi sequence. The age of the zone boundary is approximated by a U–Pb age of 17.03 ± 0.11 Ma for a tuff bed ∼2 m above the boundary. We also conducted a magnetostratigraphic analysis in which site-mean characteristic remanent magnetization directions were determined for 19 sites and showed a reverse–normal–reverse (R–N–R) stratigraphic change in polarity. The N7/N8 (M4/M5) zone boundary is within the upper reverse polarity zone. The R–N–R sequence can be correlated with a stratigraphic portion corresponding to chrons C5Dr–C5Dn–C5Cr; hence, the age of the zone boundary is within Chron C5Cr. Our results are thus compatible with the age (16.97 Ma) assigned to the boundary in ATNTS2004 rather than ATNTS2012 and should be considered when the Neogene time scale is updated. A normal polarity subzone found in the upper reverse polarity zone can also be dated at ∼17.0 Ma and is probably correlated with a short normal polarity episode in Chron C5Cr. In terms of tectonic rotation, the site-mean magnetization directions determined in this study suggest an ∼35° clockwise rotation in Ichishi with respect to the North China Block in the Asian continent since ∼17.2 Ma (the upper boundary age of Chron C5Dn). Clockwise rotation of Southwest Japan, which occurred in association with backarc opening of the Japan Sea, was probably in progress during deposition of the upper Ichishi Group sediments.

Keywords: Magnetostratigraphy Neogene time scale Planktonic foraminifera stratigraphy Tectonic rotation U–Pb dating

1. Introduction The astronomically tuned Neogene time scale (ATNTS) is an indispensable basis for correlations and geochronological interpretations of Neogene stratigraphic sequences. An astronomically tuned magnetobiostratigraphic time scale for the entire Neogene was first constructed

⁎

in 2004 (hereafter referred to as ATNTS2004; Lourens et al., 2004) and incorporated into “A Geologic Time Scale 2004” (Gradstein et al., 2004). Eight years later it was updated (referred to as ATNTS2012; Hilgen et al., 2012) and incorporated into “The Geologic Time Scale 2012” (Gradstein et al., 2012). The differences between ATNTS2004 and 2012 are minor on the whole; however, there are relatively large and non-

Corresponding author at: Department of Earth Sciences, Aichi University of Education, 1 Hirosawa, Kariya, Aichi 448-8542, Japan. E-mail address: [email protected] (H. Hoshi).

https://doi.org/10.1016/j.chemgeo.2019.119333 Received 30 April 2019; Received in revised form 22 September 2019; Accepted 9 October 2019 Available online 16 October 2019 0009-2541/ © 2019 Elsevier B.V. All rights reserved.

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Fig. 2. (a) Map of Japan and surrounding regions showing the location of Ichishi. (b) Simplified geological map of the northern Ichishi area redrawn from Geological Survey of Japan (2015). The red box shows the sampling region. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article).

Fig. 1. Age difference of the N7/N8 (M4/M5) planktonic foraminifera zone boundary (red lines) between the astronomically tuned Neogene time scales from 2004 (Lourens et al., 2004) and 2012 (Hilgen et al., 2012). The diatom biochronology is after Watanabe and Yanagisawa (2005). NPD = Neogene north Pacific diatom zone code numbers (Akiba, 1986; Yanagisawa and Akiba, 1998). (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article).

2. Stratigraphic outline of the Ichishi Group The Ichishi Group is a partly fossiliferous, clastic sedimentary sequence of early Miocene age (Shibata, 1967, 1970; Yoshida, 1991; Yoshida et al., 1995). The group is composed of three formations, the Haze, Oi, and Katada formations in ascending stratigraphic order, and each formation is further subdivided into two or more members. The sediments unconformably overlie a granitic and metamorphic basement with irregular paleotopography and coarse-grained lithofacies developed around basement highs are collectively named the Ieki Formation (Shibata, 1967). In general, the sediments gently dip (typically < 20°) to the northeast or east (Fig. 2b) except for local steeper dips near faults. Oshida et al. (2018) recently identified the LOD of P. sicana based on a detailed study of the planktonic foraminifera biostratigraphy of the upper Ichishi Group (Figs. 3 and 4). They collected mudstone samples in the Nagano River section from 23 stratigraphic horizons (sites) of the Mitsugano Member of the upper Oi Formation and the Chaya and Yakuoji members of the Katada Formation (Fig. 5) and obtained planktonic foraminifera assemblages consisting of 37 species from 17 genera at 21 sites. The stratigraphic distribution of major key species in the biostratigraphic zonation is shown in Fig. 4. The occurrence of Globorotalia birnageae, Fohsella peripheroronda, Praeorbulina curva, and P. sicana is of particular importance. The stratigraphically lowest site containing P. sicana is ∼19 m below the Oi/Katada formation boundary, and above it P. sicana successively occurs up to the uppermost site. Praeorbulina curva, one of the descendant species of P. sicana, was detected at only one site near the Oi/Katada formation boundary. Based on these observations, Oshida et al. (2018) concluded that in the Nagano River section, the LOD of P. sicana is ∼19 m below the Oi/ Katada formation boundary. Their taxonomic concept of the Globigerinoides-Praeorbulina-Orbulina lineage follows that of Jenkins et al. (1981) which is widely accepted by many stratigraphic works including ATNTS2012, although it was recently reviewed by Turco et al. (2011) who reexamined the evolutionary lineage at several Mediterranean sections. The population at the LOD of P. sicana in Ichishi is taxonomically correlated with “Globigerinoides sicanus Mophotype 3” of Turco et al. (2011) as having three apertures along the suture of the last

negligible differences in ages for some biostratigraphic datums. We focus here on the stratigraphically lowest occurrence datum (LOD) of Praeorbulina sicana, which marks the base of Zone N8 in Blow’s (1969) planktonic foraminifera N-zone scheme and the base of Zone M5 in Berggren et al.’s (1995) M-zone scheme (Fig. 1). In ATNTS2004, this datum is placed at ∼16.97 Ma in Chron C5Cr of the astronomically tuned geomagnetic polarity time scale (Lourens et al., 2004). This age assignment is based on the astronomical calibration of the planktonic foraminifera datum on the Ceara Rise in the western Atlantic (Shackleton et al., 1999). In ATNTS2012, this datum is placed at ∼16.38 Ma in Chron C5Cn.2n (Hilgen et al., 2012; Anthonissen and Ogg, 2012). This is based on the review by Wade et al. (2011) who adopted data from the Rio Grande Rise in the southwestern Atlantic instead of those from the Ceara Rise because of the relative rarity of the members of the Praeorbulina group in the Ceara Rise (Pearson and Chaisson, 1997). The ∼600-kyr difference for this datum between these two time scales is the largest among the Neogene datums of Blow’s (1969) N-zone scheme and is ∼3.5% of the assigned ages, much larger than the uncertainties of high-precision Ar/Ar and U–Pb isotopic ages. Resolving this issue is vital to the improvement of the early Miocene planktonic foraminifera biochronology and can be achieved via a detailed investigation of a stratigraphic sequence containing the Praeorbulina lineage. In this paper, we present new U–Pb and magnetostratigraphic data that constrain the age of the LOD of P. sicana. Our integrated U–Pb geochronological and paleomagnetic study was conducted on the early Miocene marine sediments of the Ichishi Group in Southwest Japan (Fig. 2a). A recent study of planktonic foraminifera biostratigraphy has shown that the LOD of P. sicana is in the upper part of the group (Oshida et al., 2018). To accurately date it, our investigation was designed to collect samples for U–Pb and magnetic measurements from the same stratigraphic section as that of Oshida et al.’s (2018) biostratigraphic study. Our new data also enables discussion of possible tectonic rotations occurring in Ichishi in association with the backarc opening of the Japan Sea. 2

