1. Introduction
There are two distinct types of jets in the Northern Hemisphere tropospheric zonal circulation: eddy-driven jets and subtropical westerly jets (WJ). Although these two jets coincide in many regions, the formation of these jets is attributed to different dynamical processes. The eddy-driven jet results from the eddy-momentum flux convergence of baroclinic waves that develop in a region of enhanced baroclinicity, while the subtropical WJ is driven by angular momentum transport from the deep tropics (Lee and Kim 2003; Li and Wettstein 2012; Merz et al. 2015). The WJ normally appears in the upper troposphere at midlatitudes as a strong wind belt, featuring unique seasonal transition, and contributes prominent climatic effects (Manney and Hegglin 2018). Typically, the intense WJ is related to midlatitude wind storms and precipitation (Messori and Caballero 2015; Messori et al. 2016). The vertical shear in the jet manifests a strong thermal wind state, which tends to form synoptic-scale disturbances and therefore makes the jet a vital source for storms (Holton and Hakim 2013). WJ can also initiate anomalous weather or climate by acting as a waveguide for the propagation of large-scale Rossby waves and teleconnection patterns (Branstator 2002; Liu and Wang 2013). Moreover, changes in the intensity and position of WJ can influence regional climate patterns. For example, various studies have proven that the evolution of East Asia summer monsoon rainfall on modern or longer time scales is modulated by the position of WJ relative to the Tibetan Plateau (Nagashima et al. 2011; Kong et al. 2017; Herzschuh et al. 2019).
Primarily, the WJ is influenced by the uneven meridional distribution of solar radiation, and thus is susceptible to the change of thermal contrasts (Holton and Hakim 2013). According to the thermal wind equation, the intensity of WJ is proportional to the temperature gradient. Climatologically, the topographic features (e.g., the Tibetan Plateau and other mountains) are responsible for the regional differences in WJ (Shi et al. 2015; Sha et al. 2020). The heating effect of Tibetan Plateau largely determines the strength and the migration of WJ over Asia by regulating the tropospheric meridional temperature gradient (Li and Liu 2015; Lei et al. 2021). Over the last decades, the subtropical WJ was found to generally shift poleward in both hemispheres under global warming (Archer and Caldeira 2008; Pena-Ortiz et al. 2013). This can be attributed to the direct radiative forcing from greenhouse gases in the troposphere (Hudson 2012). The increased CO2 also strengthens the subtropical meridional temperature gradient, which accelerates the WJ via thermal wind balance (Shaw 2019).
The Last Glacial Maximum (LGM) is the latest extreme cold period, when massive ice sheets covered much of the Northern Hemisphere continents (Clark et al. 2009). Abundant geological records indicate that the Northern Hemisphere WJ is marked by significant anomalies during LGM. As revealed by grain size, flux, and provenance records of the dust over East Asia, the westerly circulation in LGM is considered to be intensified and the core of the westerlies shifts southward (Nagashima et al. 2007, 2011). Zhu et al. (2015) interpreted the lacustrine sedimentary deposits from Lake Nam Co by using a pollen discrimination index and suggested that the Tibetan Plateau (TP) is dominated by enhanced and southward shifted zonal westerly winds from 24 to 16.5 ka, which is consistent with the results concluded from the sediments in the northern TP (An et al. 2012). Furthermore, changes in lake water balance are effectively used to trace the alteration of westerly winds (Li and Zhang 2020). The increased lake level in the central Sahara during the LGM is closely linked to the vigorous activity of subtropical WJ, the path of which is displaced more southward than today (Maley 2000). Over North America, numerous records suggest that the southwestern United States in the LGM was wetter than at the present day (Oster et al. 2015; Tabor et al. 2021; Feakins et al. 2019). During cooler glacial periods, the Searles Valley in the California’s Mojave Desert (dry in modern climate) was covered by conifer forests, and the pluvial lake systems in the southwestern United States reached a high stand (Reheis et al. 2014; Peaple et al. 2022). The hydroclimate response in the western United States at the LGM largely relates to the increased and southward shifted extratropical cyclones and storm tracks (Tabor et al. 2021).
Numerical simulations for LGM Northern Hemisphere WJ changes are diverse. During the LGM, the atmospheric baroclinicity is increased and a strong cyclonic circulation is generated in the North Atlantic Ocean and Europe, and these changes all have a close relationship with the enhanced and southward jet from North America to North Atlantic (Shin et al. 2003; Ludwig et al. 2017). Based on the results from the Paleoclimate Modeling Intercomparison Project phase 2 (PMIP2) database, Yanase and Abe-Ouchi (2007) suggested that most LGM simulations indicate a southward shift of upper WJ and precipitation maximum in the midlatitudes over East Asia and the North Pacific. The stronger and southward displaced WJ over the central North Pacific also is in accordance with the increased meridional thermal contrasts there (Kitoh et al. 2001). Moreover, the PMIP2 database shows that the amplification and the southeastward shift of the precipitation and storm track in western Europe can be partly attributed to a similar displacement of the WJ (Laîné et al. 2009). In the western Mediterranean region, the simulations from PMIP3 and phase 5 of the Coupled Model Intercomparison Project (CMIP5) illustrate that the precipitation signal during the LGM winter is characterized by a southward displacement of the North Atlantic jet stream (Beghin et al. 2016). The strengthening and southward movement of the upper WJ over North America and the North Atlantic are also the characteristics of the PMIP4–CMIP6 LGM simulations (Kageyama et al. 2021). However, some studies indicate that in response to LGM boundary conditions, the WJ over the North Atlantic region shifts poleward in winter (Pausata et al. 2011; Beghin et al. 2015). Nine available models from CMIP5 database even simulate a poleward shift of the climatological mean 200-hPa jet at the whole Northern Hemisphere scale (Wang et al. 2018). Therefore, the simulated changes in Northern Hemisphere WJ during LGM still exhibit large variability.
