Co-evolutions of terrestrial temperature and seasonal precipitation from the latest Pleistocene to the mid-Holocene in Japan: carbonate clumped isotope record of a stalagmite

doi:10.21203/rs.3.rs-993823/v1

Download PDF

Research article

Co-evolutions of terrestrial temperature and seasonal precipitation from the latest Pleistocene to the mid-Holocene in Japan: carbonate clumped isotope record of a stalagmite

https://doi.org/10.21203/rs.3.rs-993823/v1

This work is licensed under a CC BY 4.0 License

Version 1

posted

You are reading this latest preprint version

Quantitative paleotemperature reconstruction is a challenging and important issue in terrestrial paleoenvironmental studies, for which carbonate clumped isotope (Δ₄₇) thermometry is a promising approach. Here we analyzed Δ₄₇ values from 68 layers of OT02 stalagmite from Ohtaki Cave in central Japan, covering two separate time intervals (2.6–8.8 and 34.8–63.5 ka) to reconstruct temperature and meteoric d¹⁸O records. The average Δ₄₇ temperature of the Holocene portion of this stalagmite was 16.3℃ ± 5.6℃, 6.6℃ ± 7.2℃ higher than the average of the latest Pleistocene portion, which was 9.7℃ ± 4.6℃. Δ₄₇ thermometry also revealed that the coldest intervals (5℃–10℃) correspond to the Heinrich cooling events H4–6, and the warmest interval (up to 19.9℃ ± 6.0℃) in middle Holocene (approximately 6–5 ka) accompanied by the Hypsithermal climate optimum. We also reconstructed past meteoric δ¹⁸O by subtracting the temperature effect from stalagmite δ¹⁸O. Average meteoric δ¹⁸O was less negative in the Holocene (8.22‰ ± 0.99‰ VSMOW) than in the latest Pleistocene (8.81‰ ± 0.84‰). Over centennial timescales, meteoric δ¹⁸O was more negative during colder periods, such as Heinrich cooling events and the cooling event around 7 ka, and less negative in warmer periods, such as Hypsithermal warming. These relations indicated co-evolution of terrestrial paleotemperature and paleoprecipitation. A temperature dependency of ¹⁸O fractionation from water to vapor is a likely reason for the negative correlation between temperature and meteoric δ¹⁸O. Additionally, it is possible that increasing lower δ¹⁸O precipitation from East Asian winter monsoon (EAWM) has decreased the averaged meteoric δ¹⁸O in colder periods. These temperature effects on meteoric δ¹⁸O occur in opposite directions to fractionation between water and the stalagmite δ¹⁸O, explaining the small amplitudes of changes observed in the δ¹⁸O of Japanese stalagmites.

Planetary Science

Inorganic Chemistry

Stalagmite paleoclimatology

Carbonate clumped isotope

Oxygen isotope

Paleotemperature

Paleoprecipitation

Latest Pleistocene

Holocene

Heinrich events

Hypsithermal

East Asian monsoons

The oxygen isotopic composition of stalagmite calcite (δ¹⁸O_C) is a key source of information for terrestrial paleoclimates. Nevertheless, paleoclimatic interpretation of stalagmite δ¹⁸O_C should consider two main controlling factors: the temperature of calcite formation and the oxygen isotopic composition of cave drip water (δ¹⁸O_W). The latter factor also involves changes in some meteorological conditions, such as monsoon intensity, moisture trajectory, and precipitation seasonality. Although previous stalagmite studies in East Asia often emphasized the so-called “amount effect” or the change in moisture sources (e.g., Wang et al., 2001) as a cause of changes in δ¹⁸O_C, limited attention has been paid to the role of temperature changes on the stalagmite δ¹⁸O_C records (e.g., Tremaine et al., 2011; Meckler et al., 2015).

The importance of temperature signals was reaffirmed for late Pleistocene to middle Holocene stalagmites from Mie and Gifu Prefectures, central Japan, located at the eastern margin of the East Asian monsoon regime (Mori et al., 2018; Fig. 1). These stalagmites exhibit notably smaller amplitudes of variation in δ¹⁸O_C than comparable Chinese cave records; hence, the glacial/interglacial contrast within these stalagmites can be accounted for by a typical warming temperature of 9℃ from the last glacial maximum (LGM) to the mid-Holocene (Mori et al., 2018). This estimated temperature change is comparable with other paleoclimate studies using stalagmites and lake and marine deposits around the Japanese Islands (e.g., Nakagawa et al., 2002; Kawahata et al., 2011; Kigoshi et al., 2014; Uemura et al., 2016). Mori et al. (2018) concluded that climate control on δ¹⁸O_W values was insignificant for Japanese caves, perhaps because of their proximity to moisture sources in the Pacific.

Changes in temperature and precipitation are principal factors influencing terrestrial climate and are intimately connected. To reconstruct their co-evolution, their signals should be divided quantitatively. Regrettably, it is challenging to determine the relative importance of the two fundamental controls on stalagmite δ¹⁸O_C, temperature and δ¹⁸O_W, using conventional isotopic compositions alone. An independent approach to estimate paleotemperature is carbonate clumped isotope (Δ₄₇) thermometry (e.g., Ghosh et al., 2006; Eiler et al., 2007), which does not require assumptions regarding δ¹⁸O_W values.

In our previous study (Kato et al., 2021), we measured Δ₄₇ values from 50 layers of a well-dated stalagmite Hiro-1 (4.5–18.1 ka) from Maboroshi Cave in Hiroshima Prefecture, southwestern Japan (Shen et al., 2010; Hori et al., 2013, 2014; Fig. 1). We revealed terrestrial temperature changes after the LGM to the mid-Holocene, except for two hiatuses around 12.8–11.4 and 10.8–7.7 ka and some layers in which Δ₄₇ and δ¹⁸O_C were significantly disturbed by kinetic isotope effect (KIE) because of dry cave conditions. The difference between the average Δ₄₇ temperatures in the pre-Holocene (18.0–16.0 ka) and mid-Holocene (7.7–4.9 ka) was 8.3℃ ± 4.2℃, which is comparable with other estimates for the difference between the LGM and the Holocene from the two Japanese caves, Gyokusendo (8.2℃, 0 versus 26 ka; Uemura et al., 2016) and Kiriana (9℃, 7 versus 24 ka; Mori et al., 2018). The Δ₄₇ values of Hiro-1 exhibit an abrupt warming around 6.3–4.9 ka, which likely corresponds to the Hypsithermal event (e.g., Wanner et al., 2008), but the record during the corresponding Heinrich cooling event (HS1) was unavailable because of strong influences of KIE related to dry cave conditions (Kato et al., 2021). We also reconstructed changes in the meteoric water δ¹⁸O_MW from the stalagmite δ¹⁸O and the Δ₄₇ temperature, finding a correlation between δ¹⁸O_MW and Δ₄₇ temperature. We show that the δ¹⁸O_MW largely shifted positively because of changes in vapor trajectory to the region due to the Holocene transgression of the Seto Inland Sea (SIS; Fig. 1b; Kato et al., 2021).

In this study, we analyzed the Δ₄₇ values of the OT02 stalagmite (2.6–8.8 and 34.8–63.5 ka; Mori et al., 2018) from Ohtaki Cave (Fig. 1) to obtain a longer climatic record in Japan (back to 63 ka) including four periods of Heinrich cooling that have not yet been elucidated in Japan. We discuss the co-evolution of terrestrial temperature and precipitation from the Late Pleistocene to Holocene, making comparisons with the results from the Hiro-1 stalagmite (Kato et al., 2021).

2.1 Ohtaki Cave and the OT02 stalagmite

Ohtaki Cave (35°44′N, 136°59′E; 400 m asl at the entrance) is located in central Gifu Prefecture, Honshu (Fig. 1). As measured at Nagataki (altitude: 430 m asl), a nearby meteorological station, the annual average rainfall over the latest 40 years was 3081 mm. The region is wet year-round: 22.6% in spring (March–May), 36.0% in summer (June–August), 24.6% in autumn (September–November), and 16.8% in winter (December–February). The 40 year mean annual average temperature is 11.5℃, ranging from −0.3℃ in January to 23.9℃ in August (Fig. 2a).

Ohtaki Cave has a total length of more than 1000 m (Yura, 2011) into the Permian limestone of the Mino Terrane (Kajita et al., 1971). It comprises three cave levels; stalagmites are best developed in the middle level, which is located 110 m underground. The cave has been modified by artificial tunnels for tourists, but the temperature in the cave remains stable at approximately 13.0℃, which is higher than the annual mean temperature observed in the nearest weather observatory at Nagataki (11.5℃). A larger amount of water seepage during summer likely elevates the cave water temperature throughout the year.

