The small-to-insignificant CO 2 change during the short stadials may imply that AMOC perturbations happened at these events but were too short to result in a change in atmospheric CO 2 . If this were the case, we would expect to observe no CO 2 increase during the first 800–1200 years of the long stadials, because duration of the short stadials ranges 800–1200 years. However, we observe that CO 2 increases from the beginning of the long Greenlandic stadials predating DO8 and DO4 (Fig. 2a). The time lag of CO 2 relative to the isotopic signal during Greenlandic stadials predating DO12 and 17 from other ice core records appears small as well11,12. However, we cannot clearly rule out the possibility of a time lag of several centuries (Fig. 3). In addition, there is some ambiguity about the start of the long stadial predating DO4, which is conventionally defined by the end of a small temperature proxy peak (DO4.1)21. We follow convention here, but note that additional high-resolution data from other long stadial events will be needed to further address the question of when CO 2 starts to rise during events of this type. The above observation suggests that climate perturbations associated with the long and short stadials are different. Cave deposits reveal less weakening in the Asian monsoon30 and less intense South American monsoon31 during the short stadials compared with the long stadials, suggesting that the perturbation to the climate system related to the short stadial events in Greenland was weaker than for the long ones. A comparison of Antarctic ice core climate records with a thermodynamic model also indicates that the long stadials were caused by a stronger climate perturbation than short ones32. Finally, although it is not conclusive, δ13C in benthic foraminifera from North Atlantic sediment cores indicates less shoaling of AMOC during short stadials than that during the long stadials33. Thus it is likely that strength of the climate perturbation is related to change in atmospheric CO 2 during the Greenlandic stadials. Massive iceberg discharge events in the North Atlantic (Heinrich events) occurred within time intervals of the long stadials. The Heinrich events could have increased fresh water forcing into the North Atlantic and also caused large perturbations to atmospheric circulation (for example, southward movement of the ITCZ34). However, multiple studies suggest that the Heinrich events lag onsets of long stadials35,36, although exact timing of those events within the stadials is not well constrained37,38.

Figure 3: Atmospheric CO 2 and climate records from multiple ice cores. (a–d) Ice core records extended from Fig. 1. Siple Dome CO 2 and CH 4 records are compared with existing low-resolution records from EPICA EDML12,21, Byrd11 and Talos11,58 ice cores, Antarctica. Age intervals for HS2 and HS5a are not well constrained owing to chronological uncertainty in the paleoproxy records. Full size image

The control mechanisms for the two CO 2 modes may exist in oceanic processes such as AMOC reduction and consequent upwelling in the Southern Ocean. Those oceanic processes can be linked by change in vertical salinity transport and stratification in the Southern Ocean39 and/or latitudinal shift of Southern Hemisphere Westerlies14,40 and/or strength of the Southern Hemisphere Westerlies41,42,43. Although we cannot pinpoint a precise oceanic mechanism, we speculate that the weakening in AMOC during the short stadials might have not been sufficient to cause enough of a change in upwelling to impact atmospheric CO 2 . Marine proxy data for upwelling in the Southern Ocean do not clearly show strong peaks in between long stadials that bracket several short stadials14, supporting this hypothesis.

Other potential oceanic mechanisms that change CO 2 outgassing include variations in sea ice extent and changes in iron fertilization in the Southern Ocean44. Sea-salt-Na may be a proxy for sea ice extent, but Siple Dome, Dome C and EDML ice core records do not show significant differences between long and short Greenlandic stadials44,45. Proxy records for the Fe-flux (non-sea-salt Ca) from Dome C and EDML cores show highly reduced Fe-flux during several long Greenlandic stadials that predate DO8, but after DO8 the reduction during long stadials is not larger than that during short stadials44. Thus a difference in iron fertilization in the Southern Ocean is not likely the main cause of the two modes in CO 2 change.

Atmospheric CO 2 can be also controlled by exchange of land carbon. Terrestrial carbon is mostly affected by temperature and precipitation because they both control vegetation and organic carbon in soil. Compared with interstadials, paleoproxy data for both short and long stadials indicate colder and dryer conditions in the northern hemisphere, and warmer and wetter conditions in the southern hemisphere, although the magnitude of those changes depends on the type of stadials6. However, model simulations predict either a decrease46,47 or increase48 in land carbon during the stadials. Although we cannot rule out terrestrial control on the two modes of CO 2 change, we suggest that the control mechanism exists more likely in the ocean rather than on land, because we have supporting evidence for an oceanic CO 2 source during the long stadials in the last deglacial period14,15,49,50.

In principle, the lack of change in atmospheric CO 2 could also result from compensating changes in sources (for example, coincident terrestrial uptake and oceanic release) as predicted in models of AMOC shutdown and carbon cycle response46,47,51. However, the global impact of short stadials on the terrestrial biosphere was probably small, given that paleoproxy records indicate weaker terrestrial climate perturbations during the short Greenlandic stadials compared with the long ones30,31 as discussed above. Thus, terrestrial uptake balancing oceanic release during the short Greenlandic stadials is not likely the main explanation for the lack of CO 2 response.