Abstract

Using observational data and outputs from an ocean general circulation model, the growth and decay of the South Atlantic subtropical dipole (SASD) are studied. The SASD is the most dominant mode of interannual variability in the South Atlantic Ocean, and its sea surface temperature (SST) anomaly shows a dipole pattern that is oriented in the northeast–southwest direction. The positive (negative) pole develops because the warming of the mixed layer by the contribution from the climatological shortwave radiation is enhanced (suppressed) when the mixed layer is thinner (thicker) than normal. The mixed layer depth anomaly over the positive (negative) pole is due to the suppressed (enhanced) latent heat flux loss associated with the southward migration and strengthening of the subtropical high. During the decay phase, since the temperature difference between the mixed layer and the entrained water becomes anomalously large (small) as a result of the positive (negative) mixed layer temperature anomaly, the cooling of the mixed layer by the entrainment is enhanced (reduced). In addition, the cooling of the mixed layer by the contribution from the climatological latent heat flux is enhanced (suppressed) by the same thinner (thicker) mixed layer. This paper demonstrates the importance of taking into account the interannual variations of the mixed layer depth in discussing the growth and decay of SST anomalies associated with the SASD.

1. Introduction

Southern African rainfall reaches its peak during austral summer owing to the southward migration of the south Indian convergence zone (Cook 2000). Since it affects the rain-fed agriculture and the public health in southern Africa, its interannual variations are very important in the regional society. Several studies discussed its link with El Niño–Southern Oscillation (Lindesay 1988; Richard et al. 2000) and the interannual sea surface temperature (SST) variations in the southern Indian Ocean (Reason and Mulenga 1999; Behera and Yamagata 2001; Reason 2001), and some other studies focused on impacts of interannual SST variations in the South Atlantic Ocean (Walker 1990; Mason 1995; Reason 1998).

Following Richard et al. (2000), we introduce the southern African rainfall index (SARI) defined as the area-averaged land rainfall anomalies south of 10°S. Interestingly, the correlation between the SARI and SST anomalies during 1960–2008 in the South Atlantic shows a remarkable dipole pattern (Fig. 1); the SST anomalies in the southwestern (northeastern) region are positively (negatively) correlated with the SARI. Among studies devoted to the understanding of interannual SST variations in the South Atlantic, Venegas et al. (1996) was the first to discuss this SST dipole. After the similar phenomenon in the southern Indian Ocean named the Indian Ocean subtropical dipole (IOSD; Behera et al. 2000; Behera and Yamagata 2001; Suzuki et al. 2004), it is called the South Atlantic subtropical dipole (SASD). Venegas et al. (1997) showed that the atmosphere leads the ocean by 1–4 months by calculating a lag correlation between the principal components of the sea level pressure (SLP) and SST anomalies. Furthermore, Fauchereau et al. (2003) suggested that the SST anomalies are related to the latent heat flux anomalies associated with the southward migration as well as strengthening of the subtropical high. This variation in the subtropical high was suggested to have a link with the atmospheric pattern of zonal wavenumber 3 or 4 in the Southern Hemisphere. Only a few studies discussed the oceanic roles in the generation of the SASD, however. In this regard, Sterl and Hazeleger (2003) examined the mixed layer heat balance and showed that the anomalous Ekman heat transport contributes to the growth of SST anomalies even though the latent heat flux anomaly plays the dominant role. Because of the lack of ocean data, they used the climatological mean mixed layer depth derived from an ocean reanalysis in the heat balance calculation. Also, Hermes and Reason (2005) showed that the Ekman upwelling anomaly affects the evolution of SST anomalies using outputs from an ocean general circulation model (OGCM). Since the mixed layer depth undergoes significant seasonal and interannual variations in the subtropics, those studies that do not take the variations into account have serious flaws as we show later. In fact, Morioka et al. (2010) have recently discussed the importance of the interannual variations in the mixed layer depth on the growth of SST anomalies associated with the IOSD.

Fig. 1.

Correlation between the SARI and SST anomalies during 1960–2008. The contour interval is 0.1, and the values exceeding 90% confidence level in a two-tailed t test are shaded. The box over the southwestern (northeastern) pole is defined between 10° and 30°W and between 30° and 40°S (0°–20°W and 15°–25°S).

Fig. 1.

Correlation between the SARI and SST anomalies during 1960–2008. The contour interval is 0.1, and the values exceeding 90% confidence level in a two-tailed t test are shaded. The box over the southwestern (northeastern) pole is defined between 10° and 30°W and between 30° and 40°S (0°–20°W and 15°–25°S).

This paper investigates the mechanism of the SASD using outputs from an OGCM. It is organized as follows. A brief description of the observational data and an OGCM is given in the next section. In section 3, we discuss the annual cycle of the SST and the mixed layer depth in the subtropical South Atlantic. In section 4, we define the SASD events by introducing an SASD index (SASDI) and examine the growth and decay mechanisms of the SST anomalies associated with the SASD. The cause of the interannual variations in the mixed layer depth is also discussed by calculating the Monin–Obukhov depth. The final section summarizes the main results.

