1. Introduction
The Mascarene high brings a huge amount of moisture from the southern Indian Ocean over southern Africa, which helps develop the South Indian convergence zone (Cook 2000) during austral summer (December–February, hereafter the season is defined in the Southern Hemisphere) and causes most of the annual rainfall over southern Africa via synoptic-scale rainfall systems called the tropical–temperate trough (Harrison 1984; Todd and Washington 1999; Todd et al. 2004; Ratna et al. 2013). As shown in Fig. 1a, the Mascarene high during austral summer shifts eastward and intensifies to the west of Australia owing to a large land–sea thermal heat contrast (Miyasaka and Nakamura 2010). Also, it experiences a large year-to-year variation, which strongly affects the sea surface temperature (SST) in the southern Indian Ocean (Behera and Yamagata 2001) and southern African rainfall (Reason 2001, 2002; Washington and Preston 2006). Therefore, it is of great importance for the local society to accurately understand and skillfully predict the interannual variability in the summertime Mascarene high.
Climatological SLP (hPa) during the austral summer of 1982–2011 for (a) the JRA-55 and (b) CTR experiment.
Citation: Journal of Climate 28, 2; 10.1175/JCLI-D-14-00133.1
The interannual variability within the summertime Mascarene high has been widely discussed in particular its relationship to SST anomalies in the southern Indian Ocean. Behera and Yamagata (2001) first discussed a northeast–southwest-oriented dipole of SST anomalies associated with the variations in the summertime Mascarene high, referred to as Indian Ocean subtropical dipole (IOSD). Both the surface heat flux and the wind speed anomalies associated with the Mascarene high variation induce the anomalous mixed-layer thickness, and contribute to the SST anomalies by changing the warming of the mixed layer by shortwave radiation (Morioka et al. 2010, 2012). Regarding variations within the Mascarene high during the development of the IOSD, Behera and Yamagata (2001) suggested the possibility of the local SST influence because of the low correlation with major climate modes such as El Niño–Southern Oscillation (ENSO) and the Indian Ocean dipole (IOD; Saji et al. 1999).
To further investigate local SST influence, Reason (2001, 2002) performed a series of atmospheric general circulation model (AGCM) experiments by prescribing these SST anomalies in the southern Indian Ocean, which are associated with the positive IOSD. The authors claimed that, in a case where the AGCM is forced with the warm and cold SST anomalies in the southwestern and southeastern Indian Ocean, respectively, the rainfall over southern Africa remarkably increases as a result of the anomalous low pressure generated by the warm SST anomalies. However, the atmospheric response in the AGCM experiment contradicts with the reanalyses during the development of the positive IOSD, when the Mascarene high anomalously strengthens in its southern part. Rather, it represents the weakening of the Mascarene high during the decay of the positive IOSD, suggesting a potential negative feedback from the warm SST anomalies (Wang 2010).
On the other hand, Washington and Preston (2006) performed another set of AGCM experiments by specifying meridional dipole SST anomalies in the southwestern Indian Ocean. The authors discussed a significant role of the northern cold SST anomalies in the anticyclonic anomaly, which helps to advect more moisture from the southern Indian Ocean to the eastern part of southern Africa, leading to an anomalous increase in the rainfall.
The apparent differences in the Mascarene high variability of the AGCMs may stem partly from the failure of AGCMs in simulating the reanalyses, but we cannot rule out the local feedback from the warm SST anomalies on the Mascarene high variability. In fact, Nakamura (2012) found that warm SST anomalies in the Agulhas Return Current region modify the near-surface baroclinicity and, hence, affect the storm tracks associated with synoptic-scale atmospheric eddies. Since the southern part of the Mascarene high is located near the Agulhas Return Current system, it is more likely that the changes in the storm-track activity may have some effects on the quasi-stationary flow through an eddy–mean interaction (Lau 1988; Hoskins and Valdes 1990) and hence a semipermanent high pressure system like the Mascarene high.
