INTRODUCTION
The transition of the Earth's climate from full glacial to full interglacial conditions, known as the deglaciation process, is characterized by global warming and sea-level rising (ice sheet melting) (e.g., Denton et al., Reference Denton, Anderson, Toggweiler, Edwards, Schaefer and Putnam2010; IPCC, Reference Stocker, Qin, Plattner, Tignor, Allen, Boschung, Nauels, Xia, Bex and Midgley2013). During the penultimate deglaciation, massive iceberg discharges and the accompanying freshwater influx into the North Atlantic triggered a slowing or shutdown of the Atlantic Meridional Overturning Circulation (AMOC) (e.g., Grant et al., Reference Grant, Rohling, Ramsey, Cheng, Edwards, Florindo and Heslop2014; Tzedakis et al., Reference Tzedakis, Drysdale, Margari, Skinner, Menviel, Rhodes and Taschetto2018; Max et al., Reference Max, Nürnberg, Chiessi, Lenz and Mulitza2022). The reduced heat transport in the surface-ocean from the tropic to the North Atlantic resulted in the millennial-scale cooling stadial, namely Heinrich Stadial 11 (HS11), and the associated global effect.
Generally, deglaciations were often punctuated by millennial-scale events (e.g., Barker et al., Reference Barker, Diz, Vautravers, Pike, Knorr, Hall and Broecker2009; Cheng et al., Reference Cheng, Edwards, Broecker, Denton, Kong, Wang, Zhang and Wang2009, 2016) and/or significant variations in the AMOC (Marcott et al., Reference Marcott, Clark, Padman, Klinkhammer, Springer, Liu and Otto-Bliesner2011; Rasmussen et al., Reference Rasmussen, Bigler, Blockley, Blunier, Buchardt, Clausen and Cvijanovic2014; Cheng et al., Reference Cheng, Edwards, Sinha, Spötl, Yi, Chen and Kelly2016, Reference Cheng, Xu, Dong, Zhao, Li, Baker and Sinha2021; Brook and Buizert, Reference Brook and Buizert2018; Ng et al., Reference Ng, Robinson, McManus, Mohamed, Jacobel, Ivanovic, Gregoire and Chen2018; Toucanne et al., Reference Toucanne, Soulet, Vázquez Riveiros, Boswell, Dennielou, Waelbroeck and Bayon2021). The large-scale climatic fluctuations that occurred during deglaciations are virtually an extension of the recurrence of millennial climatic variations during the glacials (e.g., the Dansgaard–Oeschger [D-O] oscillations). Previous research has revealed that the D-O oscillations are closely associated with changes in the strength of the AMOC and corresponding changes in atmospheric circulation and sea ice coverage (e.g., Rasmussen and Thomsen, Reference Rasmussen and Thomsen2004; Dokken et al., Reference Dokken, Nisancioglu, Li, Battisti and Kissel2013; Menviel et al., Reference Menviel, Skinner, Tarasov and Tzedakis2020).
One of the most studied D-O-like events in the last deglaciation is the Bølling–Allerød (B/A) warm period that occurred between ca. 14.7 and 12.9 ka (thousand years before present, where present = AD 1950) in the Northern Hemisphere (NH). This event began at the end of the cold period known as the HS1 and ended at the onset of the Younger Dryas (YD). Intriguingly, a similarly abrupt warm event occurred at ca. 134.5 ka (the 134-ka event) during the penultimate deglaciation, as inferred in the Asian summer monsoon (ASM) speleothem archives and Antarctic ice core records (e.g., Cheng et al., Reference Cheng, Edwards, Wang, Kong, Ming, Kelly, Wang, Gallup and Liu2006; Wang et al., Reference Wang, Cheng, Edwards, Kong, Shao, Chen, Wu, Jiang, Wang and An2008).
The 134-ka event is a strong ASM interval centered at 134.5 ka in the penultimate deglaciation, characterized by a significantly negative oxygen isotope (δ18O) excursion in the speleothem records from the ASM domain (e.g., Cheng et al., Reference Cheng, Edwards, Wang, Kong, Ming, Kelly, Wang, Gallup and Liu2006; Kelly et al., Reference Kelly, Edwards, Cheng, Yuan, Cai, Zhang, Lin and An2006; Wang et al., Reference Wang, Wang, Shao, Liang, Zhang and Kong2018; Duan et al., Reference Duan, Cheng, Tan, Li and Edwards2019). Atmospheric methane (CH4) and nitrous oxide (N2O) also show significant peaks at approximately 134 ka, as recorded in the European Project for Ice Coring in Antarctica (EPICA) Dome C (EDC) ice core records (Schmidely et al., Reference Schmidely, Nehrbass-Ahles, Schmitt, Han, Silva, Shin, Joos, Chappellaz, Fischer and Stocker2021). The rise in CH4 concomitant with the event also can be seen in the Vostok ice core record (Delmotte et al., Reference Delmotte, Chappellaz, Brook, Yiou, Barnola, Goujon, Raynaud and Lipenkov2004), and presumably resulted from the expanded wetland due to enhanced tropical rainfall and temperature in NH (Bock et al., Reference Bock, Schmitt, Beck, Seth, Chappellaz and Fischer2017). Additionally, contemporaneous hydroclimatic changes are widely reported in different proxy records from the North Atlantic region, including a short break in ice-rafted debris (IRD) deposition and a concomitant decrease in N. pachyderma abundance in marine sediments in the eastern subpolar North Atlantic (Martrat et al., Reference Martrat, Jimenez-Amat, Zahn and Grimalt2014; Mokeddem et al., Reference Mokeddem, McManus and Oppo2014; Tzedakis et al., Reference Tzedakis, Drysdale, Margari, Skinner, Menviel, Rhodes and Taschetto2018). On the other hand, the sea surface temperature (SST) fell at this time in the South Atlantic (the South Atlantic inversion) (Scussolini et al., Reference Scussolini, Marino, Brummer and Peeters2015). Collectively, the co-occurrence of significant changes surrounding ca. 134 ka in different proxies in both hemispheres have delineated a large-scale sub-millennial climate event within HS11—the 134-ka event.