Although the changes in WJ under LGM conditions have been intensively explored, the responses of the WJ to different individual forcing are still unclear. In particular, the ice sheet serves as a special boundary condition, contributing a great influence to shaping and modulating midlatitude circulation processes during the glacial period. However, the dynamic and thermodynamic effects of ice sheet have not been well distinguished until now. Therefore, this study analyzed the seasonal variation of Northern Hemisphere WJ during LGM through a series of sensitivity experiments, and evaluated the relative contributions and dynamic mechanisms of different forcing factors, with a special focus on ice sheet topography and albedo. The model and sensitivity experiments are introduced in section 2. Section 3 describes the main results of this study. Discussion and conclusions are provided in section 4.
2. Experiment design and model validation
a. Model and experiments
The Community Atmosphere Model version 4 (CAM4) from the National Center for Atmospheric Research (NCAR) coupled to a slab ocean model (CAM4-SOM) is employed in this study to investigate the responses of Northern Hemisphere WJ to different individual LGM external forcings. The CAM4-SOM is composed of the Community Atmosphere Model version 4 (CAM4), the Community Land Model version 4 (CLM4), the Community Ice Code (CICE), and a slab ocean model (SOM). The SOM is a simplified ocean model that enables ocean prescribed with a finite mixed layer depth varying geographically. The lack of ocean dynamics in the model is compensated by prescribing an ocean heat flux, often called Q fluxes. Generally, the Q fluxes were derived from a stable and well-equilibrated climate model with an active full-depth ocean (Mahajan et al. 2013). In this study, the Q flux is estimated from the outputs of the Community Climate System Model 4 (CCSM4) experiment for preindustrial conditions and was applied to all experiments. The SOM is computationally faster and cost-effective and also can be used to evaluate the role of the ocean dynamics by comparing with the fully coupled model with an active full-depth ocean component.
Seven experiments have been conducted in the research, including the preindustrial (PI) control experiment and other six sensitivity experiments. The sensitivity experiments are referred to as the LGM full forcing experiment (LGM), the ice sheet albedo experiment (IS_albedo), the ice sheet topography experiment (IS_topo), the land–sea configuration experiment (LSC), the Earth orbital parameters experiment (ORB), and the greenhouse gases experiment (GHG). The boundary conditions for PI and LGM experiments are in accordance with the protocols given by PMIP3 (Table 1). The concentrations of CO2, CH4, and N2O are set to 284.7 ppm, 791.6 ppb, and 275.68 ppb in PI and set to 185 ppm, 350 ppb, and 200 ppb in LGM, respectively. In LGM experiment, the orbital parameters are set to their values at 21 000 years ago and a blended ice sheet configuration by using the average of three different reconstruction products (ICE-6G v2.0, MOCA, and ANU) was applied. For detailed information see the official website (http://pmip3.lsce.ipsl.fr/).
Summary of boundary conditions for the PI and LGM experiments.
The IS_albedo, IS_topo, ORB, and GHG experiments were modified based on the LGM full-forcing experiment. In the IS_albedo experiment, only the surface type of the LGM North American and European ice sheets were replaced with the specified surface type in the PI experiment, and the remaining ice sheet over the continental shelves was replaced by bare soil. Figure 1a shows the difference of ice sheet cover between the LGM and IS_albedo experiments. In the IS_topo experiment, only the topographic heights of the North American and European ice sheets were altered to their values in PI experiment. The difference of ice sheet height between the LGM and IS_topo experiments is shown in Fig. 1b. In the ORB and GHG experiments, only Earth’s orbital parameters and concentration of greenhouse gases were set to their PI values, respectively. The LSC experiment was designed based on the PI control experiment by changing the land–sea configuration to the LGM value, the difference of land–sea configuration between the LSC and PI experiments is exhibited in Fig. 1c. The detailed experimental designs for all of the experiments are summarized in Table 2. Each simulation with a horizontal resolution of 0.9° latitude × 1.25° longitude and has 26 layers in the vertical direction. The monthly mean outputs of the last 15 years of simulations were used after a spinup period of 45 years.
The difference of (a) the ice sheet cover (%) between the LGM and IS_albedo experiments, (b) the ice sheet height (m) between the LGM and IS_topo experiments, and (c) the land fraction (%) between the LSC and PI experiments.
Citation: Journal of Climate 37, 2; 10.1175/JCLI-D-22-0869.1
Experimental design of the PI control experiment and six sensitivity experiments.
b. Model validation
The ERA-Interim monthly-mean data from 1979 to 2014 with the horizontal resolution of 0.75° × 0.75° were selected to examine the feasibility of the model. In this study, the mean climate state in boreal summer is defined as an average of June, July, and August (JJA) and that in winter is defined as an average of December, January, and February (DJF). The simulated Northern Hemisphere mean surface temperature is first examined in Fig. 2. Except for Greenland and the Arctic region, the Northern Hemisphere surface temperature is positive in JJA (Figs. 2a,c), with the maximum surface temperature existing in northwestern North Africa and the Arabian Peninsula. Compared to ERA data (Fig. 2e), the simulated surface temperature is lower in northern North Africa, the Arabian Peninsula, and the Iranian Plateau, but is higher in mid-to-high latitudes of Eurasia and most of North America. In DJF (Figs. 2b,d), strong surface cooling appears in the mid-to-high latitudes, especially over the continents of northern North America and eastern Russia. The difference between PI experiment and ERA data (Fig. 2f) indicates notable positive differences in mid-to-high latitudes of Eurasia and North America, suggesting that the simulated surface temperature in these regions is higher than that in ERA data. Overall, the surface temperature in PI experiment is overestimated in the mid-to-high latitudes of Eurasia, North America, the Arctic region, and most of the oceans, but it is underestimated in the low latitudes of North Africa and Asia. Part of this difference can be attributed to the fact that the ERA data reflect present-day conditions, while the PI experiment represents the preindustrial period. Despite these differences, the PI experiment realistically reproduces the spatial distribution and seasonal variation of the Northern Hemisphere surface temperature (Figs. 2c,d).