Several recent studies discussed other possible factors responsible for the summertime Mascarene high variability such as remote influences from ENSO and the Antarctic Oscillation (AAO; Thompson and Wallace 2000) by analyzing the reanalyses (Fauchereau et al. 2003; Hermes and Reason 2005) and the coupled general circulation model (CGCM) results (Huang and Shukla 2008). Recently, Morioka et al. (2013, 2014) conducted a set of CGCM experiments to examine the relative importance of tropical SST variability and the AAO in the Mascarene high variability during the IOSD. Their CGCM results suggest that in the absence of the tropical SST variability, the AAO plays a major role in inducing the Mascarene high variability. However, it remains uncertain to what extent the variations in the summertime Mascarene high are explained by the remote influence and the local air–sea interaction.
Therefore, this study aims to investigate the potential role of the local SST variability in the southern Indian Ocean on the interannual variations in the summertime Mascarene high by performing a series of the CGCM experiments. Details of the CGCM experiments, data, and methodology for the atmospheric analysis are given in the next section. Section 3 gives a brief discussion of the climatological oceanic and atmospheric fields during austral summer. In section 4, we investigate a potential mechanism of the local ocean–atmosphere feedback on the summertime Mascarene high variability. In section 5, the possible impact of the SST anomalies as well as the Mascarene high variability on the southern African rainfall are examined. Finally, section 6 provides a summary of results and discussion on the possible impact of the local SST variability on the southern African rainfall.
2. Model, data, and methodology
a. CGCM experiments
For the CGCM experiments, we use the Scale Interaction Experiment-Frontier Research Center for Global Change 2 model (SINTEX-F2; Masson et al. 2012), which is an upgraded version of the SINTEX-F1 (Luo et al. 2003, 2005) model and has a sufficient skill of describing climate variability in the tropics (Sasaki et al. 2013) and the midlatitudes of the Southern Hemisphere (Morioka et al. 2014). The atmospheric component is based on ECHAM5 (Roeckner et al. 2003), which was originally developed at the European Centre for Medium-Range Weather Forecasts (ECMWF) and has a parameterization package developed in Hamburg, Germany. The ECHAM5 has 31 vertical levels and a spectrally truncated T106 Gaussian grid in the horizontal. The oceanic component is the Nucleus for European Modelling of the Ocean (NEMO; Madec 2008), which includes the Louvain-la-Neuve Sea Ice Model 2 (LIM2; Fichefet and Maqueda 1997) and has 0.5° × 0.5° horizontal resolution of ORCA grid configuration (ORCA05) with 31 vertical levels. The atmospheric and oceanic fields are exchanged every 2 h with no flux correction by means of the Ocean Atmosphere Sea Ice Soil 3 (OASIS3) coupler (Valcke et al. 2004). For the control run (CTR) experiment, the SINTEX-F2 was integrated for 100 years, and monthly-mean outputs from the last 80 years are analyzed.
SST nudging experiments were also performed to investigate the potential role of SST variations in the southern Indian Ocean (SIO; Table 1). In the SIO experiment, SSTs in the southern Indian Ocean (10°–55°S, 25°–125°E) are nudged to the monthly climatology of the CTR experiment. A strong negative feedback (−9600 W m−2 K−1) is added to the surface heat flux so that the temperature in the upper 200 m, the typical wintertime mixed layer in the midlatitudes, is restored within 1 day. Within 5° off the area used for the SST nudging, the restoration of the simulated SST to the model climatology gradually weakens. The SIO experiment was integrated for 50 years, and the monthly outputs for the last 30 years are used for the present analysis after removing the first 20 years for the oceanic adjustment to the interannually varying atmospheric forcing.