2. Observational data and OGCM design

We use the monthly mean observed SST data from the Hadley Centre sea ice and sea surface temperature (HadISST; Rayner et al. 2003). They are gridded data with 1° × 1° resolution, and we analyze the period of 1960–2008 because there were few observations in the southern part of the South Atlantic before the 1960s. Monthly SST anomalies are calculated by subtracting the monthly mean climatology after removing a linear trend using a least squares fit. To calculate the mixed layer depth, we use the monthly climatology of ocean temperature from the World Ocean Atlas 2009 (WOA09; http://www.nodc.noaa.gov). It has 24 levels in the vertical with 1° × 1° horizontal resolution. We also use the surface heat flux, specific humidity at 2-m height, and wind speed at 10 m above the surface from the National Centers for Environmental Prediction–National Center for Atmospheric Research (NCEP–NCAR) reanalysis dataset (Kalnay et al. 1996). It covers the same period on a T62 Gaussian grid. Since qualitatively similar results are obtained using the European Centre for Medium Range Weather Forecasts reanalysis dataset (Uppala et al. 2005), we only show results from the NCEP–NCAR reanalysis dataset in this paper. For precipitation, we use the gridded monthly rainfall data provided by the University of Delaware (Legates and Willmott 1990) from 1960 to 2008 with a horizontal resolution of 0.5° in both longitude and latitude.

To calculate the mixed layer heat balance, we use outputs from an OGCM. The ocean model is based on the Modular Ocean Model, version 3.0 (MOM 3.0), developed at the National Oceanic and Atmospheric Administration/Geophysical Fluid Dynamics Laboratory (Pacanowski and Griffies 1999) and covers the global ocean from 65°S to 30°N. It has a horizontal resolution of 0.5° in both longitude and latitude and 25 levels in the vertical with eight levels in the upper 100 m. The bottom topography and coastlines adopted in this model are based on the 5-min earth topography (ETOPO5) dataset. The vertical eddy viscosity and diffusivity are calculated using the parameterization of Pacanowski and Philander (1981), while the lateral eddy viscosity and diffusivity are based on the formula given by Smagorinsky (1963). Near the southern and northern boundaries (poleward of 62°S and 27°N), values of these coefficients are increased steadily so that the damping time scale reaches one day at 65°S and 30°N. The temperature and salinity are relaxed to the monthly mean climatology (Levitus and Boyer 1994; Levitus et al. 1994) within the sponge layer so that artificial wall effects are reduced. All measures described above are conventional.

The model is first spun up for 20 yr by the monthly mean climatology of the wind stress from the NCEP–NCAR reanalysis dataset and by the surface heat flux calculated by the bulk formula using the simulated SST and the atmospheric variables obtained from the reanalysis data. The sea surface salinity (SSS) is restored to the monthly climatology with the relaxation time scale of 30 days. The initial condition is the annual mean climatology (Levitus and Boyer 1994; Levitus et al. 1994) with no motion. Then, the model is further integrated for 59 yr from 1950 to 2008 using the daily mean wind stress from the NCEP–NCAR reanalysis data and the daily surface heat flux calculated by the bulk formula using the simulated SST and the atmospheric variables obtained from the reanalysis data. The model SSS is again restored to the monthly mean climatology. Considering the oceanic adjustment time to the interannually varying forcing, we only analyze the outputs after 1960.

3. Annual cycle of the SST and the mixed layer depth in the subtropical South Atlantic

Figures 2a and 2b show the monthly climatology of the observed and simulated SST averaged over the southwestern pole (30°–40°S, 10°–30°W) and northeastern pole (15°–25°S, 0°–20°W) as defined in Fig. 1. Both observed and simulated SSTs over the poles reach their maxima in early spring. To examine the mechanism of this seasonal variation, we calculate the tendency of mixed layer temperature Tm averaged over the boxes of positive and negative poles using

Fig. 2.

Annual march of (a),(b) the SST (°C), (c),(d) each term in Eq. (1) (10−7 °C s−1), and (e),(f) the mixed layer depth (m) at the (left) southwestern and (right) northeastern poles defined in Fig. 1. In (a) and (b), the SST from HADISST (solid line) and the model (dashed line) are shown. In (c) and (d), MLT tend., NSHF, Hor. adv., Ent., and Res. indicate the mixed-layer temperature tendency (thick solid line), net surface heat flux (thick dashed line), horizontal advection (thick dotted line), entrainment (thin solid line), and residual (thin dashed line) terms, respectively. In (e) and (f), the mixed-layer depths derived from the WOA09 (solid line) and the model (dashed line) are shown.

Fig. 2.

Annual march of (a),(b) the SST (°C), (c),(d) each term in Eq. (1) (10−7 °C s−1), and (e),(f) the mixed layer depth (m) at the (left) southwestern and (right) northeastern poles defined in Fig. 1. In (a) and (b), the SST from HADISST (solid line) and the model (dashed line) are shown. In (c) and (d), MLT tend., NSHF, Hor. adv., Ent., and Res. indicate the mixed-layer temperature tendency (thick solid line), net surface heat flux (thick dashed line), horizontal advection (thick dotted line), entrainment (thin solid line), and residual (thin dashed line) terms, respectively. In (e) and (f), the mixed-layer depths derived from the WOA09 (solid line) and the model (dashed line) are shown.

 
formula

(Qiu and Kelly 1993; Moisan and Niiler 1998). In the first term on the right-hand side, Qnet is the net surface heat flux, qd is the downward insolation penetrating through the bottom of the mixed layer (Paulson and Simpson 1977), ρ is the density of the seawater, and cp is the specific heat of the seawater. The mixed layer depth H is defined as the depth at which temperature is 0.5°C lower than the SST. Since only less than 5% of the shortwave radiation can penetrate the shallowest mixed layer during austral summer, its penetration is negligible relative to Qnet in its annual cycle. We also obtain similar results even when the mixed layer depth is defined as the depth at which the potential density is 0.125 kg m−3 larger than that at the surface. In the second term, um denotes the horizontal velocity averaged in the mixed layer. In the third term, ΔT (≡TmTH−20m) represents the temperature difference between the mixed layer and the entrained water; we use the water temperature at 20 m below the mixed layer base as the temperature of the entrained water following Yasuda et al. (2000). Also, we is the entrainment velocity defined by

 
formula

where we assume we vanishes when it becomes negative (Kraus and Turner 1967; Qiu and Kelly 1993). The residual term Res in Eq. (1) includes diffusion and other neglected oceanic processes such as roles of high-frequency variability.