We analyzed A 140-mm-long stalagmite (OT02) collected 200 m from the cave entrance. U-Th ages and δ¹⁸O values from OT02 have been reported by Mori et al. (2018). The age model of OT02 was established on the basis of U-Th ages from 11 horizons using a Bayesian statistical model (StalAge; Scholz and Hoffman, 2011). A hiatus of 26 kyr (8.8–34.8 ka) was recognized at a discontinuous surface at the 55 mm horizon (Mori et al., 2018). The age ranges of the upper and lower parts are 2.6–8.8 and 34.8–63.5 ka, respectively; however, the age model of the upper portion of OT02 includes a relatively large uncertainty (up to ±2.5 kyr; Mori et al., 2018). The lower part of OT02 stalagmite records four Heinrich cooling events (H4–6 including H5.2).

2.2 Climatic and hydrological settings of Ohtaki and Maboroshi Caves

Two major moisture sources to Japan are the Pacific Ocean and the Japan Sea, from which the East Asian summer and winter monsoons (EASM and EAWM, respectively) bring moisture mainly by warm southerly winds and cold northerly winds, respectively. The climate of Japan is characterized by clear seasonality because the East Asian summer/winter monsoons migrate across the region. Consequently, the precipitation source switches seasonally between the Pacific Ocean in summer and the Japan Sea in winter (Fukui, 1977). On the Pacific side of the central mountains in Japan, including the sites of Ohtaki Cave (OT02) and Maboroshi Cave (Hiro-1), the EASM delivers the majority of the annual precipitation to the Pacific side of Japan. During the winter, the cold and dry EAWM winds acquire moisture from the Tsushima Warm Current, which blows into the Japan Sea. Most moisture from the Japan Sea side is then released as heavy snow and rainfall on the Japan Sea coast, whereas residual moisture generates comparatively short snow and rainfall on the Pacific side. Consequently, less negative δ¹⁸O_MW values are observed during the warm season, and more negative δ¹⁸O_MW values are observed during the cold season in several areas in Japan.

Hori et al. (2009) collected meteoric water at monthly intervals at Nagaya (Okayama Prefecture), located 20 km east of the site of Hiro-1, in 2005–2007. δ¹⁸O_MW exhibited less negative values from April to October (−7.1‰ ± 1.4‰ in amount-weighted average) and more negative values from November to March (−9.7‰ ± 1.4‰; Hori et al., 2009; Fig. 2d).

Mori et al. (2018) collected 13 samples of cave drip water at stalagmite OT02 from July 2013 to November 2015, as well as 137 separate rain events from November 2013 to September 2015 at Ohgaki City (altitude 10 m), located 60 km southwest from upstream along the major moist trajectory during the EASM season. In Ohgaki, δ¹⁸O_MW from 2013 to 2015 shows a trend similar to that in Nagaya (Mori et al., 2018); however, the δ¹⁸O_MW values at Ohgaki show clearer seasonal differences, i.e., less negative values from March to November (−5.7‰ ± 2.3‰ in amount-weighted average) and more negative values from December to February (−10.8‰ ± 2.1‰; Mori et al., 2018; Fig. 2c). The drip water δ¹⁸O value at Ohtaki (−8.3‰ on average) is clearly lower than the average δ¹⁸O_MW value at Ohgaki (−6.4‰; Mori et al., 2018). This discrepancy is explained by locations along the dominant path of moisture trajectory and differences in altitude; Ohtaki Cave is further from the Pacific and at a higher altitude (400 m asl) than Ohgaki (10 m asl). Asai et al. (2014) investigated monthly δ¹⁸O_MW values at 11 locations on Mt. Ontake (1030–2750 m asl), 45 km east–northeast of Ohtaki Cave, in 2003–2005. The δ¹⁸O_MW values in the region were highest in spring (April–May; approximately −10‰ to −8‰), lowest in winter (January–February; approximately −15‰ to −14‰), and moderate during the summer (Asai et al., 2014). Although the altitudes of these sites (Nagaya, Ohgaki, and Ontake) differ significantly from that of Ohtaki Cave, there is a common trend in that the most negative δ¹⁸O_MW values were observed in winter. Since Ohgaki is located upstream of Ohtaki along the major moisture trajectory of the region, the seasonal pattern of δ¹⁸O_MW around Ohtaki Cave is likely most similar to that in Ohgaki.

A major difference between the climatic settings of Maboroshi Cave (Yuki) and Ohtaki Cave (Nagataki) is a seasonal bias of precipitation amount. In Yuki, precipitation during the warmer 7 months (April–October) accounts for 78% of the annual total, which is more than three times the precipitation in the other 5 months (November–March; 22%). Particularly, precipitation during the coldest 3 months (December–February) is limited to only 10% of the annual total and the yearly variation is also very limited in these months (Fig. 2d). In the region of Maboroshi Cave, the Chugoku Mountains (1729 m asl) obstruct the winter rain/snowfall from the Japan Sea side. Conversely, in Nagataki (near Ohtaki Cave), precipitation during the three winter months (December–February) accounts for 16.8% of annual precipitation and yearly variations in winter precipitation are far larger than those in Yuki (Fig. 2c and d). Consequently, the amount ratio of winter and annual precipitation is also an important factor controlling the annual average δ¹⁸O_MW values at Ohtaki Cave.

Figure 3 shows the standard deviation of monthly average temperature (1σ; black bold line) and the correlation coefficients between annual and monthly average temperatures (R; dashed line) in Nagataki (Fig. 3a) and Yuki (Fig. 3b) over the last 40 years. It also shows the impact of monthly temperatures on the annual average (℃/℃; gray line), i.e., the regression slope between monthly and annual temperature. In Nagataki (near Ohtaki Cave), temperature deviations tend to be large in the colder season (December–April) and small in the warmer season (May–October) (the largest deviation occurs in February and November and the smallest in June). The correlation coefficient (R) of annual and monthly temperatures is large in February–May and September–October and smallest in July–August, the warmest months (Fig. 2). This indicates that summer temperature is the least important factor influencing the annual average temperature in Nagataki. These relations will be discussed in section 5.4.

3.1 Stable oxygen isotope compositions normalized to seawater values

Mori et al. (2018) already reported the stable oxygen isotope compositions for 644 layers (at 0.2 mm intervals) from the growth axis of the OT02 stalagmite. δ¹⁸O values in the OT02 stalagmite range from −8.2‰ to −6.3‰. We used their δ¹⁸O data in which it is assumed that the oxygen isotope values of drip water and OT02 (δ¹⁸O_OT02) were influenced by seawater δ¹⁸O_SW change because the hydrological origin of moisture circulation in Japan is predominantly seawater. To correct the δ¹⁸O_OT02 values, we accounted for the influence from δ¹⁸O_SW change (Δ¹⁸O_SW) following Mori et al. (2018) as follows:

Δ¹⁸O_SW = (δ¹⁸O_foram − 3.26) * 1/(5.03 − 3.26) (1)

where δ¹⁸O_foram is the time series from deep Pacific benthic forams reported by Lisiecki and Stern (2016), which changes from +5.03‰ during the LGM to +3.26‰. Equation 1 converts δ¹⁸O_foram to Δ¹⁸O_SW (Fig. 4b) with the assumption that ¹⁸O enrichment in seawater during the LGM was around 1‰ (Schrag et al., 1996, 2002; Bintanja and van de Wal, 2008; Shakun et al., 2015). We follow Mori et al. (2018) in defining the residual variation in δ¹⁸O_OT02 by subtracting the influence from δ¹⁸O_SW change (Δ¹⁸O_SW) as follows:

Δ¹⁸O_OT02−SW = δ¹⁸O_OT02 − Δ¹⁸O_SW (in VPDB scale) (2)

We then assume that residual variation (Δ¹⁸O_OT02−SW) reflects variations in air temperature (temperature dependency of isotopic fractionation) and hydrologic processes (e.g., fractionation from seawater to vapor, amount effect, and continental effect on meteoric δ¹⁸O).