Details of the CGCM experiments performed in this study.
b. Data
Monthly-mean SSTs from the Advanced Very High Resolution Radiometer (AVHRR)-only daily optimum interpolation SST (AVHRR-only OISST, version 2; Reynolds et al. 2007) are used for validating simulated SSTs. It has a horizontal resolution of ¼° × ¼° and covers the 1982–2011 period. For the atmosphere, because of the lack of observation, both monthly and 6-hourly data from the Japanese 55-year Reanalysis (JRA-55; Ebita et al. 2011) provided by the Japan Meteorological Agency are used for comparison to the atmospheric fields simulated from the modeling experiments. It covers the same period with a horizontal resolution of 1.25° × 1.25°. Another dataset from the National Centers for Environmental Prediction–U.S. Department of Energy (NCEP–DOE) Atmospheric Model Intercomparison Project phase 2 (AMIP-II) reanalysis (NCEP2; Kanamitsu et al. 2002) is also used, but the results are qualitatively the same as in the JRA-55 (not shown). For precipitation, we use the observed data from the Climate Prediction Center (CPC) Merged Analysis of Precipitation (CMAP; Xie and Arkin 1997), which has a horizontal resolution of 2.5° × 2.5°. For all the above datasets, monthly anomalies were calculated after removing the linear trend related to climate change through the least squares method and the monthly climatology.
c. Methodology
As one of possible factors responsible for the summertime Mascarene high variability, we discuss the role of storm tracks associated with synoptic eddies. For the evaluation of the storm-track activity, we calculated the meridional heat flux (υ′T′) at 925 hPa and the meridional wind variance (υ′υ′) at 250 hPa associated with the synoptic eddies with a frequency of less than 8 days in the lower and upper troposphere, respectively (Nakamura and Shimpo 2004; Nakamura 2012). Here, we calculated the monthly anomaly of the high-frequency eddy component. First, we removed the monthly climatology from the original 6-hourly data to calculate the anomaly. Then, we applied an 8-day running mean filter to the anomaly to extract the low-frequency component. Finally, we subtracted this from the anomaly to obtain the high-frequency component less than 8 days associated with synoptic eddies. For the convenience of interpretation, the meridional heat flux in the Southern Hemisphere is shown with a minus sign to make the positive values correspond to the poleward heat flux.
3. Mean states during austral summer
The atmospheric and oceanic mean states over the analysis period of 1982–2011 are important for understanding the interannual variations. In the reanalyses, the core of the Mascarene high is located between 30°–35°S and 80°–90°E with a central pressure of 1020 hPa (Fig. 1a). The CTR experiment successfully simulates the location as well as the amplitude of the Mascarene high in the reanalyses (Fig. 1b). A good agreement between the reanalyses and the CTR experiment is found in the upper troposphere. A strong westerly jet, the polar front jet with a maximum amplitude of about 36 m s−1 between 40° and 50°S, is found in the reanalyses (Fig. 2a). The westerly jet is well reproduced in the CTR experiment (Fig. 2b), although it is located at a slightly lower latitude.
Climatological zonal wind (m s−1) at 250 hPa during the austral summer for (a) the JRA-55 and (b) CTR experiment. (c),(d) As in (a),(b), but for the SST (°C) and AVHRR-only OISST is shown in (c).
Citation: Journal of Climate 28, 2; 10.1175/JCLI-D-14-00133.1
For the oceanic mean state, the observation shows a warm SST south of the Mozambique Channel owing to the southward Agulhas Current, which creates a sharp SST front around 40°–45°S together with a strong Agulhas Return Current (Fig. 2c). Simulated SSTs in the subtropical basin show a slight cold bias compared with the observation, but the spatial SST pattern is well simulated (Fig. 2d).
In austral summer, because of the weak Hadley cell, the deep polar-front jet forms in the troposphere over the subarctic frontal zone, which affects the storm-track activity associated with synoptic eddies (Nakamura and Shimpo 2004). Figure 3 shows the climatology of the meridional SST gradient, the poleward heat flux by the synoptic eddies at 925 hPa, and the meridional wind variance associated with the synoptic eddies at 250 hPa. In the reanalyses, a strong poleward heat flux at 925 hPa is found to be associated with synoptic eddies between 40° and 50°S (Fig. 3c), which exists slightly poleward of the maximum SST front around 40°–45°S in the observation (Fig. 3a). Since this level is very close to the surface within the atmospheric boundary layer, the surface friction may play a role in the difference of the maximum latitude. In the upper troposphere, the meridional wind variance has its maximum between 40° and 50°S (Fig. 3e) and corresponds well to the exit region of the polar-front jet in Fig. 2a. In the CTR experiment, the SST front is simulated at lower latitude around 40°S (Fig. 3b). This may cause a poleward heat flux by the synoptic eddies at a lower latitude than that in the reanalyses (Fig. 3d). Also, the meridional wind variance by the synoptic eddies in the model are slightly stronger by 10% than that in the reanalyses (Fig. 3f). The NCEP2 data yield qualitatively similar results but the amplitude of the poleward heat transport as well as the meridional wind variance is weaker than that of the JRA-55. This may be partly due to the difference in the horizontal resolution between the reanalyses, although there is a difference in the model physics.