Averaged surface temperature (°C) over Northern Hemisphere in (a),(b) ERA data and (c),(d) the PI experiment in JJA and DJF, and (e),(f) their difference.
Citation: Journal of Climate 37, 2; 10.1175/JCLI-D-22-0869.1
Figure 3 exhibits the 200-hPa zonal wind fields in ERA data and PI experiment, in which the zonal wind speeds exceeding 20 and 35 m s−1 are drawn as arrows in JJA and DJF, respectively. In JJA, the WJ in ERA data (Fig. 3a) is weak and characterized by a discontinuous wind belt that emerges over the midlatitudes of North America, Eurasia, and the western North Pacific. The WJ in the PI experiment is similar to that in ERA in terms of spatial distribution (Fig. 3c), while the wind speed over western North Pacific is underestimated (Fig. 3e). In DJF (Figs. 3b,d), the WJ in both ERA data and PI experiment are apparently enhanced and move to the south of the Tibetan Plateau. There is a dramatic intensification in the WJ over Japan, which makes it the maximum wind speed center in the Northern Hemisphere (Fig. 3f). Compared to ERA data, the westerly winds in the PI experiment are stronger in the northern North Atlantic, but weaker in the western North America. Overall, the WJ in PI experiment is weak and located northward in summer whereas it becomes stronger and transits to the south in winter, which suggests the model can capture the distribution of WJ, as well as its seasonal transition.
Northern Hemisphere 200-hPa zonal wind fields (shaded; m s−1) in (a),(b) ERA data and (c),(d) the PI experiment in JJA and DJF, and (e),(f) their difference. In ERA data and PI experiment, the zonal winds with the velocity exceeding 20 and 35 m s−1 are drawn as arrows in JJA and DJF, respectively.
Citation: Journal of Climate 37, 2; 10.1175/JCLI-D-22-0869.1
The simulated Northern Hemisphere surface temperature in LGM is shown in Fig. 4. Compared to PI, the LGM is featured by distinct cooling in the continental mid-to-high latitudes, with the greatest cooling occurring in northern North America and Scandinavia (Figs. 4e,f). The maximum cooling is strongly related to the presence of continental ice sheets, which elevate the topography and increase the surface albedo. In addition, the polar amplification phenomenon induced by ice–albedo feedback is also responsible for the cooling in high latitudes. Based on the PMIP2 models and a range of proxy data, the estimates for the global surface cooling in the LGM are about 5°–6°C lower relative to the preindustrial period (Annan and Hargreaves 2015), while the simulated cooling in this study is larger than that in PMIP2. This may be because the cooling in the high latitudes is amplified in the SOM model, which results in increased sea ice and contributes to the larger decrease in surface temperature. Also, the cooling is calculated on the Northern Hemisphere scale. Abe-Ouchi et al. (2007) suggested that the strength of the cooling varies from different models, and the different feedbacks in different models largely determine the climate sensitivity of each model.
Averaged surface temperature (°C) over the Northern Hemisphere in (a),(b) the LGM and (c),(d) the PI experiments in JJA and DJF, and (e),(f) their difference.
Citation: Journal of Climate 37, 2; 10.1175/JCLI-D-22-0869.1
According to Fig. 3, the WJ over North America, central Asia, and Japan have distinct seasonal discrepancies; therefore, these three regions are selected as representative WJ regions for detailed analysis. Figure 5 displays the vertical profiles of zonal-mean zonal winds along North America (270°–290°E), central Asia (70°–90°E), and Japan (125°–145°E) in ERA data and the PI experiment, respectively. For JJA (Figs. 5a–i), the westerly winds are weak and prevailing in the troposphere north of 30°N, while the easterly winds occupy the upper troposphere in the tropical region. In the PI experiment, the strongest westerly wind center appears over central Asia (Fig. 5e), which is consistent with that in ERA data (Fig. 5d). From the difference between the ERA data and PI experiment, the westerly winds in 30°–40°N of Japan in JJA are underestimated (Fig. 5i). As for DJF (Figs. 5j–r), the strengthened westerly winds occupy the midlatitude troposphere and parts of the troposphere south of 30°N, accompanied by the maximum westerly winds center moving southward. In comparison, the westerly winds over Japan are the most powerful (Figs. 5p,q), which is in good agreement with the 200-hPa zonal wind fields (Figs. 3b,d). However, the intensity of the westerly winds in DJF is generally underestimated in the model, especially in North America (Fig. 5l). In general, the results from the PI simulation generally capture the seasonal transition and the regional differences of WJ suggested by observations, which further demonstrates the feasibility of the model.
Vertical profile of the zonal wind fields (m s−1) averaged for (a),(b) North America (270°–290°E), (d),(e) central Asia (70°–90°E), and (g),(h) Japan (125°–145°E) in (left) ERA and (center) the PI experiment in JJA, and (c),(f),(i) their difference. (j)–(r) As in (a)–(i), but for DJF. The gray shading represents the topography.
