IV. H OW DOES THE HIGH - LATITUDE THERMAL FORCING IN ONE
4.3 R ESULT
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).
Figure 4.6. Vertical cross-section of zonal-mean composite responses of temperature (shading; K), transient eddy momentum convergence (solid-dashed contour; interval = 0.07 m s-1 day-1), and meridional streamfunction (clockwise in orange and counterclockwise in purple; interval = 4×108 kg s-1) in (left) PERI and (right) FSST_20S at each time slice denoted in Fig. 1c,d. Hatched regions denote statistically insignificant temperature response at the 95% confidence level based on a Student’s t-test.
Figure 4.7. (left) Composite across the forcing cycle of zonal-mean tropospheric temperature (200-800 hPa averaged) response (shading; K) and 500 hPa mean meridional streamfunction response (clockwise in solid and counterclockwise in dashed: interval = 10 Sv). (right) Composite across the forcing cycle of zonal-mean SST response (shading; K) and 200 hPa zonal-mean zonal wind response (solid-dashed contour: interval = 0.4 m s-1) in corresponding experiments. The ITCZ is indicated by a green line (°).
Hatched regions denote the statistically insignificant temperature response at the 95% confidence level based on a Student’s t-test. Only the statistically significant values are shown in contours at the 95%
confidence level based on a Student’s t-test. The results are for PERI with (a,b) 30-yr, (c,d) 10-yr, (e,f) 5- yr, (g,h) 3-yr, and (i,j) 1-yr forcing period. Note that the magnitude of heat injected into the slab ocean over a half forcing cycle for (c-j) is one-third of that for (a,b), so that the signal-to-noise ratio is larger in (a,b).
The equatorward shift of the eddy-driven jet induces anomalous eddy momentum convergence in mid-latitude and anomalous eddy momentum divergence in high-latitude (solid-dashed contour in Fig. 4.6). The meridional wind changes induced by changes in the eddy momentum flux can be inferred from the zonal momentum budget for quasi-geostrophic motions:
o1̅ ≈ /IJ/5,<,K, (5.2)
where the overbar denotes the zonal and time mean, prime denotes a deviation thereof, and p and 1 respectively are the zonal and meridional wind. The anomalous eddy momentum
divergence in the SH high-latitude would induce anomalous northerlies, forming a counterclockwise meridional circulation (orange-purple contours in Fig. 4.6b,c). The descending motion poleward of 50°S adiabatically warms the SH polar troposphere (Fig.
4.6b-c,g-i). The strongest descent occurs near 60°S but the tropospheric temperature response is most amplified at the South Pole potentially due to the Planck feedback (e.g., Pithan and Mauritsen 2014). The tropospheric temperature in the SH polar region varies synchronously with the anomalous Hadley cell (Fig. 4.3 and 4.7a), both maximizing at year = 11.92. The surface warming over the SH high latitudes enables an equatorward propagation of warming response through the SH lower troposphere (Fig. 4.6c-e). As the eddy momentum-induced SH polar warming propagates toward the tropics, a synchronous warming is evident throughout the globe (Fig. 4.3d).
Figure 4.8. Similar to (left) Figure 4.6 and (right) Figure 4.3, but for PSEA.
Figure 4.9. Similar to Figure 4.1c, but for PSEA.