Climatology of the meridional SST gradient (10−6 °C m−1) during the austral summer for (a) the AVHRR-only OISST and (b) CTR experiment. (c),(d) As in (a),(b), but for the poleward heat flux by synoptic eddies at 925 hPa (m K s−1) and JRA-55 is shown in (c). (e),(f) As in (c),(d), but for the meridional wind variance (m2 s−2) associated with synoptic eddies at 250 hPa.
Citation: Journal of Climate 28, 2; 10.1175/JCLI-D-14-00133.1
4. Role of local SST variability in the summertime Mascarene high variability
In this study, we focus on the interannual variability of the Mascarene high intensity as well as the spatial variability of the Mascarene high. For this purpose, we performed an empirical orthogonal function (EOF) analysis on SLP anomalies in the domain during the austral summer of the analysis period (Fig. 4). As one alternative way to detect the variability, the SLP anomalies at the center of the Mascarene high can be used, but they do not necessarily represent the spatial variability in the whole Mascarene high system. The first EOF mode of the SLP anomalies discussed below is found to sufficiently separate from the second EOF mode by North et al.’s (1982) test (not shown). The first EOF mode in the reanalyses has positive values in the southern part of the southern Indian Ocean and negative values over northwestern Australia, representing a meridional dipole structure (Fig. 4a). The strengthening of the Mascarene high in its southern part is also simulated in the CTR experiment, although the simulated amplitude is slightly larger than the reanalyses and the negative values northwest of Australia are weaker (Fig. 4b). The explained variance (45.6%) is close to that in the reanalyses (40.8%). Interestingly, the first EOF mode in the SIO experiment also exhibits the strengthening of the Mascarene high in its southern part, although with larger amplitude compared to the CTR experiment (Fig. 4c). The close resemblance between the two experiments implies that the SST variability in the southern Indian Ocean may not have much influence on the Mascarene high variability during the austral summer. However, we cannot rule out the possibility that there may exist a slight difference in the two experiments for a particular month in the austral summer. Since the synoptic atmospheric disturbances cannot sustain their signals over more than one month without external forcing, the results of the EOF analysis for the seasonally averaged anomalies may be sensitive to a strong signal in a particular month.
(a) Spatial pattern of the first EOF mode for the SLP anomalies (hPa) during the austral summer of the analysis period for JRA-55 . The contour interval is 0.5 hPa. Positive values are shaded. The value on the top right of each panel indicates the explained variance. (b),(c) As in (a), but for the CTR and SIO experiments, respectively.
Citation: Journal of Climate 28, 2; 10.1175/JCLI-D-14-00133.1
To explore the possible difference between the CTR and SIO experiments, we defined the positive (negative) events as years when the principal components of the first EOF mode exceed 0.8 (−0.8) standard deviation and conducted a composite analysis for each month in the austral summer. The observed years for the positive and negative events are listed in Table 2. Note that because of the limited analysis period for the observation, reanalyses, and the SIO experiment, we decided to use this criterion of standard deviation for the event detection. Since the results for the negative events are just mirror images of those for the positive events, here we discuss only the positive cases.
Positive and negative event years for the composite analysis based on the JRA-55 data. The positive (negative) events are defined as years when the principal components of the first EOF mode of SLP anomalies during the austral summer of the analysis period exceed 0.8 (−0.8) standard deviation.
