The changed insolation in the LIG gives rise to a global annual mean warming of 0.5 K. The annual mean warming is thus relatively high compared to the recent multi-model mean estimate of 0.0 ± 0.5 K mean annual warming (Masson-Delmotte et al. 2013). Figure 2 nonetheless illustrates that the spatial pattern of the annual mean warming resembles the multi-model mean of previous LIG simulations (Lunt et al. 2013; Masson-Delmotte et al. 2013), albeit with stronger warming in the North Atlantic and stronger cooling in the tropics. The annual cycle of the insolation anomalies means that the seasonal insolation cycle is enhanced in the Northern Hemisphere (increase during summer; JJA) and reduced in the Southern Hemisphere (reduction during summer; DJF). The Northern Hemisphere annual mean warming is 0.7 K, while the Southern Hemisphere mean is 0.2 K. The Arctic region (60–90°N) experiences a substantial year-round warming with an annual mean of 2.4 K and seasonal mean warmings ranging from 1.8 K in spring (March–April–May; MAM) to 2.9 K during summer (JJA). The seasonal mean near-surface air temperature anomalies are presented in Fig. 3 alongside the zonal mean insolation changes. Both winter (DJF) and summer (MAM) anomalies resemble the Lunt et al. (2013) multi-model mean, although we simulate stronger warming at high latitudes and in the North Atlantic. Comparison of the insolation forcing and the temperature response reveals that the overall response over the continents follows the annual cycle of the insolation anomalies. Nevertheless, some continental regions stand out with temperature response that cannot be attributed to a direct warming (cooling) from increased (decreased) insolation: High northern latitudes (especially Northern Asia and Greenland) and Europe exhibit warming, and the African and Indian monsoon regions exhibit cooling throughout the year. Consequently, we will give special attention to the responses over the polar regions, the North Atlantic region, and the tropics.
The polar regions
Both polar regions experience substantial forcing from the LIG insolation changes, albeit naturally with different seasonal timing of the changes. Figure 1 reveals that the Arctic experiences a positive insolation anomaly which peaks near Northern Hemisphere summer solstice countered by a negative anomaly in the fall, i.e. an earlier onset of the polar night. The Antarctic experiences relatively similar insolation anomalies, but the insolation increase is during austral spring and the decrease during austral summer. From a cryospheric perspective, the insolation in the Arctic gives increased potential for summer melt and an earlier onset of the freezing period. In the same view, the Antarctic has decreased potential for summer melt, but an earlier onset of the melt period.
In accordance with the seasonal insolation differences, the sea ice response is varying between the two hemispheres (Fig. 4). The total Antarctic sea ice extent (i.e. the area bounded by the 15 % concentration contour) is reduced throughout the year, but exhibits only a limited decrease in late austral summer and fall. The Arctic sea ice extent exhibits a very uniform decrease throughout the year (a reduction of approximately 2 × 106 km2), but with a slightly larger decrease during late summer and early fall (July–September). The sea ice concentration in the central Arctic remains almost unchanged; the reduced extent is primarily due to a northward retreat of the sea ice edge.
From the warming over the high northern latitudes and Europe (Fig. 3), it is evident that the summertime increase overwhelms the impact of the insolation reduction during fall: Despite the reduced insolation, substantial warming is seen across high northern latitudes during fall (September–October–November; SON). The continuation of warming after the summertime insolation increase is related to the increased melt of sea ice during summer, which reduces the sea ice extent throughout the year. The loss of sea ice impacts the surface energy budget by lowering the surface albedo and reducing the insulating layer between the ocean and the atmosphere (Stroeve et al. 2012). During summer, the heat uptake by the ocean is increased following the insolation anomaly and the surface albedo feedback, which is more efficient due to the increased LIG insolation. In fall and winter, the sea ice reduction allows increased heat transfer from the ocean to the overlying atmosphere, and the heat flux is further strengthened by the anomalously warm ocean surface. The LIG experiment reveals a substantial increase of the turbulent heat flux upwards from the ocean surface during fall (SON) and winter (DJF), which peaks in the areas of sea ice loss (not shown). This combined effect of the warmer ocean and loss of sea ice causes the sustained warming through fall and winter [as found by previous studies of Arctic sea ice loss (Vihma 2014; Pedersen et al. 2016a)]. Tuenter et al. (2005) and Otto-Bliesner et al. (2013) have previously shown how increased summer insolation invokes a year-round sea ice loss that contributes to Arctic warming throughout the year.