Figures 2c and 2d show the monthly climatology of each term in Eq. (1) over the southwestern and northeastern poles. The mixed layer temperature tendency over both poles is positive (negative) from October to February (from March to September). This is explained mostly by the net surface heat flux dominated by the shortwave radiation (figure not shown). Also, the entrainment contributes to the cooling when the mixed layer deepens as shown in Figs. 2e and 2f; the mixed layer depth at both poles becomes deepest from July to September.

4. SASD events

The first empirical orthogonal function (EOF) mode of the observed SST anomalies in the South Atlantic shows a dipole pattern and explains 20.4% of the total variance (Fig. 3a). The second EOF mode with a monopole pattern explains 12.9% of the total variance and is well separated from the first mode (North et al. 1982). This dipole pattern of the first mode corresponds well to that in Fig. 1 derived from the correlation with the SARI. To capture this interesting SST variability in a simple way, we introduce the SASDI, which is defined by the difference in the SST anomalies between the southwestern pole and the northeastern pole. As expected, the SASDI shows a high correlation of 0.92 with the principal component of the first EOF mode presented in Fig. 3b.

Fig. 3.

(a) Spatial pattern (0.1°C) and (b) principal component of the first EOF mode of SST anomalies during 1960–2008. Negative values in (a) are shaded, and the boxes over the positive and negative poles are defined as in Fig. 1. The values in (b) are normalized by the standard deviation.

Fig. 3.

(a) Spatial pattern (0.1°C) and (b) principal component of the first EOF mode of SST anomalies during 1960–2008. Negative values in (a) are shaded, and the boxes over the positive and negative poles are defined as in Fig. 1. The values in (b) are normalized by the standard deviation.

The monthly standard deviation of the SASDI undergoes significant variations, with values greater than 0.8°C from December to March (Fig. 4a). This suggests that the interannual variations of the SASDI are mostly locked to the austral summer. Therefore, we define years in which the SASDI exceeds 1 standard deviation during December–March as SASD-event years (Fig. 4b). This procedure leads to 8 (11) positive (negative) SASD events (Table 1).

Fig. 4.

(a) Monthly standard deviation of the observed (bar) and simulated (line) SASDI (°C). (b) Time series of the observed (solid line) and simulated (dashed line) SASDI averaged from December to March during 1960–2008. The values in (b) are normalized by the standard deviation.

Fig. 4.

(a) Monthly standard deviation of the observed (bar) and simulated (line) SASDI (°C). (b) Time series of the observed (solid line) and simulated (dashed line) SASDI averaged from December to March during 1960–2008. The values in (b) are normalized by the standard deviation.

Table 1.

Positive and negative SASD years used in the composite analysis.

Positive and negative SASD years used in the composite analysis.
Positive and negative SASD years used in the composite analysis.

The phase-locking nature of the SASD events is well reproduced in the simulation model as seen in Fig. 4a, and the time series of the simulated SASDI are highly correlated (0.95) with the observed SASDI (Fig. 4b). Hereinafter, we analyze the positive and negative SASD events in detail using both data and simulation results.

a. Positive SASD

Figure 5 shows composites of the observed and simulated SST anomalies for the positive SASD events over four seasons from austral spring to winter: September(0)–November(0), December(0)–February(1), March(1)–May(1), and June(1)–August(1). Both positive and negative SST anomaly poles start to develop synchronously from spring, reach their peak during summer, and decay after autumn. Since the pattern correlation between the observed and simulated SST anomalies of the peak (decay) phase in summer (autumn) is 0.94 (0.94), we may conclude that the model reproduces the evolution of composite SST anomalies very well except for some slight differences in magnitude.

Fig. 5.

Evolution of (left) observed and (right) simulated composite SST anomalies for the positive SASD. The contour interval is 0.2°C. The boxes over the positive and negative poles are defined as in Fig. 1. The shading indicates anomalies exceeding 90% confidence level in a two-tailed t test.

Fig. 5.

Evolution of (left) observed and (right) simulated composite SST anomalies for the positive SASD. The contour interval is 0.2°C. The boxes over the positive and negative poles are defined as in Fig. 1. The shading indicates anomalies exceeding 90% confidence level in a two-tailed t test.

The upper panels in Fig. 6 show composite anomalies of each term in Eq. (1) over the positive and negative poles. The mixed layer temperature anomalies over both poles develop significantly from November(0) to January(1). It is mostly due to the anomalous contribution from the net surface heat flux. The anomaly in the horizontal advection also contributes to the growth of the positive pole. This is dominated by the anomalous meridional advection owing to the anomalous southward Ekman and geostrophic flows (Sterl and Hazeleger 2003). Contributions from four components of the net surface heat flux are shown in Figs. 6c and 6d. It is clear that the anomalous contribution from the shortwave radiation plays the dominant role. This is in marked contrast with the previous results from Fauchereau et al. (2003) and Hermes and Reason (2005), who suggested the importance of the latent heat flux anomaly among four components of the net surface heat flux anomalies as shown in Figs. 6e and 6f. This interesting discrepancy needs to be solved.