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Importantly, at two sites below the LOD of P. sicana, they found Crucidenticula sawamurae and Crucidenticula ikebei, both of which are agediagnostic species indicating an early Miocene age, and no occurrence of Crucidenticula kanayae and Thalassiosira fraga at the same sites. They thus concluded that the two sites in the Mitsugano Member correspond to Zone NPD 2B (= Crucidenticula sawamurae Zone; Fig. 1) of Gladenkov and Barron (1995) and Yanagisawa and Akiba (1998). This zone ranges in age from ∼18.1 Ma (near the top of Chron C5Er) to ∼17.0 Ma (in the middle of Chron C5Cr) based on the robust correlation between the diatom stratigraphy and magnetostratigraphy at ODP Site 887 in the northeast Pacific (Watanabe and Yanagisawa, 2005). Therefore, the depositional ages of the two sites in the Mitsugano Member are between ∼18.1 and ∼17.0 Ma. This interpretation does not contradict the age of the N7/N8 (M4/M5) zone boundary both in ATNTS2004 and ATNTS2012 (Fig. 1). Hayashida and Ito (1984) have reported a paleomagnetic study of the upper Ichishi Group in which they showed a magnetic polarity stratigraphy based on remanent magnetization directions obtained from 11 sites (Fig. 3). More than half of the sites were reversely magnetized while the uppermost part of the Mitsugano Member contained four normal polarity sites. Using the biostratigraphic data of Ito (1982), Hayashida and Ito (1984) suggested that the normal polarity sites correspond to Anomaly 5C of the marine magnetic anomaly (Chron C5Cn of the GPTS; Fig. 1). The LOD of P. sicana specified by Oshida et al. (2018) appears to be assigned within the interval in which the normal polarity sites are dominant. However, caution should be exercised when interpreting the magnetostratigraphy of Hayashida and Ito (1984) because the locations of their sampling sites are not presented; the magnetostratigraphy shown in Fig. 3 of the present paper is based on Fig. 2 of Hayashida and Ito (1984). Notably, the remanent magnetization directions were obtained after blanket demagnetization

Fig. 3. Lithostratigraphic divisions of the upper Ichishi Group (Shibata, 1967; Yoshida et al., 1995) and previous magnetostratigraphic and biostratigraphic data. Mag = magnetic polarities of Hayashida and Ito (1984), with the closed and open circles showing normal and reverse polarities, respectively; Foram = planktonic foraminifera stratigraphy of Oshida et al. (2018), with the N7/N8 zone boundary ∼19 m below the Oi/Katada formation boundary (see text); Diatom = diatom biostratigraphy of Oshida et al. (2018).

chamber (Oshida et al., 2018, Fig. 6, 14a–b). Oshida et al. (2018) also reported diatom assemblages from the upper Ichishi Group sediments in the Nagano River section.

Fig. 4. Stratigraphic distributions of important foraminifera species occurring from the upper Ichishi Group (data from Oshida et al., 2018) and the stratigraphic position of the N7/N8 zone boundary. 3

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Fig. 5. Locations of the Nagano River and R163 sections and sampling sites. The arrow at the end of the section denotes the stratigraphic upsection.

and more than half of them had large α95 (radius of the 95% confidence limit; Fisher, 1953) values of > 20° and small k (precision parameter; Fisher, 1953) values of < 20; three out of the four normal polarity directions corresponded to the insufficient data. For these reasons, we decided to revisit the upper Ichishi Group to obtain reliable magnetostratigraphic data. In our U–Pb and magnetostratigraphic study, we investigated two stratigraphic sections, the Nagano River and Route 163 (R163) sections, that together cover the upper Ichishi Group (Fig. 5). Lithologic logs of these sections are shown in Fig. 6 with the stratigraphic positions of the U–Pb and magnetostratigraphic sampling sites. The Nagano River section is the same as that of Oshida et al. (2018), although their biostratigraphic sampling was concentrated in its upper half (Fig. 6). At the base of this section are unsorted boulder conglomerates/breccias (< 30 m in thickness) that unconformably overlie the Cretaceous granitic basement. These coarse-grained sediments are not the basal conglomerate of the Ichishi Group but a local marginal lithology developed around a basement high. Upsection, there are sandstones ∼10 m in thickness containing granules and pebbles derived from the basement. Further upsection, there are relatively finegrained sediments (e.g., dark gray mudstones, dark gray sandy mudstones, and alternating mudstones and sandstones; in total ∼430 m in thickness) that comprise the main lithology of the Mitsugano Member. The member contains some intercalated felsic tuff beds typically several centimeters to ∼10 cm in thickness and the mudstones are partly tuffaceous. There is a white to pale cream, felsic tuff bed ∼2 m above the

LOD of P. sicana and the same tuff bed is also seen near the base of the R163 section which will be described later. This tuff bed, ∼3 cm in thickness, is laterally traceable but intermittent because of bioturbation. We sampled this tuff bed for U–Pb dating (site Tf-1). The Mitsugano Member is conformably overlain by the Chaya Member of the Katada Formation and a conformable contact can be seen between paleomagnetic sampling sites 15 and 16. The Chaya Member is mainly composed of massive sandstones and turbidites (graded beds of sandstone to mudstone) with minor debrites at some stratigraphic horizons that contain numerous mudstone rip-up clasts. The Nagano River section does not reach the upper part of the Katada Formation. The R163 section stratigraphically covers the Katada Formation, although outcrop exposures are relatively poor compared to those of the Nagano River section. The basal part of this section consists of dark gray mudstones (∼15 m in thickness) of the Mitsugano Member in which the Tf-1 marker felsic tuff (as previously described) is intercalated. The dominant lithology of the lower half of the section is massive sandstones and turbidites of the Chaya Member and that of the upper half is dark gray to gray mudstones and alternating mudstones and sandstones of the Yakuoji Member. The boundary between the Chaya and Yakuoji members is unclear in this section because of discontinuous outcrops. The uppermost part of the section, near the stratigraphic top of the Ichishi Group, consists of a 14-m-thick portion of alternating sandstones and mudstones in which we found an intercalated, 10–30-cm thick, pale blue to white, fine-grained felsic tuff bed. We sampled this tuff for U–Pb dating (site Tf-2; Fig. 6). The Ichishi

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Group is unconformably overlain or in fault contact with the Plio–Pleistocene sediments of the Tokai Group (Yoshida et al., 1995) but the contact cannot be seen in this section. 3. Materials and methods 3.1. U–Pb dating Felsic tuffs were sampled at sites Tf-1 and Tf-2 for dating. Care was taken to avoid contamination from the strata immediately above and below the tuff layers. Zircon crystals were separated from the tuff samples by crushing, sieving, water-based panning, and magnetic and heavy liquid separation techniques in the laboratory at Kyoto FissionTrack, Japan. Approximately 100 zircon grains were randomly handpicked, mounted in a PFA Teflon sheet (Danhara et al., 1993), and polished with diamond paste. Before dating using laser ablation-inductively coupled plasma mass spectrometry (LA-ICPMS), the mounts were washed in a 0.3 M HNO3 solution to reduce Pb contamination. U–Pb isotopic analyses were conducted using a multiple collectorICPMS (MC-ICPMS; Nu Plasma II, Nu Instruments Ltd., Wrexham, UK) combined with a Type-C Ti:S femtosecond laser (IFRIT, Cyber Laser Inc., Tokyo, Japan) installed in the Geochemical Research Center, University of Tokyo. The femtosecond laser pit diameter was approximately 10 μm. The use of the galvanometric optics enabled creation of a larger ablation pit via multiple laser ablations over short time intervals (Yokoyama et al., 2011). The laser was operated at a repetition rate of 10 Hz and fluence of 2.83 Jcm−2 and the ablation area was 30 μm × 30 μm. Helium was used as the carrier gas inside the ablation cell and was mixed with argon gas before entering the ICPMS. Three full-size electron multipliers were used for the simultaneous detection of the following ion signals: 202Hg, 204(Hg + Pb), and 208Pb. Signal intensities of 206 Pb, 207Pb, and 238U were monitored using multiple Daly detectors and 232Th using a faraday collector (Hattori et al., 2017; Obayashi et al., 2017). 207Pb/235U ratios were calculated based on the constant value of 137.88 for the 238U/235U natural abundance ratio. Elemental fractionation and instrumental mass bias of 206Pb/238U were corrected using the isotopic ratio of 0.17928 for the 91500 zircon reference material (Wiedenbeck et al., 1995). Mass bias of the 207Pb/206Pb and 208 Pb/206Pb ratios were normalized to those of NIST SRM612 standard glass, 0.90726 and 2.164, respectively. External reproducibilities of the 206 Pb/238U, 207Pb/207Pb, and 208Pb/206Pb ratios were determined via repeated analysis of the 91500 zircon and NIST SRM612. To confirm quality control of the dating result, OD-3 zircons were used as a secondary standard (∼33 Ma; Iwano et al., 2013; Lukács et al., 2015). Cathodoluminescence (CL) images were taken to confirm the internal structures of the analyzed zircons following LA-ICPMS analysis. To accurately determine U–Pb ages, one must consider the effect of initial disequilibria associated with intermediate nuclides in the 238U and 235U decay series. Furthermore, the contribution from non-radiogenic Pb (common Pb) to the U–Pb age is also not negligible in the dating of young or low-U zircons. To reduce both effects, we applied the modified 207Pb method of Sakata (2018). In this method, it is required to determine Th/U and Pa/U ratio partitioning between zircon and the melt system (i.e. fTh/U = (Th/U)zircon/(Th/U)melt and fPa/U = (Pa/ U)zircon/(Pa/U)melt, respectively). We used the average values of fTh/ U = 0.20 ± 0.05 and fPa/U = 3.36 ± 0.20 as estimated from previous literatures (Schmitt, 2011; Sakata et al., 2017; Sakata, 2018). Common 207 Pb/206Pb was calculated to be 0.83668 at 17 Ma based on the terrestrial Pb isotopic evolutionary model of Stacey and Kramers (1975). Mean ages were calculated using the Isoplot program (Ludwig, 2008).