Citation: Journal of Climate 37, 2; 10.1175/JCLI-D-22-0869.1
3. Results
a. Responses of surface temperature to LGM forcing
First, the responses of Northern Hemisphere mean surface temperature to the full and to different individual LGM forcing in JJA and DJF are examined and shown in Figs. 6 and 7, respectively. In JJA, influenced by the full LGM forcing (Fig. 6a), the surface cooling generally occurs in the whole Northern Hemisphere, with the largest decrease in surface temperature occurring in northern North America and Scandinavia. In the IS_albedo experiment (Fig. 6b), the responses of surface temperature closely resemble that of full LGM forcing, indicating that the ice sheet albedo plays a decisive role in the cooling of Northern Hemisphere. However, the ice sheet topography results in a general warming in boreal summer and the cooling is limited in northern North America and Scandinavia where the ice sheet is located (Fig. 6c). In the GHG experiment (Fig. 6f), the significant cooling emerges in high latitudes and eastern equatorial Pacific, and the latter contributes to La Niña–like conditions. By contrast, the cooling in the LSC (Fig. 6d) and ORB experiments (Fig. 6e) is relatively weak, suggesting that the land–sea configuration and orbital parameters have limited direct influence on the total Northern Hemisphere cooling.
Responses of Northern Hemisphere mean surface temperature (°C) to (a) LGM full forcing and to different individual LGM forcing in the (b) IS_albedo, (c) IS_topo, (d) LSC, (e) ORB, and (f) GHG sensitivity experiments in JJA. The dots indicate where the differences are statistically significant at the 5% level, as calculated by a two-sided t test.
Citation: Journal of Climate 37, 2; 10.1175/JCLI-D-22-0869.1
In DJF, the Northern Hemisphere surface cooling in response to LGM full forcing is intensified (Fig. 7a). The cooling caused by greenhouse gases which contributes the most to the decrease in surface temperature (Fig. 7f). Unlike JJA, the cooling generated by ice sheet albedo is attenuated in winter (Fig. 7b), which may be due to the increased snow cover albedo at high latitudes offsetting the effect of ice sheet albedo. In the LSC experiment (Fig. 7d), the land–sea configuration is largely responsible for the decreased surface temperature in northern Scandinavia. Similar to JJA, the ice sheet topography (Fig. 7c) and orbital parameters (Fig. 7e) have less or even opposite impacts on the Northern Hemisphere surface cooling. It should be noted that the effects of different external forcing cannot be combined linearly because the atmospheric response to the forcing sources of the boundary conditions is nonlinear and there are interactions between forcing factors.
b. Responses of WJ to LGM forcing in JJA
To investigate the change of WJ in LGM, the responses of 200-hPa wind fields to full and to individual LGM forcing are analyzed. In JJA, under the influence of full LGM forcing (Fig. 8a), there is a significant zonal wave–like difference that emerges in the mid-to-high latitudes of the Northern Hemisphere. The difference pattern indicates that during the LGM summer, the westerly winds are enhanced over the midlatitudes of North America and western North Pacific but decreased over central Asia. In addition, the easterly winds over low latitudes are generally reinforced. The differences caused by ice sheet albedo (Fig. 8b) and ice sheet topography (Fig. 8c) closely resemble the response of wind fields to full LGM forcing. Compared to the IS_topo experiment, the variation of wind fields in the IS_albedo experiment is the most similar to the total zonal wave–like difference pattern, which demonstrates that the ice sheet albedo plays an important role in atmospheric circulation at mid-to-high latitudes in summer. In the ORB and GHG experiments (Figs. 8e,f), the orbital parameters and greenhouse gases exert little impact on mid-to-high latitudes but significantly strengthen the easterly winds over the tropical region. By contrast, the change of 200-hPa wind fields in response to land–sea configuration could be neglected (Fig. 8d).
Responses of Northern Hemisphere 200-hPa wind vectors (m s−1) in JJA to (a) LGM full forcing and to different individual LGM forcing in the (b) IS_albedo, (c) IS_topo, (d) LSC, (e) ORB, and (f) GHG sensitivity experiments. Only the anomalies significant at the 5% level are shown, as calculated by a two-sided t test. The LGM ice-sheet extents are outlined in red.
Citation: Journal of Climate 37, 2; 10.1175/JCLI-D-22-0869.1
Along North America (Figs. 9a,d), the westerly winds mainly appear in the troposphere north of 30°N, and the maximum westerly winds center in LGM is more southward. The difference between LGM and PI (Fig. 9g) reveals that the westerly winds over the midlatitude of North America are stronger during LGM, which also contributes to the strengthening of WJ. In central Asia and Japan, the westerly winds in LGM (Figs. 9e,f) are coherent with that in PI (Figs. 9b,c) in terms of spatial distribution. However, the westerly winds over the upper troposphere between 30° and 50°N of central Asia are weakened in LGM (Fig. 9h), which may reflect a feeble WJ as well. Over Japan, the significant positive differences appear in the troposphere between 35° and 50°N (Fig. 9i), which indicates that the westerly winds and WJ are enhanced in LGM. In the low latitudes of the three regions negative differences exist in the upper levels, suggesting that the tropical easterly winds are increased in the LGM.
The vertical profile of the zonal wind fields (m s−1) averaged for (left) North America (270°–290°E), (center) central Asia (70°–90°E), and (right) Japan (125°–145°E) in the (a)–(c) PI and (d)–(f) LGM experiments in JJA, and (g)–(i) their differences. The gray shading represents the topography, and the dots indicate the differences are statistically significant at the 5% level, which is calculated by a two-sided t test.