Fig. 6.

Time series of composite anomalies of (a),(b) the mixed layer temperature tendency terms in Eq. (1) (10−7 °C s−1), (c),(d) components of the net surface heat flux term in Eq. (1) (10−7 °C s−1), and (e),(f) components of the net surface heat flux (W m−2) over the positive and negative SST anomaly poles for the positive SASD. A 3-month running mean is applied to smooth the time series. Filled (open) circle shows anomalies exceeding 90% (80%) confidence level in a two-tailed t test. In (a) and (b), MLT tend., NSHF, Hor. adv., Ent., and Res. indicate the mixed-layer temperature tendency (thick solid line), net surface heat flux (thick dashed line), horizontal advection (thick dotted line), entrainment (thin solid line), and residual (thin dashed line) terms, respectively. In (c) and (d), SW, LW, LH, and SH indicate the shortwave radiation (thick solid line), longwave radiation (thick dotted line), latent heat flux (thin solid line), and sensible heat flux (thin dashed line) terms, respectively. In (e) and (f), Qnet, QSW, QLW, QLH, and QSH indicate the net surface heat flux (thick dashed line), shortwave radiation (thick solid line), longwave radiation (thick dotted line), latent heat flux (thin solid line), and sensible heat flux (thin dashed line) anomalies, respectively.

Fig. 6.

Time series of composite anomalies of (a),(b) the mixed layer temperature tendency terms in Eq. (1) (10−7 °C s−1), (c),(d) components of the net surface heat flux term in Eq. (1) (10−7 °C s−1), and (e),(f) components of the net surface heat flux (W m−2) over the positive and negative SST anomaly poles for the positive SASD. A 3-month running mean is applied to smooth the time series. Filled (open) circle shows anomalies exceeding 90% (80%) confidence level in a two-tailed t test. In (a) and (b), MLT tend., NSHF, Hor. adv., Ent., and Res. indicate the mixed-layer temperature tendency (thick solid line), net surface heat flux (thick dashed line), horizontal advection (thick dotted line), entrainment (thin solid line), and residual (thin dashed line) terms, respectively. In (c) and (d), SW, LW, LH, and SH indicate the shortwave radiation (thick solid line), longwave radiation (thick dotted line), latent heat flux (thin solid line), and sensible heat flux (thin dashed line) terms, respectively. In (e) and (f), Qnet, QSW, QLW, QLH, and QSH indicate the net surface heat flux (thick dashed line), shortwave radiation (thick solid line), longwave radiation (thick dotted line), latent heat flux (thin solid line), and sensible heat flux (thin dashed line) anomalies, respectively.

The interannual anomaly in the contribution from the net surface heat flux in Eq. (1) may be decomposed as

 
formula

where an overbar indicates the monthly climatology, δ( ) means the interannual variation from the monthly climatology, and Q denotes Qnetqd. Figures 7a and 7b show contributions from each term of shortwave radiation in Eq. (3). Over the positive (negative) pole, the anomalous contribution from the shortwave radiation is dominated by the second term on the right-hand side of Eq. (3). This means that the warming of the mixed layer by the contribution from the climatological shortwave radiation is enhanced (suppressed) by the thinner (thicker) mixed layer as shown in Fig. 7c (Fig. 7d). More shortwave radiation may penetrate below the base of the mixed layer when the mixed layer becomes anomalously shallow, but this effect is negligible. The climatological latent heat flux is always negative, and the cooling of the mixed layer by the contribution from this is enhanced (suppressed) by the thinner (thicker) mixed layer. Thus, the warming (cooling) of the mixed layer by the contribution from the positive (negative) latent heat flux anomaly as shown in Fig. 6e (Fig. 6f) is canceled. The important role of the interannual variations in the mixed layer depth is missed in Fauchereau et al. (2003) and Hermes and Reason (2005) because of their assumption of the constant mixed layer depth. This intriguing aspect was discussed first by Morioka et al. (2010) for the IOSD.

Fig. 7.

Time series of composite anomalies of (a),(b) the shortwave radiation terms in Eq. (3) (10−7 °C s−1) and (c),(d) the mixed layer depth (m) at the positive and negative SST anomaly poles for the positive SASD. Filled (open) circle shows anomalies exceeding 90% (80%) confidence level in a two-tailed t test. A 3-month running mean is applied to smooth the time series. In (a) and (b), the net shortwave radiation term (thick solid line), first term (thick dashed line), second term (thick dotted line), and residual term (thin solid line) are shown.

Fig. 7.

Time series of composite anomalies of (a),(b) the shortwave radiation terms in Eq. (3) (10−7 °C s−1) and (c),(d) the mixed layer depth (m) at the positive and negative SST anomaly poles for the positive SASD. Filled (open) circle shows anomalies exceeding 90% (80%) confidence level in a two-tailed t test. A 3-month running mean is applied to smooth the time series. In (a) and (b), the net shortwave radiation term (thick solid line), first term (thick dashed line), second term (thick dotted line), and residual term (thin solid line) are shown.

