Interpreting our records in terms of atmospheric variability requires a closer inspection of the extreme values observed in the CH4 and N2O records.

The CH4 data point at ∼140 ka, previously published by , is measured in the section characterized by the widest GAD in our record. For the time interval 141–139 ka, the width of the GAD is on the order of ∼220 years, using the approach of , or ∼120 years using the approach of . At ∼140 ka, the adjacent data points are ∼160 years older and ∼170 years younger than the data point in question. Consequently, it appears unlikely that such an abrupt fluctuation in the ice core record represents an atmospheric signal. Spurious peaks could potentially result from layered bubble trapping . However, this outlier is measured in a period of otherwise stable CH4 concentrations, where neither early nor late pore closures are expected to generate an anomaly . Secondly, similarly high concentrations are not observed in our record before the 134 ka event. Such an age anomaly (between the layers enclosing gas of abnormal age and the layers enclosing gas of the expected age) would be unrealistically high compared to the characteristic age anomalies reported by for the Vostok ice core (∼200 years). Accordingly, we regard this CH4 data point as an outlier.

The large variability observed in adjacent data points during the early part of the N2O record (Fig. ) is unlikely to reflect atmospheric fluctuations given its atmospheric lifetime of 116 ± 9 years . Elevated concentrations and disproportionately high N2O variability have been observed in many instances for ice samples rich in mineral dust in both Antarctic and Greenland ice cores and are attributed to in situ production . Measurements in the Vostok ice core by demonstrated this excess N2O production to have a strong imprint in both δ15N (N2O) and δ18O (N2O) records. For some of our EDC samples we observe the same systematic isotopic deviations as for the Vostok samples of that period, with δ15N (N2O) and δ18O (N2O) values that are more enriched and depleted, respectively, than typical atmospheric values. The coupling with dust is the basis of an empirical artefact detection method applicable to EDC samples. This method considers depth intervals where dust concentrations exceed an arbitrary threshold of 300 µg kg-1 as affected by in situ production . We follow this approach and define 134.5 ka as the boundary for the section affected by artefacts. This is a slightly younger age than the last value regarded as unbiased by . Consequently, we refrain from interpreting N2O data points older than 134.5 ka as reflecting atmospheric variability.

The younger part of our records are used to study CH4 and N2O fluctuations during TII. In particular, we scrutinize our record with respect to the newly resolved variations belonging to the modes of variability presented above. To achieve this, we analyse the characteristics of the individual events as well as the background climate in which they occur.

The 129 ka event refers to the pronounced increase in CH4 and N2O concentrations to interglacial levels at the end of TII (Fig. ). For both gases, this event accounts for about half of the deglacial change in concentrations.

CH4 and N2O concentrations rose in parallel with the main resumption of the AMOC, as indicated by the evolution of the isotopic ratio of neodymium 143 and 144 (εNd) . At the same time, the evolution of the isotopic composition of speleothem calcite (δ18O (CaCO3)) indicates a northward shift of the ITCZ (Fig. ). These are fingerprints of DO-type variability. Accordingly, the abrupt CH4 rise likely reflects the response of terrestrial emissions from NH tropical wetlands. The simultaneous increase in both gases indicates that terrestrial emissions contributed, at least partly, to the N2O rise at ∼129 ka. However, because the deglacial AMOC resumption occurred at this time, it can be assumed that marine sources also played a role. A contribution from both marine and terrestrial emissions is in agreement with the findings of and for DO-type fluctuations during TI. Higher resolution measurements of the isotopic composition of N2O combined with a deconvolution, similar to the aforementioned studies, are needed to quantitatively determine the relative contribution of the sources during the 129 ka event.

The 129 ka event is preceded by a phase of rising N2O concentrations in the interval from 130.7 to 130 ka at the end of HS11 (Fig. ). This is reminiscent of the pattern of late HS-type increase, where N2O concentrations rose before the rapid DO-type Greenland temperature and CH4 increases . The timescale of the 130 ka event is in the range of the duration typically observed for these episodes (∼0.5–2 millennia) , while the N2O growth rate appears slightly larger (∼3 ppb compared to typically ∼1 ppb per century). The lack of a concomitant CH4 fluctuation suggests that only the marine source contributed to the N2O increase, in agreement with the findings of and for the late HS1 rise. The change in marine emissions may be linked to a long-term adjustment of the nitrate and oxygen inventories in the upper ocean after a period of AMOC interruption . Overall, despite the somewhat different rates of change, the 130 ka event likely constitutes an instance of a late-HS N2O fluctuation during TII. The attribution of the 130 ka feature to the pattern of late HS increase would be strengthened by additional measurements of δ15N (N2O) and δ18O (N2O), allowing the unambiguous identification of the source contributing to this event.