Since ~70% of the glacial–interglacial sea-level rise (i.e., melt water pulse 2B [MWP-2B]) occurred during a phase of the weak AMOC (HS11) (Grant et al., Reference Grant, Rohling, Bar-Matthews, Ayalon, Medina-Elizalde, Ramsey, Satow and Roberts2012, 2014), the penultimate deglaciation fundamentally differed from the previous deglaciation (Marino et al., Reference Marino, Rohling, Rodriguez-Sanz, Grant, Heslop, Roberts, Stanford and Yu2015). During the previous deglaciation, ~75% of the sea-level rise postdated the major deglacial cooling phase in the North Atlantic (HS1), and the largest meltwater pulse (MWP-1A, equivalent to 15–20% of the deglacial sea-level rise) peaked during the B/A warm period (Clark et al., Reference Clark, Mitrovica, Milne and Tamisiea2002; Carlson and Clark, Reference Carlson and Clark2012). These observations fueled arguments that Antarctic ice sheets had contributed substantially to the MWP-1A (Clark et al., Reference Clark, Mitrovica, Milne and Tamisiea2002). In contrast, the MWP-2B during the penultimate deglaciation was more than three times larger than the MWP-1A and was tied directly to circum-North Atlantic ice sheet reduction and attendant North Atlantic cooling (Marino et al., Reference Marino, Rohling, Rodriguez-Sanz, Grant, Heslop, Roberts, Stanford and Yu2015).
This fundamentally different relationship between the North Atlantic climate and sea-level change during the two deglaciations indicates that during the penultimate deglaciation, NH ice sheets collapsed earlier in the deglaciation process, possibly in response to a combination of overall stronger (and more rapidly rising) boreal summer insolation and higher atmospheric CO2 (Marino et al., Reference Marino, Rohling, Rodriguez-Sanz, Grant, Heslop, Roberts, Stanford and Yu2015). Currently, it is not clear whether the 134-ka event exhibits any similarity in terms of structure and mechanisms to the B/A event in the last deglaciation. Whether the different meltwater pulses during the two deglaciations would have caused the apparent difference between the 134-ka event and the B/A event remains unclear as well. Likewise, the bottom of the Greenland ice records reaches only to the Eemian interglacial (Marine Isotopic Stages [MIS] 5e) (Rasmussen et al., Reference Rasmussen, Bigler, Blockley, Blunier, Buchardt, Clausen and Cvijanovic2014). As such, the Greenland ice records are not useful for understanding millennial-scale climate changes beyond MIS 5e (Rasmussen et al., Reference Rasmussen, Bigler, Blockley, Blunier, Buchardt, Clausen and Cvijanovic2014), including the 134-ka event, although a detailed characterization of the 134-ka event is critical to decipher the difference and link between the last two deglaciations, as well as the underlying causal mechanism(s).
In this study, we report a newly generated speleothem δ18O record from Zhangjia Cave, southwestern China, ranging from 132.8 ± 0.3 ka to 138.8 ± 0.4 ka. On the basis of precise 230Th dating and annual-lamina counting, we characterized the detailed structure of centennial climate events in the penultimate deglaciation, particularly the 134-ka event. The comparison/correlation of our Zhangjia record with records from a number of previous speleothems, ice-core records, and marine sediment records provides new constrains on the timing of the 134-ka event globally with unprecedentedly high-precision.
CAVE LOCATION AND STALAGMITE SAMPLE
Zhangjia Cave (32°35′N, 105°58′E, 680 m above sea level) is located at northeastern Guangyuan City, on the northern edge of the Sichuan Basin in the south of the Daba Mountains (Micang Mountains) (Fig. S1). The cave, formed in the limestones of the Lower Triassic Feixianguan Formation, has an entrance of 3 × 4 m2, and its total length exceeds 1 km (Fig. 1).
Location map and cave topographic map of Zhangjia Cave. The blue star depicts the studied cave (Zhangjia Cave). Red dots indicate cave records from the Asian monsoon domain. The cave topographic map of Zhangjia Cave is also shown.
The columnar-shaped calcite stalagmite sample ZJD2020-1, which is 25.3 cm in height and ~6 cm in width (Fig. 2a), was collected in the first cave chamber, ~800 m from the cave entrance. Sample ZJD2020-1 was broken and scattered on the ground, which most likely was a result of an earthquake. According to instrumental data from the Guangyuan meteorological station (~54 km southwest of the cave), the mean annual air temperature is 16.1°C, and the mean annual precipitation is ~950 mm, ~74% of which occurs during summer (June to September) (1951–2019) (Fig. S1).
Age model of the studied speleothem record. (a) Stalagmite sample ZJD2020-1 (b) Age model of ZJD2020-1 from Zhangjia Cave. (c) Fluorescent lamina of ZJD2020-1 at 161~163 mm. Error bars show U-Th dates and errors (2σ). Age models were calculated combining the StalAge algorithm (Scholz and Hoffmann, Reference Scholz and Hoffmann2011) and lamina counting.