Fig. 6. Lithologic columns of the two mapped sections with the stratigraphic positions of sampling sites (italicized numbers = magnetic, orange Tf-1 and Tf2 = U–Pb). NR = Nagano River. Biostratigraphic ages reported by Oshida et al. (2018) are indicated by colored symbols.

3.2. Magnetic measurements The Nagano River and R163 sections contain 23 and 10 sites for

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Table 1 Site-mean remanent magnetization directions for category A and B sites. Site

Lithology

N [nl /np]

Du (°)

Iu (°)

Dir/Dip (°)

Dc (°)

Ic (°)

α95 (°)

k

Pol

7 8* 9 10 11 12* 13 14* 15 17 21 24 26B 27 28 29* 30 31* 32

Mud Mud Mud Tfc mud Fel tuff Fel tuff Sandy mud Fel tuff Mud Sandy mud Mud Fine sand Fel tuff Fel tuff Fel tuff Fel tuff Sandy mud Mud Sandy mud

4 [2/5] 4 4 4 [1/10] 4 [2/9] 3 3 4 5 6 6 6 [1/2] 6 [1/24] 6

217.7 21.7 45.6 20.4 33.9 207.7 201.4 222.1 206.7 215.5 28.0 233.3 112.6 182.0 201.1 149.4 192.9 187.5 175.3

−64.0 59.5 55.0 51.7 46.6 −57.0 −53.7 −54.2 −29.1 −67.2 42.9 −51.4 −51.0 −47.3 −59.4 −50.6 −57.6 −58.8 −51.9

99/15 99/15 91/10 90/12 89/15 108/10 108/10 93/4 67/10 53/11 65/6 88/6 53/43 113/22 113/22 93/35 99/10 115/9 115/9

237.6 42.7 53.8 32.8 44.2 221.8 214.5 226.0 209.7 220.8 31.0 237.0 171.7 207.5 234.1 199.4 207.8 202.7 186.5

−54.5 53.4 47.5 46.3 36.9 −54.1 −52.0 −51.5 −21.3 −56.6 38.0 −46.4 −51.8 −50.6 −53.6 −56.6 −55.6 −60.4 −55.6

10.6 6.7 10.2 10.8 8.4 4.9 6.3 4.0 52.5 16.8 5.1 6.6 6.1 8.3 4.0 37.0 4.4 3.4 6.4

76.3 98.5 82.6 73.8 120.8 102.4 211.8 149.6 6.6 54.8 326.9 134.0 123.1 65.5 285.5 31.7 229.9 79.5 111.7

R N N N N R R R R R N R R R R R R R R

Note. *Determined by applying the great circle method for a set of direct lines (specimen ChRM directions) and best-fit great circles. Lithology: Fel = felsic; Mud = mudstone; Sand = sandstone; Tfc = tuffaceous. N = number of independently oriented core samples. nl = number of direct lines (specimen ChRM directions). np = number of planes (best-fit great circles). Du and Iu = tilt-uncorrected (in situ) declination and inclination. Dir/Dip = direction of dip/angle of dip. Dc and Ic = tilt-corrected declination and inclination. α95 = radius of 95% confidence circle. k = precision parameter. Pol = polarity (N = normal; R = reverse).

magnetic measurements, respectively (33 sites in total; Figs. 5 and 6). The rock type of each site is summarized in Table 1. We used a batterypowered portable drill for collecting independently oriented core samples 25 mm in diameter and a magnetic compass for core orientation. Between four and eight cores were collected at each site. The cores were > 50 mm (typically c. 100 mm) in length for accurate orientation as well as to obtain two or more 22-mm cylindrical specimens from each core. Paleomagnetic measurements were conducted in the magnetically shielded room at the Center for Advanced Marine Core Research, Kochi University, Japan. Natural remanent magnetization (NRM) was measured using a 2G-760R pass-through superconducting rock magnetometer. Specimens (n = 569) were demagnetized stepwise following measurement of NRM. Alternating-field demagnetization (AFD) was performed to a maximum field of 80 mT using an in-line demagnetization system of the pass-through magnetometer and thermal demagnetization (ThD) was performed in air up to a maximum temperature of 600 °C using a Natsuhara TDS-1 thermal demagnetizer. First, as a pilot study, we selected four or more specimens per site and subjected them to a detailed stepwise demagnetization experiment. During stepwise ThD, thermal alteration of minerals was checked by measuring the initial magnetic susceptibility of specimens using a Bartington MS2B meter. Demagnetization results were analyzed by plotting them on both an orthogonal and an equal-area diagram. Then, all remaining specimens were demagnetized stepwise using fewer steps than the detailed experiment. Principal component analysis (Kirschvink, 1980) was applied using the PuffinPlot v.1.02 software (Lurcock and Wilson, 2012) for demagnetization data to determine the directions of the characteristic remanent magnetization (ChRM) components or the best-fit great circles if the ChRM components were not recognizable. We adopted ChRM directions and best-fit great circles with maximum angular deviation (MAD) values of < 15°. To infer magnetic mineralogy, we performed a rock magnetic experiment for selected specimens in which three-axis isothermal remanent magnetization (IRM) was thermally demagnetized (Lowrie, 1990). Three-axis IRM was produced by applying different steady fields

of 3.00, 0.40, and 0.12 T along the three orthogonal axes of the specimens using a Magnetic Measurements MMPM-10 pulse magnetizer. The resultant high-, intermediate- and low-coercivity components were measured using a spinner magnetometer to show the change in intensity with ThD temperature. This experiment was performed in the paleomagnetic laboratory at Aichi University of Education, Japan. 4. Results and interpretation 4.1. U–Pb ages The two tuff samples yielded abundant zircons (Tf-1 = ∼5000 grains from 0.20 kg of rock; Tf-2 = ∼300 grains from 0.50 kg of rock). The Tf-1 zircons were euhedral and pale pink and the Tf-2 zircons were small and of a higher uranium concentration compared to those from Tf-1. In both samples, most zircons were generally homogeneous in terms of age (∼17 Ma; Table 2) and had no inherited cores, as can be seen in the CL images in Supplementary Fig. A. Of the zircons measured, two from Tf-1 showed ages of ∼190 Ma (early Jurassic) and three from Tf-2 showed ages of ∼74 to ∼65 Ma (late Cretaceous). The CL images of these grains show the presence of inherited cores (Supplementary Fig. A). Clearly, they are determined as detrital grains. We finally obtained weighted mean ages of 17.03 ± 0.11 Ma for Tf-1 (95% confidence, n = 28/30, MSWD = 3.1; Fig. 7a and c) and 17.09 ± 0.06 Ma for Tf-2 (95% confidence, n = 25/27, MSWD = 1.03; Fig. 7b and d). Both mean ages were precisely determined with their 95% confidence bounds less than 1% of the ages, and this may result from the application of the modified 207Pb method (Sakata, 2018). The wider grain age distribution for Tf-1 can be attributed to the low U concentration (mostly less than 100 ppm), or lower counting statistics, or both. The two U–Pb ages of ∼17 Ma are consistent with diatom ages for two sites in the Mitsugano Member (between ∼18.1 and ∼17.0 Ma as previously described). It is thus reasonable to assume that the U–Pb ages represent the eruptive/depositional ages of the ashes of the two tuff beds.