3.2 Carbonate clumped isotope (Δ₄₇)

3.2.1 Stalagmite Δ₄₇ value

The principle of clumped isotope thermometry is based on the temperature dependency of the observed abundance anomaly of the ¹³C-¹⁸O bond in carbonate relative to the stochastic abundance calculated using bulk δ¹³C and δ¹⁸O values. The abundance anomaly of ¹³C¹⁸O¹⁶O (Δ₄₇) in CO₂ from acid digestion of carbonate minerals is negatively correlated with the temperature of carbonate precipitation (Ghosh et al., 2006; Schauble et al., 2006). The analysis of carbonates synthesized at known temperatures provides an experimental (empirical) calibration between temperature and Δ₄₇ (e.g., Ghosh et al., 2006; Dennis and Schrag, 2010; Zaarur et al., 2013; Fernandez et al., 2014; Tang et al., 2014; Defliese et al., 2015; Kluge et al., 2015; Bonifacie et al., 2017; Kelson et al., 2017; Kato et al., 2019). Nevertheless, estimating paleotemperature from stalagmite Δ₄₇ values is challenging. CO₂ degassing from cave water leads to isotopic disequilibrium in the dissolved inorganic carbon (DIC) pool and thus a smaller Δ₄₇ value of carbonate, which yields higher temperatures than predicted (e.g., Guo and Zhou., 2019). KIEs under these conditions have been reported from natural and laboratory speleothem Δ₄₇ results (Affek et al., 2008, 2014; Daëron et al., 2011; Kluge and Affek, 2012; Affek, 2013; Affek and Zaarur, 2014; Kato et al., 2021).

Our previous study (Kato et al., 2019) measured the Δ₄₇ values of two sample sets: 1) calcites synthesized at different temperatures (2.9℃–61.0℃) and 2) Japanese natural tufa collected at monthly intervals, for which the precipitation temperature was carefully estimated (5.6℃–16.0℃). Synthetic calcites were precipitated according to the “mixed-solution method” to minimize the kinetic effect of CO₂ degassing by bulk solution, and tufa samples were collected monthly at two sites by Kawai et al. (2006). Similar to stalagmite settings, CO₂ degassing induces calcite precipitation in tufa by increasing its saturation state in stream water (Ford and Pedley, 1996; Kano et al., 2003, 2019; Kawai et al., 2006). We found a clear difference in the Δ₄₇ values of synthetic calcites and tufa samples. The Δ₄₇ values of tufa were lower than those of synthetic calcites, which is in the same direction as reported in the stalagmite studies referred above (Kato et al., 2019). According to our results, the range of the Δ₄₇ offset between the tufa samples and synthetic calcite appears stable throughout the year, within a temperature range of 5.6℃–16.0℃ and a pCO₂ of 0.55–1.01 matm (Kato et al., 2019). Since these factors (temperature and pCO₂) control the rate of CO₂ degassing, our observations indicate that the degassing rate does not significantly influence the kinetic effect on the Δ₄₇ offset (Kato et al., 2019). Besides that for the synthetic calcite, Kato et al. (2019) developed an empirical temperature calibration for the natural tufa samples (Eq. 3) as follows:

Δ₄₇ = (0.0336 ± 0.0036) * 10⁶ / T² + (0.301 ± 0.048) (3)

where T is absolute temperatures in Kelvins. The evaluated temperature offset between the two calibrations was ~4℃ (Kato et al., 2019), which broadly corresponds to the temperature offset from equilibrium conditions observed in modern-day speleothems, such as Soreq Cave, Israel (Affek et al., 2014). The Δ₄₇ temperature calibration of tufa (Eq. 3; Kato et al., 2019) was applied to the Δ₄₇ values of Hiro-1 stalagmite and successfully yielded the paleotemperature (Kato et al., 2021). We therefore applied the same calibration (Eq. 3) to stalagmite Δ₄₇ to estimate paleotemperature in this study.

3.2.2 Δ₄₇ measurement and calculation

We collected subsamples for carbonate clumped isotope (Δ₄₇) measurement from 68 layers throughout OT02. To avoid strong disequilibrium effects as much as possible, subsamples were collected from clear layers excluding opaque and muddy horizons presumed to have formed under depressed and/or discontinuous calcite precipitation. The significantly large prior calcite precipitation (PCP) associated with dry conditions in Maboroshi Cave likely caused smaller Δ₆₃ of drip water DIC and smaller Δ₄₇ of Hiro-1 stalagmite (Kato et al., 2021).

Δ₄₇ measurements were performed using a dual inlet mass spectrometer (Finnigan MAT-253) at Kyushu University for four discrete periods from 2018 to 2019. Each subsample was measured once or twice, and the stability of Δ₄₇ measurements was confirmed by repeated measurements of an in-house calcite standard (Hiroshima standard; δ¹³C = −0.47‰, δ¹⁸O = −5.04‰ and Δ₄₇ = 0.671‰), which yielded 1 SD of ±0.006‰–0.009‰ (n = 5–13 for each period of measurement) corresponding to ±1.8℃–3.7℃ in the temperature range of OT02. Five milligrams of powdered carbonate were digested by phosphoric acid at 90℃ for ~5 min and the generated gas was immediately trapped by a stainless-steel tube cooled by liquid nitrogen. Moisture was separated from CO₂ gas using a liquid nitrogen and ethanol slush, after which CO₂ was introduced to a 30-m-long capillary column (Supel-Q PLOT) with a helium carrier gas. The column was cooled to −10℃ to remove organic contaminants. Purified CO₂ was analyzed by MAT-253 configured for measurements of masses 44–49 with pressure adjustment to produce an m/z 44 signal of 16 V. The integration time was 30 s. We applied the Pressure Baseline correction of He et al. (2012) with the off-peak measurement of the background intensities of masses 45–49. Each analysis comprised 5–8 acquisitions, with 4.5 off-peak (nine alternate detections; five times for the reference side and four times for the sample side), 8 on-peak, and 4 off-peak cycles per acquisition.

Further data processing to obtain Δ₄₇ values was performed using Microsoft Excel spreadsheets that first provided Δ₄₇ value versus working gas (Oztech; δ¹³C = −3.61‰ VPDB, and δ¹⁸O = 24.90‰ VSMOW). We applied the ¹⁷O correction of Brand et al. (2010) (with K = 0.01022461, λ = 0.528) to calculate Δ₄₇ values. Each Δ₄₇ value was adjusted on the absolute reference frame of Dennis et al. (2011) and expressed as Δ_47−ARF. Raw Δ_{47−[EGvsWG]} was 1) corrected using the δ⁴⁷-dependence slope (−0.000064‰ in Δ₄₇ per 1‰ of δ⁴⁷; Kato et al., 2019), 2) converted to Δ_47−RF using an empirical transfer function, and 3) converted to Δ_47−ARF by adding the acid fractionation factor of Defliese et al. (2015). We describe Δ_47−ARF simply as Δ₄₇ unless otherwise noted.

3.3 Paleometeoric δ¹⁸O reconstruction

As mentioned above, stalagmite δ¹⁸O_C value is controlled by the temperature of calcite formation and the δ¹⁸O_W value of water. Using the paleotemperature determined by carbonate clumped isotope and the temperature dependency of isotopic fractionation between calcite and water, past δ¹⁸O_W values can be calculated from the calcite δ¹⁸O_C value. We applied the temperature dependency of δ¹⁸O of Tremaine et al. (2011; Eq. 4) as follows:

1000lnα_{water−calcite} = 16.1 (10³T⁻¹) − 24.6 (4)

Eq. 4 was established for cave deposits (Tremaine et al., 2011) and is also consistent with δ¹⁸O variability in Japanese stalagmites studied by Mori et al. (2018) and Kato et al. (2021).

For the present study, we defined two types of the index of past meteoric isotopes, δ¹⁸O_MW and ε¹⁸O_MW−SW, which were calculated from δ¹⁸O_OT02 and Δ¹⁸O_OT02−SW, respectively (defined in section 3.1). δ¹⁸O_MW principally refers to cave drip water, which is assumed to be almost the average value of meteoric water δ¹⁸O. ε¹⁸O_MW−SW refers to the difference between δ¹⁸O_MW and δ¹⁸O_SW; in other words, the total fractionation in the hydrologic circulation process (from seawater to meteoric water) involving changes in climatic factors including the monsoon intensity, moisture trajectory, and the seasonality of precipitation.

4.1 Oxygen isotopes

The δ¹⁸O_OT02 (VPDB) values presented by Mori et al. (2018) vary between −8.2‰ and −6.3‰ (Fig. 4a). At 63.5–34.8 ka, the δ¹⁸O_OT02 generally becomes less negative from older to the younger layers and displays four millennial-scale events with an amplitude of 0.5‰–1‰, which likely correspond to Heinrich events. The δ¹⁸O_OT02 values of the Holocene interval (8.8–2.6 ka) mostly fall between −7‰ and −8‰; however, a slight negative shift of ~0.5‰ is observed around 6 ka (Fig. 4a). As suggested by Mori et al. (2018), the overall trend of δ¹⁸O_OT02 follows that of seawater δ¹⁸O (Fig. 4b).

Δ¹⁸O_OT02−SW shows a smaller variation between −8.3‰ and −6.7‰. Most values fall between −7‰ and −8‰ (Fig. 4c). Although no significant difference was observed between the averaged Δ¹⁸O_OT02−SW values of the latest Pleistocene (63.5–34.8 ka) and the Holocene (8.8–2.6 ka), the Δ¹⁸O_OT02−SW show a very slight increasing trend from the lower to the upper portions.