METHODS
230Th dating
Subsamples of 20–60 mg were hand-drilled along growth layers for 230Th dating. Eighteen powder subsamples were drilled from the polished slab of sample ZJD2020-1 using a carbide dental burr of 0.5 mm diameter (Table 1). The subsamples were dated using multi-collector inductively coupled plasma mass spectrometry (Neptune Plus, Thermo Scientific) in the Isotope Laboratory of Xi'an Jiaotong University. All errors are reported as two standard deviations (2σ). Standard chemistry procedures were used to separate U and Th (Edwards et al., Reference Edwards, Chen, Ku and Wasserburg1987). A triple-spike (229Th–233U–236U) isotope dilution method was employed to correct instrumental fractionation and determine U and Th isotopic ratios and concentrations. The instrumentation, standardization, and half-lives are reported in Cheng et al. (Reference Cheng, Edwards, Hoff, Gallup, Richards and Asmerom2000, Reference Cheng, Lawrence Edwards, Shen, Polyak, Asmerom, Woodhead and Hellstrom2013). All U/Th isotopes were measured in peak-jumping mode on a MasCom multiplier placed behind the retarding potential quadrupole. We followed procedures similar to those described in (Cheng et al., Reference Cheng, Edwards, Hoff, Gallup, Richards and Asmerom2000) to characterize the multiplier. Uncertainties in U/Th isotopic data were calculated offline, including corrections for blanks, multiplier dark noise, abundance sensitivity, and spike composition. 230Th ages were corrected using an initial 230Th/232Th atomic ratio of (4.4 ± 2.2) ×10−6, the values for a material at secular equilibrium with respect to the bulk Earth 232Th/238U value of 3.8. The U and Th decay constants are reported in Jaffey et al. (Reference Jaffey, Flynn, Glendenin, Bentley and Essling1971) and Cheng et al. (Reference Cheng, Lawrence Edwards, Shen, Polyak, Asmerom, Woodhead and Hellstrom2013).
230Th dating results; 2σ error.
U decay constants: λ238 = 1.55125 × 10−10 (Jaffey et al., Reference Jaffey, Flynn, Glendenin, Bentley and Essling1971) and λ234 = 2.82206 × 10−6 (Cheng et al., Reference Cheng, Lawrence Edwards, Shen, Polyak, Asmerom, Woodhead and Hellstrom2013). Th decay constant: λ230 = 9.1705 × 10−6 (Cheng et al., Reference Cheng, Lawrence Edwards, Shen, Polyak, Asmerom, Woodhead and Hellstrom2013). Corrected 230Th ages assume the initial 230Th/232Th atomic ratio of 4.4 ± 2.2 × 10−6. Those are the values for a material at secular equilibrium, with the bulk earth 232Th/238U value of 3.8. The errors are arbitrarily assumed to be 50%.
*d234U = ([234U/238U] activity − 1) × 1000.
**d234Uinitial was calculated based on 230Th age (T) (i.e., δ234Uinitial = δ234Umeasured × eλ234 × T).
Stable isotopes
For each oxygen isotope measurement, ~100 μg of powder samples were drilled from the central axis of the stalagmite. Subsamples (253 in total) were micro-milled at 1-mm increments perpendicular to the growth axes and analyzed using a Thermo Scientific MAT253 plus mass spectrometer coupled with an online carbonate preparation device (Kiel IV) at the Isotope Laboratory, Xi'an Jiaotong University, China. The results show an analytical error (1σ) for δ18O of 0.06‰. International standards were added to the analysis every 10–20 samples to check reproducibility. The results are reported relative to the Vienna Pee Dee Belemnite (VPDB) standard in δ-notation (‰).
Fluorescent lamina study
The ZJD2020-1 sample was polished and scanned with a confocal laser fluorescence microscope (CLFM) (Model: Nikon, A1) at the State Key Laboratory of Mechanical Manufacturing Systems Engineering, Xi'an Jiaotong University. The image scanning was operated with a 40 mW, 488 nm laser line. The fluorescence images were obtained using an emission filter, which allows light with wavelengths of 505–550 nm (visible, green) to pass (Zhao and Cheng, Reference Zhao and Cheng2017). The laminae counting was implemented in a section (82–170 mm) where laminae were quite clear. The results show an alternation between light and dark laminae with each light-dark pair corresponding to an annual growth cycle (Fig. 2c). Counts within the section (82–170 mm) were done for a total of five times. The laminae can be continuously identified, and the numbers of paired laminae between two consecutive 230Th dates match the 230Th age difference within uncertainty. This agreement supports our interpretation that the paired laminae indicate an annual cycle, which thus allows us to construct a precise relative age model.
RESULTS
Chronology
The 18 high-precision 230Th dates (Table 1) obtained from the sample ZJD2020-1 were used to build the age model. The results show that sample ZJD2020-1 grew continuously from 132.8 ± 0.3 ka to 138.8 ± 0.4 ka (Table 1), covering the late portion of the penultimate glacial period. All dates are in stratigraphic order within errors (2σ), ranging from 327 yr to ca. 822 yr. The samples are clean, with high 230Th/232Th atomic ratios (0.04–1.2). The 238U content is high (3–13 ppm), and the δ234UInitial values are approximately 127.
