Penultimate deglaciation Asian monsoon response to North Atlantic circulation collapse

During glacial terminations, massive iceberg discharges and meltwater pulses in the North Atlantic triggered a shutdown of the Atlantic Meridional Overturning Circulation (AMOC). Speleothem calcium carbonate oxygen isotope records (δ18OCc) indicate that the collapse of the AMOC caused dramatic changes in the distribution and variability of the East Asian and Indian monsoon rainfall. However, the mechanisms linking changes in the intensity of the AMOC and Asian monsoon δ18OCc are not fully understood. Part of the challenge arises from the fact that speleothem δ18OCc depends on not only the δ18O of precipitation but also temperature and kinetic isotope effects. Here we quantitatively deconvolve these parameters affecting δ18OCc by applying three geochemical techniques in speleothems covering the penultimate glacial termination. Our data suggest that the weakening of the AMOC during meltwater pulse 2A caused substantial cooling in East Asia and a shortening of the summer monsoon season, whereas the collapse of the AMOC during meltwater pulse 2B (133,000 years ago) also caused a dramatic decrease in the intensity of the Indian summer monsoon. These results reveal that the different modes of the AMOC produced distinct impacts on the monsoon system. The influence of meltwater pulse events on Asian monsoon systems varied in line with the degree of AMOC weakening, according to a multi-proxy analysis of speleothems from China covering the penultimate glacial termination.

T he Asian monsoon provides fresh water for a densely populated region and plays a major role in global climate as a conveyor of latent and sensible heat 1 . Changes in the cross-equatorial ocean energy transport driven by the Atlantic Meridional Overturning Circulation (AMOC) play a fundamental role in the monsoon system through its impact on the interhemispheric temperature gradient and the latitudinal position of the intertropical convergence zone (ITCZ) 2 . A weaker AMOC also decreases sub-polar temperatures, which affect the course and strength of the westerly winds that interfere with southerly monsoon winds 3 . Palaeo-climate records offer a unique opportunity to study the links between changes in AMOC strength and the Asian monsoon system during periods of the past when the AMOC was substantially weaker or even completely collapsed 4,5 , but there is no general consensus about the dominant mechanisms. Understanding the response of the monsoon system to rapid changes in the strength of the AMOC is particularly relevant in the context of recent observational evidence suggesting a weakening of the AMOC over the past decades [6][7][8][9] .
The oxygen isotopic composition of speleothem calcite from East Asia suggests that rapid changes in the strength of the AMOC and the Asian monsoon were indeed tightly linked 3,[10][11][12][13][14][15] . During the coldest parts of the last glacial, that is, during Heinrich events, a weakening or collapse of the AMOC 4 was systematically associated with higher speleothem δ 18 O Cc (refs. 12,13,16 ). Classically, higher speleothem δ 18 O Cc values have been interpreted as 'weak monsoon intervals' and explained by a lower proportion of low δ 18 O P (Indian summer monsoon (ISM) and East Asian summer monsoon (EASM)) rainfall in annual totals 11,16,17 and or decreased upstream depletion of 18 O in either the ISM or East Asian monsoon (EAM) regions 11,14,17,18 , and vice versa for low speleothem δ 18 O Cc values, which are interpreted as 'strong monsoon intervals' . However, with the available data, it is not possible to distinguish how the different components of the Asian monsoon system (for example, temperature gradients, the position of the ITCZ, the length of the monsoon season or the intensity of the ISM or EASM) are affected by changes in the strength of the AMOC.

Deconvolving calcite δ 18 O
Part of the challenge arises from the fact that δ 18 O Cc depends on not only mean annual drip water δ 18 O but also temperature-controlled isotope fractionation between water and calcite 19 , and potential kinetic isotope effects 20 . Here, we use an innovative approach to deconvolve the effects of kinetic isotope fractionation, temperature and drip water δ 18 O by combining: (i) the TEX 86 speleothem thermometer to reconstruct changes in cave air temperatures 21,22 , (ii) the dual clumped isotope (Δ 47 and Δ 48 ) thermometer that allows to assess both changes in cave temperature and the extent of kinetic isotope fractionation 23,24 and (iii) the hydrogen and oxygen isotope composition of speleothem fluid inclusions, that is, δ 2 H Fl and δ 18 O Fl , as a proxy for the isotopic composition of precipitation, that is, δ 2 H P and δ 18 O P (refs. 25,26 ). We applied these techniques to two south-west Chinese speleothems from Jiangjun Cave spanning the penultimate deglaciation 27 ( Fig. 1 and Extended Data Figs. 1 and 2), a time period that was characterized by two successive episodes of vast freshwater discharge into the North Atlantic 28 (meltwater pulses (MWPs)) that caused a weakening 29 (MWP 2A) and full collapse (MWP 2B) of the AMOC 4 .
