Sea ice loss

This study makes use of large-ensemble atmospheric model simulations with perturbed sea ice conditions to isolate the influence of Arctic sea ice loss on the negative phase of the NAO (NAO−). It focuses on the NAO− for two key reasons. First, the climatological winter mid-tropospheric circulation response to sea ice loss in these simulations projects onto the NAO− (Supplementary Fig. 1), prompting a closer look at NAO− events specifically. Second, considering the wider literature, the one dynamical change that appears commonplace (if not ubiquitous) in response to sea ice loss is a tendency towards NAO−. The main analyses are based upon two 502-member ensembles, one with below-average sea ice cover and the other with above-average sea ice cover (see Methods for further details), hereafter referred to as the low ice (LI) and high ice (HI) ensembles. Figure 1a,b show the differences in sea ice concentration between LI and HI during early winter (November–December) and midwinter (January-–February), respectively. There are reduced sea ice concentrations, in LI compared with HI, along the sea ice edge and in the sub-polar seas. The difference patterns are largely similar between early and midwinter, except for larger sea ice reductions in Hudson Bay and the Chukchi Sea in the former, and larger reductions in the Sea of Okhotsk and Labrador in the latter. The Barents–Kara Sea is a region where sea ice reductions are understood to be especially effective at influencing the NAO18,33,36,46,56. Decreased ice cover in this region is evident in both early and midwinter.

Figure 1: Arctic sea ice loss prescribed in the model simulations. (a) Early winter (November–December) sea ice concentration differenced between the low ice (LI) and high ice (HI) simulations (LI minus HI). (b) As a but for midwinter (January–February). (c,d) As a,b but for the difference between twenty-first century (C21) and twentieth century (C20) simulations (C21 minus C20). The red and green lines mark the sea ice edge (15% contour) in LI and HI, respectively, in both a,b and in C21 and C20, respectively, in both c,d. Full size image

Figure 1c,d show differences in sea ice concentration between two additional experiments, referred to as the twenty-first century (C21) and twentieth century (C20) ensembles (see Methods). The differences in sea ice cover between C21 and C20 are more spatially extensive than between LI and HI, and the difference in sea ice area is roughly twice as large (−8.9 versus −4.8 million km2 in early winter; −10.2 versus −5.2 million km2 in midwinter). However, in some regions, the differences in sea ice concentration are larger between LI and HI than between C21 and C20 (Supplementary Fig. 2). This is especially the case near to the observed climatological sea ice edge (stemming from the fact that C20 has less sea ice than HI). The C21 and C20 ensembles will be utilized later, but first the focus is on the differences in the NAO between LI and HI.

NAO response

Unlike previous studies that have examined seasonal mean changes of the NAO in response to Arctic sea ice loss, the large ensembles used here allow in-depth analysis of strongly negative NAO events specifically. This distinction is important as changes in extreme events (in this case, low surface NAO index values) may not simply follow changes in mean climate and society is arguably more sensitive to the former. This study focuses on seasonal-scale (January–February mean) NAO− events, which are associated with prolonged periods of anomalous weather and significant impacts on society. Simulated NAO− events (defined here as when the midwinter surface NAO index is >1 s.d. below its mean; see Methods) are characterized by raised mid-tropospheric (500 hPa) geopotential heights centred over Greenland and lowered heights over the North Atlantic (Fig. 2a, contours). These anomalies are vertically coherent and extend from the surface into the stratosphere, as illustrated by a vertical cross-section along the 40° W meridian (Fig. 2b, contours).