We used the least square method to anchor the results from annual band counting to the 230Th dates (Domínguez-Villar et al., Reference Domínguez-Villar, Baker, Fairchild and Edwards2012), and establish a high-precision age model for the 82–170 mm section (Fig. 2c). In general, chronological models based on annual laminae counts are more accurate. In principle, there are two main sources of uncertainty associated with a floating chronology obtained by annual lamina counting: (A) lamina counting (LC) uncertainty (~113 [2σ] resulted from counting five times for ZJD2020-1), and (b) the uncertainty from anchoring (A) the LC chronologies to 230Th dates (ca. 70 yr for ZJD2020-1). The combined error (ɛ) of annual-lamina counts in the 82–170 mm section is 132.5 yr (2σ), calculated by the following equation:
The age models for the 0–82 mm and 170–253 mm segments were built using the StalAge Monte-Carlo simulation (Fig. 2c), and the 95% confidence limit was calculated from the distribution of the simulated fits (Scholz and Hoffmann, Reference Scholz and Hoffmann2011). The ages are expressed in years before the present (AD 1950).
δ18O and δ13C records
The replication test between the ZJ2020-1 δ18O record and different speleothem records suggests that the ZJD2020-1 δ18O record reflects mainly large-scale regional precipitation δ18O (Fig. 4), hence with insignificant influence from disequilibrium effects during carbonate precipitation. The ZJD2020-1 record contains ~253 δ18O and δ13C data points with a mean temporal resolution of ca. 24 yr. The δ18O record spans 132.8–138.8 ka with δ18O values ranging from −4.3‰ to −8.3‰, with an average of −6.3‰, and δ13C data values varying between −6.2‰ to −10.2‰. The δ18O data recorded clear structures across a large portion of HS11, including the 134-ka event.
DISCUSSION
Interpretation of cave speleothem δ18O and δ13C
Statistically significant negative correlations are seen between speleothem δ18O values and rainfall amount proxies or wet/dry indexes from some regions in the ASM domain over the past ca. 60 yr (Zhang et al., Reference Zhang, Cheng, Spötl, Cai, Sinha, Tan and Yi2018; Cheng et al., Reference Cheng, Zhang, Zhao, Li, Ning and Kathayat2019; Zhao et al., Reference Zhao, Cheng, Yang, Tan, Spötl, Ning and Zhang2019). Moreover, the results based on the Experimental Climate Prediction Center's Isotope-incorporated Global Spectral Model (IsoGSM) (Yoshimura et al., Reference Yoshimura, Kanamitsu, Noone and Oki2008) and water isotope-enabled Community Earth System Model (iCESM) (Hurrell et al., Reference Hurrell, Holland, Gent, Ghan, Kay, Kushner and Lamarque2013) show a significant negative correlation between the simulated δ18O of precipitation and rainfall amount on large regional scales (Fig. S2). IsoGSM and iCESM have been widely used for both modern and past climate simulations and have been proven to be in good agreement with observations of precipitation δ18O from the Global Network of Isotopes in Precipitation (GNIP: https://www.iaea.org/services/networks/gnip) (Yoshimura et al., Reference Yoshimura, Kanamitsu, Noone and Oki2008). Additionally, the significant influence of upstream isotope depletion of δ18O (“rainout effect”) in East Asian speleothem δ18O records was attributed to variations in the summer (southerly) monsoon wind (or spatial scales of the summer monsoon circulation) and related rainout-effect changes in the moisture trajectory (Pausata et al., Reference Pausata, Battisti, Nisancioglu and Bitz2011; Liu et al., Reference Liu, Wen, Brady, Otto-Bliesner, Yu, Lu and Cheng2014). In that regard, we suggest that the integrated oxygen isotopic fractionation from rainfall drives the δ18O variability observed in the Zhangjia δ18O record, including both rainfall at the cave site and in the upstream region of the moisture transport. The latter is sensitive to monsoon intensity, or the spatial scale of summer monsoon circulation, thus includes changes in the source of moisture. Recently, summer monsoon intensity as the main control of speleothem δ18O variability in the ASM domain has been suggested (e.g., Cheng et al., Reference Cheng, Edwards, Sinha, Spötl, Yi, Chen and Kelly2016, 2019, 2022; Zhao et al., Reference Zhao, Cheng, Cao, Sinha, Dong, Pan and Pérez-Mejías2023). The consistency in the variations of the ASM speleothem δ18O records on various timescales implies a common climatic control of the overall monsoon intensity (e.g., Cheng et al., Reference Cheng, Sinha, Wang, Cruz and Edwards2012, 2016, 2019; Tan, Reference Tan2014; Zhao et al., Reference Zhao, Cheng, Yang, Tan, Spötl, Ning and Zhang2019; Liang et al., Reference Liang, Zhao, Edwards, Wang, Shao, Zhang, Zhao, Wang, Cheng and Kong2020). More broadly, the changes in large-scale monsoon circulation or monsoon intensity, including the moisture sources and integrated rainfall, would be the major factors affecting the speleothem δ18O variation with large spatial consistency in ASM regions (e.g., Cheng et al., Reference Cheng, Sinha, Wang, Cruz and Edwards2012, 2019; Zhao et al., Reference Zhao, Cheng, Yang, Tan, Spötl, Ning and Zhang2019), corresponding to more (less) incorporated oxygen isotope fractionation of moisture from the remote (proximate) tropical oceanic sources to cave sites (e.g., Tan, Reference Tan2014; Cheng et al., Reference Cheng, Edwards, Sinha, Spötl, Yi, Chen and Kelly2016, 2019).
Climate environment can affect the δ13C values of stalagmites by one or several mechanisms, and the main controlling factors of δ13C can vary with different historical periods and/or timescales. For example, δ13C may reflect a variation of C3/C4 at glacial–interglacial timescales (Dorale et al., Reference Dorale, Edwards, Ito and Gonzalez1998; Denniston et al., Reference Denniston, Asmerom, Lachniet, Polyak, Hope, An, Rodzinyak and Humphreys2013), but also may reflect variation of vegetation density and the yield of CO2 at the millennial scale. Changes in the amplitude of temperature/precipitation at various timescales are main factors of the above phenomenon (Genty et al., Reference Genty, Blamart, Ouahdi, Gilmour, Baker, Jouzel and Van-Exter2003). In this study, we suggest that millennial-scale positive excursions in our δ13C record are linked to increasing summer monsoon rainfall (Wang et al., Reference Wang, Wang, Shao, Liang, Zhang and Kong2018).