4.2 Carbonate clumped isotope (Δ₄₇)

The Δ₄₇ values of 68 layers in OT02 range from 0.670‰ to 0.747‰ (Fig. 4d), corresponding to 1.4℃–28.6℃ using the equation 3 (Table 1). The overall trend of Δ₄₇ variations broadly agrees with trends of δ¹⁸O_OT02 and Δ¹⁸O_OT02−SW. Comparatively high Δ₄₇ values (low temperatures) were observed in portions in which less negative δ¹⁸O_OT02 and Δ¹⁸O_OT02−SW values were observed.

The average Δ₄₇ temperature of the Holocene portion (16.3℃, 8.8–2.6 ka) is 6.6℃ higher than the average of the lower portion (9.7℃, 63.5–34.8 ka). The observed temperature drops corresponding to the Heinrich cooling events are 3℃–5℃.

4.3 δ¹⁸O_MW and ε¹⁸O_MW−SW

δ¹⁸O_MW and ε¹⁸O_MW−SW values vary between −10.3‰ and −6.2‰ and −11.0‰ and −6.2‰, respectively (Table 1). Variations in δ¹⁸O_MW show a flatter trend than those of δ¹⁸O_OT02 and Δ¹⁸O_OT02−SW. The averaged δ¹⁸O_MW and ε¹⁸O_MW−SW values of the Holocene (8.8–2.6 ka) portion are less negative than those of the latest Pleistocene (63.5–34.8 ka) portion, as is the crosscurrent of δ¹⁸O_OT02 and Δ¹⁸O_OT02−SW values.

5.1 Discrimination of temperature and KIEs

Before paleoclimatic reconstruction, we inspected the reliability of our results as paleoclimatic records. In principle, both calcite δ¹⁸O and Δ₄₇ values depend on temperature; nevertheless, this temperature dependency can be complicated by the KIE due to CO₂ degassing from the parent water. It has been suggested that CO₂ degassing elevates δ¹⁸O values and lowers Δ₄₇ values of carbonate minerals by incomplete isotopic exchange among water and DIC species (Guo, 2008; Daëron et al., 2011; Kluge and Affek, 2012; Guo and Zhou, 2019) and can generate isotopic offsets (differences between actual and equilibrium values). Some speleothems show a negative correlation between δ¹⁸O offset and Δ₄₇ offset due to the KIE (Daëron et al., 2011; Wainer et al., 2011; Kluge and Affek, 2012; Kluge et al., 2013; Affek et al., 2014). This type of covariation (negative correlation) between δ¹⁸O and Δ₄₇ offsets develops when KIE fractionation significantly affects δ¹⁸O and Δ₄₇. In such cases, Δ₄₇ temperature is unusable for paleoclimatic reconstruction.

Kato et al. (2021) examined the covariation between Δ¹⁸O_{Hiro−1−SW} and Δ₄₇ values of the Hiro-1 stalagmite, which were calculated using the same method as in the present study, to distinguish the influence from KIE. The positive covariation expected for equilibrium conditions was found in most parts of the Hiro-1 stalagmite, supporting the interpretation that both Δ₄₇ and δ¹⁸O_Hiro−1 changed mainly as a function of temperature. Several layers of low growth rate and high δ¹³C in Hiro-1 which imply dry conditions and high PCP (Hori et al., 2013), however exhibited features of the strong influence of KIE disequilibrium: less negative δ¹⁸O values and small Δ₄₇ values. In these layers of the Hiro-1 stalagmite, according to our estimate (Kato et al., 2021), significantly large PCP likely caused small Δ₆₃ of DIC and small stalagmite Δ₄₇, and the KIE was likely further induced by rapid supersaturation and precipitation in the evaporating water film, in which both δ¹⁸O and δ¹³C increased.

In the OT02 stalagmite, a weak but positive covariation between Δ¹⁸O_OT02−SW and Δ₄₇ values was noted throughout (Fig. 5a, r = 0.36, n = 78). This supports the interpretation that temperature change is the principal factor controlling both Δ₄₇ and Δ¹⁸O_OT02−SW and that the signals (Δ₄₇, δ¹⁸O_OT02, and Δ¹⁸O_OT02−SW) seem not to be disturbed by strong KIE disequilibrium. The growth rate of the OT02 stalagmite is higher and more stable than that of the Hiro-1 stalagmite. Additionally, subsamples for Δ₄₇ measurements were collected from clear layers of OT02 excluding opaque/muddy horizons to avoid the effect of KIE as much as possible, as mentioned in section 3.2.2.

5.2 Features in the correlation between Δ₄₇ vs Δ¹⁸O_OT02−SW

The relationship between the two temperature-controlled factors (Δ₄₇ and δ¹⁸O_C) is not simple. In the case of the Hiro-1 stalagmite, different regressions of Δ₄₇ versus Δ¹⁸O_{Hiro−1−SW} were obtained from three discrete periods, 18.0–16.0, 14.2–12.6, and 7.7–4.9 ka (Fig. 5a, Kato et al., 2021). A especially large difference was noted between regressions in the pre-Holocene and mid-Holocene. Kato et al. (2021) explained this discrepancy in regressions by differences in meteoric water δ¹⁸O_MW (ε¹⁸O_MW−SW) due to the presence or absence of the SIS (Fig. 1b), which is an important vapor source for Maboroshi Cave in Hiroshima. The SIS is currently shallow (mostly <50 m deep) and was almost exposed during the glacial sea-level low. Because of the absence of this inland sea, moisture was predominantly transported over longer distances from the marine vapor source. Less negative δ¹⁸O_MW from the SIS was established by the Holocene glacial retreat (Kato et al., 2021). By contrast, the influence of the SIS is negligible for Ohtaki Cave in Gifu prefecture (Fig. 1b). This is likely the reason for the unified regression between Δ₄₇ and Δ¹⁸O_OT02−SW for Ohtaki Cave (Fig. 5a).

Comparing to Hiro-1 values, another characteristic in the regression of OT02 Δ₄₇ vs Δ¹⁸O_OT02−SW becomes apparent (Fig. 5a), which is expressed as follows:

Δ¹⁸O_OT02−SW = 5.05 Δ₄₇ − 11.15 (Eq. 5; n = 78, R² = 0.132)

With a slope of 5.05, this is clearly shallower than that of the Hiro-1 stalagmite, for which the corresponding value is approximately 40 (37.71–44.15; Kato et al., 2021; Fig. 5a). Where both stalagmite δ¹⁸O_C and Δ₄₇ are solely controlled by depositional temperature (in other words, assuming a constant drip water δ¹⁸O_W), the relationship between δ¹⁸O_C and Δ₄₇ can be theoretically calculated. We applied the temperature dependency of Δ₄₇ (Eq. 3) and δ¹⁸O (Eq. 4; Tremaine et al., 2011). The theoretical relationship between Δ₄₇ and Δ¹⁸O_{stalagmtie−SW} is almost linear, with a slope around 70, but is slightly dependent on the Δ₄₇ value (i.e., temperature) (72.5 at Δ₄₇ = 0.69, 68.5 at Δ₄₇ = 0.74).

For Hiro-1, Kato et al. (2021) interpreted that a possible process for the shallower slopes was the temperature dependency of meteoric water δ¹⁸O_MW; thus, ¹⁸O depletion during vapor generation decreases with increasing temperature (e.g., Horita and Wesolowski, 1994; Clark and Fritz, 1997). To evaluate this interpretation, Kato et al. (2021) defined the temperature-dependent fraction from seawater to meteoric water as FT_SW−MW (in ‰) and expressed the relationships between Δ₄₇ and Δ¹⁸O_{C (Hiro−1 or OT02)−SW} as shown in Eqs. 6 and 7:

f(Δ₄₇) = AΔ₄₇ + B = Δ¹⁸O_{C (Hiro−1 or OT02)−SW} − FT_SW−MW (6)

dFT_SW−MW/dT = − aFT_SW−MW = − aT + b (7)

where T is the temperature in ℃ and FT_SW−MW is assumed to be a linear function of temperature. The temperature-dependent coefficient of FT_SW−MW is expressed by a (‰/℃). The slope A in Eq. 6 increases with a of Eq. 7 and attains a theoretical slope of Δ₄₇ versus Δ¹⁸O_{stalagmtie−SW} (approximately 70) when a takes an appropriate value.