Figure 2: Changes to NAO− events induced by Arctic sea ice loss. (a) Midwinter (January–February) 500 hPa geopotential height during NAO− events, differenced between the low ice (LI) and high ice (HI) simulations (shading; LI minus HI). (b) Midwinter geopotential height along the 40° W meridian (marked by a grey line in a) during NAO− events, differenced between LI and HI (shading). (c) Polar cap (north of 65° N) averaged geopotential height for the 4 months preceding (September–December; leftmost), during (January–February; centre) and 4 months following (March–June; rightmost) midwinter NAO− events, differenced between LI and HI (shading). Green hatching (a–c) denotes differences that are statistically significant at the 95% (P=0.05) confidence level. Black contours (a–c) show the average geopotential height for NAO− events relative to climatology (average of both LI and HI; solid for positive; dashed for negative; drawn from −200 to 200 at intervals of 20 m, excluding zero). The solid grey line in b marks the tropopause. Full size image

To estimate the influence of Arctic sea ice loss on NAO− events, the difference is taken between a composite mean of NAO− events in LI and that in HI. Mid-tropospheric height differences between NAO− events in LI and HI (Fig. 2a, shading) project strongly onto the climatological NAO− pattern (Fig. 2a, contours). NAO− events are associated with raised heights over Greenland and depressed heights over the North Atlantic in LI compared with HI. The vertical profile of height differences between NAO− events in LI and HI (Fig. 2b, shading) also closely resembles the vertical structure of the climatological NAO− (Fig. 2b, contours). These differences imply that midwinter NAO− events are amplified (intensified) by Arctic sea ice loss. This intensification can also be seen as significant (P<0.001) increase in the s.d. of the surface NAO index (1.36 hPa (95% confidence intervals: 0.86 to 1.86)) but no significant (P=0.96) change in its mean (0.02 hPa (−0.64 to 0.68); Supplementary Fig. 3). It is noteworthy that the climatological midwinter circulation response is NAO like in the mid-troposphere but not at the surface, hence no mean shift in the surface NAO index. This study focuses on midwinter (January–February) as the intensification of NAO− events is most pronounced in these months (Supplementary Fig. 4); however, other studies have found NAO responses to be maximal in late winter19,20. The timing of the NAO response may be dependent of the atmospheric model used and/or the sea ice conditions prescribed.

Troposphere–stratosphere interaction

The temporal evolution of polar cap (>65° N) height (PCH) is a commonly used metric to infer the evolution of the NAO (or AO) through time19,20,46. Figure 2c shows the evolution of PCH in the months preceding, during and following midwinter NAO− events, and how this differs between LI and HI. Typically, midwinter NAO− events are preceded by increases in stratospheric PCH in late autumn and early winter months (Fig. 2c, contours). These positive PCH anomalies descend through time and become apparent in the troposphere by midwinter. Although the stratospheric PCH anomalies persist into early spring following the midwinter NAO− event, the tropospheric anomalies dissipate. Comparing LI and HI, PCH is enhanced in the stratosphere from October to March in LI with positive tropospheric anomalies emerging a few months later in December and persisting until April. This familiar response pattern to Arctic sea ice loss19,20,46 strongly suggests a warming and weakening of the stratospheric vortex, followed by a downward propagation of circulation anomalies into the troposphere with a lag of around 1–2 months. Negative PCH differences in the stratosphere during spring may be linked to delayed final breakdown of the polar vortex, which often follows the recovery from a weakened winter vortex.

