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Southerly Surges on Submonthly Time Scales over the Eastern Indian Ocean during
the Southern Hemisphere WinterYOSHIKI FUKUTOMI
Frontier Research Center for Global Change, JAMSTEC, Yokohama, JapanTETSUZO YASUNARI
Hydrospheric Atmospheric Research Center, Nagoya University, Nagoya, and Frontier Research Center for Global Change,
JAMSTEC, Yokohama, Japan(Manuscript received 5 June 2004, in final form 18 October 2004)
ABSTRACT
Meridional wind surges from the extratropics into the Tropics strongly regulate tropical convective
activity. This paper confirms that extratropical forcing manifested as a meridional surge does modulate thetropical atmosphere over the eastern Indian Ocean, and it describes the tropical–extratropical connectionin the region.
Surges in the lower atmosphere on submonthly (6–25 days) time scales over the eastern Indian Ocean
were examined in tandem with associated tropical convection and large-scale atmospheric fields during theSouthern Hemisphere (SH) winter (June–August). Data used in this study are NCEP-2 reanalyses and daily
NOAA/Climate Diagnostics Center (CDC) outgoing longwave radiation (OLR) data for 23 yr, from 1979
to 2001. A low-level surge index was calculated using the 850-hPa meridional wind component (␷) averaged
over a region where sub–monthly scale ␷ variance shows a local maximum (17.5°–2.5°S, 87.5°–97.5°E). The
surge index defines 62 different surge events. Composites of various components were generated based on
the index to define relationships between surge events and large-scale fields.
Low-level southerly surges over the eastern Indian Ocean originate from midlatitude Rossby waves withstrong baroclinic development in the entrance region of a subtropical jet core off Australia’s west coast.
Strengthened low-level wind surges cause cross-equatorial flow stretching from the subtropical eastern
Indian Ocean to the southern Bay of Bengal. Surges are accompanied by the advection of cold, dry air frommidlatitudes into the Tropics. A cold and dry front develops at the leading surge edge during the surge
period. Two to four days later, as the surge peaks, negative OLR anomalies develop near the key region.
The OLR anomalies indicate a local blow up of convection over the tropical eastern Indian Ocean. Convection reflects increased instability in the surge region, which is caused by low-level dry air advection andnear-surface moistening that is forced by enhanced sea surface evaporation associated with the surge. The
southerly surge on submonthly time scales is an important bridge linking the Tropics and midlatitudes overthe Indian Ocean.

1. Introduction
Meridional wind surges from the extratropics into the
Tropics strongly regulate tropical convective activity.
Over the last two decades, numerous studies have described the influence of cold surges on large-scale
Corresponding author address: Dr. Yoshiki Fukutomi, Frontier
Research Center for Global Change, Yokohama Institute for
Earth Sciences, Japan Agency for Marine-Earth Science andTechnology (JAMSTEC), 3173-25 Showa-machi, Kanazawa-ku,
Yokohama, Kanagawa 236-0001, Japan.
E-mail: fukutomi@jamstec.go.jp© 2005 American Meteorological Society

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fields. These studies include pioneering works after theWinter Monsoon Experiment (WMONEX) in 1978–79,
which established the fundamental relationships between the east Asian–North Pacific cold surges andconvective activity over the tropical Pacific. Midlatitude wave propagation and cold surges have beenshown to modulate tropical convective activity in theAsia–Pacific region during the Northern Hemisphere(NH) winter (e.g., Murakami and Sumi 1981; Pan andZhou 1985; Kiladis et al. 1994; Meehl et al. 1996; Slingo
1998; Compo et al. 1999). These studies have shownthat a high-frequency intraseasonal time scale shorter

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than 30 days is a dominant time scale of tropical–extratropical interactions associated with east Asian cold
surges. This time scale is conventionally called submonthly. Similar cold surges linking the Tropics andextratropics also occur over the eastern South Pacificand South America during the Southern Hemisphere(SH) winter and summer (e.g., