III. T HE T ROPICAL R ESPONSE TO P ERIODIC E XTRATROPICAL T HERMAL
3.4 S UMMARY AND D ISCUSSION
suppressed WES feedback (labeled E10_WEF), which ignores wind speed variations in the bulk aerodynamic formulation of surface turbulent fluxes (Mahajan et al. 2011; Kang et al.
2012; Kang et al. 2014). If the WES feedback were to be a key propagation mechanism, the WES feedback suppression would increase the time lag and/or reduce the amplitude of the tropical response. However, the transient behavior of climate response to the periodic extratropical forcing depends little on the WES feedback (Fig. 3.10), suggesting that the WES feedback is not essential in the equatorward propagation of extratropical signals.
Figure 3.10. Same as (a) Fig. 2.3, (b) Fig. 3.3a, and (c) Fig. 3.5, but for the E10_WEF experiment.
Recent studies point to an important role of dynamic ocean in the quasi-equilibrium tropical response to extratropical radiative forcing. Dynamic ocean adjustments associated with Pacific subtropical cells and AMOC dampen the meridional ITCZ shift response to extratropical thermal forcing (Green and Marshall 2017; Yu et al. 2019) and global response to Arctic sea ice loss (Deser et al 2015). A more complete picture of extratropics-to-tropics teleconnection should be explored in a full hierarchy of coupled atmosphere-ocean-land models, but we believe the key propagation mechanism will operate through the surface, as revealed by our aquaplanet mixed layer ocean experiments.
Acknowledgements
We thank Doyeon Kim, Hanjun Kim, Kiwoong Park, Yen-Ting Hwang, Dargan Frierson, Malte Stuecker, and Young-Min Yang for insightful comments. Ken Takahashi provided solid and thoughtful suggestion for theoratical modeling. Comments from four reviewers, including Aaron Donohoe and Spencer Hill, greatly improved the paper. Y. S. was funded Global Ph.D. fellowship (2017H1A2A1044044) by the National Research Foundation of Korea (NRF).
Chapter 4
How does the high-latitude thermal forcing in one hemisphere affect the other hemisphere?
Abstract
Significant progress has been made in our understanding of extratropical impacts on the tropical climate via energetics framework. It is of question whether the impact of extratropical thermal forcing in one hemisphere would extend far into high-latitudes of the other hemisphere. We examine the possibility of the pole-to-pole linkage via atmospheric teleconnections by imposing a cyclic surface thermal forcing in the northern extratropics of an aquaplanet slab ocean model. We reveal a synchronous temperature response between the two poles mediated by zonal-mean atmospheric dynamics. A warming in one polar region leads to a strengthened Hadley circulation of the unforced hemisphere, fluxing more momentum toward the subtropics, thereby pulling the eddy-driven jet equatorward. A consequent anomalous descent over the polar region causes warming. The polar surface warming in the unforced hemisphere reaches 30% of that in the forced hemisphere, inferring a significance of the pole-to-pole connection.
4.1 Introduction
Regional climate change effect does not remain local but can extend globally. In particular, the polar regions are suggested to be most effective at inducing remote climate changes (Park et al. 2018). Even the two polar regions are inferred to be connected by paleoclimatic evidence. Ice cores show that cold periods of the northern high-latitudes coincided with warm periods of the southern high-latitudes and vice versa during the last glacial period on timescales of thousands of years (e.g., Blunier and Brook 2001; EPICA Community Members 2006). This asynchronous coupling of the Arctic and Antarctic temperature variations is called the bipolar seesaw. Variability in the Atlantic Meridional Overturning Circulation (AMOC) is thought to be important in formulating the bipolar seesaw (e.g., Stocker and Johnsen 2003; Pedro et al. 2018). Newly-established proxies further confirm the central role of oceanic connection in regulating the interhemispheric heat redistribution (WAIS Divide Project Members 2015).
In the meantime, paleoclimate studies on the last glacial termination invoke an atmospheric pathway for the pole-to-pole linkage (e.g., Denton et al. 2010). A prolonged cold period invariably terminated when Northern Hemisphere (NH) ice sheets were close to their greatest area and volume. The excess ice collapses into the North Atlantic, resulting in a weakening of AMOC and a southward ITCZ shift. The resultant poleward-shifted or strengthened southern westerlies enhance the upwelling of a CO2-fertilized deep-water in the Southern Ocean (e.g., Sigman and Boyle 2000), thereby increasing the atmospheric CO2
concentration, which ultimately drives a termination of the ice age (Anderson et al. 2009;
Togweiler and Lea 2010). This hypothesis is corroborated by idealized modeling studies, which demonstrate a prescribed NH cooling displaces the ITCZ southward and as a
consequence the southern mid-latitude jet shifts poleward (e.g., Lee et al. 2011; Ceppi et al.
2013). A poleward-shifted Southern Hemisphere (SH) jet may be accompanied by an increased poleward atmospheric energy transport, leading to Antarctic warming (e.g., Yim et al. 2016). Hence a collection of studies points to a plausible pole-to-pole linkage via atmospheric teleconnections.
The atmospheric connection between the two poles is more relevant for the ongoing climate change, as it operates on a much shorter—decadal and less—timescale than the millennial-scale ocean-driven bipolar seesaw. The two polar regions over the past century exhibit multi-decadal variations in surface temperatures that are seemingly connected. Over the 20th century, multi-decadal fluctuations of Arctic and Antarctic surface temperatures are anti-correlated (Chylek et al. 2010; Wang et al. 2015). However, the observed multi-decadal surface temperature fluctuations at the two poles are recently shown to be caused by distinct drivers (e.g., Deser et al. 2020; England 2021). The Arctic signal is forced by anthropogenic aerosols and greenhouse gases, whereas the Antarctic signal is driven by Pacific decadal variability. Moreover, multiple large ensembles of climate simulations point to the positively correlated climate of the two poles over the past century (England 2021), counter to an anti- correlation suggested by Chylek et al. (2010). This calls for a careful examination of atmospheric processes responsible for a pole-to-pole linkage.
Hence the primary purpose of this study is to outline the atmospheric mechanism by which the high-latitude thermal forcing in one hemisphere affects the other hemisphere. We use an aquaplanet configuration to explore the first-order zonal-mean atmospheric dynamics.
An atmospheric pathway connecting the two poles is revealed by examining the momentum balance. This study expands our understanding of polar climate change and its global impacts.
4.2 Methodology
We examine how extratropical thermal forcing confined to one hemisphere affects the other hemisphere using the Geophysical Fluid Dynamics Laboratory (GFDL) AM2 (The GFDL Global Atmospheric Model Development Team 2004) coupled to an aquaplanet slab ocean model. The heat capacity of the slab mixed-layer ocean is specified as 2 × 108 J K-1 m-2, approximately corresponding to a 50 m depth. The model is run at a resolution of 2° latitude by 2.5° longitude with 24 vertical levels.
To investigate the temporal evolution of forced response, we impose a time-varying energy flux (Qflux) into the slab ocean as in Shin et al. (2021):
j(V, k) = N sin EGH1 VG *(k − k+). (4.1)
The Heaviside function *(n) allows the Qflux to be limited to the region poleward of k+ and the amplitude of Qflux varies sinusoidally with time. We set the parameters as: k+ = 40°N, N = 10 W m-2, and # = 30 years. The 30-year period is chosen to qualitatively mimic the high-latitude multi-decadal variability such as the Atlantic Multidecadal Variability. We impose the Qflux under perpetual equinox conditions (denoted PERI) and integrate it for 330 years (i.e., 11 cycles). The periodic Qflux itself neither adds nor subtracts energy from the global system over one forcing cycle, but the global-mean surface temperature slightly decreases by 0.4 K during the initial 30-year adjustment period (Shin et al. 2021). This drift occurs because the surface temperature is more sensitive to an imposed cooling than to an imposed warming associated with nonlinear cloud radiative effect (CRE) (e.g., Shaw et al.