The mixed layer depth anomaly may be explained by the anomalous wind stirring and surface heat flux. To quantify each effect, we calculate a diagnostic value of the mixed layer depth during a shoaling phase, which is given by the Monin–Obukhov depth:

 
formula

(Kraus and Turner 1967; Qiu and Kelly 1993). Here, m0 is a coefficient that indicates the efficiency of the wind stirring, and, following Davis et al. (1981), we use m0 = 0.5. The frictional velocity u* is defined by u*, where ρa is the density of the air, CD is the drag coefficient (=0.001 25), and u10 is the wind speed at 10-m height. Also, α is the thermal expansion coefficient (=0.000 25), and q(z) is the downward insolation (Paulson and Simpson 1977). Now, we examine the variation of the Monin–Obukhov depth anomaly by decomposing it into three components as

 
formula

where

 
formula

is the effective buoyancy forcing, and

 
formula

is the effective penetrative shortwave radiation. Figure 8 shows contributions from the wind-stirring anomaly (the first term), the surface heat flux anomaly (the second and third terms), and the residual in November(0)–December(0) and January(1)–February(1). Although there exists a slight difference in the amplitude between the actual mixed layer depth anomaly (Figs. 7c,d) and the Monin–Obukhov depth anomaly, the latter captures a quick shoaling phase from November(0)–December(0) to January(1)–February(1) at both poles. This may validate the use of the Monin–Obukhov depth as a measure of the mixed layer depth. The Monin–Obukhov depth anomaly at both poles of November(0)–December(0) is mostly due to the effect of the surface heat flux anomaly (Fig. 8), in which the latent heat flux anomaly plays a dominant role (Figs. 6e,f). Since the evaporation increases (decreases) as the positive (negative) SST anomaly grows, the contribution from the latent heat flux anomaly quickly decays from November(0)–December(0) to January(1)–February(1), leading to the quick decay of the Monin–Obukhov depth anomaly (Fig. 8). This situation may be linked with the southward migration and strengthening of the subtropical high in the South Atlantic (Fig. 9), as discussed by Fauchereau et al. (2003).

Fig. 8.

Composite anomalies of the contributions from the wind stirring (black bar), surface heat flux (gray bar), and residual (white bar) terms in Eq. (5) at the (left) positive and (right) negative SST anomaly poles for the positive SASD (m). Here, ND(0) represents the average from November(0) to December(0), and JF(1) represents the average from January(1) to February(1).

Fig. 8.

Composite anomalies of the contributions from the wind stirring (black bar), surface heat flux (gray bar), and residual (white bar) terms in Eq. (5) at the (left) positive and (right) negative SST anomaly poles for the positive SASD (m). Here, ND(0) represents the average from November(0) to December(0), and JF(1) represents the average from January(1) to February(1).

Fig. 9.

Composite anomalies of sea level pressure (thick contours with interval of 0.5 hPa) and latent heat flux (thin contours with interval of 3 W m−2) averaged in (left) November(0)–December(0) and (right) January(1)–February(1) of the positive SASD. Latent heat flux anomalies exceeding 90% confidence level in a two-tailed t test are shaded.

Fig. 9.

Composite anomalies of sea level pressure (thick contours with interval of 0.5 hPa) and latent heat flux (thin contours with interval of 3 W m−2) averaged in (left) November(0)–December(0) and (right) January(1)–February(1) of the positive SASD. Latent heat flux anomalies exceeding 90% confidence level in a two-tailed t test are shaded.

Furthermore, a significant difference in the mixed layer depth (Monin–Obukhov depth) anomaly between two poles is found especially in November(0)–December(0) as shown in Figs. 7c and 7d (Fig. 8). Besides the fact that the specific humidity difference near the surface over the northeastern pole is slightly larger than that over the southwestern pole, the wind speed at 10-m height over the northeastern flank of the subtropical high is nearly 2 times that over the southwestern flank (figure not shown). This causes larger latent heat loss over the negative pole in November(0)–December(0), leading to the smaller net surface heat flux in Eq. (5) than that over the positive pole. In addition, the absolute value of the latent heat anomaly over the negative pole is 2 times that over the positive pole, leading to the larger absolute value of the net surface heat flux anomaly (Figs. 6e,f). Both of these contribute to the asymmetry of the mixed layer depth (Monin–Obukhov depth) anomaly between the poles.

Now let us consider the decay phase. Both positive and negative SST anomaly poles decay during early (late) autumn owing to the anomalous contribution from the entrainment (net surface heat flux) in Figs. 6a and 6b. To investigate the anomalous contribution from the entrainment in Eq. (1) in more detail, we decompose it as

 
formula

In early autumn, the first term on the right-hand side plays a major role (Figs. 10a,b), indicating that, over the positive (negative) SST anomaly pole, the cooling of the mixed layer by the entrainment is enhanced (suppressed) by the positive (negative) anomaly of the temperature difference between the mixed layer and the entrained water. Since changes in the entrained water temperature are relatively small, the anomaly in the temperature difference is mostly due to the positive (negative) mixed layer temperature anomaly. In late autumn, the anomalous contribution from the net surface heat flux is mostly due to that from the latent heat flux (Figs. 6c,d). Figures 10c and 10d show time series of composite anomalies of each term in Eq. (3) for the latent heat flux. Those anomalies in late autumn are mostly due to the effect of the second term on the right-hand side of Eq. (3), except for the anomalous contribution from the latent heat flux in May(1) when the first term also contributes. Thus, over the positive (negative) pole, the cooling of the mixed layer by the contribution from the climatological latent heat flux is enhanced (suppressed) by the thinner (thicker) mixed layer, as indicated in Fig. 7c (Fig. 7d). Although the anomalous contribution from the latent heat flux in late autumn plays a major role in the decay of the SASD (Figs. 6c,d), that from the shortwave radiation continues to maintain the SASD. This opposite contribution from the shortwave radiation may cause the anomalies in the contribution from the net surface heat flux to lag behind those from the entrainment by 2 months.