Citation: Journal of Climate 37, 2; 10.1175/JCLI-D-22-0869.1
From sensitivity experiments, the strengthening of midlatitude westerly winds in North America is largely attributed to the effects of ice sheet albedo and topography (Figs. 10a,b), and the change of westerly winds in the IS_albedo experiment is closer to the total difference (Fig. 10g). The land–sea configuration (Fig. 10c), orbital parameters (Fig. 10d), and greenhouse gases (Fig. 10e) have no contribution to the intensification in westerly winds over North America due to the negative differences. In central Asia (Figs. 10f–j), the attenuated upper westerly winds have the closest relationship with the effect of ice sheet albedo (Fig. 10f). Although there are negative differences in the ORB and GHG experiments (Figs. 10i,j), it suggests the great influences of orbital parameters and greenhouse gases on the enhancement of easterly winds over low latitudes. In Japan (Figs. 10k–o), the change of wind fields in response to ice sheet albedo (Fig. 10l) is most similar to the total difference (Fig. 9i). Owing to the positive differences in IS_topo (Fig. 10l) experiments, the effect of ice sheet topography slightly accounts for the increment of westerly winds. By contrast, the enhancement resulted from the land–sea configuration (Fig. 10m), orbital parameters (Fig. 10n) and greenhouse gases (Fig. 10o) are insignificant.
Responses of vertical profile of the zonal wind fields (m s−1) averaged for North America (270°–290°E) to different individual LGM forcing in the (a) IS_albedo, (b) IS_topo, (c) LSC, (d) ORB, and (e) GHG sensitivity experiments in JJA. (f)–(j) and (k)–(o) As in (a)–(e), but for central Asia (70°–90°E) and Japan (125°–145°E), respectively. The gray shading represents the topography, and the dots indicate the differences are statistically significant at the 5% level, which is calculated by a two-sided t test.
Citation: Journal of Climate 37, 2; 10.1175/JCLI-D-22-0869.1
c. Responses of WJ to LGM forcing in DJF
The responses of 200-hPa wind fields in DJF to full and to individual LGM forcing are presented in Fig. 11. In LGM (Fig. 11a), the change of wind fields over North America to North Atlantic features a northwest–southeast-oriented tripole difference pattern. The westerly winds in LGM are significantly strengthened over southeastern North America and central North Atlantic, while weakened over northern North America and southern North Atlantic. Over Eurasia and North Pacific, the midlatitude westerly winds are also apparently weakened in the LGM. In DJF, the ice sheet topography (Fig. 11c), rather than the ice sheet albedo (Fig. 11b), exerts a decisive role in modulating the westerly winds over North America, whereas the ice sheet albedo weakens the westerly winds over Asia and North Pacific. In the ORB and GHG experiments (Figs. 11e,f), the westerly winds are dramatically decreased over the midlatitudes of Eurasia and North Pacific. In comparison, the differences in the ORB and GHG experiments are stronger than that in the IS_albedo experiment, which illustrates the dominance of orbital parameters and greenhouse gases in wind fields over Eurasia and the North Pacific in winter. Due to the discrepancy of westerly winds in the LSC experiment (Fig. 11d), the land–sea configuration has a significant positive effect on westerly winds over the North Pacific.
Relative to JJA, the westerly winds in the vertical profile of North America, central Asia, and Japan are consistently intensified in DJF, with the maximum westerly winds center generally moving to the south of 40°N (Figs. 12a–f). In North America, the significant positive wind differences exhibited in the troposphere around 30°N (Fig. 12g) indicate that the midlatitude westerly winds, as well as the WJ, are enhanced during the LGM. The difference between LGM and PI in the vertical profile of central Asia (Fig. 12h) is largely coherent with that in Japan (Fig. 12i), which is expressed by negative wind differences covering most of the mid-to-high troposphere between 30° and 50°N. The negative differences coincide with decreased westerly winds in these regions, further suggesting an attenuated WJ over central Asia and Japan.
The responses of zonal-mean zonal winds in the vertical profile of North America, central Asia, and Japan to individual LGM forcing in DJF are exhibited in Fig. 13. In North America (Figs. 13a–e), the responses of zonal wind fields in the IS_topo experiment (Fig. 13b) contribute the most to the total zonal winds change in LGM. The greatest positive differences around 30°N demonstrate that the ice sheet topography plays a crucial role in the strengthening of WJ over North America. In the IS_albedo (Fig. 13a), LSC (Fig. 13c), and GHG experiments (Fig. 13e), the effects of ice sheet albedo, land–sea configuration, and greenhouse gases slightly enhance the WJ over North America. Comparatively, the orbital parameters (Fig. 13d) have less impact on the WJ over North America. With regard to central Asia (Figs. 13f–j), the significant negative differences are found in the upper troposphere around 30°N in the IS_albedo (Fig. 13f), LSC (Fig. 13h), ORB (Fig. 13i), and GHG experiments (Fig. 13j). By comparing these differences, the weakened WJ in LGM is mainly attributed to the effects of orbital parameters and greenhouse gases. As for Japan (Figs. 13k–o), the wind fields in response to ice sheet albedo (Fig. 13k), orbital parameters (Fig. 13n), and greenhouse gases (Fig. 13o) are decreased in the upper troposphere around 30°N, which contributes to the weakening of WJ over Japan. Similar to central Asia, the change in WJ over Japan is attributed to the effects of orbital parameters and greenhouse gases due to the significant negative differences.
d. Responses of 500-hPa temperature to LGM forcing in JJA
To analyze the variation of WJ, the responses of 500-hPa temperature deviation fields to full and to individual LGM forcing are investigated in Fig. 14. The deviation of each grid means the zonal-mean values removed. In response to the full LGM forcing (Fig. 14a), there is a wavelike difference pattern that emerges in the 500-hPa temperature deviation fields at mid-to-high latitudes of the Northern Hemisphere. The changes in 500-hPa temperature deviation correspond well to the response of 200-hPa wind fields, suggesting that the behaviors of westerly winds in the upper troposphere are well explained by the temperature anomalies in the midtroposphere. Due to the similar zonal wave–like responses in the IS_albedo (Fig. 14b) and IS_topo experiments (Fig. 14c), the ice sheet albedo and topography significantly influence the atmospheric temperature in the Northern Hemisphere. Over Eurasia and the North Pacific, the differences triggered by ice sheet albedo closely match the total difference. In contrast, the differences in the LSC (Fig. 14d), ORB (Fig. 14e), and GHG experiments (Fig. 14f) contribute little. The orbital parameters and greenhouse gases are expected to adjust the upper easterly winds in low latitudes through reducing the temperature deviation over Southeast Asia and western North Pacific.