Reduction of the sea ice thickness could also lead to increased heat flux from the ocean through reduced insulation effect (Gerdes 2006). However, while the sea ice thickness is substantially decreased in the Arctic in LIG, the heat flux over the sea ice covered areas is largely unchanged. Despite the large thinning, the sea ice thickness in the central Arctic remains about 2–4 m in LIG (4–6 m in PI) and thus still efficiently insulates the ocean from the atmosphere.
In the Antarctic, the mean temperature change follows the annual cycle of the insolation anomalies (Fig. 3). Over the continent, temperatures decrease about 1–3 K during summer (DJF), and increase by a similar magnitude (regionally more than 3 K) during spring (SON). The Southern Ocean and Antarctic coastal seas show warming in selected regions throughout the year. The warming coincides geographically with regions of sea ice loss in all seasons except summer (cf. Figs. 3, 4) and is, again, accompanied by increases in the upward turbulent heat fluxes from the surface (not shown). The sea ice related warming does, however, only appear to have a limited impact over the Antarctic continent.
The North Atlantic region
Substantial warming is evident in the North Atlantic region throughout the year. Part of this warming is related to the northward retreat of the sea ice edge (Fig. 4) described above, and increased absorption of incoming sunlight during summer. Additionally, the maximum strength of the Atlantic Meridional Overturning Circulation (AMOC) is increased in LIG compared to PI, increasing the heat transport toward the North Atlantic; especially in the colder seasons when deep convection is active. The maximum annual mean overturning strength increases by 37 % from 15.8 Sv in PI to 21.6 Sv in LIG (with standard deviations 1.03 and 1.27 Sv, respectively). This increase is in line with the increase of approximately 30 % found in previous, similar GCM simulations of 125 ka (Govin et al. 2012; Langebroek and Nisancioglu 2014).
Using the mixed layer depth as a proxy for deep water formation, Fig. 5 indicates that the main AMOC changes are related to the activation of convection in the Labrador Sea, where no convection occurs in PI, and an increased sinking in the Greenland Sea (south of Svalbard). The lack of Labrador Sea convection in PI is likely related to model biases, as marine proxy records indicate active convection since the early part of the current interglacial (Hillaire-Marcel et al. 2001; Solignac et al. 2004).
Besides the changes in the Labrador Sea, the locations of deep water formation appear unchanged, but we observe an expansion of the convection areas following the northward retreat of the sea ice edge (as indicated by the mixed layer depths). The deep water formation is active from December to April with no substantial differences in the timing between PI and LIG (the northernmost areas have deepening of the mixed layer depth starting from November and ending in May).
The largest convection anomalies occur in connection with northward sea ice retreat in the North Atlantic region. The convection and the sea ice are interconnected, as the sea ice cover inhibits convection by insulating the ocean from the atmosphere, while the convection strength affects sea ice growth (retreat) through regional cooling (warming). Unfortunately, our experiments do not allow determination of the causality of the changes in the Labrador Sea region, i.e. whether the expanded sea ice cover in PI is the cause or the result of the convection shutdown.
We observe a general increase of the sea surface salinity (SSS) in the LIG experiment that is also consistent with the increased AMOC strength. The SSS affects the convection, which itself drives the circulation that brings saline water into the region (Stommel 1961; Kuhlbrodt et al. 2007). Hence, the increased AMOC would also favor increased salinity in the North Atlantic region, making it difficult to assess the causality pattern. One factor affecting SSS in the North Atlantic is the sea ice export southward from the Arctic Ocean (SSS and sea ice drift anomalies are shown in Fig. 6). In the LIG simulation the ice edge in the North Atlantic has retreated northward, and the changed sea ice drift indicates that the southward sea ice export through the Fram Strait and along the Greenland east coast is substantially decreased especially in summer and fall (as indicated by the northward drift anomaly vectors along the Greenland east coast in Fig. 6). The decreased sea ice export coincides with an increased SSS, e.g. along the Greenland east coast, in Baffin Bay and in the Labrador Sea. Sea ice export through the Fram Strait constitutes a substantial freshwater export southward from the Arctic Ocean (Serreze et al. 2006), and previous studies have illustrated that the sea ice export affects the salinity in the North Atlantic, the Labrador Sea convection and AMOC strength (Born et al. 2010; Govin et al. 2012).
The tropics (30°S–30°N) experience an annual mean insolation reduction, but exhibit an almost unchanged annual mean temperature (0.02 K increase). Some tropical regions exhibit substantial cooling even in seasons without reduced insolation (Fig. 3), namely the sub-Saharan/Sahel region in Africa and, to lesser extent, India and parts of southeast Asia. The cooling effect is due to changes in cloudiness, soil moisture, and precipitation related to the monsoonal systems.