The present-day local climate at Jiangjun Cave is characterized by warm/wet summers and cold/dry winters. The summer monsoon peaks from May to September 30 , contributing ~73% to annual rainfall (~1,320 mm per year), and the mean annual air temperature is ~23 °C. Wet-season (summer) precipitation in the region of Jiangjun correlates positively with precipitation in north-east India, which, in turn, depends on the strength of the ISM 31 . Consequently, Jiangjun δ 18 O P is about 10‰ lower in July-August compared with January-March (Extended Data Fig. 3).   isotope data ( Fig. 2b and Supplementary Data Tables 1c-f). In addition, our dual clumped isotope data show that calcite precipitated indistinguishable from isotope equilibrium (Extended Data Fig.  5). Applying a temperature-dependent water-calcite isotope fractionation of 0.21‰/1 °C (ref. 32 Table 1g). One potential explanation for these differences could be that the fluid inclusions are not well preserved and do not represent palaeo-rainfall. A key criterion to assess this is that, if unaltered, the fluid inclusion isotope data should plot close to the local and global meteoric water line 33 . For Jiangjun Cave, the most representative Global Network of Isotopes in Precipitation data are from Kunming, which are highly consistent with the fluid inclusion data (Fig. 3a). Furthermore, Jiangjun's interglacial δ 18 O Fl values are very similar to values calculated for parent drip water based on Jiangjun's TEX 86 and dual clumped isotope temperatures and the δ 18 O temperature calibration from ref. 34 , which bolsters the quality of our dataset (Extended Data Fig. 6). The fluid inclusions are thus well preserved and represent the isotopic composition of the parent water.

Structure of Termination ii
Closer examination of our dataset revealed a ~1-3‰ offset between parent water δ 18 O (calculated by subtracting the temperature contribution from calcite δ 18 O) and δ 18 O Fl from ~139 to ~128 ka, with δ 18 O Fl being lower than calculated parent water values (Extended Data Fig. 6). Since this offset occurs in the overlap zone of the two speleothems where δ 18 O Cc , δ 18 O Fl and TEX 86 proxy records replicate well, we do not believe that this reflects kinetic fractionation or diagenetic processes. We propose instead that, in this interval, the difference arises because δ 18  . These data suggest that, during the driest part of the glacial period, characterized by low speleothem growth rates (Fig. 2d) and a less developed soil on top of the cave, strong monsoon rainfall events infiltrated rapidly and were captured in fluid inclusion-rich calcite laminae. In the dry season, fluid inclusion-poor laminae were formed, causing a distinct seasonal bias in fluid inclusion incorporation, while the calcite grew in both wet and dry seasons 35 (Fig. 3b). This effectively decouples δ 18 O Fl and δ 18 O Cc records, and highlights that δ 18 O Fl predominantly captures the strength or 'intensity' of the ISM that integrates the rainfall history along the moisture pathway, while δ 18 O Cc values are more likely to reflect the annually averaged hydroclimate signal.