Two-stages of weak monsoon intervals
Our Zhangjia δ18O record spans the weak monsoon interval-II (WMI-II) (Cheng et al., Reference Cheng, Edwards, Wang, Kong, Ming, Kelly, Wang, Gallup and Liu2006, 2009; Kelly et al., Reference Kelly, Edwards, Cheng, Yuan, Cai, Zhang, Lin and An2006; Wang et al., Reference Wang, Wang, Shao, Liang, Zhang and Kong2018), which is characterized by higher δ18O values with an average of approximately −5.8‰. The Zhangjia record revealed two stages of the WMI-II: WMI-IIa (132.8–134.1 ka) and WMI-IIb (134.4–136.4 ka) (Fig. 4), separated by a short, strong ASM interval (134.1–134.4 ka). The WMI-IIa and WMI-IIb correspond to the cold events H11.2 and H11.1, respectively, as revealed by the cold stadials identified in the North Atlantic sediments (i.e., ice-rafted debris [IRD] events) recorded in cores MD03-2664 and ODP984 (Fig. 5i), benthic δ18O events in core ODP 983, and Uk’37 SST record in core MD01-2444 (Fig. 5e) (Channell et al., Reference Channell, Hodell and Lehman1997; Mokeddem et al., Reference Mokeddem, McManus and Oppo2014; Marino et al., Reference Marino, Rohling, Rodriguez-Sanz, Grant, Heslop, Roberts, Stanford and Yu2015; Irvalı et al., Reference Irvalı, Ninnemann, Kleiven, Galaasen, Morley and Rosenthal2016; Tzedakis et al., Reference Tzedakis, Drysdale, Margari, Skinner, Menviel, Rhodes and Taschetto2018). Thus, the WMIs identified in the δ18O record in southern China speleothems are presumably the counterparts of the low North Atlantic SST stadials (Jiménez-Amat and Zahn, Reference Jiménez-Amat and Zahn2015).
On the other hand, the North Atlantic stadials and Chinese WMIs coincide with a warming trend in the Southern Hemisphere, the increasing trend of atmospheric CO2 and CH4 (Petit et al., Reference Petit, Jouzel, Raynaud, Barkov, Barnola, Basile and Bender1999; Cheng et al., Reference Cheng, Edwards, Sinha, Spötl, Yi, Chen and Kelly2016; Schmidely et al., Reference Schmidely, Nehrbass-Ahles, Schmitt, Han, Silva, Shin, Joos, Chappellaz, Fischer and Stocker2021), and a southward shift and/or strengthening of the Southern Ocean westerlies (Toggweiler et al., Reference Toggweiler, Russell and Carson2006). The antiphase pattern between hemispheres is consistent with the “bipolar seesaw” (Stocker et al., Reference Stocker and Johnsen2003; Marino et al., Reference Marino, Rohling, Rodriguez-Sanz, Grant, Heslop, Roberts, Stanford and Yu2015), which was accompanied by collapse of the AMOC (Deaney et al., Reference Deaney, Barker and van de Flierdt2017), a southward shift of the ITCZ (Jacobel et al., Reference Jacobel, McManus, Anderson and Winckler2016), and ASM weakening (Cheng et al., Reference Cheng, Edwards, Sinha, Spötl, Yi, Chen and Kelly2016). In addition, the two stadials in NH appear to resemble the HS1 and YD (or HS0) during the last deglaciation, which apparently supports the assumption that the past two deglaciations are broadly similar in their structures (Cheng et al., Reference Cheng, Edwards, Broecker, Denton, Kong, Wang, Zhang and Wang2009; Broecker et al., Reference Broecker, Denton, Edwards, Cheng, Alley and Putnam2010).
Notably, the mean value of the Zhangjia δ18O record is higher in the WMI-IIb (−5.1‰) than that in WMI-IIa (−6‰) (Fig. 5c), suggesting a weaker monsoon interval over the WMI-IIb in the penultimate deglaciation. This interval apparently correlates the North Atlantic stadial, as inferred by the low SST recorded both in the alkenone-based record in the ODP 976 core (Fig. 5d) from the Alboran Basin (Martrat et al., Reference Martrat, Jimenez-Amat, Zahn and Grimalt2014) and the alkenone-based and Mg/Ca SST reconstructions from MD01-2444 core at the Iberian Margin (Fig. 5e) (Martrat et al., Reference Martrat, Grimalt, Shackleton, de Abreu, Hutterli and Stocker2007, 2014; Tzedakis et al., Reference Tzedakis, Drysdale, Margari, Skinner, Menviel, Rhodes and Taschetto2018). These observations suggest an extremely cold H11.1 in H11, corresponding to an abrupt weakening of the ASM. However, this feature is not evident in some of the speleothem records from southern China. For instance, there is no obvious difference regarding the mean δ18O values between WMI-IIa and WMI-IIb in the Dongge and Sanbao records (Fig. 4a, b). This discrepancy may be attributed to the regional differences in the ASM domain in response to the H11.1, which requires further investigation.