2015; Shin et al. 2017). Hence, the first 30-year is regarded as a spin-up period. The composite response is obtained by subtracting the time-mean climatology of PERI over 31-
with a 10-year cutoff period to extract the forced periodic response. The reference experiment, conducted with no imposed heat flux (i.e., j = 0), is integrated for 120 years. The internal variability in the reference experiment is used to examine the statistical significance of the climate response.
In order to evaluate the importance of SH extratropical SST adjustments for perturbing the southern tropics, we repeat PERI with the SSTs poleward of 20°S prescribed to the time-mean climatology over 31-330 years of PERI (denoted FSST_20S). Hence, the equatorward teleconnection from the southern extratropics is removed while allowing the SSTs equatorward of 20°S to freely evolve with an interactive slab ocean boundary. The FSST_20S is also integrated for 330 years with a 30-year spin-up and the composite response is calculated with the last 10 cycles.
Additionally, we use the identical 50 m aquaplanet slab ocean model but forced by seasonally varying solar insolation with zero eccentricity and 23.5° obliquity for conducting the same perturbed experiment as PERI (denoted PSEA). Given the inclusion of the seasonal cycle requires a longer adjustment period, PSEA is integrated for 450 years, corresponding to 15 cycles. All the main features that we discuss here are robust irrespective of the seasonal cycle (contrast Figs. 4.1,3,6 vs 4.8-9). Consistent results can be also found in a more complex model that includes realistic land-sea contrast and orography coupled to the 1.5 layer reduced gravity ocean model (Lee et al. 2011). Therefore, we mainly examine the perpetual equinox experiment, PERI, to focus on the first-order atmospheric dynamics.
4.3 Result
Figure 4.1a shows the zonal-mean sea surface temperature (SST) response to the cyclic northern extratropical thermal forcing. The SST response propagates into the tropics with a 1.67-year lag, measured by the lag time at which the correlation is maximized between the SST anomaly at the forcing edge (40°N) and that at the equator. This timescale is consistent with the previously suggested timescale of the tropical response to an extratropical forcing (e.g., Fig. 3b in Woelfle et al. 2015; Fig. 3 in Shin et al. 2021). The response timescale in the latitude-pressure domain indicates that the equatorward propagation preferentially occurs through the lower atmosphere (Fig. 4.2), hence mediated by SST response (e.g., Voigt et al.
2017). The ITCZ, defined as the precipitation centroid between 20°S and 20°N (e.g., Roberts et al., 2017), oscillates following the relatively warmer hemisphere (green line in Fig. 4.1a).
The composite ITCZ response shows a smooth oscillation (Fig. 4.1c). Since the forced response to warming and cooling is nearly symmetric, we only discuss the warming response.
The composite clearly shows that the effect of the northern extratropical forcing reaches far into the opposite hemisphere (Fig. 4.1c). However, the equatorward SST propagation is disconnected near the equator, which coincides with the latitude of climatological ITCZ. This phenomenon is reported as the "blocking effect" by the climatological ITCZ (Kang et al. 2020). As the lower branch of the Hadley cell is responsible for the equatorward SST propagation in the tropics (Shin et al. 2021), its reversed direction at the climatological ITCZ hampers further propagation. The ITCZ blocking effect is evident in Figure 4.2, which shows a sharp jump in the response timescale at the equator below 800 hPa. The tropospheric temperature responses in Figure 4.3b-d also show that the boundary layer temperature response is disconnected at the equator.
Figure 4.1. (a,b) Hovmöller diagram of zonal-mean SST response (shading; K) and ITCZ latitude (green line; °) for the last 4 cycles (i.e., 120 years) and (c,d) composite across the forcing cycle of zonal-mean SST response (shading; K), TOA clear-sky radiation response (solid-dashed contour: interval = 1.5 W m-
2), and ITCZ (green line; °) in (upper) PERI and (lower) FSST_20S. The equatorial edge of the forcing region (φ0 = 40°N) is indicated by a cyan line. In FSST_20S, the SSTs poleward of 20°S, demarcated by a magenta line, are prescribed to the 31-330 year average of PERI. Vertical gray lines denote the time slices analyzed in Figs. 2 and 3. Hatched regions denote statistically insignificant values at the 95%
confidence level based on a Student’s t-test. (e) Composite across the forcing cycle of the meridional ITCZ displacement in PERI (black) and FSST_20S (red). Faded lines indicate the ITCZ response in each forcing cycle. Thin smooth lines indicate a composite response, with thicker lines showing statistically significant values at the 95% confidence level based on a two-side Student’s t-test. The vertical lines denote the timing at which the ITCZ crosses the equator.