implications for monsoon dynamics
Our data indicate that MWP 2A did not affect δ 18 O Fl but did affect temperature and δ 18 O Cc , so how can this be reconciled? MWP 2A was associated with a peak in ice-rafted debris 36 (IRD) in the North Atlantic and caused a weakening of the AMOC as observed in benthic δ 13 C values (ref. 29 ) (Fig. 4). Two additional IRD peaks of similar magnitude have been observed before and after MWP 2A, also associated with benthic δ 13 C minima, suggesting multiple periods of weakening and reinvigoration of the AMOC during this time period. A weakening of the AMOC leads to cooling in the subpolar North Atlantic, which increases the equator-to-pole temperature gradient and triggers a southward shift of the westerly jet, as shown by climate simulations of HS1 (ref. 3 ). Interestingly, our speleothem TEX 86 temperature record shows three cooling events (of ~1 °C) in south-west China that occur within age uncertainties of these North Atlantic IRD events (Fig. 4). However, MWP 2A, as well as the two additional IRD events, were likely too small to induce a complete shutdown of the AMOC. Consequently, ocean-to-continent temperature gradients were likely not affected to the extent that they impacted the strength of the ISM, as suggested by the relatively stable Jiangjun δ 18  is supported by the high correlation (r = −0.74, P < 0.001, n = 17) between δ 18 O Fl and independent marine reconstructions of the ISM wind-induced upper ocean stratification in the east equatorial Indian Ocean 37 , which suggest a weak response of the ISM wind strength to MWP 2A and surrounding MWP events (Fig. 4). Despite this apparent lack of response of the ISM intensity, a clear signal is observed in the δ 18 O Cc records during MWP 2A and the onset of MWP 2B that cannot be explained by the observed temperature change alone, suggesting a change in annual mean δ 18 O P . We hypothesize that these observations can be reconciled if the timing of the northward retreat of the westerly jet in the Asian monsoon region was delayed in response to the enhanced equator-to-pole temperature gradient, shortening the monsoon season in a similar fashion as during HS1 (ref. 3 ). In agreement with climate simulations 3 and Jiangjun speleothem growth rate (Fig. 2), this scenario would have resulted in a decrease in annual mean precipitation at Jiangjun during this interval (Fig. 2d). The consequences of a delayed northward retreat of the westerly jet would have been similar in northern India and in the Arabian Peninsula, causing also a shortening of the monsoon season and thus more positive annual mean parent water δ 18 (Fig. 4) is in agreement with a moderate southward shift of the westerly jet. Thus, our interpretation represents a viable mechanism to explain the consistency of speleothem δ 18 O Cc across the Asian monsoon region, as well as their apparent discrepancy with the available ISM records. In contrast to MWP 2A, MWP 2B was associated with massive iceberg discharge events in the North Atlantic and a full shutdown of the AMOC, as shown by the kinematic AMOC tracer 231 Pa/ 230 Th 4 (Fig. 4e). Our data reveal that MWP 2B was characterized by colder temperatures and a dramatic decrease in the intensity of the ISM (reflected in a 2-3‰ increase in δ 18 O Fl ; Figs. 2 and 4). These changes coincide with a maximum in planktonic foraminifera surface-to-thermocline δ 18 O gradients in the southernmost Bay of Bengal (Fig. 4b). The role of sea surface temperature in the Bay of Bengal during Heinrich events has not been resolved yet 40 , but climate model simulations of HS1 suggest that, as opposed to the Arabian Sea, sea surface temperature did not change significantly in that region (see fig. S4 in ref. 15 ), suggesting that the increase in the planktonic foraminifera gradients reflects a strengthening of ocean stratification in response to a weaker ISM wind strength. Our data provide empirical support for this interpretation. In addition, the timing and structure of our δ 18 O Fl mirror the Mulu cave δ 18 O Cc record, suggesting that the weakening of the ISM during MWP 2B was closely coupled to the migration of the ITCZ towards its southernmost position (Fig. 4a). Thus, our data suggest that the strong Northern Hemisphere cooling associated with the collapse of the AMOC during MWP 2B affected both the ocean-to-continent and the equator-to-pole temperature gradients, impacting both the intensity of the ISM and the length of the monsoon season.
These results underpin the potential of using multiple analytical techniques to quantitatively deconvolve the different components affecting speleothem δ 18 O Cc . Our δ 18 O Fl record suggests that the intensity of the ISM only decreased substantially with the full shutdown of the AMOC during the peak of MWP 2B, which was associated with an extreme southward migration of the ITCZ. In contrast, smaller AMOC reductions associated with earlier North Atlantic IRD events, including MWP 2A, appeared to have impacted atmospheric temperatures and reduced the length of the monsoon rain season (as suggested by δ 18 O Cc records), but did not affect the intensity of the ISM substantially. Thus, our data indicate a distinct response of the Asian monsoon system to changes in the strength of the AMOC.