The 134-ka event during the penultimate deglaciation
One of the noticeable features of the Zhangjia δ18O record is a centennial-scale abrupt strengthening of the ASM at 134.1–134.4 ka, which divides the WMI-II into two parts: the WMI-IIa and WMI-IIb (Fig. 4e). This strong ASM event, which is centered at 134.3 ± 0.13 ka, is constrained by the combination of the 230Th dates and the fluorescence lamina counting results. This is consistent with the Hulu record, in which the corresponding strong ASM event occurred at 134.56 ± 1.0 ka, namely the Chinese Interstadial B.1 (CI B.1) (Cheng et al., Reference Cheng, Edwards, Wang, Kong, Ming, Kelly, Wang, Gallup and Liu2006) (Fig. 4d). Based on our annual-lamina counting results of sample ZJD2020-1, the duration of the 134-ka event is 349 ± 20 yr. The onset of the event inferred by the δ18O decrease (~3‰) endures for ca. 149 yr, and the end of the event, which is inferred by the δ18O increase (~3‰), spans ca. 200 yr (Fig. 3). Our precise chronology confirms a previous notion that this event (or CI B.1) lasts no longer than several hundred years (Cheng et al., Reference Cheng, Edwards, Broecker, Denton, Kong, Wang, Zhang and Wang2009). In the central and southern Chinese speleothem δ18O records (Sanbao and Dongge records), the 134-ka event is also remarkable, although the amplitudes (~2‰) are smaller than those in southwestern and northeastern China (Zhangjia and Hulu) (~3‰) (Fig. 4). In addition, an abrupt negative excursion of the Zhangjia δ13C (~3‰) co-occurred at the onset of the 134-ka event inferred by the δ18O change (Fig. 3a), suggesting an abrupt change of vegetation type/coverage and soil microbial activities in response to the onset of the event (Xue et al., Reference Xue, Cai, Ma, Cheng, Cheng, Edwards, Li and Tan2019).
δ18O and δ13C records of Zhangjia Cave. From bottom to top: (a) ZJD2020-1 δ13C record; (b) ZJD2020-1 δ18O record. (c) Insolation at 65°N on 21 July. (d) Fluorescent lamina of ZJD2020-1 at 90–100 mm. Please note that the shifts in the 134-ka event shown on the Figure 3 have been calculated in the section of the stalagmite where lamina counting was conducted. This strong ASM event, which is centered at 134.3 ± 0.13 ka BP, is constrained by the combination of 230Th dates and results of the fluorescence lamina counting.
The 134-ka event in speleothem δ18O from the ASM domain. From bottom to top: (a) δ18O record from Dongge Cave speleothem D4 (Kelly et al., Reference Kelly, Edwards, Cheng, Yuan, Cai, Zhang, Lin and An2006); (b) speleothem SB-25(dark green) and SB-11(light green) δ18O records from Sanbao Cave (Cheng et al., Reference Cheng, Edwards, Broecker, Denton, Kong, Wang, Zhang and Wang2009); (c, d) speleothem HuLu162 (light blue), MSP (pink), and MSX (navy blue) δ18O records from Hulu Cave (Cheng et al., Reference Cheng, Edwards, Wang, Kong, Ming, Kelly, Wang, Gallup and Liu2006; Wang et al., Reference Wang, Wang, Shao, Liang, Zhang and Kong2018); (e) speleothem δ18O record from Zhangjia Cave (this study). The 230Th dates with uncertainties (± 2σ) are shown in each of the error bars. The vertical blue fields indicate weak monsoon intervals, and the vertical yellow interval indicates the 134-ka event. The horizontal red-dashed lines and blue-dashed lines represent the average levels of δ18O during the time intervals. CIS = Chinese Interstadials.
Climatic variations across the 134-ka event are also widely documented in a wide range of records in the North Atlantic region that are comparable to the ASM δ18O records within age uncertainties; for example, the short cessation of IRD deposition with a concomitant decrease in N. pachyderma abundance in the ODP 984 core records from south of Iceland (Mokeddem et al., Reference Mokeddem, McManus and Oppo2014; Tzedakis et al., Reference Tzedakis, Drysdale, Margari, Skinner, Menviel, Rhodes and Taschetto2018), the rise both in the alkenone-based SST in the ODP 976 core record (Fig. 5d) from the Alboran Basin (Martrat et al., Reference Martrat, Jimenez-Amat, Zahn and Grimalt2014) and the Mg/Ca-based SST in the MD01-2444 core record from the Iberian Margin (Fig. 4e) (Martrat et al., Reference Martrat, Grimalt, Shackleton, de Abreu, Hutterli and Stocker2007, 2014; Tzedakis et al., Reference Tzedakis, Drysdale, Margari, Skinner, Menviel, Rhodes and Taschetto2018). In the Southern Hemisphere, the SST in the South Atlantic dropped around the same time (the South Atlantic inversion), which is consistent with the “bipolar seesaw” scenario (Scussolini et al., Reference Scussolini, Marino, Brummer and Peeters2015). Previously, the Vostok ice core methane (CH4) record from Antarctica (Delmotte et al., Reference Delmotte, Chappellaz, Brook, Yiou, Barnola, Goujon, Raynaud and Lipenkov2004) showed a CH4 peak near the 134-ka event. Recently, high-resolution records from the EPICA Dome C (EDC) ice-core in Antarctica (Schmidely et al., Reference Schmidely, Nehrbass-Ahles, Schmitt, Han, Silva, Shin, Joos, Chappellaz, Fischer and Stocker2021) further characterized the significant CH4 and nitrous oxide (N2O) peaks at ca. 134.5 ka with durations similar to that in the Zhangjia δ18O record (Fig. 5a, b). This abrupt increase in atmospheric CH4 may be casually linked to the increased CH4 emission due to enhanced tropical rainfall in the source regions (Bock et al., Reference Bock, Schmitt, Beck, Seth, Chappellaz and Fischer2017), as well as NH temperature increase and monsoon intensification. Taken together, the 134-ka event has clear imprints in both hemispheres, with a global pattern similar to the D-O events, Greenland interstadials, and Chinese interstadials that are well defined in the last glacial–deglacial period (e.g., Cheng et al., Reference Cheng, Edwards, Wang, Kong, Ming, Kelly, Wang, Gallup and Liu2006, 2009; Rasmussen et al., Reference Rasmussen, Bigler, Blockley, Blunier, Buchardt, Clausen and Cvijanovic2014; Tzedakis et al., Reference Tzedakis, Drysdale, Margari, Skinner, Menviel, Rhodes and Taschetto2018; Duan et al., Reference Duan, Cheng, Tan, Li and Edwards2019; Schmidely et al., Reference Schmidely, Nehrbass-Ahles, Schmitt, Han, Silva, Shin, Joos, Chappellaz, Fischer and Stocker2021).