In the Hiro-1 stalagmite, slope A in Eq. 6 attains the theoretical slope when a is at 0.077‰–0.098‰/℃ (black lines in Fig. 5b; Kato et al., 2021). The range of a_Hiro−1 is comparable with the temperature dependency of ¹⁸O fractionation between water and vapor, which is 0.094‰–0.102‰/℃ at 10℃, 0.085‰–0.090‰/℃ at 20℃, and 0.076‰–0.079‰/℃ at 30℃ (Horita and Wesolowski, 1994; Clark and Fritz, 1997). Figure 5b also shows the calculation results for the OT02 stalagmite (red line for the Holocene and blue line for latest Pleistocene). Slope A attains the theoretical slope (approximately 70) when a_OT02 is at 0.18‰/℃ for both the Holocene and latest Pleistocene, which is clearly larger than a_Hiro−1.

5.3 Reason for the larger temperature-dependent coefficient of FT_SW−MW (a) in OT02

We assume that the divergence in a values for OT02 and Hiro-1 are ascribed to regional differences in precipitation characteristics. As described in section 2.2, in both Nagataki (near Ohtaki Cave; Fig. 2c) and Yuki (near Maboroshi Cave; Fig. 2d), less negative δ¹⁸O_MW values are observed during the warm season, whereas more negative δ¹⁸O_MW values are observed in the cold season. However, in Nagataki, winter precipitation accounts for a larger proportion of the annual total, and the yearly variation of winter precipitation is far larger than in Yuki (Fig. 2a and b). Hence, the amount ratio of winter/annual precipitations is an important factor influencing the annual average δ¹⁸O_MW values at Ohtaki Cave.

The major control of winter precipitation from the Japan Sea side is the strength of the EAWM. Hirose and Fukudome (2006) found a significant correlation (R² = 0.85) between the average winter precipitation on the Japan Sea side of Honshu and Hokkaido islands in 1907–2006 and the winter monsoon index of Hanawa et al. (1988), which reflects differences in the sea-level pressures between Irkutsk and Nemuro in winter and represents the strength of the EAWM. Hirose and Fukudome (2006) also found that the Japan Sea SST in winter (November–January) exhibited a significant correlation with winter precipitation on the Japan Sea side (R² = 0.82). This relation occurs because the high SST of the Japan Sea increases evaporation from seawater, which is the supply source of winter snow/rainfall to Japan. In the Northwest Pacific and East Asia, the land–sea thermal contrast is a driving force behind both the EASM and EAWM. As described above, the proportion of winter/summer precipitation is more variable in the Ohtaki region and the proportion is largely dependent on winter precipitation. Cooling of the land area results in increased winter precipitation and decreased annual average δ¹⁸O_MW values in Ohtaki. This is likely the reason for the apparently larger temperature dependency of meteoric water δ¹⁸O_MW (a; ‰/℃) in the Ohtaki region.

5.4 Terrestrial paleotemperature

We interpret temperature to be the dominant control on the OT02 Δ₄₇ record, on the basis of the discussion presented in section 5.1. The Holocene Δ₄₇ temperature appears consistent with the present cave temperature of 13℃, although the growth of OT02 was suspended in 2.6 ka and present Δ₄₇ data are absent. The Δ₄₇ temperature record from OT02 (Table 1) is also broadly consistent with known climatic stages, such as cooling during Heinrich events, warming during the Holocene, and a warming peak around the Hypsithermal event (9 to 6–5 ka; e.g., Wanner et al., 2008; Fig. 6). In the period common to the OT02 and Hiro-1 stalagmites (7.7–4.5 ka), the two stalagmites exhibit very similar patterns in Δ₄₇ temperature (Fig. 4 and 6); a rapid warming in 7.0–6.0 ka after a temporal cooling of ~5℃ at 7.0 ka from around 15℃ at 8.0–7.5 ka and a gradual cooling to the cessation of stalagmite growth. Three intervals of distinct cooling by 3–5℃ are likely correlated to Heinrich events H6, H5.2, and H5 (Fig. 6). Although the averaged Δ₄₇ temperature of the latest Pleistocene portion of OT02 (63–35 ka) is lower than that of the Holocene portion, the Δ₄₇ temperature in warm periods between each Heinrich event reached 10℃–15℃, which is as high as the present temperature and the average of the middle Holocene.

Our Δ₄₇ temperature equation (Eq. 3) was calibrated using the Δ₄₇ values of natural tufa deposited at 5.6℃–16.0℃. The high Δ₄₇ temperatures over 20℃ around 5.4–2.9 ka therefore somewhat deviate from the temperature range covered by Eq. 3. Because of the nature of carbonate clumped isotope thermometry, the uncertainty of Δ₄₇ temperatures is greater at higher temperatures, because the Δ₄₇ value exhibits an inverse proportionality to the square of the temperature in Kelvins. Even after considering these larger errors and uncertainties, the average temperature during the period from 5.4 to 3.8 ka (considering the larger dating error for OT02 mentioned above, which likely corresponds to the warm maximum of 6.3–4.9 ka recorded in Hiro-1) is 19.9℃ ± 6.0℃ and is clearly higher than the present cave temperature of 13℃.

Seasonal temperature variation is currently very limited in the deeper part of Ohtaki Cave, and it is thus likely that the past cave temperature was also stable on seasonal timescales. We do not presume, however, that the high-average Δ₄₇ temperature during the Hypsithermal means uniform warming of all seasons. In Nagataki (near Ohtaki Cave), the temperature variation tends to be largest in the colder season (December–April) and smaller in the warmer season (May–October; Fig. 3a). The correlation coefficient (R) of annual and monthly temperatures is large in February–May and September–October and smallest in the warmest months (July–August; Fig. 3a). Together, this indicates that summer temperature is not an important control in the determination of annual average temperatures. Variability in annual average temperature is largely dependent on temperature variability in spring and autumn and is also somewhat dependent on winter temperature. This indicates that the annual average temperature in Nagataki is characterized by the length of the summer (warmer interval) and the winter (colder interval). We presume that the higher cave temperature of 19.9℃ ± 6.0℃ observed in the 5.4–3.8 ka period of OT02 (corresponding 6.3–4.9 ka of Hiro-1) was induced by longer high summers and shorter winters than at present, accompanied by the warm climate optimum of the Hypsithermal event.

5.5 Paleoprecipitation history

In section 5.4, we explained changes in cave temperature arising from interactions between the summer and winter seasons. In Japan, summer and winter climates are characterized by the EASM and EAWM winds, respectively, where rainfall from EASM winds has a less negative δ¹⁸O_MW value and rain/snowfall from EAWM winds has a more negative δ¹⁸O_MW value. Changes in summer and winter durations might therefore affect the precipitation balance from EASM and EAWM winds and the annual average of δ¹⁸O_MW value. Based on this assumption, paleometeoric water should have a more negative δ¹⁸O_MW value in colder periods and a less negative δ¹⁸O_MW value in warmer periods. However, the relation between EASM and EAWM is not simple. Previous paleoclimatic studies have indicated numerous relationships between the intensities of EASM and EAWM: an inverse correlation (Xiao et al., 1995; Yancheva et al., 2007; Liu et al., 2009), a positive correlation (Zhang and Lu, 2007), and both positive and inverse correlations depending on the period and timescale concerned (Steinke et al., 2011; Ge et al., 2017). Recently, Yan et al. (2020) presented simulation results to investigate the relationship between time series changes in EASM and EAWM, showing that their intensities are positively correlated at orbital timescale due to seasonal insolation forcing but are negatively correlated over multidecadal to millennial timescales, primarily as a result of internal variability in the Atlantic Meridional Overturning Circulation and its subsequent teleconnection to East Asia via land–sea thermal contrasts.

Focusing on climatic changes at the centennial–millennial timescale, i.e., Heinrich cooling events and Hypsithermal warming, the result of our meteoric water δ¹⁸O_MW reconstruction is consistent with the assumption of a more negative (less negative) δ¹⁸O_MW value in colder (warmer) periods (Fig. 6). The ε¹⁸O_MW−SW value (and accordingly the δ¹⁸O_MW value) was more negative in colder periods such as Heinrich cooling events and the period of cooling around 7 ka and less negative in warmer periods.