The North African monsoon has previously been shown to be sensitive to insolation changes (de Noblet et al. 1996; Braconnot et al. 2008; Govin et al. 2014; Bosmans et al. 2015). Specifically, the strength of the summer monsoon is increasing with anomalous high northern hemisphere summertime insolation (e.g. during LIG). The proposed mechanism (Braconnot et al. 2008; Bosmans et al. 2015) is that warming of the continents (by insolation during summer) increases ocean–land thermal and pressure gradients. The increased gradients and strengthened thermal low systems over land drive increased winds and moisture transport from the ocean to the continent. The precipitation increase is set up by a combination of increased moisture transport and local recycling (from evaporation); the latter being a minor contribution. The consequence is that the continental temperatures are expected to decrease due to increased cloud cover and evaporation (de Noblet et al. 1996; Montoya et al. 2000; Bosmans et al. 2015). Bosmans et al. (2015) investigate the links between insolation changes and the North African monsoon using the previous version of the EC-Earth model. Looking at the summer (JJA) means, the authors conclude that both low precession (summer solstice near perihelion) and high obliquity strengthens the monsoon by inducing a low pressure anomaly over Northern Africa which increases winds and moisture transport from the tropical Atlantic.
Consistent with the previous studies, the summertime (JJA) precipitation (Fig. 7) is increased substantially over the Indian and North African monsoon regions in the LIG. The precipitation increase coincides with increased cloud cover (not shown) and cooling (Fig. 3), consistent with an increased strength of the summer monsoonal systems. The total cloud fraction is increased by more than 0.2 in a broad band over the African continent covering approximately 5–25°N, and the cloud increase is seen throughout the atmospheric column (i.e. both high, mid, and low clouds are increased). The atmospheric circulation anomaly (not shown) is very similar to the response outlined by Bosmans et al. (2015): Sea level pressure is substantially decreased over Northern Africa, and the anomalous circulation increases the flow from the tropical Atlantic towards the interior continent south of Sahara.
The circulation anomaly is only evident during summer (JJA), while the precipitation increase remains through the fall (SON). Some of the changes related to the summer monsoon even appear to persist throughout the year: near-surface cooling, increased surface evaporation, and increased cloud cover are dominant in the region all year, albeit in varying meridional extents. The increased cloud cover means that the down welling shortwave radiation at the surface is decreased throughout the year, regardless of the insolation anomaly (not shown). Thus, the monsoonal changes affect the hydrological cycle in the region and impact the climate the entire year. This is illustrated by the turbulent fluxes from the surface in the region: Fig. 8 shows the annual cycle of latent and sensible heat flux in the North African monsoonal region (here defined as 5°S–25°N and 20°W–40°E). The latent heat flux is increased gradually during the summer monsoon season, and remains higher than PI through the fall. Evidently, the region is in a new regime where increased soil moisture allows more efficient cooling of the surface through evaporation, as illustrated by the increased latent heat flux. The colder surface consequently leads to a reduced sensible heat flux throughout the year. The wetter surface conditions combined with the increased cloud cover negates the impact of the summertime insolation increase on the surface.
Separation of contributions
The series of AGCM simulations is designed to investigate the mechanisms behind the simulated changes, and compare the direct and indirect effects of the insolation anomalies. Figure 9 displays the seasonal mean near-surface air temperature anomalies relative to iP + oP (PI conditions) in the three simulations (Table 2): iL + oL (LIG conditions), iL + oP (LIG insolation, PI SST and sea ice), iP + oL (PI insolation, LIG SST and sea ice), and iP + oP-ice (PI insolation, PI SST and LIG sea ice). As desired, the temperature anomaly in iL + oL closely resembles the anomaly from the coupled experiment (compare with Fig. 3) over oceans as well as continents. The response in iL + oP is limited to the continents (and sea ice covered areas), as the near-surface air temperature over the ocean is largely determined by the prescribed SSTs. Conversely, the iP + oL experiment reveals that the changed oceanic conditions have impacts across all continents even with unchanged insolation. The iP + oP-ice experiment reveals that the sea ice reduction can explain a substantial part of the Arctic and Southern Ocean wintertime warming, even with unchanged SSTs.