Responses of Northern Hemisphere 500-hPa temperature deviation (°C) in JJA to (a) LGM full forcing and to different individual LGM forcing in the (b) IS_albedo, (c) IS_topo, (d) LSC, (e) ORB, and (f) GHG sensitivity experiments. The dots indicate the differences are statistically significant at the 5% level, which is calculated by a two-sided t test. The LGM ice-sheet extents are outlined in red.
Citation: Journal of Climate 37, 2; 10.1175/JCLI-D-22-0869.1
Since the intensity of WJ is proportional to the tropospheric meridional temperature gradient (MTG), the responses of Northern Hemisphere 500-hPa MTG to full and to different individual LGM forcing in JJA are shown in Fig. 15. Also, the responses of 500-hPa MTG in 30°–60°N of North America (250°–300°E), central Asia (60°–100°E), and Japan (120°–150°E) to full and to different individual LGM forcing in JJA are calculated and presented in Table 3. In response to the full LGM forcing, the 500-hPa MTG is generally enhanced in the midlatitudes of the Northern Hemisphere (Fig. 15a). The increased 500-hPa MTG in North America and Japan corresponds to the intensified WJ in these regions, consistent with the thermal wind balance. Through sensitivity experiments, the enhanced 500-hPa MTG in North America is mainly related to the effects of ice sheet albedo and ice sheet height (Figs. 15b,c), with values of 1.59° and 0.90°C (103 km)−1, respectively. Especially, the ice sheet albedo contributes the most to the increase of 500-hPa MTG by causing cooling over North America (Fig. 14b). In Japan, although the positive differences are observed in the IS_albedo (Fig. 15b), LSC (Fig. 15d), ORB (Fig. 15e), and GHG (Fig. 15f) experiments, the strengthening of 500-hPa MTG is primarily attributed to the effect of ice sheet albedo, leading to warming in the North Pacific (Fig. 14b). In central Asia, the increased 500-hPa MTG in LGM is also mainly controlled by the effect of ice sheet albedo (Fig. 15a), with a value of 0.47°C (103 km)−1. This enhanced 500-hPa MTG fails to explain the changes in WJ via thermal wind relationship, since the WJ in central Asia is weakened in LGM (Fig. 8h).
Responses of Northern Hemisphere 500-hPa meridional temperature gradient [°C (103 km)−1] in JJA to (a) LGM full forcing and to different individual LGM forcing in the (b) IS_albedo, (c) IS_topo, (d) LSC, (e) ORB and (f) GHG sensitivity experiments. The dots indicate the differences are statistically significant at the 5% level which is calculated by a two-sided t test. The LGM ice-sheet extents are outlined in red.
Citation: Journal of Climate 37, 2; 10.1175/JCLI-D-22-0869.1
Responses of 500-hPa meridional temperature gradient [°C (103 km)−1] in 30°–60°N of North America (250°–300°E), central Asia (60°–100°E), and Japan (120°–150°E) to LGM full forcing and to ice-sheet albedo, ice-sheet topography, land–sea configuration, orbital parameters, and greenhouse gas forcings in JJA.
e. Responses of 500-hPa temperature to LGM forcing in DJF
The influences of full and individual LGM forcing on 500-hPa temperature deviation in DJF are shown in Fig. 16. Influenced by the full LGM forcing (Fig. 16a), the temperature deviation intensely decreased in central and eastern North America, while increased to its northwest and southeast, shaping a tripole difference pattern. Moreover, the positive differences occupy most of Asia and the western North Pacific. The total difference of temperature deviation is highly correlated with the response of 200-hPa wind fields and is therefore responsible for the WJ change. In DJF, the negative differences over North America are the strongest in the IS_topo experiment (Fig. 16c). The positive differences in northwestern North America and central North Atlantic, which are caused by ice sheet topography, are also similar to the total difference. The tripole difference in the IS_topo experiment suggests that the ice sheet topography dominates the wind fields over North America and North Atlantic in winter. Rather than ice sheet topography, the positive differences over Asia and western North Pacific are mainly attributed to the ice sheet albedo (Fig. 16b), orbital parameters (Fig. 16e), and greenhouse gases (Fig. 16f). Especially in the ORB and GHG experiments, the significant meridional dipole difference over western North Pacific indicates that the orbital parameters and greenhouse gases are primarily responsible for the weakening of westerly winds over Asia and North Pacific.
Figure 17 illustrates the responses of 500-hPa MTG to full and to individual LGM forcing in DJF. As the WJ shifts southward in winter, the responses of 500-hPa MTG in 20°–50°N of North America, central Asia, and Japan to full and to different individual LGM forcing are calculated and presented in Table 4. In DJF (Fig. 17a), the 500-hPa MTG is significantly strengthened in central North America and central North Atlantic, while weakened in northern North America, the Arctic region, and the Northern Hemisphere around 30°N. Based on the thermal wind relationship, the enhanced 500-hPa MTG in North America is consistent with the stronger WJ and the reduced 500-hPa MTG in central Asia and Japan also in accordance with the weakened WJ in these regions. Through comparison, the effect of ice sheet topography takes the major responsibility for the intensified 500-hPa MTG in North America (Fig. 17c) by causing cooling over northeastern North America and warming over central North Atlantic (Fig. 16c), with a value of 1.56°C (103 km)−1. In the IS_albedo (Fig. 17b), ORB (Fig. 17e), and GHG (Fig. 17f) experiments, the weakened 500-hPa MTG appears in Asia and the North Pacific south of 40°N. The effects of ice sheet albedo, orbital parameters, and greenhouse gases significantly reduce the 500-hPa MTG in central Asia and Japan by causing warming in the midlatitudes of eastern Asia and North Pacific and cooling in its southern region (Figs. 16b,e,f). Overall, the variation of WJ is closely linked to the changes in the meridional temperature gradient at midtroposphere, which resulted from the different external forcings.