6

7

81 97 50 70 48 113 84 77 110 101 96 76

Tf-2 Selected Tf-2-1 Tf-2-2 Tf-2-3 Tf-2-4 Tf-2-5 Tf-2-6 Tf-2-8 Tf-2-9 Tf-2-10 Tf-2-11 Tf-2-12 Tf-2-14

0.85 0.91 0.77 0.87 0.76 0.93 0.87 0.83 0.89 0.99 0.65 0.88

0.61 0.47

184 135

grains 96 107 64 80 63 122 97 93 124 102 149 87

1.08 1.30 1.03 1.09 1.09 0.90 0.89 1.09 0.85 1.04 1.11 0.85 1.21 1.12 1.03 1.13 1.06 1.09 1.01 1.05 0.98 1.02 1.12 1.02 1.02 0.89 1.03 1.04 1.00 1.06

36 183 43 49 56 25 34 49 34 54 63 36 143 63 38 56 65 52 45 32 38 45 59 36 48 26 30 46 44 103

0.01664 0.01689 0.01685 0.01823 0.01636 0.01709 0.01609 0.01829 0.01773 0.01650 0.01689 0.01769

0.21959 0.20719

0.01708 0.01721 0.01786 0.01712 0.01641 0.02079 0.01783 0.01706 0.02068 0.01663 0.01903 0.02006 0.01788 0.01755 0.01940 0.01919 0.01734 0.01966 0.01886 0.01702 0.01752 0.01807 0.01655 0.01715 0.01704 0.01800 0.01804 0.01664 0.01900 0.01752

0.00098 0.00094 0.00126 0.00113 0.00126 0.00087 0.00096 0.00103 0.00088 0.00094 0.00078 0.00106

0.00349 0.00334

0.00194 0.00080 0.00166 0.00160 0.00145 0.00233 0.00177 0.00161 0.00181 0.00144 0.00140 0.00169 0.00086 0.00134 0.00181 0.00148 0.00126 0.00154 0.00159 0.00199 0.00169 0.00156 0.00135 0.00185 0.00146 0.00209 0.00210 0.00153 0.00159 0.00095

2σ

0.00267 0.00261 0.00265 0.00265 0.00267 0.00263 0.00265 0.00263 0.00265 0.00259 0.00264 0.00262

0.02986 0.02992

0.00266 0.00266 0.00268 0.00263 0.00260 0.00262 0.00262 0.00256 0.00269 0.00269 0.00270 0.00267 0.00263 0.00267 0.00257 0.00263 0.00262 0.00275 0.00264 0.00254 0.00275 0.00267 0.00261 0.00270 0.00263 0.00259 0.00258 0.00250 0.00267 0.00264

Pb/238U

206

0.00005 0.00005 0.00006 0.00005 0.00006 0.00005 0.00005 0.00005 0.00005 0.00005 0.00005 0.00005

0.00041 0.00041

0.00006 0.00004 0.00006 0.00006 0.00006 0.00007 0.00006 0.00006 0.00006 0.00006 0.00005 0.00005 0.00003 0.00004 0.00006 0.00005 0.00004 0.00005 0.00005 0.00006 0.00005 0.00005 0.00005 0.00006 0.00005 0.00006 0.00006 0.00005 0.00005 0.00004

2σ

Pb/206Pb

207

2σ

0.04517 0.04698 0.04604 0.04992 0.04448 0.04711 0.04410 0.05043 0.04844 0.04619 0.04633 0.04901

0.05334 0.05023

0.04654 0.04695 0.04824 0.04714 0.04579 0.05749 0.04942 0.04841 0.05573 0.04478 0.05104 0.05446 0.04935 0.04774 0.05472 0.05295 0.04800 0.05189 0.05190 0.04861 0.04629 0.04911 0.04592 0.04599 0.04696 0.05032 0.05069 0.04821 0.05167 0.04821

0.00250 0.00244 0.00330 0.00292 0.00331 0.00223 0.00249 0.00267 0.00221 0.00248 0.00196 0.00278

0.00042 0.00042

0.00517 0.00205 0.00436 0.00429 0.00393 0.00626 0.00477 0.00443 0.00471 0.00376 0.00362 0.00447 0.00230 0.00356 0.00497 0.00398 0.00341 0.00395 0.00425 0.00559 0.00438 0.00414 0.00367 0.00486 0.00393 0.00573 0.00579 0.00433 0.00422 0.00253

16.76 17.01 16.97 18.35 16.48 17.21 16.21 18.40 17.84 16.62 17.00 17.80

201.57 191.20

17.19 17.32 17.97 17.24 16.52 20.89 17.94 17.17 20.79 16.75 19.14 20.17 17.99 17.67 19.51 19.30 17.46 19.76 18.98 17.14 17.64 18.19 16.67 17.27 17.16 18.12 18.16 16.76 19.11 17.63

U-207Pb

235

Pb/235U

207

Th/U

U

Th

Age (Ma)

Isotope ratios

Elemental concentration (μg/g) and ratio

Tf-1 Selected grains Tf-1-156 34 Tf-1-157 141 Tf-1-158 42 Tf-1-159 45 Tf-1-160 51 Tf-1-164 28 Tf-1-165 39 Tf-1-166 45 Tf-1-167 40 Tf-1-168 52 Tf-1-170 57 Tf-1-171 42 Tf-1-172 119 Tf-1-173 56 Tf-1-174 37 Tf-1-175 50 Tf-1-176 61 Tf-1-177 47 Tf-1-178 45 Tf-1-181 31 Tf-1-185 39 Tf-1-186 45 Tf-1-187 52 Tf-1-188 35 Tf-1-191 47 Tf-1-192 29 Tf-1-199 29 Tf-1-202 44 Tf-1-203 44 Tf-1-204 97 Excluded detrital grains Tf-1-162 300 Tf-1-169 286

Sample

Table 2 U–Th–Pb isotopic data for two tuff samples (Tf-1 and Tf-2) and a secondary standard (OD-3).

0.98 0.93 1.26 1.13 1.26 0.87 0.96 1.03 0.87 0.94 0.78 1.06

2.91 2.81

1.94 0.80 1.66 1.60 1.45 2.31 1.77 1.60 1.80 1.44 1.40 1.68 0.86 1.34 1.81 1.48 1.26 1.53 1.58 1.99 1.69 1.55 1.35 1.84 1.46 2.08 2.10 1.53 1.59 0.95

2σ

17.20 16.79 17.09 17.06 17.17 16.94 17.04 16.93 17.09 16.68 17.02 16.85

189.67 190.04

17.13 17.11 17.29 16.96 16.73 16.89 16.84 16.45 17.33 17.34 17.41 17.20 16.91 17.17 16.55 16.92 16.87 17.68 16.97 16.35 17.67 17.19 16.83 17.41 16.94 16.71 16.62 16.12 17.17 16.97

U-206Pb

238

*

*

0.33 0.32 0.36 0.34 0.36 0.32 0.33 0.33 0.32 0.32 0.31 0.33

2.58 2.58 Wtd. mean ± 2σ MSWD n Rejected (*)

0.41 0.28 0.38 0.37 0.35 0.44 0.39 0.37 0.39 0.36 0.35 0.33 0.22 0.29 0.36 0.30 0.28 0.32 0.32 0.37 0.34 0.32 0.29 0.36 0.31 0.38 0.38 0.32 0.32 0.23

2σ

17.29 16.88 17.18 17.08 17.26 17.02 17.13 16.94 17.14 16.77 17.10 16.89

17.03 ± 0.11 3.1 28 2

17.22 17.20 17.34 17.04 16.82 16.75 16.88 16.51 17.22 17.43 17.41 17.12 16.95 17.24 16.47 16.88 16.93 17.66 16.95 16.40 17.76 17.22 16.92 17.50 17.03 16.72 16.63 16.18 17.15 17.02

U-Pb age (Ma)

(continued on next page)

0.34 0.32 0.36 0.34 0.36 0.32 0.33 0.33 0.32 0.32 0.31 0.33

0.41 0.28 0.39 0.37 0.36 0.44 0.39 0.37 0.39 0.36 0.35 0.33 0.22 0.29 0.36 0.31 0.28 0.32 0.32 0.38 0.34 0.32 0.30 0.36 0.31 0.39 0.39 0.32 0.33 0.24