The continental warming during the insolation maximum in Northern Hemisphere summer (JJA) is dominated by the direct impact of the insolation. The oceanic changes do, however, contribute to temperature increase over high northern latitudes and over Europe. The SST and sea ice changes appear to dominate the response over the same regions during fall and winter, where widespread warming occurs despite the lower insolation. Part of this all-year warming in the high northern latitudes, especially in the North Atlantic region including Greenland and Europe, can be ascribed to the AMOC increase and a seasonal memory of sea ice retreat [as described by Otto-Bliesner et al. (2013)]. In these regions, the oceanic changes more than outweigh the direct impact of the fall (SON) insolation decrease. The isolated impact of the sea ice loss (in iP + oP-ice) is largely limited to the areas of sea ice retreat. One exception occurs during fall and winter where a more widespread Arctic warming is simulated. As the ice thickness is fixed in these simulations and no substantial concentration changes occur in the Arctic Ocean (Fig. 4), the widespread warming indicates that warming is advected from the ice loss regions. This is also manifested in the warming over Greenland. During winter (DJF) iP + oP-ice indicates that a substantial contribution to Greenland warming is related directly to the sea ice retreat (a detailed analysis of the Greenland response presented in Pedersen et al. 2016b). During summer (JJA in the Arctic, DJF in the Antarctic) there is no temperature response to the sea ice loss in iP + oP-ice. This is a direct consequence of the experiment design. In summer, the most dominant impact of sea ice loss is an albedo feedback that favors increased absorption of solar radiation that drives accelerated melt. As the SST is fixed, this effect is not captured in this setup.
In the Antarctic, reduced insolation dominates austral summer (DJF) temperatures that decrease over the entire continent. The warming from the oceanic changes spreads over most of the continent during both winter (JJA) and spring (SON). The strongest impact is during winter, with warming of more than 1 K over the majority of East Antarctica and about 0.5 K over West Antarctica. Increased spring (SON) insolation causes warming over the entire continent, while oceanic changes primarily impact the near-coastal regions in the vicinity of the sea ice loss in the Ross Sea (150°E–90°W) and east of the Weddell Sea (20°W–60°E). The resemblance of iP + oL and iP + oP-ice (Fig. 9) illustrates that the sea ice loss is the dominant cause of the warming near Antarctica.
Figure 9 shows high resemblance between the temperature response over the tropical region in iL + oL and iL + oP (insolation only). The contribution from the oceanic changes in iP + oL is warming Northern Africa through most of the year (except MAM), while also contributing to cooling over the Sahel region (except DJF). During SON the oceanic changes in iP + oL lead to significant warming over all continents, except Northern Australia and, again, the Sahel region in Africa.
The precipitation changes reveal that both insolation and the related oceanic changes cause substantial changes in tropical precipitation patterns. Figure 10 shows the precipitation changes in JJA, where the seasonal anomalies over the continents are largest. As for the temperature anomalies, the precipitation anomalies in iL + oL resemble the response in the coupled simulation. The hybrid simulations reveal that while the affected areas are similar in iL + oP and iP + oL, the changes show large contrasts. In southeastern Asia, iL + oP exhibits wetter conditions over the continent at the expense of the near-coastal waters in the Indian and Pacific Oceans. The iP + oL simulation exhibits precipitation increase over the same ocean regions (especially in the Indian Ocean) while no substantial changes are seen over the continent. Similar results are found in this region by previous studies of the Mid-Holocene, where the oceanic changes were even found to limit the continental precipitation increase driven by insolation changes (Liu et al. 2004; Braconnot et al. 2007). Substantial evaporation changes over the Indian Ocean (not shown) contribute to the varying precipitation changes between iL + oP and iP + oL: The insolation change in iL + oP reduces the JJA mean evaporation in the region, while it is increased in iP + oL coincident with the higher SSTs.
On both the Pacific and Atlantic sides of Central America, contrasting changes are seen in iL + oP and iP + oL: The insolation changes cause decreased rainfall, while the oceanic changes alone cause a substantial increase. The result from iL + oL indicates that the combined response is close to the sum of the two, with only a slight increase on the Pacific side. In the Atlantic, similar contrasting changes are evident off the North American coast (co-located with the North Atlantic current), where the drier conditions from iL + oP dominate in iL + oL. In the tropical Atlantic both iL + oP and iP + oL show a belt of drier conditions from South America to Africa, but iL + oP also has a substantial precipitation increase off the African coast.
The strengthening of the Northern African monsoon discussed previously appears to be related directly to the insolation changes in iL + oP rather than the oceanic changes in iP + oL. Nevertheless, the oceanic changes in iP + oL do seem to contribute to the increased precipitation, albeit in smaller extent and magnitude. This is in line with the mechanism described above (suggested by Braconnot et al. 2008; Bosmans et al. 2015). Figure 11 presents the circulation changes in terms of mean sea level pressure and 10 m winds, and reveals high resemblance between the full response in iL + oL and iL + oP. As described above, the insolation changes set up a low pressure anomaly over Northern Africa, which increases the flow and thus the moisture transport from the tropical Atlantic across the continent. These experiments indicate that the changes in Atlantic SSTs do not play a major role in shaping the monsoon response.