This is consistent with the assumption in section 5.4 that the longer summer durations and larger amounts of EASM rainfall caused the average δ¹⁸O_MW value in the Hypsithermal warm period. By contrast, it is presumed that EAWM brought a larger amount of snow and rainfall during long winters to Ohtaki Cave during Heinrich cooling events. Very few previous studies have reported terrestrial climatic records in Heinrich events from Japan. Nakamura et al. (2013) performed stratigraphic analyses of acoustic records of sediments in Lake Nojiri, the Japan Sea side of central Japan (Fig. 1b) where winter snowfall accounts for approximately 30% of annual precipitation. They revealed lake level fluctuation during the past 45,000 years and found eight sets of rising and falling levels, although their results contain large dating errors of ±1,000–15,000 years. They also found that peaks of lake level corresponded to cold climate stages such as Heinrich cooling events 1–4 and the Younger Dryas. They presumed that lake level rises were caused by increased snowfall from an enhanced winter monsoon (EAWM). Light blue vertical bands in Fig. 6 indicate the periods of peak lake level in Nojiri identified by Nakamura et al. (2013). Unfortunately, the record of Nakamura et al. (2013) does not include older Heinrich events H5–6, for which particularly negative ε¹⁸O_MW−SW values were reconstructed from OT02 (Fig. 6). However, the durations of high lake level periods in Nojiri seem to correspond to the periods of colder temperature and more negative ε¹⁸O_MW−SW values. Increased precipitation from EAWM was likely also bought into the Ohtaki region and caused a decrease in the annual average of ε¹⁸O_MW−SW during these cold stages.

As described here, changes in the amount ratio of summer and winter precipitations from EASM and EAWM caused the divergence in ε¹⁸O_MW−SW values in the Ohtaki region, although the temperature dependency of the fractionation between seawater and meteoric water (as discussed in section 5.3) also accounts for the relationship between temperature and ε¹⁸O_MW−SW.

As described in section 2.2, precipitation from EAWM is minor in the region of Maboroshi Cave. We suggest that a strong EAWM and weak EASM during cold periods caused the dry conditions of Maboroshi Cave, which might have resulted in the interruptions of the growth of Hiro-1 and KIE disturbance of Hiro-1 records (Fig. 6).

We analyzed carbonate clumped isotopes (Δ₄₇) of the OT02 stalagmite from Ohtaki Cave for the age intervals of 2.6–8.8 and 34.8–63.5 ka and revealed changes in terrestrial temperature and meteoric δ¹⁸O. The average Δ₄₇ temperature of the Holocene portion (16.3℃, 8.8–2.6 ka) is 6.6℃ higher than that of the latest Pleistocene portion (9.7℃, 63.5–34.8 ka). Reductions in Δ₄₇ temperature correspond to Heinrich cooling events and are approximately 3℃–5℃. We presume that higher cave temperatures of 19.9℃ ± 6.0℃ observed during the middle Holocene (5.4–3.8 ka of OT02) were induced by longer high summers and shorter winters than at present, accompanied by the Hypsithermal warm climate optimum.

We also reconstructed two indices of meteoric oxygen isotopes, δ¹⁸O_MW and ε¹⁸O_MW−SW, by subtracting the temperature effect from δ¹⁸O_OT02 and Δ¹⁸O_OT02−SW. The averaged δ¹⁸O_MW and ε¹⁸O_MW−SW values through the Holocene (8.8–2.6 ka) portion are less negative than those of the latest Pleistocene (63.5–34.8 ka) portion. Focusing on climatic changes at centennial–millennial timescale, ε¹⁸O_MW−SW values (and accordingly δ¹⁸O_MW values) were more negative in colder periods, such as Heinrich cooling events and the cooling event around 7 ka, and less negative in warmer periods such as the Hypsithermal. These relationships indicate the co-evolution of terrestrial paleotemperature and paleoprecipitation. In Japan, summer and winter climates are characterized by EASM and EAWM winds, respectively. Rainfall from EASM winds has less negative δ¹⁸O_MW values and rain/snowfall from EAWM winds has more negative δ¹⁸O_MW values. Increased precipitation brought by EASM has likely increased the average δ¹⁸O_MW in warmer periods such as the Hypsithermal, whereas increased precipitation brought by EAWM has decreased the averaged δ¹⁸O_MW in colder periods such as Heinrich cooling events.

Hence, we revealed a trend of more/less negative meteoric water δ¹⁸O_MW in warmer/colder climate stages. This trend is the opposite of that assumed in conventional paleoclimate studies using Japanese stalagmites, which suggest that meteoric δ¹⁸O_MW becomes more/less negative in warm–humid/cold–dry climates due to the “amount effect.” Our results also do not follow previous interpretations that variation in meteoric δ¹⁸O_MW values is the dominant controlling factor of stalagmite δ¹⁸O value. In our study regions, major factors determining the average δ¹⁸O_MW value are 1) the proportion between summer and winter precipitation and 2) the temperature dependency of the fractionation from seawater to meteoric water. Additionally, 3) seawater δ¹⁸O_SW change (Δ¹⁸O_SW) at orbital scales also influences δ¹⁸O_MW. These effects occur due to the hydrological setting in Japan in which moisture is brought from surrounding seas. Factor 1 is related to climatic stages on centennial scales, which is also deeply related to terrestrial temperature. Besides these three factors controlling δ¹⁸O_MW, 4) temperature-dependent fractionation at the time of stalagmite deposition (FT_{water−stalagmite}) is an important control on stalagmite δ¹⁸O_C. These temperature effects on δ¹⁸O_MW and δ¹⁸O_C are, however, in opposite directions and the negative influence of temperature on δ¹⁸O_C (FT_{water−stalagmite}) exceeds the information available for past δ¹⁸O_MW. This explains the small amplitude of δ¹⁸O change in Japanese stalagmites (Mori et al., 2018; Kato et al., 2021) and has complicated the interpretation of stalagmite δ¹⁸O_C records using only conventional methods of stalagmite climatology depending only on δ¹⁸O analysis.

Availability of data and material

Table 1 contains all data presented in this study.

Competing interests

The authors declare that they have no competing interest.

Funding

This study was supported by Grants-in-Aid from the Japan Society for the Promotion of Science [70782019, 20J00843 and 21K18393 for HK, and 16H02235 and 20H00191 for AK] and partially supported by grants from the Science Vanguard Research Program of the Ministry of Science and Technology (MOST) (110-2123-M-002-009), the National Taiwan University (109L8926 to C.-C.S.), the Higher Education Sprout Project of the Ministry of Education (110L901001 and 110L8907) Taiwan ROC for C.-C.S..

Authors' contributions

HK proposed the topic, conceived and designed the study. HK, TM, AK, CCW, and CCS carried out the experimental work. HK and AK wrote the manuscript of this study. AK and CCS supervised the study.The authors read and approved the final manuscript.

Acknowledgment

Weather data and bathymetric data in this study are from the Japan Meteorological Agency (http://www.jma.go.jp) and the NOAA National Geophysical Data Center, 2009: ETOPO1 1 Arc-Minute Global Relief Model (Accessed in July 2021). We thank Yoshihiro Kuwahara and Ryoko Senda (Kyushu University) for supporting the Δ₄₇ analysis.

Affek H. P. (2013) Clumped isotopic equilibrium and the rate of isotope exchange between CO₂ and water. Amer. Jour. Sci. 313, 309–325.

Affek H. P., Bar-Matthews M., Ayalon A., Matthews A., Eiler J. M. (2008) Glacial/interglacial temperature variations in Soreq cave speleothems as recorded by “clumped isotope” thermometry. Geochim. Cosmochim. Acta 72, 5351–5360.

Affek H. P., Matthews A., Ayalon A., Bar-Matthews M., Burstyn Y., Zaarur S., Zilberman T. (2014) Accounting for kinetic isotope effects in Soreq Cave (Israel) speleothems. Geochim. Cosmochim. Acta 143, 303–318.

Affek H. P., Zaarur S. (2014) Kinetic isotope effect in CO₂ degassing: insight from clumped and oxygen isotopes in laboratory precipitation experiments. Geochim. Cosmochim. Acta 143, 319–330.

Asai K., Tsujimura M., Fantong W. Y. (2014) Temporal variation of stable isotope ratios in precipitation on Chubu-mountainous areas: a case study of Mt. Ontake, Japan (in Japanese). Journal of Japanese Association of Hydrological Sciences 44, 67−77.

Bintanja R. and van de Wal R. S. W. (2008) North American ice-sheet dynamics and the onset of 100,000-year glacial cycles. Nature 454, 869–872.

Bonifacie M., Calmels D., Eiler J. M., Horita J., Chaduteau C., Vasconcelos C., Agrinter P., Katz A., Passey B. H., Ferry J. M., Bourrand J.-J. (2017) Calibration of the dolomite clumped isotope thermometer from 25 to 350 °C, and implications for a universal calibration for all (Ca, Mg, Fe)CO₃ carbonates. Geochim. Cosmochim. Acta 200, 255–279.

Brand W. A., Assonov S. S., Coplen T. B. (2010) Correction for the ¹⁷O interference in δ(¹³C) measurements when analyzing CO₂ with stable isotope mass spectrometry (IUPAC Technical Report). Pure Appl. Chem. 82, 1719–1733.

Clark I., Fritz P. (1997) Environmental Isotopes in Hydrology. Lewis Publishers, New York.

