Influence of tropical Atlantic meridional dipole of sea surface temperature anomalies on Antarctic autumn sea ice

To examine the influence of tropical Atlantic SST on Antarctic SIC throughout the year, we apply the MCA as a function of time lag in month for anomalies of SST and SIC. Figure 1(a) shows the SCs of the leading mode of MCA, and its ordinate is the SIC calendar month and abscissa is the time lag in month, with a negative value indicating that SST leads SIC. It displays a strong seasonal dependence, i.e. the SCs are large when SST leads May SIC by 0–2 months, and small but insignificant when SST leads SIC in January and September. This indicates a strong influence of austral autumn tropical Atlantic SST on Antarctic SIC. The SCFs and COs are consistent with SCs, as shown in figure S2.

Zoom In Zoom Out Reset image size

Figures 1(b) and (c) display spatial patterns of May SIC anomalies and March SSTAs associated with the leading MCA mode at lag −2 months (SCF, 56%), depicted by regressions of SIC and SST onto the normalized SST time coefficients. The SIC anomalies show a tri-pole structure (figure 1(b)), accompanied by the aggregated SIC increase near the Antarctic Peninsula, and decrease in the Ross Sea and east of the Weddell Sea. This structure resembles the leading mode of the EOF of Antarctic SIC in May, accounting for 20% of the total variance (figure S3(a)). The CO coefficient between the MCA1 SIC time coefficients at lag −2 months and the PC1 of May SIC EOF is 0.81, suggesting large impact of the austral autumn tropical Atlantic SST on SIC anomalies.

Figure 1(c) shows the March SSTAs regressed against the MCA1 SST index (over the Atlantic and Pacific). SSTAs in the tropical Atlantic show a meridional-dipole-like pattern (hereinafter SST-dipole), with positive values north of the equator and relatively small negative values south of the equator. The SSTAs obtained by our MCA highly resemble the Atlantic meridional mode (AMM, Amaya et al 2016, Chiang and Vimont 2004) both in spatial distribution (figure S4) and interannual variation (r = 0.77). The AMM is defined as the leading MCA mode of the SST and the zonal and meridional components of the 10 m winds over the tropical Atlantic (32° N–22° S, 74° W to the African coast; Chiang and Vimont 2004). The dynamic mechanisms of the AMM have been extensively studied, including the trade wind variations (Nobre and Shukla 1996) and wind-evaporation-SST (WES) feedback (Chang et al 1997). The AMM modulates the interannual variability of the coupled ocean-atmosphere system over the tropical Atlantic, and is often considered independent of the ENSO (Chiang and Vimont 2004). SSTAs in the tropical Pacific, on the other hand, are much weaker than in the Atlantic (recall figure 1(c)), highlighting the role of tropical Atlantic SSTAs in modulating the SIC. The CO coefficient between the normalized SST and SIC time coefficients (figure 1(d)) is 0.66, and the power spectrums (figure S5) of them both exhibit a spectral peak around five years. At lag −1 and 0 month, the SIC anomalies (figures S6(b) and (d)) show a similar distribution as in figure 1(b). The dipole-like structure become more pronounced as the negative anomalies in the tropical South Atlantic intensifies from March to May (recall figures 1(c), S6(a) and (c)).

We define an index as the difference between the area-averaged SSTA (with ENSO influence removed) in north equatorial Atlantic (NEA, 0.5° N–20.5° N, 84.5° W–16.5° E) and south equatorial Atlantic (SEA, 0.5° S–20.5° S, 84.5° W–16.5° E), which is widely adopted in previous studies (Servain 1991, Melice and Servain 2003). The NEA-minus-SEA index is highly correlated with the MCA1 SST time coefficients at lag −2 months at 0.95, well above the 99.9% confidence level, and the regression of SIC anomalies against the NEA-minus-SEA index shows a similar tri-polar SIC anomalies as the one obtained by the MCA in figure 1(b). This result is independent from the MCA, reassuring the robustness of the MCA analysis. The time coefficients defined by the pan-tropical Atlantic area-averaged SST (the combined NEA and SEA region) have been used to study the impacts of the AMO and warming trends in the tropical Atlantic on long-term changes of Antarctic sea ice, especially in austral winter and spring (Li et al 2014, Simpinks et al 2014). Note that the NEA-minus-SEA denotes the regional difference (i.e. the meridional dipole mode), whereas the pan-tropical Atlantic region denotes area-averaged SST (i.e. similar to the equatorial zonal mode). Different from the tri-pole SIC anomalies (recall figure 1(b)), regression of Antarctic SIC on the pan-tropical SST displays less significant positive SIC anomalies in the Ross Sea. Thus, what we report in this study is an additional component, i.e. the tropical Atlantic meridional dipole-Antarctica teleconnections that operates on interannual timescales.