Fig. 10.

Time series of composite anomalies of (a),(b) the entrainment terms (10−7 °C s−1) in Eq. (6) and (c),(d) the latent heat flux terms (10−7 °C s−1) in Eq. (3) over the (left) positive and (right) negative SST anomaly poles for the positive SASD. Filled (open) circle shows anomalies exceeding 90% (80%) confidence level in a two-tailed t test. A 3-month running mean is applied to smooth the time series. In (a) and (b), the net entrainment term (thick solid line), first term (thick dashed line), second term (thick dotted line), third term (thin solid line), and residual term (thin dashed line) are shown. In (c) and (d), the net latent heat flux term (thick solid line), first term (thick dashed line), second term (thick dotted line), and residual term (thin solid line) are shown.

Fig. 10.

Time series of composite anomalies of (a),(b) the entrainment terms (10−7 °C s−1) in Eq. (6) and (c),(d) the latent heat flux terms (10−7 °C s−1) in Eq. (3) over the (left) positive and (right) negative SST anomaly poles for the positive SASD. Filled (open) circle shows anomalies exceeding 90% (80%) confidence level in a two-tailed t test. A 3-month running mean is applied to smooth the time series. In (a) and (b), the net entrainment term (thick solid line), first term (thick dashed line), second term (thick dotted line), third term (thin solid line), and residual term (thin dashed line) are shown. In (c) and (d), the net latent heat flux term (thick solid line), first term (thick dashed line), second term (thick dotted line), and residual term (thin solid line) are shown.

b. Negative SASD

The evolution of the SST anomalies for the negative SASD is shown in Fig. 11. Both positive and negative SST anomaly poles start to develop synchronously from austral spring, reach their peaks during summer, and decay after autumn. The pattern correlation between the observed and simulated SST anomalies during the peak (decay) phase of summer (autumn) is 0.95 (0.93), and the evolution of the negative SASD is also well reproduced in the model.

Fig. 11.

As in Fig. 5, but for the negative SASD.

Fig. 11.

As in Fig. 5, but for the negative SASD.

The mixed layer temperature anomalies over both poles develop significantly from November(0) to January(1) mostly owing to the anomalous contribution from the net surface heat flux (Figs. 12a,b). In particular, the contribution from the shortwave radiation is dominant (Figs. 12c,d). This contribution is mostly explained by the second term on the right-hand side of Eq. (3) (Figs. 13a,b), indicating that, over the positive (negative) pole, the warming of the mixed layer by the contribution from the climatological shortwave radiation is enhanced (suppressed) by the thinner (thicker) mixed layer as shown in Fig. 13c (Fig. 13d). As indicated in Fig. 14, the quick decay as well as the amplitude of the mixed layer depth anomaly is well diagnosed by those of the Monin–Obukhov depth anomaly, which is mostly due to the contribution from the surface heat flux anomaly [the second and third terms on the right-hand side of Eq. (5)]. In particular, the latent heat flux anomaly plays the dominant role (Figs. 12e,f). This latent heat flux anomaly may also be linked with the weakening of the subtropical high in the South Atlantic (Fig. 15).

Fig. 12.

As in Fig. 6, but for the negative SASD.

Fig. 12.

As in Fig. 6, but for the negative SASD.

Fig. 13.

As in Fig. 7, but for the negative SASD.

Fig. 13.

As in Fig. 7, but for the negative SASD.

Fig. 14.

As in Fig. 8, but for the negative SASD.

Fig. 14.

As in Fig. 8, but for the negative SASD.

Fig. 15.

As in Fig. 9, but for the negative SASD.

Fig. 15.

As in Fig. 9, but for the negative SASD.

The mixed layer temperature anomalies in both poles decay during early (late) autumn owing to the anomalous contribution from the entrainment (net surface heat flux) in Figs. 12a and 12b. Over the positive (negative) pole, the cooling of the mixed layer by the entrainment is enhanced (suppressed) by the larger (smaller) temperature difference between the mixed layer and the entrained water in early autumn as shown in Fig. 16a (Fig. 16b). In particular, the positive (negative) mixed layer temperature anomaly mostly contributes to the temperature difference. In late autumn, the anomalous contribution from the net surface heat flux is mostly explained by that from the latent heat flux (Figs. 12c,d). The cooling of the mixed layer by the contribution from the climatological latent heat flux is enhanced by the thinner mixed layer at the positive pole (Fig. 16c). On the other hand, over the negative pole, the contribution from the positive latent heat flux anomaly contributes mostly to the damping of the negative mixed layer temperature anomaly (Fig. 16d). Although the mean mixed layer depth and latent heat flux over the negative pole are similar to those over the positive pole, the mixed layer depth anomaly at the negative pole is 0.2 m, much smaller than −1.8 m at the positive pole (Figs. 13c,d). This is why the contribution from the climatological latent heat flux over the negative pole is much smaller than that over the positive pole.