2σ

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2σ

Pb/238U

206

2σ

Pb/206Pb

207

2σ

0.81 0.72 0.86 0.71 0.60 0.81 0.64 0.86 0.78

0.62 0.71 0.67

68 78 70

OD-3: Secondary standard OD3 1-1 172 139 OD3 1-2 127 92 OD3 2-1 145 124 OD3 3-1 142 100 OD3 3-2 92 56 OD3 4-1 199 161 OD3 7-1 120 76 OD3 8-1 278 240 OD3 9-1 187 145

0.96 1.07 0.88 0.89 0.91 0.93 0.99 0.93 0.94 0.92 0.87 0.93 0.67 0.90 0.96

83 242 112 69 140 70 132 78 144 85 77 130 52 89 139

0.03461 0.03381 0.03277 0.03301 0.03261 0.03411 0.03826 0.03354 0.03345

0.07634 0.06740 0.06801

0.01823 0.01762 0.01722 0.01816 0.01855 0.01836 0.01675 0.01778 0.01714 0.01699 0.01734 0.01670 0.01625 0.01750 0.01728

0.00110 0.00122 0.00113 0.00113 0.00138 0.00098 0.00132 0.00077 0.00092

0.00214 0.00197 0.00193

0.00107 0.00064 0.00086 0.00117 0.00081 0.00120 0.00083 0.00111 0.00078 0.00103 0.00106 0.00081 0.00113 0.00100 0.00080

0.00518 0.00504 0.00505 0.00514 0.00507 0.00515 0.00523 0.00515 0.00515

0.01148 0.01018 0.01052

0.00268 0.00272 0.00265 0.00267 0.00276 0.00263 0.00266 0.00263 0.00267 0.00267 0.00260 0.00266 0.00263 0.00265 0.00265

0.00009 0.00009 0.00009 0.00007 0.00007 0.00006 0.00008 0.00005 0.00005

0.00019 0.00017 0.00013

0.00005 0.00004 0.00004 0.00005 0.00004 0.00004 0.00004 0.00004 0.00004 0.00004 0.00004 0.00004 0.00004 0.00004 0.00004

0.04842 0.04863 0.04706 0.04660 0.04667 0.04805 0.05303 0.04724 0.04707

0.04825 0.04803 0.04688

0.04941 0.04701 0.04711 0.04931 0.04869 0.05061 0.04568 0.04908 0.04662 0.04619 0.04837 0.04561 0.04485 0.04786 0.04736

0.00129 0.00153 0.00140 0.00147 0.00186 0.00125 0.00163 0.00099 0.00122

0.00108 0.00115 0.00120

0.00272 0.00160 0.00224 0.00307 0.00201 0.00318 0.00216 0.00296 0.00201 0.00269 0.00285 0.00211 0.00302 0.00264 0.00208 2.02 1.88 1.84

1.06 0.64 0.85 1.17 0.81 1.19 0.83 1.11 0.78 1.03 1.06 0.81 1.13 1.00 0.80

2σ

34.55 1.08 33.76 1.20 32.74 1.11 32.98 1.11 32.58 1.36 34.06 0.96 38.13 1.29 33.50 0.75 33.41 0.91 Wtd. mean ± 2σ MSWD n

74.69 66.23 66.81

18.34 17.74 17.34 18.27 18.66 18.47 16.87 17.89 17.26 17.11 17.46 16.82 16.36 17.61 17.39

U-207Pb

235

Pb/235U

207

Th/U

U

Th

Age (Ma)

Isotope ratios

Elemental concentration (μg/g) and ratio

Tf-2-15 86 Tf-2-16 227 Tf-2-17 128 Tf-2-18 78 Tf-2-19 153 Tf-2-21 75 Tf-2-22 134 Tf-2-23 84 Tf-2-24 154 Tf-2-25 92 Tf-2-26 88 Tf-2-27 140 Tf-2-28 78 Tf-2-29 99 Tf-2-30 145 Excluded detrital grains Tf-2-7 109 Tf-2-13 109 Tf-2-20 104

Sample

Table 2 (continued)

8

*

*

1.23 1.10 0.83 Wtd. mean ± 2σ MSWD n Rejected (*)

0.34 0.23 0.25 0.29 0.25 0.29 0.25 0.28 0.24 0.28 0.27 0.25 0.29 0.27 0.25

2σ

33.34 0.57 32.42 0.57 32.47 0.56 33.04 0.43 32.58 0.45 33.10 0.41 33.65 0.52 33.11 0.30 33.14 0.32 33.03 ± 0.26 2.50 9

73.55 65.28 67.47

17.22 17.50 17.07 17.20 17.79 16.94 17.12 16.91 17.17 17.18 16.74 17.10 16.91 17.07 17.04

U-206Pb

238

17.09 ± 0.06 1.03 25 2

17.26 17.59 17.15 17.23 17.84 16.95 17.21 16.95 17.25 17.26 16.80 17.19 17.00 17.14 17.11

U-Pb age (Ma) 0.34 0.23 0.25 0.29 0.25 0.29 0.25 0.28 0.25 0.28 0.28 0.25 0.29 0.27 0.25

2σ

Disequilibrium & common Pb-corrected

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Fig. 7. U–Pb zircon dating results from two tuff samples (Tf-1 and Tf-2). (a, b) Concordia diagrams showing U–Pb data from selected grains without older detrital grains. Data are uncorrected for the initial disequilibria and common Pb effects. (c, d) Grain ages and their weighted means following application of the modified 207 Pb method. Diagrams were created using Isoplot v.4.15.

4.2. Demagnetization experiments

components predominated from room temperature to 600 °C and had a maximum unblocking temperature of 550–600 °C. In addition, magnetic iron sulfide, probably greigite, may also be present in the latter group, as suggested from a bend in the curve of the low- and intermediate-coercivity components between 300 and 350 °C (Fig. 9b). Such a bend was not seen in the IRM demagnetization results for the former group (Fig. 9a). Susceptibility curves characterized by a gentle or rapid increase above ∼375–400 °C were observed in all specimens of the latter group and several (not all) of the former group (Figs. 8c–f and 9). This may represent the alteration of magnetic iron sulfide to a strongly magnetic phase (probably magnetite). In summary, we interpret that the ChRM is mainly carried by magnetite and also partially by magnetic iron sulfide. The coexistence of magnetite and magnetic iron sulfide (particularly greigite) is not uncommon in marine clastic sediments (e.g., Hoshi and Yamada, 2016; Okada et al., 2017). Importantly, there was no recognizable change in the magnetization direction in the specimens of the former group that showed the increase in susceptibility (Fig. 8c–e). This observation suggests little or no difference between the direction carried by the magnetite and that by the magnetic iron sulfide. The magnetic iron sulfide in the Ichishi Group sediments is believed to have been formed by early diagenetic processes immediately following deposition.

The NRM intensity of ∼90% of the measured specimens ranged from ∼1.0 × 10−4 to ∼6.5 × 10−3 A/m, with a maximum of ∼8.0 × 10−2 A/m (sandy mudstone at site 13) and a minimum of ∼2.2 × 10−5 A/m (mudstone at site 5). The directions of ChRMs were determined for 169 specimens (∼30% of total) and best-fit great circles were determined for 156 specimens (∼27% of total). At sites where the ChRM directions were obtained through both AFD and ThD, the directions using the two methods were consistent (Fig. 8). Soft magnetic overprints believed to be of viscous origin were removed by AFD to 20 mT or by ThD to 200 °C prior to the appearance of ChRMs, as best shown in Fig. 8c and e. In AFD, the median destructive fields were generally 20–30 mT, indicating the moderate coercivities of the magnetic minerals. ThD results yielding ChRMs from 53 specimens were divided into two main groups: (1) those showing a maximum unblocking temperature of ∼500 °C or higher, produced from ∼70% of the specimens (Fig. 8a–e), and (2) those showing a magnetization that mostly decayed by 325–350 °C, produced from ∼30% of the specimens (Fig. 8f). ThD curves of the composite IRM were basically similar between both groups (Fig. 9). Magnetite is suggested to be the main IRM carrier for both groups on the basis of the observation that low-coercivity

9

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Fig. 8. Orthogonal plots showing representative demagnetization data showing ChRM components. Six pairs of AFD and ThD results from the same sites are presented as follows: (a) mudstones from site 9, (b) sandy mudstones from site 13, (c) felsic fine tuffs from site 11, (d) sandy mudstones from site 30, (e) fine sandstones from site 24, and (f) mudstones from site 21. Solid (open) circles denote vector endpoints on the horizontal (vertical) plane and those that are red indicate data used to calculate ChRM directions. Values in square brackets are NRM intensities in 10−3 A/m. Small graphs attached to the ThD plots show normalized variations in the initial magnetic susceptibility with temperature. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article).