Online content
Any methods, additional references, Nature Research reporting summaries, source data, extended data, supplementary information, acknowledgements, peer review information; details of author contributions and competing interests; and statements of data and code availability are available at https://doi.org/10.1038/ s41561-021-00851-9.
For the MS set-up, the water vapour is transported with He carrier gas to a cold trap at −105 °C. Water is trapped for 6 min, then flash-heated and transported to the TC-EA 26,48 . We used a standard bracketing approach with an evaluation procedure to correct for signal intensity, memory and non-linearity effects 26 . Memory and non-linearity effects of the set-up were assessed on a weekly basis following ref. 26 . By choosing a water standard with an isotope composition that is within a few ‰ of the fluid inclusions, the memory effect was minimized. To correct for signal intensity, a series of at least three standard waters were analysed before the crush, with water amounts ranging from 0.28 to 0.12 µL. After the crush, two or three injections of 0.4 µL of standard water were injected to flush the system, which was followed by one standard water analysis, which was not used for the data evaluation because it may still be affected by a potential memory effect related to the fluid inclusions from the crushed calcite. A minimum of another two water standards with water amounts similar to the crush were analysed to assess the stability of the system and verify that it remained leak-free during and after the crush. We applied two outlier criteria: signal intensity and replication. Only samples with a water yield resulting in a peak area larger than 35 V-s for δ 2 H were considered for the dataset, while any samples with a lower water yield were discarded. For 22 out of 31 samples, it was possible to do two or more analyses with the TC-EA MS technique. Only one sample did not replicate within 1‰ for δ 18 O and was discarded (Supplementary Data Table 1g).
For the CRDS set-up, an online preparation line detailed in ref. 26 is coupled to a Picarro L2140i water isotope analyser. A stable water background of known isotope composition is generated and subtracted from the sample peak, in a similar way to refs. 49,50 . For correction purposes, we analysed seven water standards of 0.3 µL before the first sample and five water standards with volumes corresponding to the water yield from the crushed calcites after the last sample. Each of the four samples measured with the CRDS technique was measured twice, replicating within 0.1‰ for δ 18 O (Supplementary Data Table 1g).
For fluid inclusion isotope analysis conducted on both the MS and CRDS systems, the uncertainties are based on the reproducibility reported by ref. 26 . In that study, fluid inclusion isotope analysis was conducted on the same TC/ EA-MS and CRDS systems that produced the fluid inclusion isotope data for this manuscript. The reproducibility is based on repeated measurements of five samples. For the IRMS system, the reproducibility is 0.4‰ for δ 18 Table 1c-f) was performed at Goethe University Frankfurt using the analytical set-up described in ref. 23 . Δ 47 and Δ 48 data are reported on the Carbon Dioxide Equilibrium Scale at a reaction temperature of 90 °C (CDES 90). For data processing, we followed the CDES 90 correction protocol detailed in ref. 24 . It includes (1) correction of m/z 47, 48 and 49 raw intensities for negative backgrounds using scaling factors and intensities measured on the m/z 47.5 cup, (2) correction for scale compression using CO 2 gases equilibrated at 25 °C and 1000 °C, respectively, and (3) correction for subtle long-term variations in Δ 47 values of reference carbonates ETH 1 and ETH 2. The two samples (JJ0406-312 and JJ0406-690) were analysed between May and August 2019, that is, covering the same period as session 2 in ref. 24 . Equilibrated gas data, reference carbonate data, empirically determined scaling factors, empirical transfer functions and temporal Δ 47 offset functions relevant for sample data correction are, therefore, identical to those reported for session 2 in ref. 24  Analysis of GDGTs. The procedure to analyse the distribution of glycerol dibiphytanyl glycerol tetraether (GDGT) lipids in the stalagmites from Jiangjun Cave was based on an adaptation of the methods proposed by ref. 22 . GDGT measurements (Supplementary Data Table 1b) were performed in the same samples used for fluid inclusion measurements, to maximize the comparability Methods Stalagmite samples. The speleothems were collected from Jiangjun's main passage, which has a length of 4 km, height of 1.5-20 m and width of 3-25 m. The speleothems were collected ~2.8 km from the main entrance. The surface above the cave is densely vegetated during the summer season. The main entrance is covered by ~80 m of host rock.