Figure 4.2. The lag time for which the correlation coefficient is maximized between the tropospheric zonal- mean MSE response and the imposed forcing ( as a function of latitude and height (shading; year) in PERI. The contours show the time-mean meridional stream function (clockwise circulation in solid and counter-clockwise circulation in dashed; interval = 0.5×1011 kg s-1). White areas indicate locations where the imposed forcing explains less than 25% of MSE variance (R2 < 0.25). The equatorial edge of the forcing region (40°N) is indicated by a vertical cyan line.
Given that the ITCZ acts as an effective barrier, one may expect the extratropical influence to be confined to the forced hemisphere. However, the southern high-latitudes exhibit a significant SST response, which in turn propagates into the lower latitudes (Fig.
4.1a,c). The southern extratropical SST oscillates at a period of about 30 years. The agreement of the response cycle between the two hemispheres indicates that the SH response is caused by the cyclic forcing imposed in the northern extratropics, indicative of an interhemispheric pole-to-pole teleconnection.
As the southern extratropical response propagates into the tropics, the southern tropics exhibit a warming response 11 years after the emergence of northern tropical warming (Fig. 4.1a,c). This suggests that the tropics of the unforced hemisphere is the region that
atmospheric processes (e.g., Fig. 4.2). One may suspect that either the ITCZ shift or diffusive processes across the equator could be important in forming the delayed southern tropical warming. We hence utilize FSST_20S where the response chain in the southern extratropics is disabled. Compared to PERI, the southern tropical SST response becomes completely out- of-phase with the northern tropical SST response (Fig. 4.1b,d). As the vapor-rich ITCZ is shifted toward a warmer hemisphere, water vapor amount increases in the northern tropics and decreases in the southern tropics, which respectively enhances and reduces the absorption of the top-of-atmosphere clear-sky radiation (black contour in Fig. 4.1d). Hence, the water vapor feedback associated with the ITCZ shift is responsible for the out-of-phase tropical SST response in the two hemispheres (Clark et al. 2018). The distinct tropical SST response pattern between PERI and FSST_20S implies that the equatorward propagation of the southern extratropical SST response is important for shaping the overall tropical climate response. As the lagged southern extratropical SST response is of the same sign as in the NH, the SH response chain acts to damp the ITCZ response by 10%p while shortening the response timescale by 3 years (Fig. 4.1e). This suggests the importance of an unforced hemisphere for setting the interhemispheric thermal contrast and ITCZ displacement.
To better understand how the northern extratropical forcing perturbs the southern extratropics, we show in Figure 4.3 the vertical cross-section of zonal-mean temperature response composite: 5.67, 9.42, 13.17, 16.92 and 20.67 years after the heating is prescribed poleward of 40°N, which are chosen based on the temporal evolution of the SST anomaly averaged between 3°N-5°N (vertical lines in Fig. 4.1c,d). The prescribed northern extratropical heating propagates into the tropics preferentially through the lower troposphere (Fig. 4.3a). The near-surface warming signal is trapped within the northern side of the equator, consistent with the SST response (Fig, 4.1c). The southern tropics instead exhibit a near-
surface cooling response, of the opposite sign to the northern tropical response (Fig, 4.3a), reminiscent of a previous forcing cycle (refer to Fig. 4.1c).
Figure 4.3. Vertical cross-section of zonal-mean composite responses of temperature (shading; K), zonal wind (positive in solid and negative in dashed contours; interval = 0.5 m s-1), and meridional streamfunction (clockwise in orange and counterclockwise in purple; interval = 3×1010 kg s-1) in (left) PERI and (right) FSST_20S at each time slice denoted in Fig. 4.1c-d. Hatched regions denote statistically insignificant temperature response at the 95% confidence level based on a Student’s t-test.