As large differences in stratospheric PCH emerge in November preceding midwinter NAO− events (Fig. 2c), attention now turns to this month and potential causes of the weakened polar stratospheric vortex. Previous work has suggested that sea ice loss increases vertical wave propagation into the stratosphere in early winter and leads to a weakened polar vortex46. Such increases in vertical wave activity are understood to relate to amplification of the climatological planetary waves, in particular the zonal wavenumber 1 component46. The concept of linear interference—how the forced response interacts with the climatological waves—appears a powerful paradigm to explain the effect of extratropical surface forcing, such as sea ice loss, on vertical wave activity59,60,61. Figure 3 shows the zonal wavenumber 1 component of the difference in geopotential height between LI and HI for Novembers preceding midwinter NAO− events. The differences display a westward tilt with altitude, indicative of vertical wave propagation62, and are tightly in phase with the climatological wavenumber 1. Thus, these simulations support the notion that the planetary wave response to Arctic sea ice loss interferes constructively with the climatological wave pattern in November and enhances vertical wave propagation into the stratosphere, consistent with ref. 46. In other months, the wave response does not project so well onto the climatological wave pattern (Fig. 3b), which suggests enhanced vertical wave propagation in November (and to a lesser extent October) triggered by sea ice loss is especially relevant for winter weather, as proposed by others33,37,46. It is worth noting that despite using a ‘low-top’ model (that is, with a model lid at 10 hPa and relatively poor vertical resolution of the stratosphere) in this study, the results are strongly consistent with those from ‘high-top’ models19,20,46. In summary, these simulations display a robust intensification of NAO− events in response to Arctic sea ice loss, through a mechanism whereby enhanced tropospheric wave activity leads to a weaker stratospheric polar vortex and precedes more intense NAO− events.

Figure 3: Planetary wave interference induced by Arctic sea ice loss. (a) Zonal wavenumber 1 component of November geopotential height averaged between 40 and 60° N during NAO− events, differenced between the low ice (LI) and high ice (HI) simulations (shading; LI minus HI). The black contours show the climatological wavenumber 1 (average of both LI and HI; solid for positive, dashed for negative; drawn from −300 to 300 at intervals of 50 m, excluding zero). Values below 700 hPa are omitted due to pressure levels intersecting elevated topography. (b) Pattern correlation (50–700 hPa) between the forced wavenumber 1 response and the climatological wave for the months preceding and during midwinter NAO− events. Full size image

NAO-related temperature response

Focus now shifts to the effects of NAO− events on near-surface (1.5 m) temperature. The NAO explains the largest percentage of midwinter temperature variance over a region covering 15° W–40° E 50–65° N (Fig. 4a), encompassing the British Isles, Belgium, the Netherlands, northern Germany and Poland, the Baltic States and southern Scandinavia, and hereafter referred to as Northern Europe. One-third (33.2%) of the simulated variance in midwinter Northern European 1.5 m temperature is explained by the NAO. Northern European temperature is 2.26 °C (1.84−2.68) colder than average during NAO− events (Table 1). This cooling is understood to be largely related to easterly wind anomalies and enhanced advection of cold continental air masses into Northern Europe. Averaged over Northern Europe, NAO− events are associated with mean westerlies of 0.61 m s−1 compared with an average of 1.98 m s−1 (Table 1). Figure 4b shows the spatial pattern of temperature anomalies during NAO− events. Cooler temperatures also occur over Siberia, East Asia and North America; however, in these regions the NAO explains a smaller fraction of the total variance than over Northern Europe.

Figure 4: Effects of the NAO on near-surface temperature. (a) Percentage of midwinter (January–February) 1.5 m temperature variance explained by the NAO (average of both low ice (LI) and high ice (HI) simulations). (b) Midwinter 1.5 m temperature during NAO− events (average of both LI and HI) relative to climatology. (c) Midwinter 1.5 m temperature during NAO− events, differenced between LI and HI (LI minus HI). The black box in a marks the Northern European domain. Green hatching (b,c) denotes differences that are statistically significant at the 95% (P=0.05) confidence level. Full size image

Table 1 Changes in Northern European winter climate. Full size table

One might expect that the temperature difference between NAO− events in LI and HI (Fig. 4c) would resemble an amplified NAO− temperature pattern (Fig. 4b), given the intensification of NAO− events by sea ice loss. However, this is not the case. Over Siberia, NAO− events are associated with warmer temperatures in LI compared with HI (Fig. 4c), rather than cooler temperatures that would be expected from more intense NAO− events. Over most of Europe, there is little change in temperature associated with NAO− events, despite the intensification of these events. Specifically for Northern Europe, there is a marginally significant (P=0.06) warming despite a highly significant (P<0.001) decrease in zonal wind (Table 1), the latter implying a more easterly flow regime typically linked to colder winter temperatures. This is called the missing cooling response, referring to the fact that midwinter Northern European temperature is unaffected by sea ice loss despite the marked intensification of NAO− events that would be expected to yield cooling. A lack of Northern European cooling is also apparent in the climatological midwinter response to sea ice loss (that is, including all midwinters not just NAO− ones; Supplementary Fig. 1).