Stalagmite JJ0406 from Jiangjun Cave is 860 mm long and grew 460 mm from 139 to 124 ka, JJ0403 is 485 mm long and grew 348 mm from 147 to 134 ka. JJ0406 shows a pool-type structure at the growth axis, which we avoided when sampling for fluid inclusions and 230 Th/U dating. We show a thin section of primary calcite with clear continuous lamination that corresponds to the fabric from which fluid inclusion samples were taken (Extended Data Fig. 1). The speleothem calcite δ 18 O records were already published 27 . There is a second Jiangjun Cave in south-west China 41 , which should not be confused with the Jiangjun Cave from this manuscript.
Uncrushed samples for fluid inclusion isotope analysis encompass ~5-20 mm of stratigraphy at the growth axis and thus represent about 50-200 years with a speleothem growth rate of 100 µm per year, or more with lower growth rates. Therefore, the fluid inclusion isotope record tracks ISM strength on multi-decadal to centennial timescales.  Table 1a). Per sample, between 200 and 300 mg carbonate powder was used. Uranium and Th were chemically separated following a protocol adapted after ref. 42 and were measured using the technique described in ref. 43 . We used U decay constants as in refs. 43,44 . JJ0406 and JJ0403 contain only 20 ppb U, and 234 U/ 238 U initial atomic ratios are low. These stalagmites are thus rather difficult to date, which resulted in relatively large age uncertainties (Supplementary Data Table 1a). Part of the U series ages of speleothem JJ0406 were recently published 27 , being consistent with new U-series ages produced for this study. Age-depth modelling was performed using the StalAge algorithm 45 (Extended Data Fig. 2).
Carbonate oxygen isotope analysis. The oxygen and carbon isotope records of stalagmites JJ0403 and JJ0406 were constructed in the Xi'an Jiaotong Isotope Laboratory. Oxygen isotopes were published in ref. 27 Table 1a.
The crushed calcite powders were analysed at the Max Planck Institute for Chemistry in Mainz, on a Thermo Delta V mass spectrometer equipped with a GASBENCH preparation device. A CaCO 3 sample (~20-50 μg) in a He-filled 12 ml exetainer vial was digested in >99% H 3 PO 4 at 70 °C. Subsequently the CO 2 -He gas mixture was transported to the GASBENCH in helium carrier gas. In the GASBENCH, water vapour and various gaseous compounds were separated from the He-CO 2 mixture before sending it to the mass spectrometer in nine separate peaks. Isotope values of these individual peaks were averaged and are reported as δ 18 O relative to VPDB. A total of 20 replicates of two in-house CaCO 3 standards were analysed in each run of 55 samples. CaCO 3 standard weights were chosen such that they span the entire range of sample weights of the samples. The reproducibility of these standards is typically better than 0.1‰ (1 s.d.) for δ 18 O. The oxygen isotope data of the crush residues are presented in Supplementary Data Table 1g.
High-resolution sampling on the thick section of sample JJ0406-690 was performed with a micromill at 20 and 40 µm resolution using a pointy-tipped drill bit. Approximately 3-5 µg of carbonate was analysed with a cold-trap method connected to the setup described above (see ref. 46 for more information). The oxygen isotope data are presented in Supplementary Data Table 1a.