The 134-ka event in speleothem δ18O, marine-core, and Antarctic ice-core N2O and CH4 records. From bottom to top: (a) Antarctic EDC ice-core CH4 record (Loulergue et al., Reference Loulergue, Schilt, Spahni, Masson-Delmotte, Blunier, Lemieux, Barnola, Raynaud, Stocker and Chappellaz2008; Schmidely et al., Reference Schmidely, Nehrbass-Ahles, Schmitt, Han, Silva, Shin, Joos, Chappellaz, Fischer and Stocker2021); (b) Antarctic EDC ice-core N2O record (Schmidely et al., Reference Schmidely, Nehrbass-Ahles, Schmitt, Han, Silva, Shin, Joos, Chappellaz, Fischer and Stocker2021); (c) speleothem δ18O record from Zhangjia Cave (this study). The 230Th dates with uncertainties (± 2σ) are shown in each of the error bars. (d, e) Uk’37 SST from ODP 976 (gray) and Uk’37 (purple) and Mg/Ca (gray blue) SST from MD01-2444 (Martrat et al., Reference Martrat, Grimalt, Shackleton, de Abreu, Hutterli and Stocker2007, 2014; Tzedakis et al., Reference Tzedakis, Drysdale, Margari, Skinner, Menviel, Rhodes and Taschetto2018). (f) Freshwater fluxes calculated from rates of sea-level change (Grant et al., Reference Grant, Rohling, Ramsey, Cheng, Edwards, Florindo and Heslop2014); MWP = melt water pulse. (g) Global sea level rates from benthic δ18Obenthic (Stoll et al., Reference Stoll, Cacho, Gasson, Sliwinski, Kost, Moreno and Iglesias2022); (h) Bermuda Rise 231Pa/230Th (Böhm et al., Reference Böhm, Lippold, Gutjahr, Frank, Blaser, Antz, Fohlmeister, Frank, Andersen and Deininger2015). (i) IRD (ice-rafted debris) records from ODP984 (dark) (Mokeddem et al., Reference Mokeddem, McManus and Oppo2014) and MD03-2664 (gold) (Irvalı et al., Reference Irvalı, Ninnemann, Kleiven, Galaasen, Morley and Rosenthal2016). The vertical blue fields indicate weak monsoon intervals, and the vertical yellow interval indicates the 134-ka event.
It is notable that while the durations of the 134-ka event are similar within age uncertainties among Zhangjia, Hulu, Dongge, and Sanbao speleothem δ18O records, the event in the Xinglong and Shangxiaofeng speleothem δ18O records appears to be much longer: >1600 yr (Fig. S3d; Duan et al., Reference Duan, Cheng, Tan, Li and Edwards2019) and ca. 1800 yr (Fig. S3e; Xue et al., Reference Xue, Cai, Ma, Cheng, Cheng, Edwards, Li and Tan2019), respectively. This observation could imply a regional difference. However, given the broad consistency in terms of timing and structure of the CIs during the last glacial–deglacial period (e.g., Cheng et al., Reference Cheng, Edwards, Hoff, Gallup, Richards and Asmerom2020; Li et al., Reference Li, Rao, Xu, Zhang, Liu, Wang, Cheng, Edwards and Chen2020; Duan et al., 2022), it is likely that the sample resolution and age uncertainty of the Xinglong and Shangxiaofeng records might partially explain the observed disparity. Additionally, the high-resolution CH4 and N2O records from the Antarctic EDC ice core records show a rather similar duration of a few hundred years for the 134-ka event (Fig. 5a, b), providing a strong constraint on the event duration globally. As such, we suggest that the unusually long duration of the apparent CIs around ca. 134 ka might be a unique feature in North China that resulted from the different regional climate response to the climate variations surrounding the event and/or sample resolution/dating uncertainty. Further studies are needed to resolve the discrepancy.
