Abstract
Our current understanding of ocean–atmosphere–cryosphere interactions at ice-age terminations relies largely on assessments of the most recent (last) glacial–interglacial transition1,2,3, Termination I (T-I). But the extent to which T-I is representative of previous terminations remains unclear. Testing the consistency of termination processes requires comparison of time series of critical climate parameters with detailed absolute and relative age control. However, such age control has been lacking for even the penultimate glacial termination (T-II), which culminated in a sea-level highstand during the last interglacial period that was several metres above present4. Here we show that Heinrich Stadial 11 (HS11), a prominent North Atlantic cold episode5,6, occurred between 135 ± 1 and 130 ± 2 thousand years ago and was linked with rapid sea-level rise during T-II. Our conclusions are based on new and existing6,7,8,9 data for T-II and the last interglacial that we collate onto a single, radiometrically constrained chronology. The HS11 cold episode5,6 punctuated T-II and coincided directly with a major deglacial meltwater pulse, which predominantly entered the North Atlantic Ocean and accounted for about 70 per cent of the glacial–interglacial sea-level rise8,9. We conclude that, possibly in response to stronger insolation and CO2 forcing earlier in T-II, the relationship between climate and ice-volume changes differed fundamentally from that of T-I. In T-I, the major sea-level rise clearly post-dates3,10,11 Heinrich Stadial 1. We also find that HS11 coincided with sustained Antarctic warming, probably through a bipolar seesaw temperature response12, and propose that this heat gain at high southern latitudes promoted Antarctic ice-sheet melting that fuelled the last interglacial sea-level peak.
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Change history
12 August 2015
A Correction to this paper has been published: https://doi.org/10.1038/nature14960
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Acknowledgements
We thank the International Ocean Discovery Program for providing samples from Leg 161 and W. Hale for sampling assistance. M. Charidemou contributed to a pilot study on ODP975, J. Amies to preliminary stratigraphic assessment, and J. Cali to stable isotope analyses. We thank M. Bar-Matthews for providing the updated version of the Soreq Cave chronology table, R. N. Drysdale and J. Hellstrom for helpful discussions, and B. Martrat, P. Jimenez-Amat, R. Zahn, and J. O. Grimalt for the ODP976 records, L. Skinner for the MD01-2444 data, and other colleagues who made data available through the NOAA National Climatic Data Center and/or Pangaea. This study was supported by Australian Research Council Australian Laureate Fellowship FL120100050 and by UK-NERC project NE/I009906/1 (E.J.R.).
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G.M. and E.J.R. designed the study. G.M. prepared the foraminiferal samples for stable isotope analyses and together with E.J.R. and D.H. performed the statistical analysis. L.R.S. performed the stable isotope analyses. K.M.G. and J.D.S. oversaw early stratigraphic assessment. All authors contributed to interpretation and preparation of the final manuscript.
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Extended data figures and tables
Extended Data Figure 1 Orbital parameters, insolation forcing, and atmospheric CO2 concentrations during the last two glacial–interglacial transitions.
a, f, Incoming solar radiation on 21 June and 21 December at 65° N (orange) and 65° S (blue), for T-I (a) and T-II (f)31. b, g, Precession of the Earth’s axis31 for T-I (b) and T-II (g). c, h, as b, g, but for eccentricity of the Earth’s orbit31. Note the different axis scales in c and h. d, i, Obliquity of the Earth’s axis31 for T-I (d) and T-II (i). Different orbital configurations during glacial terminations I (T-I) and II (T-II) led to markedly different insolation forcing31,71,72,73, with the summer 65° N insolation increase during T-II larger by ∼35% (∼23 W m−2), and occurring at faster rates18, than that during T-I. e, j, atmospheric CO2 concentrations during T-I (refs 74, 75, 76) and T-II (ref. 56). Yellow bands illustrate the timing of Heinrich stadial 1 (HS1) and 11 (HS11), following previous literature1 and this study, respectively. Dashed blue lines highlight that atmospheric CO2 concentrations were systematically higher during T-II than during T-I. Symbols indicate the maxima and/or minima in insolation, precession, eccentricity, and obliquity cycles.
Extended Data Figure 2 Flowchart of the approach used to construct and validate chronologies of the various records.
See main text and Methods for details.