The free troposphere of the southern tropics, by contrast, exhibits a warming response associated with the weak temperature gradient criteria following the smallness of the Coriolis parameter in the tropics (Sobel et al. 2001). Simultaneously, a tropospheric warming response emerges in the southern high latitudes (Fig. 4.3b), which consequently warms the surface (Fig. 4.3c,d) via downward clear-sky longwave radiation (Fig. 4.4; also
Figure 4.4. (upper) Composite across the forcing cycle of temperature (shading; K) and vertical motion (rising in solid and sinking in dashed; 2 cm hr-1) averaged over Antarctic (65°S-90°S). (lower) Composite across the forcing cycle of net surface heat flux (black), net surface shortwave (cyan), net surface longwave (red), downward clear-sky longwave at the surface (red-dashed), latent heat flux (green), and sensible heat flux (blue) averaged over Antarctic (65°S-90°S). The sign convention is that positive heats the surface.
(left) PERI and (right) FSST_20S.
Figure 4.5. (a) Changes in ITCZ (°) vs changes in southern subtropical jet strength (maximum zonal wind at 200 hPa; m s-1). (b) Changes in southern subtropical jet strength (m s-1) vs changes in southern eddy- driven jet location (a latitude of zonal wind maximum at 850 hPa; °) in PERI (black) and FSST_20S (red).
The southern high-latitude warming response then propagates into the tropics through the lower troposphere (Fig. 4.3d,e), similar to the near-surface equatorward progression pathway in the forced NH. As the warming response becomes widespread in the unforced SH, the global temperature response finally exhibits the so-called mini-global warming pattern, characterized by the upper-troposphere amplified tropical warming and the near-surface amplified polar warming (Fig. 4.3d), consistent with the canonical pattern resulting from either Arctic or Antarctic sea-ice loss (e.g., Deser et al. 2015; England et al.
2020a). The mid-tropospheric warming over the south pole is also evident in FSST_20S (Fig.
4.3g), suggestive of the atmospheric origin. However, the SH warming is confined to the high-latitude troposphere as the southern extratropical SSTs are inhibited to respond in FSST_20S, obscuring the mini-global warming pattern (Fig. 4.3i). The results imply that the SST responses are critical for the equatorward propagation whether the extratropical perturbation originates from the troposphere or the surface.
Then, what causes the unforced SH polar region to warm as soon as the warming response initiated from the northern extratropics crosses the equator and reaches the southern tropical troposphere (Fig. 4.3b,g)? Under realistic boundary conditions with land-sea contrast, a Rossby wave train originating in the tropical Pacific may connect the equatorial region with the unforced high-latitudes (e.g., Li et al. 2014; Ding et al. 2014; England et al. 2020b).
However, we invoke the zonal-mean dynamics relevant to our experiments under aquaplanet configurations. As the ascending branch of the Hadley cell is shifted northward, the SH Hadley cell becomes stronger, consequently transporting more angular momentum into the subtropics and accelerating the subtropical jet (Fig. 4.3b,g; Lindzen and Hou 1988; Ceppi et al. 2013).
zonal-mean zonal wind speed at 200 hPa, are indeed strongly correlated at 0.71 (Fig. 4.5a).
Moreover, the subtropical jet strength is tightly coupled with the eddy-driven jet position (Lee and Kim, 2003). A strengthening of the subtropical jet reinforces vertical wind shears, and the enhanced baroclinicity attracts the eddy-driven jet equatorward (Fig. 4.3b-c,g-h). A stronger SH subtropical jet is associated with an equatorward eddy-driven jet location, defined as the latitude of the maximum zonal-mean zonal wind at 850 hPa (Fig. 4.5b), consistent with previous studies (e.g., Brayshaw et al. 2008; Shin et al. 2017). That is, the eddy-driven jet location is dependent on the subtropical jet strength, which in turn depends on the Hadley cell strength. The intensified Hadley cell tends to pull the eddy-driven jet equatorward while the weakened Hadley cell tends to push the eddy-driven jet poleward (Ceppi et al. 2013).