It should be noted that lagged response of the SIC primarily reflects the persistent SSTAs rather than the occurrence of the SST forcing in advance, as pointed out by Czaja and Frankignoul (2002). To explore the persistence of ocean and atmosphere variables, we take convective precipitation as a representative of atmosphere variables, and calculate the autocorrelations of NEA-minus-SEA index and area-averaged convective precipitation over the equatorial Atlantic within 5.5° N–5.5° S, 84.5° W–16.5° E from March to August (figure S7). The SST-dipole can persist for more than four months, which explains why the influence of SSTAs on SIC in late austral autumn can be detected using SSTAs in March. Indeed, MCA results also show that SST-dipole in the tropical Atlantic in March (figure 1(c)) persists into May (figures S6(a) and (c)). For convective precipitation, however, the autocorrelation between March and May is much lower and insignificant, indicating that signals of tropical atmosphere anomalies cannot persist into the subsequent months.

The teleconnection from the tropical Atlantic to the Amundsen–Bellingshausen seas low is built upon a stationary Rossby wave mechanism (Li et al 2014, 2015a, 2015b). Here we aim to establish a dynamical link between season-dependent SST dipole and SIC anomalies on interannual timescales. We start by analyzing regressions of various tropical anomalies in May with the normalized MCA1 SST time coefficients at lag −2 months. For the regressions in the following paper, either SST or SIC index can be used, as these two time coefficients represent the synergistic anomalies of SST and SIC (Zhang et al 2018, 2021).

In response to the SST-dipole, both the tropical Atlantic and Pacific undergo a series of variations. For the Atlantic (see red box in figure 2), convective precipitation response shows a meridional-dipole-like pattern, with decreased precipitation over the equatorial Atlantic (figure 2(a), brown shading), especially on the east coast of tropical South America, and increased precipitation over 5° N–10° N (figure 2(a), green shading). This decreased precipitation is associated with the intensified low-level divergence (figure 2(a), vectors) and the corresponding downward vertical motion over the equatorial Atlantic (figure 2(b), red shading). Figure 2(c) displays the structure of upper troposphere divergence (shading) and 200 hPa divergent winds (vectors). The anomalous winds converge over the equatorial Atlantic, and diverge over South America at ∼30° S, suggesting the altered Hadley cell.

Zoom In Zoom Out Reset image size

The SST-dipole in the tropical Atlantic also modulates the convective activity over the tropical Pacific (see blue box in figure 2). There are significant positive convective precipitation anomalies over the equatorial Pacific despite much weaker SSTAs compared to the Atlantic (recall figure 1(c)). It is known that the Walker cell acts as an atmospheric bridge between the tropical Atlantic and Pacific. Here, the anomaly Walker cell is displayed as descending motion over the equatorial Atlantic and ascending motion over the equatorial Pacific (figure 2(b), shading), which forms a clockwise circulation together with the zonal wind anomalies (figure 2(b), contours). The increased precipitation over the equatorial Pacific is due to the ascending motion associated with the anomalous Walker cell in response to the SSTAs in the tropical Atlantic. The resultant anomalous divergence in the upper troposphere over the equatorial Pacific then induces anomalous convergence over the South Pacific via anomalous Hadley cell.

We suggest that anomalous warming of NEA in March promotes the development of SST dipole mode and the northward movement of convergence zone, including decreased convection over the equatorial Atlantic, and these anomalies peak in May. Such decreased convection drives enhanced convection over the equatorial Pacific. Thus, the tropical Atlantic-induced teleconnection can be expressed as two pathways: the one through perturbations of the local Hadley cell over the Atlantic and South America; and the other one that arises from the Atlantic–Pacific interaction. Both driven by the SST-dipole in the tropical Atlantic, the atmospheric responses over subtropical South America and South Pacific are opposite, which is key to generate the collaborative Rossby waves into mid- to high-latitude Southern Hemisphere (SH) as discussed in the next subsection. The atmospheric responses at lag −1 and 0 month are consistent with that at lag −2 months (not shown). Similar Walker cell response has also been found in preceding study of the long-term warming trends in the tropical Atlantic SST in austral spring (Simpkins et al 2016).

To consider the responses in subtropic regions, we first estimate the activity of the Rossby wave source (RWS). According to Sardeshmukh and Hoskins (1988), the RWS is given by $S = - {\zeta _a}D - {{{\boldsymbol{v}}}_\chi } \cdot \nabla {\zeta _a}$. The first term on the right-hand side accounts for the vortex stretching, and the second term represents the absolute vorticity gradient advection by the divergent flow. The specific mathematical procedure of RWS is introduced in the supplemental material text S2. Figure 3(a) (shading) shows the structure of anomalous RWS in response to the SST-dipole in the tropical Atlantic. These positive (negative) anomalies represent anticyclonic (cyclonic) RWS anomalies (RWSs), which drive anticyclonic (cyclonic) eddy stream function anomalies in the SH (Wang et al 2019). Further inspection shows that the total RWS is dominated by the vortex stretching (figure S8). Due to the dependence on absolute vorticity, the RWSs are small within the tropics, while RWS appears active in the midlatitude of the SH. As pointed by Sardeshmukh and Hoskins (1988), Rossby wave activity in the upper troposphere is generated by divergent flow in the presence of nonzero absolute vorticity. Specifically, positive RWSs located near east coast of Australia and subtropical South America (figure 3(a), R1 and R4) are attributed to the corresponding anomalous divergent out-flow (recall figure 2(c), D1 and D4). Similarly, the pronounced negative anomalous RWS region emerges (figure 3(a), R2 and R3) over the subtropical central and eastern South Pacific, corresponding to the upper-level anomalous convergence (recall figure 2(c), blue shading) and the anomalous divergent in-flow (recall figure 2(c), D2 and D3). The Rossby waves produced by the RWSs manifest as alternating negative and positive 200 hPa geopotential height anomalies (figure 3(a), contours).