Better understanding of the reasons for this missing cooling response can be obtained by considering the anatomy of NAO− events using simulated daily data. Figure 5a compares histograms of Northern European daily zonal wind for all midwinters and for NAO− events in LI and HI. During NAO− events (in both LI and HI), there are more frequent days of easterly zonal wind compared with climatology and, conversely, fewer days of westerly zonal flow. Comparing NAO− events in LI and HI, there are more easterly days and fewer westerly days in the former than the latter (Fig. 5a). Thus, the reduction in midwinter mean zonal wind over Northern Europe induced by sea ice loss is associated with more days of easterly flow. Comparable histograms for surface temperature (expressed as anomalies from the daily climatology in HI to remove the effects of the seasonal cycle) reveal more frequent days of below-average temperature during NAO− events (in both LI and HI) and, in particular, more frequent occurrences of cold extremes (Fig. 5b). There are notable differences in the histograms of daily temperature during NAO− events between LI and HI, despite the small change in mean temperature. There are fewer occurrences of temperature anomalies lower than −3 °C in LI compared with HI, but more occurrences of anomalies in the range −3 to 7 °C (Fig. 5b). In other words, although sea ice loss increases the number of moderate cold and warm anomalies, the largest cold anomalies decrease in number. These opposing differences result in only a small change in mean temperature.

Figure 5: Anatomy of NAO− events changed by Arctic sea ice loss. (a) Histograms of daily 10 m zonal wind averaged over Northern Europe (black box in Fig. 4a) for all midwinters (January–February; grey bars; both low ice (LI) and high ice (HI) simulations) and for NAO− events in LI and HI (red and blue lines, respectively). (b) As a, but for daily 1.5 m temperature anomalies (relative to the daily climatology in HI). (c) Relationship between daily 10 m zonal wind and 1.5 m temperature anomalies during NAO− midwinters. Each cross corresponds to a Northern European and bin average (classified by zonal wind with a bin size of 1 m s−1) in LI (red) and HI (blue). The solid lines show linear relationships, referred to in the main text with the blue line used to predict the expected temperature change (dT) due to the simulated change in zonal wind (dU; LI minus HI). (d) Dynamical (blue), thermo-dynamical (red) and net (black) contributions to the midwinter mean difference in Northern European 1.5 m temperature during NAO− events between LI and HI. The line graph shows the contributions as a function of daily 10 m zonal wind and the crosses show the total contribution. Full size image

There is a strong linear relationship between Northern European daily zonal wind and temperature during NAO− midwinters, evident in both LI and HI (Fig. 5c). As mentioned earlier, easterlies tend to be associated with colder conditions and vice versa. This linear relationship can be used to estimate the temperature change that one would expect for a given change in zonal wind: the decrease in zonal wind of −0.58 m s−1, between LI and HI, yields an anticipated cooling of 0.50 °C. This contradicts the simulated warming of 0.68 °C (−0.02 to 1.39).