Daëron M., Guo W., Eiler J., Genty D., Blamart D., Boch R., Drysdale R., Maire R., Wainer K., Zanchetta G. (2011) ¹³C¹⁸O clumping in speleothems: observations from natural caves and precipitation experiments. Geochim. Cosmochim. Acta 75, 3303–3317.

Defliese W. F., Hren M. T., Lohmann K. C. (2015) Compositional and temperature effects of phosphoric acid fractionation on Δ₄₇ analysis and implications for discrepant calibrations. Chem. Geol. 396, 51–60.

Dennis K. J., Affek H. P., Passey B. H., Schrag D. P., Eiler J. M. (2011) Defining an absolute reference frame for “clumped” isotope studies of CO₂. Geochim. Cosmochim. Acta 75, 7117–7131.

Dennis K. J., Schrag D. P. (2010) Clumped isotope thermometry of carbonatites as an indicator of diagenetic alteration. Geochim. Cosmochim. Acta 74, 4110–4122.

Eiler J. M. (2007) “Clumped-isotope” geochemistry—The study of naturally-occurring, multiply-substituted isotopologues. Earth Planet. Sci. Lett. 262, 309–327.

Fernandez A., Tang J., Rosenheim B. E. (2014) Siderite “clumped” isotope thermometry: a new paleoclimate proxy for humid continental environments. Geochim. Cosmochim. Acta 126, 411–421.

Ford T. D., Pedley H. M. (1996) A review of tufa and travertine deposits of the world. Earth-Sci. Rev. 41, 117–175.

Fukui T. (Ed.) (1977) The Climate of Japan. Elsevier, Amsterdam.

Ge Q., Xue Z., Yao Z., Zang Z., Chu F. (2017) Anti‐phase relationship between the East Asian winter monsoon and summer monsoon during the Holocene? Journal of Ocean University of China 16, 175−183.

Ghosh P., Adkins J., Affek H., Balta B., Guo W. F., Schauble E. A., Schrag D., Eiler J. M. (2006) ¹³C–¹⁸O bonds in carbonate minerals: a new kind of paleothermometer. Geochim. Cosmochim. Acta 70, 1439–1456.

Guo W. (2008) Carbonate clumped isotope thermometry: application to carbonaceous chondrites and effects of kinetic isotope fractionation. Ph. D. thesis, Caltech.

Guo W., Zhou C. (2019) Patterns and controls of disequilibrium isotope effects in speleothems: Insights from an isotope-enabled diffusion-reaction model and implications for quantitative thermometry. Geochim. Cosmochim Acta 267, 196–226.

Hanawa K., Watanabe T., Iwasaka N., Suga T., Toba Y. (1988) Surface thermal conditions in the Western North Pacific during the ENSO events. Journal of the Meteorological Society of Japan 66, 445–456.

He B., Olack G. A., Colman A. S. (2012) Pressure baseline correction and high-precision CO₂ clumped-isotope (Δ₄₇) measurements in bellows and micro-volume modes. Rapid Commun. Mass Spectrom. 26, 2837–2853.

Hirose N., Fukudome K. (2006) Monitoring the Tsushima Warm Current improves seasonal prediction of the regional snowfall. SOLA 2, 61−63.

Hori M., Ishikawa T., Nagaishi K., Lin K., Wang B.-S., You C.-F., Shen C.-C., Kano A. (2013) Prior calcite precipitation and source mixing process influence Sr/Ca, Ba/Ca and ⁸⁷Sr/⁸⁶Sr of a stalagmite developed in southwestern Japan during 18.0 4.5 ka. Chem. Geol. 347, 190−198.

Hori M., Ishikawa T., Nagaishi K., You C.-F., Huang K.-F., Shen C.-C., Kano A. (2014) Rare earth elements in a stalagmite from southwestern Japan: A potential proxy for chemical weathering. Geochem. Jour. 48, 73−84.

Hori M., Kawai T., Matsuoka J., Kano A. (2009) Intra-annual perturbations of stable isotopes in tufas: Effects of hydrological processes. Geochim. Cosmochim. Acta 73, 1684–1695.

Horita J., Wesolowski D. J. (1994) Liquid-vapor fractionation of oxygen and hydrogen isotopes of water from the freezing to the critical temperature. Geochim. Cosmochim. Acta 58, 3425–3437.

Kajita S., Aoyama S., Kitamura T., Hibino M. (1971) The catalogue of limestone caves, Gifu Prefecture, Central Japan 2 (in Japanese). Science report of the Faculty of Education, Gifu University. Natural science 4, 379−386.

Kano A., Matsuoka J., Kojo T., Fujii H. (2003) Origin of annual laminations in tufa deposits, southwest Japan. Palaeogeogr. Palaeoclimatol. Palaeoecol. 191, 243–262.

Kano A., Okumura T., Takashima C. Shiraishi F. (2019) Geobiochenimal Properties of Travertine with Focus on Japanese Sites. Springer, Singapore.

Kawai T., Kano A., Matsuoka J., Ihara T. (2006) Seasonal variation in water chemistry and depositional processes in a tufa-bearing stream in SW-Japan, based on 5 years of monthly observations. Chem. Geol. 232, 33–53.

Kato H., Amekawa S., Kano A., Mori T., Kuwahara Y., Quade J. (2019) Seasonal temperature changes obtained from carbonate clumped isotopes of annually laminated tufas from Japan: Discrepancy between natural and synthetic calcites. Geochim. Cosmochim. Acta 244, 548–564.

Kato H., Amekawa S., Hori M., Shen C.-C., Kuwahara Y., Senda R., Kano A. (2021) Influences of temperature and the meteoric water δ¹⁸O value on a stalagmite record in the last deglacial to middle Holocene period from southwestern Japan. Quat. Sci. Rev. 253, 106746.

Kelson J., Huntington K. W., Schauer A., Saenger C., Lechler A. R. (2017) Toward a universal carbonate clumped isotope calibration: Diverse synthesis and preparatory methods suggest a single temperature relationship. Geochim. Cosmochim. Acta 197, 104–131.

Kawahata H., Ohshima H., Kuroyanagi A. (2011) Terrestrial–Ocean environmental change in the northwestern Pacific from the glacial times to Holocene. J. Asian Earth Sci. 40, 1189–1202.

Kigoshi T., Kumon F., Hayashi R., Kuriyama M., Yamada K., Takemura K. (2014) Climate changes for the past 52 ka clarified by total organic carbon concentrations and pollen composition in Lake Biwa, Japan. Quat. Int. 333, 2–12.

Kluge T., Affek H. P. (2012) Quantifying kinetic fractionation in Bunker Cave speleothems using Δ₄₇. Quat. Sci. Rev. 49, 82–94.

Kluge T., Affek H. P., Marx T., Aeschbach-Hertig W., Riechelmann D. F. C., Scholz D., Riechelmann S., Immenhauser A., Richter D. K., Fohlmeister J., Wackerbarth A., Mangini A., Spötl C. (2013) Reconstruction of drip-water δ¹⁸O based on calcite oxygen and clumped isotopes of speleothems from Bunker Cave (Germany). Climate of the Past 9, 377–391.

Kluge T., John C. M., Jourdan A.-L., Davis S., Crawshaw J. (2015) Laboratory calibration of the calcium carbonate clumped isotope thermometer in the 25–250 ºC temperature range. Geochim. Cosmochim. Acta 157, 213–227.

Liu X., Dong H., Yang X., Herzschuh U., Zhang E., Stuut J.-B.W., Wang Y. (2009) Late Holocene forcing of the Asian winter and summer monsoon as evidenced by proxy records from the northern Qinghai−Tibetan Plateau. Earth and Planetary Science Letters 280, 276−284.

Meckler A. N., Affolter S., Dublyansky Y. V., Krüger Y., Vogel N., Bernasconi S. M., Frenz M., Kipfer R., Leuenberger M., Spötl C., Carolin S., Cobb K. M., Moerman J., Adkins J. F., Fleitman D. (2015) Glacial－interglacial temperature change in the tropical West Pacific: A comparison of stalagmite-based paleo-thermometers. Quat. Sci. Rev. 127, 90–116.

Mori T., Kashiwagi K., Amekawa S., Kato H., Okumura T., Takashima C., Wu C.-C., Shen C.-C., Quade J., Kano A. (2018) Temperature and seawater isotopic controls on two stalagmite records since 83 ka from maritime Japan. Quat. Sci. Rev. 192, 47–58.

Nakagawa T., Tarasov P. E. Nishida K., Gotanda K., Yasuda Y. (2002) Quantitative pollen-based climate reconstruction in central Japan: application to surface and Late Quaternary spectra. Quat. Sci. Rev. 21, 2099–2113.