Zoom In Zoom Out Reset image size

Next, we consider the evolution and propagation of the Rossby waves. Following Takaya and Nakamura (1997, 2001), the Rossby wave activity flux (TN flux) is used to investigate the Rossby wave propagation (specific mathematical procedure in the supplemental material text S3). This diagnostic tool illustrates the evolution and propagation of stationary Rossby wave disturbances on a zonally varying basic flow. For slowly varying basic flows, the TN flux is parallel to the local group velocity. Figure 3(b) (vectors) gives the TN flux anomalies in the SH. The RWS anomaly R3 and R4 (figure 3(a), shading) are the source of the wave train that arcs poleward to the midlatitude, and then eastward across the Antarctic Peninsula along the waveguide of the subtropical and subpolar jets (pink line in figure 3(a)). Meanwhile, there is another Rossby wave train originating over the central-western Pacific (recall R1and R2 in figure 3(a)), that propagates southward and eastward (green line in figure 3(a)). This suggests that these RWSs over subtropical South America and the central-western South Pacific are effective to trigger stationary Rossby waves. The height anomalies at 500 hPa (figure 3(b), contours) show an almost identical pattern to the one at 200 hPa (recall figure 3(a), contours), suggesting the barotropic nature of the Rossby wave trains. These two wave trains meet over the Antarctic Peninsula, and then propagate together over the Southern Ocean in the form of alternating negative and positive height anomalies. At lag −1 and 0 month, the planetary-scale atmospheric responses are similar to lag −2 months (not shown). We suggest that the superposition of teleconnection from the Atlantic and Pacific constitutes the total atmospheric responses in the SH high latitudes.

The subtropical and subpolar jets play a critical role in the teleconnection, creating a Rossby waveguide. The climatological mean of Rossby wavenumber $K$ (figure 3(c), shading) and meridional vorticity gradient ${\beta ^*}$ in May (figure S9, details of $K$ and ${\beta ^*}$ in the supplemental material Text S4) illustrates the underlying cause of the waveguide and reflecting surfaces. Regions where $K$ (figure 3(c)) tends toward zero (the boundaries of the white regions) mark reflecting surfaces, which exist because of the curvature of the flow along the polar flank of the subtropical jet and equatorward flank of the midlatitude background jet stream (figure 3(c), contours). The polar reflecting surface is a consequence of the decline in ${\beta ^*}$ (figure S9, associated with the decline in $\beta$) as one approaches the pole. The subtropical reflecting surface, however, owes its existence to the curvature of the flow (${U_{yy}}$) along the polar flank of the subtropical jet and equatorward flank of the midlatitude jet stream, which is sometimes sufficiently strong to overcome the gradient in planetary vorticity $\beta$. Regions of negative ${\beta ^*}$ are marked in dark blue, and are most prominent over the west Pacific. This reversal in the vorticity gradient forms a 'barrier' to Rossby wave propagation, keeping the wave activity trapped in the extratropic.

Figure 4(a) shows the anomalies of sea level pressure (SLP) and 10 m winds with the normalized SIC time coefficients at lag −2 months. Positive SLP anomalies are centered over the Antarctic Peninsula, accompanied with anticyclonic circulations. The anomalous southerlies induce cold air advection and offshore drifts of sea ice, contributing to the increase of SIC near the Antarctic Peninsula (figure 1(b)). In contrast, the negative SLP centers near 150° W and 10° E over 50° S–60° S induce warm air advection and onshore drifts to the Ross–Amundsen seas and Atlantic–Indian basins near Antarctica, respectively, and lead to the decrease of SIC therein (figure 1(b)).

Zoom In Zoom Out Reset image size

We then verify SAT anomalies over the SH high latitudes. There are warming indications in the Ross–Amundsen seas and east of the Weddell Sea, while an apparent cooling signal around the Antarctic Peninsula, agrees with the regression results of SIC (figure 1(b)). It has been reported that SAT datasets in reanalysis may be biased over Antarctica to some degree (Bracegirdle and Marshall 2012), so we further compare the SAT data measured by 40 stations in the Antarctic. The COs between SAT observations from 40 stations and normalized SIC time coefficients at lag −2 months are in good agreement with the reanalysis, especially in the Antarctic Peninsula (figure 4(c)). Positive COs with a high confidence level can also be seen near the Ross Sea and east Antarctica between 0° and 30°E, further supporting our analysis.