Fluid inclusion isotope analysis. Stalagmites were sampled for fluid inclusion isotope analysis (Supplementary Data Table 1g) using a handheld device (Dremel Micro), equipped with a 20-38 mm diameter diamond-covered circular saw. Depending on the water content of the calcite, we cut small calcite blocks of 0.1-1.1 g per analysis. Fluid inclusion isotope analysis was performed following the protocols adapted after ref. 26 using a Delta V advantage mass spectrometer (MS) coupled to a Conflow IV and a Thermo Conversion Element Analyser (TC-EA) from Thermo Scientific or an online preparation line coupled to a Picarro L2140i cavity ring-down spectrometer (CRDS). The calcite was crushed by a downward rotary movement in a Potsdam-type crusher 47 at a temperature of 120 °C. In ref. 26 , aliquots of the same samples were analysed with both techniques repeatedly. The results were consistent, proving that the techniques can be used interchangeably. In between samples, the crusher was cleaned with acetic acid and milliQ water and subsequently dried with pressurized air. of temperature estimates with calcite and fluid inclusion δ 18 O. After the fluid inclusion measurement, 0.55-1.54 g pulverized material was weighed in a 60 mL cylindrical glass vial. Calcite was subsequently dissolved in 20 mL of 6 N HCl and digested on a heating plate at 100 °C for 2 h. After cooling, the digested sample was extracted three times using 20 mL dichloromethane (DCM) in each extraction. The DCM fraction was collected in a 60 mL vial using a 50 mL separation funnel. After the third extraction, 40 µL of C46 GDGT internal standard was added to the extract for GDGT quantification purposes 54 . The solvent was evaporated to complete dryness under nitrogen gas on a heating plate at 45 °C. To remove potential traces of acid and other impurities, the extract was re-dissolved in a mixture of DCM-methanol and passed through a 5% deactivated silica column. The eluate was collected in 4 mL vials, dried, and filtered into a 1 mL vial through a 0.2 µm polytetrafluoroethylene membrane filter using a 1.8% mixture of isopropanol in n-hexane. The filtered extracts were analysed using Agilent 1260 Infinity high-performance liquid chromatography (HPLC) coupled to an Agilent 6130 single-quadrupole MS, using a slight adaptation of the method proposed by ref. 55 , as described in ref. 56 .
Cave temperature estimates were obtained by using the recent calibration proposed by ref. 21 Cave air T = −7.34 + 34.64 TEX86 The external analytical precision of the method was estimated by analysing a speleothem standard sample prepared by homogenizing around 1 kg of flowstone material from Scladina Cave, Belgium. This standard was analysed at least in duplicate in each batch of samples analysed, obtaining an average cave T of 14.57 °C with a standard deviation of 0.28 °C (1 s.d., n = 38). Uncertainties shown in Fig. 2 are 2 s.d. based on the external analytical precision derived from repeated measurements of the speleothem standard (0.56 °C). Due to the limited sample availability, full replicate analysis was only possible in one sample. The s.d. obtained (0.20 °C) is consistent with our estimate from our long-term in-house standard. Calibration uncertainty for speleothem TEX 86 is ~1 °C (2 s.d.) in the temperature range of the samples presented in this study 21 . However, it is important to note that calibration uncertainties mainly affect absolute values not sample-to-sample variability.
The potential effect of crushing and/or heating (to 120 °C) during the fluid inclusion measurement and the cleaning procedure of the crusher on estimated GDGT temperatures was evaluated by performing the fluid inclusion measurement protocol on three samples of our standard speleothem sample from Scladina Cave. The analysis of GDGTs after the fluid inclusion treatment yielded an average T of 14.61 °C with a standard deviation of 0.10 °C (1 s.d., n = 3), which is statistically indistinguishable from the T obtained for untreated samples (t test, P < 0.01). These results are consistent with the reported stability of TEX 86 estimates in samples subjected to hydrous pyrolysis at temperatures below 220 °C (ref. 57 ).
Calculation of δ 18 O P . Parent water δ 18 O was calculated using water-to-calcite isotope fractionation factors based on the temperature calibration from ref. 34 , our TEX 86 temperature data and dual clumped isotope temperatures.
Data smoothing. Jiangjun δ18OFl and TEX86 smoothed loess curve in Fig. 4 was modelled with a smoothing factor of 0.17, and 95% confidence limits were determined by a bootstrapping technique in PAST software 59 .

Data availability
All data produced for this study are available in Supplementary Data 1 and are stored in the data repository of the National Oceanic and Atmospheric Administration. Other source data for Figs. 2-4 and Extended Data Figs. 3-4 are referenced in the Source data provided with this paper.

Code availability
This research paper has no associated code. Fig. 1 | Sampling positions. Positions of 230 Th/U dating, oxygen isotope transect (magenta), fluid inclusions (blue) and TEX 86 analysis for which the crush residue did not provide enough material for one analysis (red) are indicated. Thin section under polarization microscope of calcite with fluid inclusions is shown for speleothem JJ0406. The bottom part of JJ0406 zoomed in on the right shows the most pronounced pool structure at the growth axis and dissolution features indicating the lowest drip water saturation indices.