Dynamical and thermodynamical effects

The temperature difference between NAO− events in LI and HI (0.68 °C (−0.02 to 1.39)) can be partitioned into contributions coming from days of differing daily zonal wind strength (Fig. 5d, black line). Further, it is possible to estimate a contribution, owing to the change in frequency of days in each wind category, assuming no change in the average temperature of days in each wind category, and a contribution owing to a change in the average temperature associated with each wind category, assuming no change in the frequency of days in each wind category. As the former describes a change of wind in the absence of a mean temperature change and the latter, a mean temperature change in the absence of changes in circulation, these are referred to as dynamical and thermodynamical components, respectively. The dynamical component (Fig. 5d, blue line) is dominated by a cooling contribution on days of zonal wind in the range −5 to 2 m s−1. These are more frequent in LI relative to HI and are associated with cold temperature anomalies; hence, they act to lower the midwinter temperature in LI relative to HI. It is noteworthy that although days of zonal wind less than −4 m s−1 are also increased, these are few in number so contribute less to the midwinter mean change. There is a smaller dynamical cooling contribution from days of zonal wind in the range 4–8 m s−1, owing to fewer of these typically warmer days in LI compared with HI. The dynamical contribution is small for days of zonal wind in the range 0–3 m s−1. In contrast, the thermodynamical contribution (Fig. 5d, red line) is largest in this range, but positive for all categories. As all wind categories are warmer in LI compared with HI (Fig. 3c), the magnitude of the thermodynamical contribution is largely dictated by the mean frequency of each category, with more frequent categories making a larger contribution to the midwinter mean temperature difference. The net contribution (Fig. 5d, black line) shows cooling (dynamically driven) on days of strong easterly flow (<−5 m s−1) and on days of strong (>5 m s−1) westerly flow, and warming (thermodynamically driven) on days of moderate (−5 to 5 m s−1) easterly and westerly flow. Summed over all days in midwinter NAO− events (that is, over all wind categories), the cooling effect of intensified NAO− is missing, owing to a larger and opposite warming effect.

Cold extremes

Figure 5b shows a reduction in the frequency of daily cold extremes. This reduction in cold extremes is caused in part by mean warming, but also by decreased daily temperature variability, consistent with previous work20,57. Reduced variability is a physical consequence of weakened horizontal temperature gradients63,64. Northern European cold extremes tend to be associated with advection of cold subpolar air from northern Eurasia, a region that is warmed by Arctic sea ice loss (Fig. 4c). It is worth noting that cold extremes decrease in frequency (Fig. 5b), despite a net cooling on days of strongest easterlies (<−5 m s−1; Fig. 5d). This can be understood by the fact that the zonal wind is only one factor of many that influences temperature, meaning that not all the coldest days are coincident with strong easterlies. Some of the coldest days fall into wind categories that are warmed by sea ice loss (for example, −5 to 5 m s−1; Fig. 5d), which explains the reduction in cold extremes.

Robustness of the response

Past work has suggested that the atmospheric response to Arctic sea ice loss can be dependent on the magnitude and spatial pattern of sea ice loss17,18,19,20,21 and the background climatic state22,23. Therefore, a pertinent question to ask is: is the missing Northern European cooling response a feature specific to these simulations or a consistent feature of the atmospheric response to Arctic sea ice loss? To begin to explore this question, the large ensemble was sub-sampled into four smaller ensembles corresponding to the four different background states (see Methods). Both the climatological NAO− response (that is, across all midwinters; there are too few NAO− events in each of the smaller ensembles to allow reliable comparison of solely NAO− midwinters) and the absence of Northern European cooling are robustly simulated in all four cases (Supplementary Figs 5 and 6), suggesting little sensitivity to the background state (this is in contrast to other aspects of the response to sea ice loss23). To further explore potential sensitivities, it is useful to attempt to reproduce the results using another set of ensemble simulations with the same model, but very different prescribed sea ice concentrations. These additional ensembles were briefly introduced earlier (C21 and C20). Recall that the difference in sea ice area between C21 and C20 is approximately twice as large as between LI and HI, and the spatial patterns of sea ice loss are very different (Fig. 1 and Supplementary Fig. 2). Despite these differences in sea ice forcing (and in the background state), there is very high consistency in the simulated atmospheric response to sea ice loss. Mid-tropospheric heights differenced between C21 and C20, show a very similar amplification of NAO− events (Fig. 6a) to that discussed previously (Fig. 2a), suggesting this is a robust feature of the atmospheric response to sea ice loss (at least in this model). Another consistent feature is the absence of European cooling associated with these more intense NAO− events (c.f. Figs 4c and 6b).