Extended Data Figure 3 Synchronization of western Mediterranean Sea ODP975 to ODP977 and Iberian Margin core MD01-2444.
a, Synchronization of G. bulloides δ18O from western Mediterranean ODP977 to its counterpart from nearby ODP975. Circles and error bars depict tie points and 2σ synchronization errors, respectively (Methods, Extended Data Table 3). The G. bulloides δ18O from ODP976 is only shown for comparison and is not used to transfer our new, radiometrically constrained chronology, to ODP977 (Methods). b, Validation of the synchronization exercise shown in a by comparing the alkenone-based sea surface temperature (SST) records from ODP977 and ODP976, on their new, radiometrically constrained chronology. The 95% confidence intervals (light grey envelope) and probability maximum (heavier grey line) and its associated 95% confidence intervals (heavier grey envelope) of the ODP976 SST data (grey circles) are based on a Monte Carlo analysis of chronological and SST uncertainties, employing a 0.2 kyr Gaussian filter (Methods). The synthetic record of Greenland climate variability16 is also shown. c, Synchronization of the G. bulloides δ18O from Iberian Margin (Atlantic Ocean) core MD01-2444 to its counterpart from ODP975 (Western Mediterranean, Methods). Circles and error bars depict tie points and 2σ synchronization errors, respectively (Methods, Extended Data Table 3). The G. bulloides δ18O from ODP976 is shown for comparison and is not used to transfer the MD01-2444 records to our new, radiometrically constrained chronology (Methods). d, Validation of the synchronization exercise shown in c by comparing alkenone-based SST records from core MD01-2444 and ODP976 (shaded envelopes and circles as in b), on their new radiometrically constrained chronology. The synthetic record of Greenland climate variability16 is also shown.
Extended Data Figure 4 Relationship between the duration of the North Atlantic cold phases (stadials) and the magnitude of Antarctic warming.
a, δ18Oice time series from North Greenland Ice Core Project (NGRIP)77(light blue). The probability maximum (solid blue line) and associated 95% confidence bounds (shaded blue envelope) of the δ18Oice record result from 10,000 Monte Carlo simulations, employing a 0.15 kyr Gaussian filter through the data and their chronological and δ18Oice uncertainties (see Methods). b, EPICA Dome C (EDC) temperature reconstructions (ΔT) based on δDice data13 (light red). The probability maximum (solid red line) and associated 95% confidence bounds (shaded red envelope) of the temperature record result from 10,000 Monte Carlo simulations, employing a 0.2 kyr Gaussian filter through the data and their chronological and ΔT uncertainties. c, Comparison between Antarctic temperature reconstructions from EDC using the methods of Jouzel et al. (ΔT, ref. 13) and Stenni et al. (TSITE, ref. 78), revealing no difference in the amplitude of the estimated Antarctic temperature shifts. d, Relationship between duration of Greenland stadials and Antarctic warming for bipolar seesaw events highlighted in a and b by yellow bands. Greenland and Antarctic records are on the latest ice core chronology AICC2012 (ref. 55). Error bars depict 1σ uncertainty in the magnitude and duration of North Atlantic cooling.
Extended Data Figure 5 Synchronization of western Mediterranean Sea ODP975 to eastern Mediterranean core LC21.
a, Synchronization of the N. pachyderma (d) δ18O from ODP975 to its counterpart from eastern Mediterranean core LC21 (Methods), which was placed on a radiometric chronology by ref. 8. Red filled circles and error bars depict the tie points used to synchronize ODP975 to LC21 and the 2σ synchronization uncertainties, respectively. b, Validation of the synchronization exercise shown in a by comparing the respective co-registered N. pachyderma (d) δ13C profiles from the same cores.
Extended Data Figure 6 Glacial 18O-enrichment in oceanic and Mediterranean seawater.
Blue (open ocean), black (eastern Mediterranean Sea) and green (western Mediterranean Sea) lines illustrate the relationship between seawater δ18O (δ18Oseawater) and relative sea level (RSL). The open ocean relationship has been re-calculated here using the latest assessment (grey symbol) of the sea-level lowstand11 of the Last Glacial Maximum (26.5–19 ka, ref. 79) and the mean ocean δ18Oseawater value derived from porewater analyses in a suite of deep-sea cores63. The eastern Mediterranean relationship is derived from the model presented in ref. 60. The western Mediterranean relationship (green line) reflects linear mixing between eastern Mediterranean (black line) and open ocean (blue line) end-members, assuming that the residence time effect60,62 on δ18Oseawater (and salinity) in the western Mediterranean is 70 ± 5% of that in eastern Mediterranean64. Shaded envelopes are 1σ confidence bounds, derived from probabilistic analysis of uncertainties.
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This file contains the MATLAB code used to calculate the freshwater fluxes associated with ice sheet melting during Termination II. (PDF 83 kb)
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Marino, G., Rohling, E., Rodríguez-Sanz, L. et al. Bipolar seesaw control on last interglacial sea level. Nature 522, 197–201 (2015). https://doi.org/10.1038/nature14499
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DOI: https://doi.org/10.1038/nature14499
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