Responses of 500-hPa meridional temperature gradient [°C (103 km)−1] in 20°–50°N of North America (250°–300°E), central Asia (60°–100°E), and Japan (120°–150°E) to LGM full forcing and to ice-sheet albedo, ice-sheet topography, land–sea configuration, orbital parameters, and greenhouse gas forcings in DJF.
4. Discussion and conclusions
In this study, the influences of the different LGM external forcing on the variability of the Northern Hemisphere WJ are investigated. Our simulations suggest that the seasonal variation of Northern Hemisphere WJ during LGM is marked by distinct regional differences. Over North America, the WJ is generally intensified and displaced southward in both summer and winter. Records from the southwestern North America indicate that the southward shift of the westerlies in glacial period have brought more precipitation and caused the wet condition in this region (Kirby et al. 2013; Oster et al. 2015). Because of the southward displacement of westerlies and their enhanced persistence through the year, much of western North America was covered in large lakes and the coverage of open conifer woodland was expanded relative to today (Thompson and Anderson 2000; Fu 2023). Changes in Asia are more complicated: the WJ over central Asia is weakened and the summer and winter mean position of WJ is northward, with the northward trends being significant in June and December but very weak in the rest of the months (not shown). In contrast, the WJ over Japan is strengthened and moves southward in summer but is attenuated and slightly shifts northward in winter. The aeolian dust from the sediments of Japan Sea provides valuable information about the path and intensity variations in Asian westerlies, suggesting that the WJ is located to the south of the Tibetan Plateau throughout most of the year in cold stadials, resulting in a prolonged spring and a dustier East Asia (Nagashima et al. 2011; Chiang et al. 2015). Dong et al. (2017) further proposed that in the LGM, the WJ from Japan Sea to North Pacific moves southward by about 3°–5° in winter and spring, while the reconstruction by using the vegetation and permafrost suggested a southward migration of WJ over Japan by around 3°–5° in summer (Ono and Irino 2004), which agrees with our results. Over central Asia, the weakened WJ may dynamically coincide with the cold and dry conditions in this region (Herzschuh 2006). However, this result is not supported by the grain-size records from Lake Qinghai and Chinese Loess Plateau, since they often reflected an intensified and southward displacement of the westerlies in glacial stages (An et al. 2012; Sun 2004).
The discrepancy of the WJ over central Asia between the simulation and geological records may be because the changes in WJ indicated by the geological records are often discussed on the glacial–interglacial scale or millennial scale, whereas the variation of WJ in this study is analyzed on seasonal scale. Furthermore, the geological records usually reflect the overall movement of the WJ across the entire Asian region, while our findings demonstrate that the variation of the Asian WJ in LGM is characterized by distinct regional discrepancies. Additionally, while the upper WJ is weakened, the mid-to-low level jet in the North Pacific is intensified and shifts southward, especially in LGM winter (Yanase and Abe-Ouchi 2007; Laîné et al. 2009; Fu 2023). The variation in the low-level jet can well interpret the dust records in Asia and the North Pacific, which also manifests that most proxy records may reflect the trends in the low-level jet, as the dust and moisture transport predominantly occur in there rather than in the upper levels.
According to our sensitivity experiments, the boundary conditions from LGM are found to substantially modify the climate in the Northern Hemisphere. In terms of the surface temperature, our study suggests that the expanded continental ice sheet and reduced greenhouse gases contribute the most to the cooling in Northern Hemisphere (Figs. 6 and 7), which corresponds to previous simulations (Kim 2004; Erb et al. 2015). The thermodynamic effect of ice sheet plays a vital role in summer cooling (Fig. 6b), while the reduced greenhouse gas has potential implications for the strong cooling in winter (Fig. 7f). Broccoli and Manabe (1987) have also demonstrated that the polar amplification associated with the reduced CO2 may contribute to the larger cooling at the high latitudes in winter. Although the effects of greenhouse gases and ice sheets have major responsibility for and contribute nearly equally to global mean cooling, it is worth noting that the effects of different external forcing cannot be simply combined because the atmospheric response to different forcing is nonlinear. Erb et al. (2015) evaluate the nonlinearities in the climate response to multiple LGM forcings by using a linear reconstruction methodology and suggest the temperature nonlinearity is larger near the areas of sea ice, emphasizing potential nonlinear responses related to sea ice. However, the nonlinear responses are hard to quantify in our study given the complex interactions among individual external forcings. Therefore, the magnitude of the contributions of each forcing calculated in this study should be taken as approximate.
As for wind fields, the alterations in tropospheric westerly circulation result from the ice sheet. In summer, the increased ice sheet albedo significantly influences the westerly winds at the hemispheric scale by inducing a wavelike response of winds at midlatitudes (Fig. 8b), resulting in an enhanced jet in North America and Japan (Figs. 10a,k) and a weakened jet in central Asia (Fig. 10f). In winter, the stronger and southward shift of the WJ in North America and North Atlantic is mainly dominated by the ice sheet topography (Fig. 13b). The ice sheet albedo, orbital parameters, and greenhouse gases are the main factors for WJ in Asia (Fig. 13). By contrast, the effect of land–sea configuration on upper westerly winds is simulated relatively weaker than other forcing, although some studies found that it prominently influences the circulation and precipitation in equatorial regions by stimulating a series of ocean–atmosphere feedbacks (DiNezio et al. 2018; Cao et al. 2019). The results are consistent with other modeling studies, which indicate that as a response to LGM ice sheet topography, the WJ over North Atlantic region shifted southward in winter and the midlatitude westerly winds have substantially enhanced across North America and Europe (Hofer et al. 2012; Oster et al. 2015; Kageyama et al. 2021). According to thermal wind balance, the variation of WJ results from the different individual forcing-induced MTG anomalies at the midtroposphere. However, the response of WJ over central Asia in JJA is inconsistent with the variation in MTG, and the complex topographic features in central Asia may also exert potential influences on the variation of WJ (Shi et al. 2015; Sha et al. 2020; Shi et al. 2021).