Climate dynamics underlying the 134-ka event
According to a common forcing mechanism for the deglaciation processes, the 134-ka event in the penultimate deglaciation, which is characterized by high amplitude and a short duration (300–400 yr) (Fig. 3), can be considered as a B/A-like event (B/A-II) in the last deglaciation. In contrast, the 134-ka event in the penultimate deglaciation had a much shorter duration (ca. 346 ± 20 yr) and smaller amplitude compared with the B/A interstadial in the last deglaciation. The B/A interstadial lasted more than 1800 yr (ca. 14.7 to 12.9 ka) (Fig. S4), during which the AMOC recovered to a level close to Holocene levels, as inferred by the 231Pa/230Th proxy from marine sediments (Fig. S4e) (Böhm et al., Reference Böhm, Lippold, Gutjahr, Frank, Blaser, Antz, Fohlmeister, Frank, Andersen and Deininger2015). Although the 134-ka event occurred in the penultimate deglaciation with a much short duration compared with the B/A interstadial, the AMOC was also at a mode equivalent to that during the B/A interstadial, as inferred by the 231Pa/230Th record (Figure 5h) (Böhm et al., Reference Böhm, Lippold, Gutjahr, Frank, Blaser, Antz, Fohlmeister, Frank, Andersen and Deininger2015). The end of the event might have been caused by a much larger freshwater forcing, the MWP-2B (Fig. 5f, g) (Schmidely et al., Reference Schmidely, Nehrbass-Ahles, Schmitt, Han, Silva, Shin, Joos, Chappellaz, Fischer and Stocker2021; Stoll et al., Reference Stoll, Cacho, Gasson, Sliwinski, Kost, Moreno and Iglesias2022), which turned the AMOC to an “off” or “Heinrich” mode (Fig. 5h) (Böhm et al., Reference Böhm, Lippold, Gutjahr, Frank, Blaser, Antz, Fohlmeister, Frank, Andersen and Deininger2015).
The MWP-2B, which is a major meltwater event during the penultimate deglaciation, contributed 70 m of sea-level rise (nearly 70% of the glacial–interglacial change) (Grant et al., Reference Grant, Rohling, Bar-Matthews, Ayalon, Medina-Elizalde, Ramsey, Satow and Roberts2012, 2014), with most meltwater discharged into the North Atlantic, as indicated by a large amount of IRD (Skinner and Shackleton, Reference Skinner and Shackleton2006; Grant et al., Reference Grant, Rohling, Ramsey, Cheng, Edwards, Florindo and Heslop2014; Marino et al., Reference Marino, Rohling, Rodriguez-Sanz, Grant, Heslop, Roberts, Stanford and Yu2015). This ice sheet melting process broadly agrees with a large set of observations, such as reduction of the surrounding North Atlantic ice sheet and concomitant long-term cooling in the North Atlantic (Marino et al., Reference Marino, Rohling, Rodriguez-Sanz, Grant, Heslop, Roberts, Stanford and Yu2015). On the other hand, the large MWP-2B apparently induced a long-term (ca. 5 ka) cold period, corresponding to the WMI-IIa (Fig. 5), which is much longer than the YD cold period (ca. 1.2 ka) in the last deglaciation (Fig. S4). Intriguingly, the MWP-1A, which was the largest meltwater pulse during the last deglaciation, occurred at an early stage of the B/A warm period (a strong ASM interval) and accounted for 15–20% of the deglacial sea-level rise mostly from Antarctic ice sheet meltwater (e.g., Weber et al., Reference Weber, Clark, Kuhn, Timmermann, Sprenk, Gladstone and Zhang2014). If the AMOC modes of the MWP-2B and MWP-1A periods are fundamentally different (Böhm et al., Reference Böhm, Lippold, Gutjahr, Frank, Blaser, Antz, Fohlmeister, Frank, Andersen and Deininger2015), they correspond to weak and strong ASM, respectively (Cheng et al., Reference Cheng, Edwards, Broecker, Denton, Kong, Wang, Zhang and Wang2009). These observations support the notion that ice volume is less effective in driving ASM changes (Cheng et al., Reference Cheng, Li, Sha, Sinha, Shi, Yin and Lu2022). In other words, the “ice volume effect” on the ASM would lie in its influence on the AMOC through North Atlantic meltwater forcing, and therefore the AMOC mode, not the ice volume, per se, would be more critical for low-latitude monsoons (Cheng et al., Reference Cheng, Li, Sha, Sinha, Shi, Yin and Lu2022). Additionally, the short duration (ca. 346 ± 20 yr) ASM event, the 134-ka event, appears to correlate with an abrupt AMOC change in a similar duration. If so, this will call for modern work to further understand the large and fast switch of the AMOC mode on centennial timescales.
CONCLUSIONS
The Zhangjia high-resolution δ18O record from southwestern China has provided unprecedented absolute and relative age precision, spanning 132.8–138.8 ka. This record precisely characterizes an abrupt ASM event, the 134-ka event, namely the CIS B.1 of previous Hulu records. The event occurred between 134.1–134.4 ka with a duration of ca. 349 ± 20 yr and δ18O amplitude of ~3‰. As inferred by our δ18O excursions, the onset of the 134-ka event endures for ca. 149 yr, and the end ca. 200 yr. This event separates the WMI-II in our record into two stages, WMI-IIa (132.8–134.1 ka) and WMI-IIb (134.4–136.4 ka). In comparison with North Atlantic climate records, we suggest that the 134-ka event, as a strong ASM event during the WMI-II that corresponded with the North Atlantic HS11, essentially corresponds to the millennial-scale events in the last glacial–deglacial period with a similar climatic pattern globally, including atmospheric CH4 and N2O jumps, the higher SST, and less IRD in the North Atlantic. It appears that the observed weak-strong-weak ASM rhythm from 138.8 ka to 132.8 ka in our record is largely controlled by the AMOC mode switches that were forced mainly by meltwater from northern high-latitude ice sheets. In this interpretative framework, the 134-ka event implies a centennial-scale faster AMOC mode or a break of the meltwater forcing in the North Atlantic, calling for further modeling study. Additionally, our results support the notion that the AMOC, rather than the ice volume, is more critical to ASM variations during deglacial processes.
Supplementary Material
The supplementary material for this article can be found at https://doi.org/10.1017/qua.2023.43
Acknowledgments
This study is supported by National Natural Science Foundation of China (NSFC 41888101, 42002199, and 42150710534).