Figure 5 shows composite SLP anomalies during the austral summer of the positive events. In the reanalyses, the positive SLP anomalies in the southern Indian Ocean significantly persist over three months until February of the year following the event [i.e., Feb(1)] (Fig. 5a). They concur with the negative SLP anomalies northwest of Australia as expected from the results of the EOF analysis (Fig. 4a). The persistence of the positive SLP anomalies throughout the austral summer is also simulated in the CTR experiment (Fig. 5b), resembling the spatial pattern found in the reanalyses to the exception of the negative SLP anomalies northwest of Australia, which are weaker in the modeling experiments. On the other hand, the SIO experiment shows that the positive SLP anomalies in the southern Indian Ocean persist over December of the event year [Dec(0)] and January following the event year [Jan(1)], but sharply weaken in Feb(1). Indeed, the pattern correlations between composite anomalies of the CTR and SIO experiments in Dec(0) and Jan(1) are significantly high at 0.83 and 0.75, respectively, but it becomes very low at 0.33 in Feb(1), not significant at the 95% confidence level using a two-tailed t test. This remarkable pattern difference between the CTR and SIO experiments in Feb(1) implies the possible influence of the local SST variability on the persistence of the positive SLP anomalies in the CTR experiment.
(a) Composite SLP anomalies (hPa) during the austral summer of the positive events for JRA-55. (b),(c) As in (a), but for the CTR and SIO experiments, respectively. Values exceeding the 90% (95%) confidence level using a two-tailed t test are colored for the JRA-55 and SIO experiment (CTR experiment).
Citation: Journal of Climate 28, 2; 10.1175/JCLI-D-14-00133.1
For further examination of the potential oceanic role, we computed composite SST anomalies during the positive events (Fig. 6). The composite SST anomalies in the observation show positive anomalies in the southern part of the southern Indian Ocean and negative anomalies in its northern part, representing a meridional dipole pattern (Fig. 6a). They are associated with strong positive SST anomalies northwest of Australia, which may be related to a coastal climate phenomenon recently identified as Ningaloo Niño (Feng et al. 2013; Kataoka et al. 2014). Indeed, three out of six events in 1996, 1999, and 2010 are associated with the Ningaloo Niño and this may have a close link with the negative SLP anomalies northwest of Australia in Fig. 6a. It should be noted that the meridional dipole of the SST anomalies reaches the maximum in their amplitudes in Feb(1), then decays during the austral autumn (not shown). The CTR experiment reasonably simulates the observed meridional dipole structure of the SST anomalies with their peak in Feb(1) (Fig. 6b), although the model seems to fail in reproducing the observed positive SST anomalies northwest of Australia, which may be responsible for the failure of simulating the negative SLP anomalies in the reanalyses (Figs. 5a,b). Because of their strong peak, the meridional dipole of SST anomalies is more likely to have an impact on the atmosphere in Feb(1).
As in Fig. 5, but for the SST anomalies (°C) and AVHRR-only OISST is shown in (a).
Citation: Journal of Climate 28, 2; 10.1175/JCLI-D-14-00133.1
One of the possible physical processes responsible for the ocean–atmosphere feedback is through potential changes in the surface heat flux. In the peak month of Feb(1), composite anomalies of the net surface heat flux for the reanalyses and the CTR experiment are weakly positive over the positive SST anomaly pole, acting to cool the atmosphere (Figs. 7a,b). This partly helps to sustain the positive SLP anomalies in Feb(1). Another possible factor is the meridional SST gradient, through which the atmospheric synoptic eddies develop and affect the stationary eddies with a frequency longer than one month via an eddy–mean interaction (Lau 1988; Hoskins and Valdes 1990). Therefore, we calculated composite anomalies of the meridional SST gradient during the positive events (Fig. 8). As expected, the meridional SST gradient in the observation reaches the minimum in Feb(1) in between the positive and negative SST anomalies and the maximum to the south of the positive SST anomalies (Fig. 8a), representing a meridional dipole pattern of the meridional SST gradient anomaly. The observed meridional dipole pattern is reasonably simulated in the CTR experiment (Fig. 8b), although it is located at a slightly lower latitude.
As in Fig. 5, but for the net surface heat flux anomalies (W m−2). The positive values indicate warming the ocean.