10

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Fig. 9. Examples of ThD of composite IRM . (a) Fine sandstone from site 24. (b) Mudstone from site 21. The variation in initial magnetic susceptibility with temperature is also shown. Fig. 10. Equal-area projections of site-mean directions. The solid (open) symbol shows the direction in the lower (upper) hemisphere with the 95% confidence limit shown by the circle around the direction. (a, b) Tilt-uncorrected (in situ) and tilt-corrected directions for 18 sites of categories A and B (site 15 is excluded because of its large α95 value). Red and blue symbols indicate the category A and category B directions, respectively. The dotted circle around the center of the projections corresponds to the inclination I = ± 54.2° expected at the representative location of Ichishi (34.7 °N, 136.4 °E) when assuming the geocentric axial dipole (GAD) field and the cross is the normal polarity GAD field direction. (c, d) Tilt-corrected category A directions of the R1–N1 and R2 polarity zones. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article).

11

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4.3. Site-mean directions

In summary, we conclude that the 13 category A site-mean directions are a paleomagnetic record acquired before tilting and have been stable over a geologically long period. It would be reasonable to regard them as a depositional remanent magnetization (DRM) or a chemical remanent magnetization (CRM) acquired at or soon after deposition because they form a cleanly defined magnetostratigraphic zonation (Fig. 11a), as will be described. The same argument can also be made for the five category B directions as they have a polarity that is stratigraphically consistent with that of the category A directions (Fig. 11a).

The paleomagnetic sites examined in this study (n = 33) were classified into three categories (A, B, and C) as described in the following and site-mean directions were determined for category A and B sites (Table 1). Category A (n = 14; sites 7, 9, 10, 11, 13, 15, 17, 21, 24, 26B, 27, 28, 30, and 32): These sites have ChRM directions from more than two independently oriented cores (= samples). When two or more specimens cut from a single sample yielded ChRM directions, we calculated a mean direction for that sample to average the possible internal magnetic inhomogeneity. Each sample direction was given a unit weight and sample directions were used to calculate site-mean directions and their statistical parameters (α95 and k). Category B (n = 5; sites 8, 12, 14, 29, and 31): These sites have only one or two ChRM directions as well as more than one best-fit great circle. By applying the great circle method of McFadden and McElhinny (1988), site-mean directions and their statistical parameters were calculated using these ChRM directions and great circles with the unit weight for each. Category C (n = 14; sites 1, 2, 3, 4, 5, 6, 16, 18, 19, 20, 22, 23, 25, and 26A): No ChRM directions were obtained for these sites although some specimens provided best-fit great circles. We excluded these sites from consideration. The site-mean directions of categories A and B in geographic coordinates (= tilt-uncorrected) and those in stratigraphic coordinates (= tilt-corrected) are shown in Fig. 10a and b, respectively. For tilt correction, site-mean directions were corrected for tilt using strike/dip data measured at or near the individual sites. In Fig. 10a and b, the red symbols illustrate category A directions except for that of site 15 which has a very large α95. The blue symbols indicate category B directions. Among the tilt-corrected directions (Fig. 10b), five have a positive inclination and a northeast declination and 13 have a negative inclination and a broadly south to southwest declination. We regard the former and the latter as normal and reverse polarity directions, respectively. The normal polarity directions are not a recently acquired secondary magnetization because, before tilt correction, they are significantly different from the geocentric axial dipole field direction. The tilt-corrected directions of the dual polarities are broadly deflected clockwise of the north–south direction. This finding is compatible with previous results (Hayashida and Ito, 1984) and would be a product of clockwise rotation, as will be discussed later. Although the normal and reverse polarity directions are broadly antipodal, tilt correction does not improve the antipodality. We performed bootstrap reversals test (Tauxe et al., 1991) on 13 tilt-corrected category A directions (except site 15) using the PmagPy program (Tauxe et al., 2016) and obtained a negative result, suggesting a difference in the 95% confidence level between the two directional groups. The most likely cause of this is the small number of normal polarity directions compared to those that are reverse. However, tilt correction improves the clustering of the 13 category A directions (except for site 15). After inverting the reverse polarity directions to the corresponding normal polarity directions, the tilt-uncorrected overall mean direction is D = 19.7°, I = 57.2°, α95 = 9.5°, and k = 19.9 and the tilt-corrected overall mean direction is D = 36.0°, I = 51.1°, α95 = 7.1°, and k = 34.7. The improvement in clustering is mainly because of the significant change in the direction for site 26B (Fig. 10a and b) where the bedding dip is steep compared to that of other sites (Table 1). Notably, when considering category A directions excluding site 26B, the clustering is still better following tilt correction than that prior (D = 25.0°, I = 55.2°, α95 = 6.6°, and k = 43.7 before tilt correction; D = 39.3°, I = 50.4°, α95 = 6.3°, and k = 48.6 after tilt correction). The bootstrap fold test of Tauxe and Watson (1994) provides an inconclusive result regardless of whether site 26B is included or excluded, and this is most likely because of the lack of variation in shallow bedding dips.

4.4. Magnetostratigraphy The polarity sequence of the 19 site-mean directions of categories A and B is straightforward (Fig. 11a) and allows definition of a magnetostratigraphic zonation consisting of three magnetic polarity zones: Zones R1, N1, and R2 in ascending stratigraphic order. Zone R1 is a ∼250 m portion of the Mitsugano Member with its lower limit defined by site 30 and its upper limit between sites 12 and 11. Zone N1 is a ∼70 m portion of the Mutsugano Member and its upper limit is between sites 8 and 7. Zone R2 is a ∼450 m sequence spanning from the upper part of the Mitsugano Member to the uppermost part of the Yakuoji Member with site 26B defining its upper limit. Zone R2 contains the normal polarity horizon of site 21, for which we assign a normal polarity subzone termed Subzone R2-n1. The robustness of the normal polarity for this site is confirmed by the fact that the site-mean direction is precisely determined using only ChRM directions and has the same directional character as that of the normal polarity sites of Zone N1 (Fig. 10a and b). Our data show that the LOD of P. sicana, marking the base of Zone N8, is in the lower part of Zone R2. This is in contrast to a combination of published data (Hayashida and Ito, 1984; Oshida et al., 2018) that suggests the datum within a normal polarity-dominated interval (Fig. 3). 5. Discussion 5.1. Age of the N7/N8 (M4/M5) zone boundary The U–Pb ages determined for two tuff beds suggest that the LOD of P. sicana is at ∼17.0 Ma. They are indistinguishable within the uncertainties, suggesting a high sedimentation rate (mostly of the Katada Formation) between the two tuff horizons. The age for Tf-2 provides an upper age limit not only for the group but also for the LOD of P. sicana. The age for Tf-1 can be interpreted as approximating the age of the datum because site Tf-1 is quite near it, with a stratigraphic separation of only ∼2 m. We have found no evidence of a sedimentary discontinuity, such as an erosional surface or a glauconite bed, within the ∼2-m mudstone column between the LOD of P. sicana and site Tf-1. Using the two U–Pb ages, the R1–N1–R2 magnetostratigraphic polarity sequence can be easily correlated with a sequence of geomagnetic polarity chrons at approximately 17 Ma (Fig. 11b). Zone R2, which contains two dated tuff beds, can be confidently correlated with Chron C5Cr. This correlation enables Zones N1 and R1 to be correlated with Chrons C5Dn and C5Dr, respectively. This magnetostratigraphic correlation suggests that the composite section we investigated is older than the upper boundary age of Chron C5Cr (∼16.7 Ma) and younger than the lower boundary age of Chron C5Dr (∼18.1 Ma). The N7/N8 (M4/M5) boundary age should therefore be older than ∼16.7 Ma. Here, we need to address the possible influence of magma residence time because the U–Pb zircon age has a very high closure temperature (e.g., Cherniak and Watson, 2001). In a strict sense, the U–Pb ages for the Tf-1 and Tf-2 tuffs may only represent the lower age limits of the eruptive/depositional ages, owing to the possible presence of unknown magma residence times. In practice, it is generally difficult to elucidate the residence times. However, the magma residence time is generally estimated to be < 100 kyr in the case of eruptive volumes of 12