Our work highlights both the dynamic and thermodynamics effects of the ice sheet on the atmospheric circulation during the LGM. Especially, the effect of ice sheet topography has been emphasized to play an important role in wintertime, leading to a southward shift of the jet and thereby generating precipitation anomalies in the North Atlantic and northern Europe (Kageyama and Valdes 2000; Hofer et al. 2012). During the LGM, the North American ice sheet is elevated to a height of over 3 km, and the midlatitude jet stream in North America and the North Atlantic assumes a strong, stable, and zonal disposition (Löfverström and Lora 2017), which is consistent with our results (Fig. 10c). Oster et al. (2015) pointed out that the enhanced WJ across the continent of North America is squeezed and steered by the high pressure system developing over the Laurentide ice sheet, which was intensified during the LGM. The dynamical effect of the ice sheet not only changes the spatial distribution characteristics of the WJ, but also modifies the local mean circulation by triggering the stationary wave changes (Löfverström et al. 2014). Furthermore, the upstream and downstream effects associated with the topographically induced stationary waves have also led to climate anomalies in distant regions (Justino et al. 2005; Ullman et al. 2014). Abe-Ouchi et al. (2007) deemed that the nonlocal topography effect can be more important than its local effects through the influences on the atmospheric circulation and/or cloud distribution. Although the ice sheet topography dominates the atmospheric circulation in winter, our study indicates that the ice sheet albedo plays a major role in summer. Yanase and Abe-Ouchi (2010) proved that it is the ice sheet albedo over North America that resulted in the North Pacific cyclonic anomalies in both summer and winter, and only the effect of ice sheet albedo can explain the North Pacific cyclonic anomalies in summer. The high albedo over the ice sheet cools North America and increases the equator-to-pole surface temperature gradient, which drives the strengthening and southward shift of WJ in North America (Li and Battisti 2008; Bhattacharya et al. 2017). This southward movement of the jet can further weaken the North American summer monsoon by diluting the energy fluxes required for convection (Bhattacharya et al. 2018). In addition, the effect of ice sheet albedo can amplify the atmospheric responses through its prominent diabatic cooling (Roberts et al. 2019). Overall, the ice sheet is also a source of thermal forcing on the atmosphere. The perturbations induced by the ice sheet are proportional to its thermal effect rather than its height when the thermal effect is significant (Roe and Lindzen 2001). Given the significance and complexity of ice sheets, more studies should be carried out to deeply explore the climate effect of ice sheets.
Raible and Blender (2004) have suggested that the storm tracks in the North Pacific shifts significantly northward in a mixed layer ocean model. Since the changes in storm tracks are linked to the WJ, the choice of such a simplified ocean model may influence the simulated WJ. In our study, it can be seen from Fig. 3 that, compared to ERA data, the simulated WJ in the North Pacific is northward in JJA while it is consistent with ERA data in DJF. As a simplified model, the SOM only contains the ocean thermodynamics and lacks any ocean dynamics, this disparity may result in the altered interactions with dynamics and subsequently affect the sensitivity simulations (Shell 2013). Bitz et al. (2012) found that the pattern of ocean heat uptake is possibly a factor that affects the climate sensitivity between the SOM model and fully coupled model. In addition, Watterson (2003) attributed the smaller climate sensitivity in SOM model to reduced poleward heat transport. Although the applicability of SOM to study the climate sensitivity is limited, the SOM can reflect a quick response process of WJ due to the ocean evolves slower than the atmosphere. The SOM used in this study has undergone substantial revisions relative to its earlier version and is intended to more closely approximate the results of the full-depth, ocean general circulation model (Bitz et al. 2012). Another point worth noting is that all experiments used the same Q flux, which is representative of the PI climate condition. In the LGM experiment, using PI Q flux may lead to certain biases in the simulated sea surface temperature, and further influence the evaporation and atmospheric humidity, potentially affecting the cloud formation and precipitation patterns. Although these potential biases might exist, the difference of Q flux between LGM and PI seems difficult to obviously affect our attribution analysis. Our results agree well with the previous studies in terms of the surface temperature, which allows us to further analyze the response of the WJ. Given the difference between the SOM and fully coupled model, care should be taken when using the SOM model, and similar experiments with a coupled model are needed in the future to refine the conclusions.
Acknowledgments.
We sincerely thank the three anonymous reviewers for their constructive comments and suggestions, which helped to improve the manuscript. This work is jointly supported by the Fund of Shandong Province (No. LSKJ202203300), the National Natural Science Foundation of China (41888101, 41977382), and the Strategic Priority Research Program of Chinese Academy of Sciences (XDB40030000). Z. Shi also acknowledges the support of Youth Innovation Promotion Association CAS and Natural Science Basic Research Program of Shaanxi Province (2022JC-17).
Data availability statement.
The ERA-Interim monthly-mean data are publicly available at https://cds.climate.copernicus.eu/cdsapp#!/dataset/reanalysis-era-interim?tab=form. The raw data used in the manuscript are available from https://doi.org/10.5281/zenodo.7340633.
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