Fig. 16.

As in Fig. 10, but for the negative SASD.

Fig. 16.

As in Fig. 10, but for the negative SASD.

5. Conclusions

Using outputs from observational data and OGCM simulation results, the interannual SST variations in the South Atlantic are investigated. The new mechanism of the positive SASD demonstrated in this study is summarized schematically in Fig. 17. During the growth phase (Fig. 17a), the anomalous southward migration and strengthening of the subtropical high in late spring cause the positive (negative) latent heat flux anomaly over the positive (negative) SST anomaly pole. This leads to the anomalous shoaling (deepening) of the mixed layer in early summer. As a result, the warming of the mixed layer by the contribution from the climatological shortwave radiation is enhanced (suppressed) by the thinner (thicker) mixed layer and the positive (negative) SST anomaly pole develops. Thus, the latent heat flux anomaly contributes to the growth of both poles through its influence on the mixed layer depth. This is in contrast to previous studies on the SASD, which suggested that the latent heat flux anomaly directly generates the SST anomaly. This important influence by the interannual mixed layer depth anomaly on the contribution from the climatological shortwave radiation is also discussed to explain the evolution of the IOSD (Morioka et al. 2010).

Fig. 17.

Schematic of (a) growth and (b) decay mechanisms of the positive and negative SST anomaly poles for the positive SASD. Red (blue) color indicates the positive (negative) mixed layer temperature anomaly.

Fig. 17.

Schematic of (a) growth and (b) decay mechanisms of the positive and negative SST anomaly poles for the positive SASD. Red (blue) color indicates the positive (negative) mixed layer temperature anomaly.

On the other hand, during the decay phase (Fig. 17b) the anomalous contributions from the net surface heat flux and entrainment damp the above SST anomalies. In early autumn, the cooling of the mixed layer by the entrainment is enhanced (suppressed) by the anomalously large (small) temperature difference between the mixed layer and the entrained water. In particular, the positive (negative) temperature anomaly in the mixed layer is responsible for this anomalous temperature difference. In addition, the cooling of the mixed layer by the contribution from the climatogical latent heat flux is enhanced (suppressed) by the thinner (thicker) mixed layer in late autumn. The almost similar mechanism is obtained for the growth and decay of the negative SASD.

This study reveals the important roles of interannual variation of the mixed layer depth on the formation as well as the decay of the SASD. The above results are based on the outputs from the OGCM, however, in which the surface heat flux is calculated by the bulk formula using the atmospheric reanalysis data and the simulated SST. Further studies using an ocean–atmosphere coupled model are necessary to investigate the air–sea interaction processes involving the SASD in more detail.

Acknowledgments

The authors thank Dr. Yukio Masumoto for his helpful comments. They also thank two anonymous reviewers for their helpful comments. The OGCM was run on the HITACHI SR11000/J1 of the Information Technology Center at the University of Tokyo as part of cooperative research with the Atmosphere and Ocean Research Institute of the University of Tokyo. This research is supported by the Japan Science and Technology Agency/Japan International Cooperation Agency through the Science and Technology Research Partnership for Sustainable Development (SATREPS). Also, the first author is supported by both the Sasakawa Scientific Research Grant from the Japan Science Society and the Research Fellowship of the Japan Society for the Promotion of Science.