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Fig. 11. (a) Synthesis of the lithostratigraphy, biostratigraphy, magnetostratigraphy, and U–Pb ages of the upper Ichishi Group. See Fig. 6 for the key to the lithologic columns. Solid (open) circles show normal (reverse) polarity. Foraminifera and diatom data are after Oshida et al. (2018). (b) Correlation of the magnetostratigraphy with the astronomically tuned geomagnetic polarity time scale and an age-depth model (sedimentation rate curve). The gray band shows the correlation between Zone N1 and Chron C5Dn.

approximately 10 km3 (Costa, 2008), with the possible exception of caldera-forming supereruptions with eruptive volumes exceeding 100 km3. Even if the Tf-1 and Tf-2 tuffs are the products of a gigantic eruption and the magma residence times are assumed to be ∼200 kyr (corresponding to the Huckleberry Tuff eruption of the Yellowstone supervolcano) or ∼300 kyr (approximately corresponding to the Fish Canyon Tuff supereruption), Zone R2 containing the two dated tuff horizons can be correlated with Chron C5Cr. Thus, even if the U–Pb ages for the Tf-1 and Tf-2 tuffs are affected by the magma residence time, the depositional ages of these tuffs are unlikely to be younger than the top of Chron C5Cr (∼16.7 Ma). This suggests that the N7/N8 (M4/ M5) boundary age is included within Chron C5Cr, even considering the magma residence time. We attempt to assess the validity of the estimated N7/N8 (M4/M5) boundary age in terms of sedimentation rates. The sedimentation rate curve for the composite section is shown in Fig. 11b. The average sedimentation rate for each polarity zone can be calculated using its thickness and the duration of the correlative chron. Here we focus attention on the sedimentation rates around the LOD of P. sicana. Zone N1 has a relatively low sedimentation rate of ∼24 cm/kyr and this is compatible with the fact that the dominant lithology is mudstone. Similarly, mudstone predominates in the ∼44-m interval between the N1/R2 zone boundary and site Tf-1, though approximately one-half of

the interval is not exposed; hence, the expectation that this interval has a similar sedimentation rate to that of Zone N1. By adopting the U–Pb age for Tf-1, we obtain sedimentation rates of 14–46 cm/kyr for this interval, which is in accord with that for Zone N1. Even though the sedimentation rate for the ∼2-m mudstone column between the LOD of P. sicana and site Tf-1 is as low as ∼14 cm/kyr, the ∼2-m column corresponds to a negligibly short period of ∼14 kyr. Therefore, our interpretation that the Tf-1 age (17.03 ± 0.11 Ma) approximates the N7/N8 (M4/M5) boundary age is compatible with sedimentation rates. Note that the sedimentation rate of the interval between sites Tf-1 and Tf-2 is very high, reaching > 370 cm/kyr. This is simply obtained by dividing the thickness of that interval by the difference in the two U–Pb ages (considering the age uncertainties) and is compatible with the lithology of the Katada Formation characterized by massive sandstones and turbidites particularly in its lower half. Rapidly increased sediment supply probably occurred during the depositional period of the Katada Formation, resulting in the high sedimentation rate. In summary, our data agree with ATNTS2004 that suggests the N7/ N8 (M4/M5) zone boundary age is 16.97 Ma. The early Miocene calcareous biochronology in ATNTS2004 is based on data from a large calcareous sediment core ∼5.5 km in length in total recovered from ODP Leg 154 sites on the Ceara Rise. Unfortunately, reliable magnetostratigraphic records were not obtained because of severe 13

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remagnetization during coring ( Curry et al., 1995); hence, no direct correlation of planktonic foraminifera and calcareous nannofossil datums to the magnetostratigraphy was possible. However, a wealth of biostratigraphic data were employed to determine astronomically tuned ages for several key datums (Shackleton and Crowhurst, 1997; Weedon et al., 1997; Shackleton et al., 1999). Shackleton et al. (1999) obtained an astronomically calibrated age of 17.0 Ma for the LOD of P. sicana in the Site 925 core, using the La93 astronomical solution (Laskar et al., 1993) as an insolation target. During the compilation of ATNTS2004, early to middle Miocene astronomically calibrated ages of Shackleton et al. (1999) were retuned by Lourens et al. (2004) using the La2004 solution (Laskar et al., 2004), which resulted in a slight modification of the LOD age of P. sicana to 16.97 Ma. Our new results from the Ichishi Group support the validity of the following: (1) the biostratigraphic assignment by Pearson and Chaisson (1997) of the LOD of P. sicana in the core at Site 925 of ODP Leg 154, (2) astronomical tuning by Shackleton et al. (1999) of the cyclicity in the lower Miocene portion of the same core and the resultant biostratigraphic calibration, and (3) retuning by Lourens et al. (2004) of the astronomical time scale for the Site 925 core with the La2004 solution. However, our results indicate that an age of 16.38 Ma for the N7/N8 (M4/M5) zone boundary in ATNTS2012 is problematic. In this time scale, ages for some of the early to middle Miocene planktonic foraminifera datums including the LOD age of P. sicana follow the compilation of Wade et al. (2011). In their compilation, they did not include the astronomically calibrated ages for the taxa of the genus Praeorbulina because of their rare occurrence at the Leg 154 sites on the Ceara Rise and for these taxa adopted the ages described by Berggren et al. (1995). The age for the LOD of P. sicana in Berggren et al. (1995) is essentially based on data from the Rio Grande Rise, where the LOD of P. sicana can be directly correlated with the magnetostratigraphy in the same core at DSDP Site 516 (Barker et al., 1983). According to Berggren et al. (1983), the LOD of P. sicana is at the upper part of a normal polaritydominated zone which they interpreted to be correlative to the later part of Anomaly 5C. Berggren et al. (1985, 1995) reassessed the magneto-biochronology of the LOD of P. sicana and Berggren et al. (1995) located its age in the middle of Chron C5Cn.2n on the geomagnetic polarity time scale of Cande and Kent (1995). However, caution is advised because, at Site 516, normal polarity occupies ∼75% of the stratigraphic interval, which Berggren et al. (1983) correlated with Zones N5 to N7 (∼85 m in thickness excluding portions of undetermined polarity). This appears to be inconsistent with the fact that the period from ∼20 to ∼16 Ma has approximately equal proportions of normal and reverse polarity intervals (Cande and Kent, 1995; Gradstein et al., 2004, 2012). Thus, one possibility is that secondary overprints severely affected the unknown lengths of the normal polarity portions including the normal polarity zone to which the LOD of P. sicana has been correlated. Finally, normal polarity Subzone R2-n1 can be dated at ∼17.0 Ma based on our U–Pb ages and is probably correlated with a short normal polarity episode in Chron C5Cr, although such a short polarity interval has not been listed on the geomagnetic polarity time scale. Interestingly, normal or intermediate polarity episodes in Chron C5Cr have been documented from marine sediment cores throughout the world (Channell et al., 2003; Lanci et al., 2004; Acton et al., 2008; Channell et al., 2013). For example, at ODP Site 1218 in the equatorial Pacific, Lanci et al. (2004) found such a record which they termed “Event C” and interpreted it to correspond to a very short subchron or field excursion. A normal or intermediate polarity horizon in Chron C5Cr has also been reported from on-land sedimentary exposures on the Loess Plateau in China (Hao and Guo, 2007g), in the western Ebro basin in Spain (Larrasoaña et al., 2006), and in the Engelswies section in Germany (Böhme et al., 2011).