Figure 6: Replication of main results in simulations with different Arctic sea ice loss. (a) Midwinter (January–February) 500 hPa geopotential height during NAO− events, differenced between the twenty-first century (C21) and twentieth century (C20) simulations (shading; C21 minus C20). (b) As a, but for 1.5 m temperature. Green hatching (a,b) denotes differences that are statistically significant at the 95% (P=0.05) confidence level. Black contours (a) show the average geopotential height for NAO− events relative to climatology (average of both C21 and C20; solid for positive; dashed for negative; drawn from −200 to 200 at intervals of 20 m, excluding zero). (c) Histograms of daily 10 m zonal wind averaged over Northern Europe (black box in Fig. 4a) for all midwinters (January–February; grey bars; both C21 and C20) and for NAO− events in C21 and C20 (red and blue lines, respectively). (d) As c, but for daily 1.5 m temperature anomalies (relative to the daily climatology in C20). (e) Relationship between daily 10 m zonal wind and 1.5 m temperature anomalies during NAO− midwinters. Each cross corresponds to a Northern European and bin average (classified by zonal wind with a bin size of 1 m s−1) in C21 (red) and C20 (blue). The solid lines show linear relationships, referred to in the main text with the blue line used to predict the expected temperature change (dT) due to the simulated change in zonal wind (dU; C21 minus C20). (f) Dynamical (blue), thermo-dynamical (red) and net (black) contributions to the midwinter mean difference in Northern European 1.5 m temperature during NAO− events between C21 and C20. The line graph shows the contributions as a function of daily 10 m zonal wind and the crosses show the total contribution. Full size image

Histograms of Northern European daily zonal wind and temperature during NAO− events reveal very similar changes between C21 and C20 (Fig. 6c,d) to those reported earlier between LI and HI (Fig. 5a,b). Namely, an increase in days of easterly flow, an increase in days of moderate cold anomalies and a decrease in days of large cold anomalies. The difference in midwinter Northern European zonal wind between NAO− events in C21 and C20 is of comparable magnitude (−0.60 m s−1 (−0.95 to −0.26); Fig. 6e) to that between LI and HI (−0.58 m s−1 (−0.88to −0.27); Fig. 5c), despite much larger sea ice differences between C21 and C20 (Fig. 1 and Supplementary Fig. 2), emphasizing that the dynamical response does not scale linearly with the magnitude of sea ice loss15,16,17. The enhanced easterlies exert a dynamical cooling contribution but, as found before, this is offset by thermodynamical warming (Fig. 6f), leading to a small and insignificant (P=0.81) mean temperature response (0.10 °C (−0.66 to 0.87)). Despite an increase in frequency (as well as intensity) of NAO− midwinters in C21 relative to C20 (see Methods), there is no evidence of Northern European cooling in the climatological temperature response to sea ice loss (that is, across all midwinters not just NAO− ones; Supplementary Fig. 7), even though the climatological circulation response is similar to the NAO−.

One noteworthy difference between the two sets of simulations is that the slope of the wind–temperature relationship is reduced in C21 compared with C20 (Fig. 6e), whereas it stays roughly constant in HI and LI (Fig. 5c). This difference in slope is attributable to larger (sub-) polar warming in the future sea ice loss scenario and further weakened horizontal temperature gradients. In summary, the intensification of NAO− events and the missing Northern European cooling are common features of the response to Arctic sea ice loss in both sets of simulations. The absence of NAO-like surface temperature change, despite a circulation response to sea ice loss reminiscent of the NAO−, was also evident in ref. 24 (although it was not explored in any detail). However, both sets of simulations analysed in this study and the simulations analysed in ref. 24 were conducted with the same model, so model dependence cannot be ruled out.