Citation: Journal of Climate 28, 2; 10.1175/JCLI-D-14-00133.1
As in Fig. 6, but for the anomalies of the meridional SST gradient (10−6 °C m−1).
Citation: Journal of Climate 28, 2; 10.1175/JCLI-D-14-00133.1
To qualitatively estimate the impact of the anomalous meridional SST gradient on the atmospheric synoptic eddies, composite anomalies of the poleward heat flux by the synoptic eddies at 925 hPa are computed (Fig. 9). During the austral summer, the poleward heat flux in the reanalyses tends to decrease with a northwest–southeast orientation over the weakened meridional SST gradient (Fig. 9a), whereas the poleward heat flux increases over the strengthened meridional SST gradient. This indicates an anomalous southward shift of near-surface storm tracks. The northwest–southeast-tilted dipole structure of the poleward heat flux anomaly is more dominant in the CTR experiment throughout the austral summer (Fig. 9b). However, in the SIO experiment, the northwest–southeast-oriented dipole pattern of the poleward heat flux anomaly in Feb(1) does not appear as clearly as in the Jan(1) (Fig. 9c).
As in Fig. 5, but for the anomalies of the poleward heat flux by synoptic eddies at 925 hPa (m K s−1).
Citation: Journal of Climate 28, 2; 10.1175/JCLI-D-14-00133.1
The anomalous southward shift of the storm tracks associated with the synoptic eddies is also found in the upper troposphere, where composite anomalies of the meridional wind variance by the synoptic eddies at 250 hPa are calculated (Fig. 10). Both the reanalyses and the CTR experiment results show a steady decrease and increase in the meridional wind variance poleward of the weakened and strengthened meridional SST gradient, respectively (Figs. 10a and 10b). This suggests that throughout the summer season, the storm tracks in the upper troposphere anomalously shift southward in a similar manner as those near the surface at 925 hPa. In the SIO experiment, the maximum amplitude of the negative anomalies in Jan(1) exceeds −120 m2 s−2, but becomes between −80 and −60 m2 s−2 in Feb(1), approximately half as strong as in Jan(1). However, the positive and negative anomalies of the meridional wind variance in Feb(1) suggest that the impact of the SST anomalies on the storm-track activity in the upper troposphere may be weaker than that in the lower troposphere.
As in Fig. 5, but for the anomalies of the meridional wind variance associated with synoptic eddies at 250 hPa (m2 s−2).
Citation: Journal of Climate 28, 2; 10.1175/JCLI-D-14-00133.1
The location and amplitude of the storm tracks in the upper troposphere are strongly related to those of the westerly jet through the eddy–mean interaction (Lau 1988; Hoskins and Valdes 1990). Hence, it is worthwhile to examine to what extent the change in the storm-track activity is associated with that in the westerly jet. Figure 11 shows composite anomalies of the zonal wind at 250 hPa and the E vector associated with the synoptic eddies. Both the reanalyses and the CTR experiment show a weakening of the westerly jet in the region of the weaker meridional wind variance, which is associated with the convergence of the E vector (Fig. 11a). To the south of the weakened westerly jet, the westerly wind strengthens in association with the divergence of the E vector, corresponding well to the stronger meridional wind variance. This indicates an apparent link between variations in the westerly jet and an anomalous southward shift of the storm tracks associated with the synoptic eddies. A similar feature is found in the SIO experiment except in Feb(1) when the westerly jet does not significantly weaken between 35° and 45°S and the associated E vector shows anomalous divergence in the region.
As in Fig. 5, but for the zonal wind anomalies at 250 hPa (color, m s−1) and the E vector (arrows, m2 s−2). The thick arrows indicates anomalies exceeding the 90% (95%) confidence level by a two-tailed t test for the JRA-55 and SIO experiment (CTR experiment).