5.2. Tectonic rotation Hayashida and Ito (1984) suggested clockwise tectonic rotation in Ichishi, based on an overall mean direction (D = 45.1°, I = 48.8°, and α95 = 11.9° for 11 sites) which they obtained from the upper Ichishi Group. They calculated the amount of rotation relative to the Asian continent to be 45 ± 18° and attributed this rotation to the clockwise rotation of entire Southwest Japan associated with the backarc opening of the Japan Sea. Their work has been quoted by later studies in assessing how and when the Japan Sea opened (Otofuji et al., 1985; Hayashida et al., 1991; Jolivet et al., 1995; Martin, 2011; Hoshi et al., 2015; Hoshi, 2018). Our new results enable discussion of the clockwise rotation in more detail. Reliable site-mean directions of polarity Zones R1–N1 and those of Zone R2 are shown in Fig. 10c and d, respectively. The clustering of the directions is better for Zones R1–N1 than for Zone R2. This is compatible with the difference in the average sedimentation rates (Fig. 11); site-mean directions of the relatively slowly deposited sediments of Zones R1–N1 have smoothed out the geomagnetic secular variation while those of the rapidly deposited sediments of Zone R2 have been influenced by the secular variation. To determine the amount of rotation and its uncertainty (R ± ΔR), we calculated the mean directions of each of the two polarity units and compared them to an expected paleomagnetic direction that was computed from 18–14 Ma paleomagnetic data for the North China Block in the Asian continent (Zhao et al., 1994). The mean direction of Zones R1–N1 is D = 39.5°, I = 49.4°, and α95 = 7.2° (7 sites; Fig. 10c) and 33.9° ± 10.9° (R ± ΔR) clockwise of the early Miocene expected direction (D = 5.6°, I = 53.0°, and α95 = 4.7°). The mean direction of Zone R2 is D = 31.5°, I = 53.0°, and α95 = 15.1° (6 sites; Fig. 10d), and R ± ΔR = 25.9° ± 21.5°. Here, two interesting points are noted. One is that the angle of rotation is smaller than that presented by Hayashida and Ito (1984). Our results suggest that in Ichishi, an ∼35° clockwise rotation with respect to the Asian continent has occurred since ∼17.2 Ma (= upper boundary age of Chron C5Dn). This is compatible with the recent argument of Hoshi (2018) that the clockwise rotation of Southwest Japan occurred between 18 and 16 Ma with the upper limit for the amount of rotation being 41.7° ± 5.4°. Another point is that the amount of rotation for Zone R2 is smaller than that for Zones R1–N1, although the difference is not significant at the 95% confidence level. This possibly indicates the progress of clockwise rotation during deposition of the upper Ichishi Group. A future paleomagnetic study of sediments from the lower part of the group may provide data to be used to help explore this possibility as well as establish the magnetostratigraphy of the entire group. 6. Conclusions Based on new U–Pb and magnetostratigraphic results from the upper Ichishi Group sediments, we demonstrated that the age of the N7/N8 (M4/M5) planktonic foraminifera zone boundary is approximated by a U–Pb zircon age of 17.03 ± 0.11 Ma. Correlation of the Ichishi Group magnetostratigraphy with the astronomically tuned geomagnetic polarity time scale provides a constraint showing that the zone boundary is within Chron C5Cr and older than the upper boundary age of that chron (∼16.7 Ma). These results are compatible with ATNTS2004 in which the zone boundary is assigned an age of 16.97 Ma. In contrast, a significantly younger age (16.38 Ma) assigned to the zone boundary in ATNTS2012 is questionable. The normal polarity Subzone R2-n1 of the Ichishi Group magnetostratigraphic zonation can also be dated at ∼17.0 Ma and is probably correlated with a short normal polarity episode in Chron C5Cr. Our new paleomagnetic data also suggest an ∼35° clockwise rotation of Southwest Japan with respect to the North China Block of the

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Asian continent since ∼17.2 Ma. It is possible that clockwise rotation associated with backarc opening of the Japan Sea was in progress during deposition of the upper Ichishi Group.

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Declaration of Competing Interest The authors declare that they have no known competing financial interests or personal relationships that could have appeared to influence the work reported in this paper. Acknowledgments We would like to thank Ryutaro Kuroki for assistance with magnetic measurements, Yoshihisa Hiroki for providing information on the stratigraphy of the Ichishi Group, Takafumi Hirata and Shuhei Sakata for technical support with U–Pb age determinations including age corrections, and Jun Hosoi for assistance with cathodoluminescence work. We would also like to thank Conall Mac Niocaill and an anonymous reviewer for helpful reviews that greatly improved the manuscript. This study was performed under a cooperative research program of the Center for Advanced Marine Core Research of Kochi University (16A003, 16B003, 17A001, 17B001) and we are grateful for the assistance of Yuhji Yamamoto during the program. This study was supported by Grants-in-Aid for Scientific Research from the Japan Society for the Promotion of Science (26400488, 17K05680). Appendix A. Supplementary data Supplementary material related to this article can be found, in the online version, at doi:https://doi.org/10.1016/j.chemgeo.2019. 119333. References Acton, G.D., Florindo, F., Jovane, L., Lum, B., Ohneiser, C., Sagnotti, L., Strada, E., Verosub, K.L., Wilson, G.S., The ANDRILL-SMS Science Team, 2008. Palaeomagnetism of the AND-2A core, ANDRILL Southern McMurdo Sound Project, Antarctica. Terra Antart. 15, 193–210. Akiba, F., et al., 1986. Middle Miocene to Quaternary diatom biostratigraphy in the Nankai trough and Japan trench, and modified lower Miocene through Quaternary diatom zones for middle-to-high latitudes of the north Pacific. In: In: Kagami, H., Karig, D.E., Coulbourn, W.T. (Eds.), Initial Reports of the Deep Sea Drilling Project 87. U.S. Government Printing Office, Washington, DC, pp. 393–480. https://doi.org/ 10.2973/dsdp.proc.87.106.1986. Anthonissen, D.E., Ogg, J.G., 2012. Cenozoic and Cretaceous biochronology of planktonic foraminifera and calcareous nannofossils. In: Gradstein, F.M., Ogg, J.G., Schmitz, M.D., Ogg, G.M. (Eds.), The Geologic Time Scale 2012. Elsevier, Amsterdam, pp. 1083–1127. https://doi.org/10.1016/B978-0-444-59425-9.15003-6. Barker, P.F., Carlson, R.L., Johnson, D.A., Čepek, P., Coulbourn, W.T., Gamboa, L.A., Hamilton, N., de Melo, U., Pujol, C., Shor, A.N., Suzyumov, A.E., Tjalsma, R.C., Walton, W.H., et al., 1983. Site 516: Rio Grande Rise. In: In: Barker, P.F., Carlson, R.L., Johnson, D.A. (Eds.), Initial Reports of the Deep Sea Drilling Project 72. US Government Printing Office, Washington, DC, pp. 155–338. https://doi.org/10. 2973/dsdp.proc.72.105.1983. Berggren, W.A., Aubry, M.P., Hamilton, N., et al., 1983. Neogene magnetobiostratigraphy of Deep Sea Drilling Project Site 516 (Rio Grande Rise, South Atlantic). In: In: Barker, P.F., Carlson, R.L., Johnson, D.A. (Eds.), Initial Reports of the Deep Sea Drilling Project 72. US Government Printing Office, Washington, DC, pp. 675–713. https:// doi.org/10.2973/dsdp.proc.72.130.1983. Berggren, W.A., Kent, D.V., van Couvering, J.A., 1985. The Neogene: Part 2 Neogene geochronology and chronostratigraphy. Geol. Soc. London, Mem. 10, 211–260. https://doi.org/10.1144/GSL.MEM.1985.010.01.18. Berggren, W.A., Kent, D.V., Swisher, C.C., Aubry, M.-P., 1995. A revised Cenozoic geochronology and chronostratigraphy. In: Berggren, W.A., Kent, D.V., Aubry, M.-P., Hardenbol, J. (Eds.), Geochronology, Time Scales, and Global Stratigraphic Correlation. SEPM (Society for Sedimentary Geology), Tulsa, OK, pp. 129–212. https://doi.org/10.2110/pec.95.04.0129. Blow, W.H., 1969. Late Middle Eocene to Recent planktonic foraminiferal biostratigraphy. In: In: Brönniman, P., Renz, H.H. (Eds.), Proceedings of the First International Conference on Planktonic Microfossils, Geneva, 1967 Vol. 1. E. J. Brill, Leiden, pp. 199–422. Böhme, M., Aziz, H.A., Prieto, J., Bachtadse, V., Schweigert, G., 2011. Bio-magnetostratigraphy and environment of the oldest Eurasian hominoid from the Early Miocene of Engelswies (Germany). J. Hum. Evol. 61, 332–339. https://doi.org/10.1016/j. jhevol.2011.04.012.

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