Turning to the 134 ka event, its occurrence within HS11 and the properties of the CH4 increase (duration and amplitude) are reminiscent of the HS-type pattern of variability evidenced by . HS-type fluctuations resolved in the WAIS Divide (WD) ice core fully developed within 90–190 years . The broader GAD of the EDC ice core implies that any such features would appear strongly dampened in our data. This is supported by continuous CH4 measurements in the Vostok ice core, demonstrating the absence of the characteristic overshoot resolved in the WD ice core during HS4 . Therefore, the sharpness and amplitude of the 134 ka event is not consistent with the picture of a substantially smoothed version of an HS-type fluctuation. The ∼70 ppb rise observed in the EDC record would translate into a WD signal exceeding by far the amplitude range of the HS1, HS2, HS4, and HS5 fluctuations (32–53 ppb) . Secondly, HS-type CH4 variability is also characterized by abrupt CO2 jumps, millennial-scale increases in δ18O (O2), enrichment in speleothem δ18O (CaCO3), and the absence of concomitant N2O variability . The simultaneous occurrence of the CH4 and N2O pulses at ∼134 ka, the depletion in speleothem δ18O (CaCO3), and the lack of any fluctuation in the δ18O (O2) record contradict these observations (Fig. ). The 134 ka event is therefore likely to have resulted from different mechanisms and does not constitute an instance of an HS-type CH4 fluctuation during TII. On the other hand, simultaneous rises of atmospheric CH4 and N2O concentrations are observed during DO-type fluctuations. During the last glacial cycle, the amplitude of DO-type CH4 increases ranges from ∼50 to 220 ppb in Greenland ice cores and elevated concentrations lasted for centuries to millennia . Accordingly, the magnitude of the 134 ka event fits well within the amplitude range observed for the well-resolved DO-type fluctuations in the EDC record of (∼50–185 ppb). Going further, results from Greenland ice cores have shown that the amplitude of DO-type CH4 rises tracks the variations in low- to mid-latitude NH summer insolation, where the most substantial changes in concentration are associated with higher insolation. This indicates that the amount or seasonal distribution of solar radiation has modulated the response of CH4 sources during interstadials . The 134 ka event occurs during a time period of low NH summer insolation (Fig. ), comparable to this at DO2. Yet, the magnitude of the 134 ka event appears larger (∼70 ppb compared to ∼50 ppb for DO2 in the EDC record).

Associating our event with a DO-type pattern of variability requires evidence for a northward shift of the ITCZ and reinvigoration of the AMOC. Synchronous with the 134 ka event, within dating uncertainty, we observe a short-lived negative excursion in speleothem δ18O (CaCO3) records (Fig. ) as well as fluctuations in proxies reflecting salinity and runoff intensity in the Bay of Bengal . These data indicate a transient strengthening of the NH tropical monsoon systems consistent with a northward shift of the ITCZ. Concerning the behaviour of the AMOC, we are not aware of studies reporting on a potential reinvigoration at this time. Nevertheless, oceanic tracers exhibit a small and short-lived fluctuation in the time interval 133–132 ka (on the timescale of ) before reaching their maximum HS11 values . On the updated chronology of the sediment core ODP Site 1063 , the negative excursion in the εNd record of coincides with a comparably small value in the data of (Fig. ). However, the revised chronology places the excursion in εNd and 231Pa/230Th substantially earlier than the 134 ka event (Fig. ). The onset of this excursion is at ∼137.4 ka, corresponding to a shift by ∼4900 years with respect to the timescale of . Since the timescale of the sediment core is tuned to the Antarctic Ice Core Chronology (AICC2012) within an estimated uncertainty of 400 years, the fluctuations resolved in the ice and marine cores should, in principle, not be regarded as synchronous. However, the alignment with the AICC2012 is performed using only one tie point between CH4 and the isotopic composition of planktonic foraminifera (at ∼128.7 ka, corresponding to the abrupt CH4 increase into the LIG) . Therefore, it may be possible that the two timescales are less tightly aligned away from this tie point. If this would be the case, it could be hypothesized that the excursions in oceanic proxies are synchronous with the fluctuations of δ18O (CaCO3), CH4, and N2O.

DO-type variability is typically also imprinted in the temperature record from Antarctic ice cores as well as in the δ18O (O2) record. However, the responses of these proxies are generally attenuated or occur on longer timescales compared to those of CH4 and δ18O (CaCO3). Accordingly, a short and small-amplitude interstadial at ∼134 ka might not have left discernible imprints in the respective records of and (Figs.  and ).

Taking the proxy evidences together, we speculate that a brief and small-scale resumption of the AMOC might have occurred within HS11, accounting for the northward ITCZ shift and the rise in atmospheric CH4 and N2O concentrations. We acknowledge that our interpretation is limited by the relatively coarse resolution of εNd and 231Pa/230Th as well as by the uncertainty arising from cross-dating sediment and ice core records. Concerning N2O, we finally note that the duration of the rise and the simultaneity with the excursion in CH4 suggest that terrestrial N2O emissions contributed by far the most to the 134 ka event .

Should our interpretation hold, the 134 ka event can be regarded as a short DO-type fluctuation. We speculate that the hypothesized AMOC reinvigoration might have been perturbed by Meltwater Pulse 2B (MWP-2B). MWP-2B represents ∼70 % of the deglacial sea level change and coincides with the 134 ka event within dating uncertainty (Fig. ) . The disruption of the oceanic circulation by freshwater forcing might have been enabled by the high susceptibility of the AMOC to perturbations at a time when the high-latitude SH was particularly cold , as was the case during HS11 (Fig. ). This situation favours the occurrence of relatively short, centennial, interstadials such as those appearing near the Last Glacial Maximum in Greenland ice cores.