Citation: Journal of Climate 28, 2; 10.1175/JCLI-D-14-00133.1
To further quantitatively evaluate the contribution from the synoptic eddies to the strengthening of the Mascarene high in its southern part, we calculated composite anomalies for the high-frequency eddy term in Eq. (3) for the upper troposphere, and the term 
As in Fig. 5, but for the anomalous contribution from the synoptic eddies to the tendency anomaly in the geopotential height at 250 hPa (m month−1). The term 
Citation: Journal of Climate 28, 2; 10.1175/JCLI-D-14-00133.1
This local feedback process may operate in Dec(0) and Jan(1), but because of the weak development of the SST anomalies, it appears to be not as strong as the remote influences by other climate variability. To explore the relative importance of the local and remote effects on the Mascarene high variability, composite anomalies of the geopotential height at 250 hPa and the associated wave activity flux were calculated (Fig. 13). All of the reanalyses and the model experiment results show that the positive anomalies in the southern Indian Ocean are associated with the zonally elongated positive anomalies in the midlatitudes and the negative anomalies over Antarctica, indicating a potential link with the AAO, as discussed in previous studies (Hermes and Reason 2005; Huang and Shukla 2008; Morioka et al. 2013, 2014). In particular, the link with the AAO may be stronger in Dec(0) and Jan(1) when the stationary Rossby waves propagate from the negative anomalies over Antarctica to the positive anomalies in the southern Indian Ocean. The wave propagation in Feb(1) does not appear as clearly as in Dec(0) and Jan(1), although the remote influence by the AAO may play a certain role in sustaining the positive geopotential height anomalies in the southern Indian Ocean. These results suggest that in Dec(0) and Jan(1), the meridional SST gradient anomaly does not develop strongly enough to impact the Mascarene high variability compared to the remote effect, but in Feb(1), the SST anomalies reach their peak because of the shallow mixed layer and those could play an important role in the Mascarene high variability.
As in Fig. 5, but for the geopotential height anomalies (color, m) and the wave activity flux at 250 hPa (arrows, m2 s−2).
Citation: Journal of Climate 28, 2; 10.1175/JCLI-D-14-00133.1
5. Impact on the southern African rainfall
As discussed in previous literature (Reason 2001, 2002; Washington and Preston 2006), the SST variability in the southern Indian Ocean plays an important role in the southern African rainfall variations. For this investigation, composite rainfall anomalies during the positive events are calculated in Fig. 14. Both the observation and the CTR experiment show a significant rainfall increase over southern Africa throughout austral summer (Figs. 14a and 14b). In contrast, the SIO experiment shows a significant rainfall decrease in Dec(0) and Jan(1) and a slight increase in Feb(1) (Fig. 14c). There exists a remarkable difference in early summer rainfall between the CTR and SIO experiments.
As in Fig. 5, but for the rainfall anomalies (color, mm month−1) and CMAP data are shown in (a). Anomalies exceeding the 90% (95%) confidence level by a two-tailed t test are shown for the CMAP and SIO experiment (CTR experiment).
Citation: Journal of Climate 28, 2; 10.1175/JCLI-D-14-00133.1
To examine the difference in the early summer rainfall anomalies between the model experiments, composite anomalies of the moisture flux integrated from 300 hPa to the surface and their divergence/convergence are computed (Fig. 15). Although both the CTR and SIO experiments show anomalous strengthening of the Mascarene high in its southern part in Dec(0) and Jan(1) (Figs. 5b and 15c), the moisture flux toward the southern Indian Ocean in the CTR experiment (Fig. 15b) is stronger than that in the SIO experiment (Fig. 15c) and causes anomalous convergence over southern Africa. This suggests a potential impact of the local SST anomalies on the southern African rainfall. Indeed, in Dec(0) and Jan(1) of the CTR experiment, the moisture flux anomalies emanate from the southern Indian Ocean through the south of Madagascar, where the positive SST anomalies are located (Fig. 6b). However, in Feb(1), both the CTR and SIO experiments show an anomalous convergence of the moisture flux over southern Africa corresponding to the anomalous rainfall increase (Figs. 14b and 14c). Thus, the anomalous moisture supply from the positive SST anomalies as well as the Mascarene high variability could be important for the early summer rainfall increase over southern Africa.