Nakamura Y., Inouchi Y., Inoue T., Kondo Y., Kumon F., Nagahashi Y. (2013) Lake-level changes and their factors during the last 45,000 years in Lake Nojiri, Central Japan (in Japanese). The Quaternary Research (Daiyonki-Kenkyu) 52, 203−212.

Schauble E. A., Ghosh P., Eiler J. M. (2006) Preferential formation of ¹³C–¹⁸O bonds in carbonate minerals, estimated using first-principles lattice dynamics. Geochim. Cosmochim. Acta 70, 2510–2529.

Scholz D., Hoffmann D. L. (2011) StalAge – An algorithm designed for construction of speleothem age models. Quat. Geochronol. 6, 369–382.

Schrag D. P., Adkins J. F., McIntyre K., Alexander J. L., Hodell D. A., Charles C. D., McManus J. F. (2002) The oxygen isotopic composition of seawater during the Last Glacial Maximum. Quat. Sci. Rev. 21, 331–342.

Schrag D. P., Hampt G., Murray D. W. (1996) Pore Fluid Constraints on the Temperature and Oxygen Isotopic Composition of the Glacial Ocean. Science 272, 1930–1932

Shakun J. D., Lea D. W., Lisiecki L. E., Raymo M. E. (2015) An 800-kyr record of global surface ocean δ¹⁸O and implications for ice volume-temperature coupling. Earth Planet. Sci. Lett. 426 58–68.

Shen C.-C., Kano A., Hori M., Lin K., Chiu T.-C., Burr G. S. (2010) East Asian monsoon evolution and reconciliation of climate records from Japan and Greenland during the last deglaciation. Quat. Sci. Rev. 29, 3327–3335.

Steinke S., Glatz C., Mohtadi M., Groeneveld J., Li Q., Jian Z. (2011) Past dynamics of the East Asian monsoon: No inverse behaviour between the summer and winter monsoon during the Holocene. Global and Planetary Change 78, 170−177.

Tang J., Dietzel M., Fernandez A., Tripati A. K., Rosenheim B. E. (2014) Evaluation of kinetic effects on clumped isotope fractionation (Δ₄₇) during inorganic calcite precipitation. Geochim. Cosmochim. Acta 134, 120–136.

Tremaine D. M., Froelich P. N., Wang Y. (2011) Speleothem calcite farmed in situ: modern calibration of δ¹⁸O and δ¹³C paleoclimate proxies in a continuously-monitored natural cave system. Geochim. Cosmochim. Acta 75, 4929–4950.

Uemura R., Nakamoto M., Asami R., Mishima S., Gibo M., Msaka K., Chen J.-P., Wu C.-C., Chang Y.-W., Shen C.-C. (2016) Precise oxygen and hydrogen isotope determination in nanoliter quantities of speleothem inclusion water by cavity ring-down spectroscopic techniques. Geochim. Cosmochim. Acta 172, 159–176.

Wainer K., Genty D., Blamart D., Daëron M., Bar-Matthews M., Vonhof H., Dublyansky Y., Pons-Branchu E., Thomas L., van Calsteren P., Quinif Y., Caillon N. (2011). Speleothem record of the last 180 ka in Villars cave (SW France): Investigation of a large δ¹⁸O shift between MIS6 and MIS5. Quat. Sci. Rev. 30, 130–146.

Wang Y. J., Cheng H., Edwards R. L., An Z. S., Wu J. Y., Shen C.-C. and Dorale J. A. (2001) A High-Resolution Absolute-Dated Late Pleistocene Monsoon Record from Hulu Cave, China. Science 451, 1090–1093.

Wanner H., Beer J., Bütikofer J., Crowley T. J., Cubasch U., Flückiger J., Goosse H., Grosjean M., Joos F., Kaplan J. O., Küttel M., Müller S. A., Prentice I. C., Solomina O., Stocker T. F., Tarasov P., Wagner M., Widmann M. (2008) Mid‐ to Late Holocene climate change: An overview. Quat. Sci. Rev. 27, 1791–1828.

Yura K. (2011) Visit and consideration of Ohtaki Limestone Cave (tourism cave), Gujo City, Gifu Prefecture (in Japanese). Cave Environmental NET Society 2, 3−10.

Xiao J., Porter S. C., An Z., Kumai H., Yoshikawa S. (1995) Grain size of quartz as an indicator of winter monsoon strength on the Loess Plateau of central China during the last 130,000 yr. Quaternary Research 43, 22−29.

Yan M., Liu Z. Y., Ning L., Liu J. (2020) Holocene EASM‐EAWM relationship across different timescales in CCSM3. Geophys. Res. Lett. 47, e2020GL088451.

Yancheva G., Nowaczyk N. R., Mingram J., Dulski P., Schettler G., Negendank J. F. W., Liu J., Sigman D. M., Peterson L. C., Haug G. H. (2007) Influence of the intertropical convergence zone on the East Asian monsoon. Nature 445, 74−77.

Zaarur S., Affek H. P., Brandon M. T. (2013) A revised calibration of the clumped isotope thermometer. Earth Planet. Sci. Lett. 383, 47–57.

Zhang, D., Lu, L. (2007). Anti‐correlation of summer/winter monsoons? Nature, 450, E7−E8.

Due to technical limitations, table 1 is only available as a download in the Supplemental Files section.

OT02graphicalabstract.pdf
Table1.xlsx
Table 1. δ18OOT02 and Δ47 values of OT02, Δ18OOT02−SW, Δ47 temperatures calculated from a tufa calibration (Kato et al., 2019), and reconstructed meteoric water δ18OMW and ε18OMW−SW values.

Download PDF

Version 1

posted

You are reading this latest preprint version

Co-evolutions of terrestrial temperature and seasonal precipitation from the latest Pleistocene to the mid-Holocene in Japan: carbonate clumped isotope record of a stalagmite

Status:

Version 1

Abstract

Figures

1. Introduction

2. Study Area And Material

2.1 Ohtaki Cave and the OT02 stalagmite

2.2 Climatic and hydrological settings of Ohtaki and Maboroshi Caves

3. Method

3.1 Stable oxygen isotope compositions normalized to seawater values

3.2 Carbonate clumped isotope (Δ₄₇)

3.2.1 Stalagmite Δ₄₇ value

3.2.2 Δ₄₇ measurement and calculation

3.3 Paleometeoric δ¹⁸O reconstruction

4. Results

4.1 Oxygen isotopes

4.2 Carbonate clumped isotope (Δ₄₇)

4.3 δ¹⁸O_MW and ε¹⁸O_MW−SW

5 Discussion

5.1 Discrimination of temperature and KIEs

5.2 Features in the correlation between Δ₄₇ vs Δ¹⁸O_OT02−SW

5.3 Reason for the larger temperature-dependent coefficient of FT_SW−MW (a) in OT02

5.4 Terrestrial paleotemperature

5.5 Paleoprecipitation history

6. Summary

Declarations

References

Tables

Supplementary Files

Status:

Version 1

Co-evolutions of terrestrial temperature and seasonal precipitation from the latest Pleistocene to the mid-Holocene in Japan: carbonate clumped isotope record of a stalagmite

Status:

Version 1

Abstract

Figures

1. Introduction

2. Study Area And Material

2.1 Ohtaki Cave and the OT02 stalagmite

2.2 Climatic and hydrological settings of Ohtaki and Maboroshi Caves

3. Method

3.1 Stable oxygen isotope compositions normalized to seawater values

3.2 Carbonate clumped isotope (Δ47)

3.2.1 Stalagmite Δ47 value

3.2.2 Δ47 measurement and calculation

3.3 Paleometeoric δ18O reconstruction

4. Results

4.1 Oxygen isotopes

4.2 Carbonate clumped isotope (Δ47)

4.3 δ18OMW and ε18OMW−SW

5 Discussion

5.1 Discrimination of temperature and KIEs

5.2 Features in the correlation between Δ47 vs Δ18OOT02−SW

5.3 Reason for the larger temperature-dependent coefficient of FTSW−MW (a) in OT02

5.4 Terrestrial paleotemperature

5.5 Paleoprecipitation history

6. Summary

Declarations

References

Tables

Supplementary Files

Status:

Version 1

3.2 Carbonate clumped isotope (Δ₄₇)

3.2.1 Stalagmite Δ₄₇ value

3.2.2 Δ₄₇ measurement and calculation

3.3 Paleometeoric δ¹⁸O reconstruction

4.2 Carbonate clumped isotope (Δ₄₇)

4.3 δ¹⁸O_MW and ε¹⁸O_MW−SW

5.2 Features in the correlation between Δ₄₇ vs Δ¹⁸O_OT02−SW

5.3 Reason for the larger temperature-dependent coefficient of FT_SW−MW (a) in OT02