REFERENCES

REFERENCES
Behera
,
S. K.
, and
T.
Yamagata
,
2001
:
Subtropical SST dipole events in the southern Indian Ocean
.
Geophys. Res. Lett.
,
28
,
327
330
.
Behera
,
S. K.
,
P. S.
Salvekar
, and
T.
Yamagata
,
2000
:
Simulation of interannual SST variability in the tropical Indian Ocean
.
J. Climate
,
13
,
3487
3499
.
Cook
,
K. H.
,
2000
:
The South Indian convergence zone and interannual rainfall variability over southern Africa
.
J. Climate
,
13
,
3789
3804
.
Davis
,
R. E.
,
R.
de Szoeke
, and
P.
Niiler
,
1981
:
Variability in the upper ocean during MILE. Part II: Modeling the mixed layer response
.
Deep-Sea Res.
,
28
,
1453
1475
.
Fauchereau
,
N.
,
S.
Trzaska
,
Y.
Richard
,
P.
Roucou
, and
P.
Camberlin
,
2003
:
Sea-surface temperature co-variability in the southern Atlantic and Indian Oceans and its connections with the atmospheric circulation in the Southern Hemisphere
.
Int. J. Climatol.
,
23
,
663
677
.
Hermes
,
J. C.
, and
C. J. C.
Reason
,
2005
:
Ocean model diagnosis of interannual coevolving SST variability in the South Indian and South Atlantic Oceans
.
J. Climate
,
18
,
2864
2882
.
Kalnay
,
E.
, and
Coauthors
,
1996
:
The NCEP/NCAR 40-Year Reanalysis Project
.
Bull. Amer. Meteor. Soc.
,
77
,
437
471
.
Kraus
,
E. B.
, and
J. S.
Turner
,
1967
:
A one-dimensional model of the seasonal thermocline: II. The general theory and its consequences
.
Tellus
,
19
,
98
106
.
Legates
,
D. R.
, and
C. J.
Willmott
,
1990
:
Mean seasonal and spatial variability in gauge-corrected, global precipitation
.
Int. J. Climatol.
,
10
,
111
127
.
Levitus
,
S.
, and
T. P.
Boyer
,
1994
:
Temperature
.
Vol. 4, World Ocean Atlas 1994, NOAA Atlas NESDIS 4, 117 pp
.
Levitus
,
S.
,
R.
Burgett
, and
T. P.
Boyer
,
1994
:
Salinity
.
Vol. 3, World Ocean Atlas 1994, NOAA Atlas NESDIS 3, 99 pp
.
Lindesay
,
J. A.
,
1988
:
South African rainfall, the Southern Oscillation, and a Southern Hemisphere semi-annual cycle
.
J. Climatol.
,
8
,
17
30
.
Mason
,
S. J.
,
1995
:
Sea-surface temperature—South African rainfall associations, 1910–89
.
Int. J. Climatol.
,
15
,
119
135
.
Moisan
,
J. R.
, and
P. P.
Niiler
,
1998
:
The seasonal heat budget of the North Pacific: Net heat flux and heat storage rates (1950–90)
.
J. Phys. Oceanogr.
,
28
,
401
421
.
Morioka
,
Y.
,
T.
Tozuka
, and
T.
Yamagata
,
2010
:
Climate variability in the southern Indian Ocean as revealed by self-organizing maps
.
Climate Dyn.
,
35
,
1059
1072
.
North
,
G. R.
,
T. L.
Bell
,
R. F.
Cahalan
, and
F. J.
Moeng
,
1982
:
Sampling errors in the estimation of empirical orthogonal functions
.
Mon. Wea. Rev.
,
110
,
699
706
.
Pacanowski
,
R. C.
, and
S. G. H.
Philander
,
1981
:
Parameterization of vertical mixing in numerical models of tropical oceans
.
J. Phys. Oceanogr.
,
11
,
1443
1451
.
Pacanowski
,
R. C.
, and
S. M.
Griffies
,
1999
:
MOM 3.0 manual
.
NOAA/GFDL, 680 pp
.
Paulson
,
C. A.
, and
J. J.
Simpson
,
1977
:
Irradiance measurements in the upper ocean
.
J. Phys. Oceanogr.
,
7
,
952
956
.
Qiu
,
B.
, and
K. A.
Kelly
,
1993
:
Upper-ocean heat balance in the Kuroshio Extension region
.
J. Phys. Oceanogr.
,
23
,
2027
2041
.
Rayner
,
N. A.
,
D. E.
Parker
,
E. B.
Horton
,
C. K.
Folland
,
L. V.
Alexander
,
D. P.
Rowell
,
E. C.
Kent
, and
A.
Kaplan
,
2003
:
Global analysis of sea surface temperature, sea ice, and night marine air temperature since the late nineteenth century
.
J. Geophys. Res.
,
108
,
4407
,
doi:10.1029/2002JD002670
.
Reason
,
C. J. C.
,
1998
:
Warm and cold events in the southeast Atlantic/southwest Indian Ocean region and potential impacts on circulation and rainfall over southern Africa
.
Meteor. Atmos. Phys.
,
69
,
49
65
.
Reason
,
C. J. C.
,
2001
:
Subtropical Indian Ocean SST dipole events and southern African rainfall
.
Geophys. Res. Lett.
,
28
,
2225
2227
.
Reason
,
C. J. C.
, and
H.
Mulenga
,
1999
:
Relationships between South African rainfall and SST anomalies in the southwest Indian Ocean
.
Int. J. Climatol.
,
19
,
1651
1673
.
Richard
,
Y.
,
S.
Trzaska
,
P.
Roucou
, and
M.
Rouault
,
2000
:
Modification of the southern African rainfall variability/ENSO relationship since the late 1960s
.
Climate Dyn.
,
16
,
883
895
.
Smagorinsky
,
J.
,
1963
:
General circulation experiments with the primitive equations. Part I: The basic experiment
.
Mon. Wea. Rev.
,
91
,
99
164
.
Sterl
,
A.
, and
W.
Hazeleger
,
2003
:
Coupled variability and air–sea interaction in the South Atlantic Ocean
.
Climate Dyn.
,
21
,
559
571
.
Suzuki
,
R.
,
S. K.
Behera
,
S.
Iizuka
, and
T.
Yamagata
,
2004
:
Indian Ocean subtropical dipole simulated using a coupled general circulation model
.
J. Geophys. Res.
,
109
,
C09001
,
doi:10.1029/2003JC001974
.
Uppala
,
S. M.
, and
Coauthors
,
2005
:
The ERA-40 Re-Analysis
.
Quart. J. Roy. Meteor. Soc.
,
131
,
2961
3012
.
Venegas
,
S. A.
,
L. A.
Mysak
, and
D. N.
Straub
,
1996
:
Evidence for interannual and interdecadal climate variability in the South Atlantic
.
Geophys. Res. Lett.
,
23
,
2673
2676
.
Venegas
,
S. A.
,
L. A.
Mysak
, and
D. N.
Straub
,
1997
:
Atmosphere–ocean coupled variability in the South Atlantic
.
J. Climate
,
10
,
2904
2920
.
Walker
,
N. D.
,
1990
:
Links between South African summer rainfall and temperature variability of the Agulhas and Benguela current systems
.
J. Geophys. Res.
,
95
,
3297
3319
.
Yasuda
,
I.
,
T.
Tozuka
,
M.
Noto
, and
S.
Kouketsu
,
2000
:
Heat balance and regime shifts of the mixed layer in the Kuroshio Extension
.
Prog. Oceanogr.
,
47
,
257
278
.