As in Fig. 5, but for the moisture flux anomalies (arrows, kg m kg−1 s−1) integrated from 300 hPa to the surface and their divergence/convergence (color, 10−7 kg kg−1 s−1). Anomalies exceeding the 90% (95%) confidence level by a two-tailed t test are shown for the JRA-55 and SIO experiment (CTR experiment).
Citation: Journal of Climate 28, 2; 10.1175/JCLI-D-14-00133.1
6. Summary and discussion
A set of CGCM experiments are performed to investigate the role of the local SST variability in the summertime Mascarene high variability, which strongly affects rainfall variability over southern Africa (Behera and Yamagata 2001; Reason 2001, 2002; Washington and Preston 2006). The control (CTR) experiment reasonably simulates the dominant variability of the Mascarene high with the persistence of the SLP anomalies in its southern part throughout the austral summer. On the other hand, in the southern Indian Ocean (SIO) experiment, where the interannual SST variability in the southern Indian Ocean is suppressed with the model climatology, the SLP anomalies persist over December–January, but significantly weaken in February.
A careful comparison between CTR and SIO experiments reveals a potential SST feedback supporting the persistence of the SLP anomalies as summarized in a schematic diagram (Fig. 16). In the case of the positive events, the SST anomalies in the southern Indian Ocean represent a meridional dipole pattern with the positive SST anomalies to the south of the southern Indian Ocean and the negative SST anomalies to the north. This leads to the weakening of the meridional SST gradient to the north of the positive SST anomalies and its strengthening to the south. As a result, the storm-track activity associated with the synoptic eddies in the lower and upper troposphere weakens poleward of the weaker meridional SST gradient and it strengthens poleward of the stronger meridional SST gradient, typical of the anomalous southward shift of the storm tracks. This is a favorable condition for strengthening the Mascarene high in its southern part, but this local ocean–atmosphere feedback effectively operates in February when the meridional dipole of the SST anomalies reaches its peak. In contrast, during December and January, the remote effect arising from the AAO may play a major role in strengthening the Mascarene high in its southern part as discussed in previous studies (Hermes and Reason 2005; Huang and Shukla 2008; Morioka et al. 2013, 2014).
Schematic diagram showing the sequence of the local SST impacts on the summertime Mascarene high variability. The deviation from the monthly climatology is described. The local ocean–atmosphere feedback effectively operates in February, when the meridional dipole of the SST anomalies reaches its peak.
Citation: Journal of Climate 28, 2; 10.1175/JCLI-D-14-00133.1
Furthermore, the possible impact of the local SST anomalies in the austral summer rainfall over southern Africa is found. In early summer, the strengthening of the Mascarene high in its southern part and the associated SST anomalies provide more moisture from the southern Indian Ocean to southern Africa and enhance the southern African rainfall. Our CGCM result in the western part of the southern Indian Ocean is qualitatively consistent with the AGCM experiment by Washington and Preston (2006), in which the meridional SST anomalies prescribed in their AGCM experiment are located mainly in the western part.
Since the rainfall variability over southern Africa during the austral summer has huge socioeconomic impacts such as changes in major crop yields, it is strongly required for local society to skillfully predict the rainfall variability in advance. For societal needs, an accurate prediction of the summertime Mascarene high variability is indispensable, but the results here provide further requirements for precisely predicting the SST anomalies in the southern Indian Ocean. However, most previous studies, to date, paid more attention to remote influences arising from ENSO and the AAO for inducing the Mascarene high variability, which generates the dipole SST anomalies associated with the IOSD (Huang and Shukla 2008; Morioka et al. 2013, 2014; Yuan et al. 2014). In this regard, this study provides a useful insight into the important role of the local SST variability in the summertime Mascarene high variability, and, hence, its unique influence on southern African climate.
Acknowledgments
The SINTEX-F2 was run on the Earth Simulator 2 at the Japan Agency for Marine-Earth Science and Technology (JAMSTEC). Constructive comments from Drs. Toshio Yamagata and Mototaka Nakamura, and Prof. Hisashi Nakamura helped improve the manuscript. Also, the authors thank three anonymous reviewers for their helpful comments. The present 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). The first author is supported by a research fellowship from the Japan Society for the Promotion of Science (JSPS).
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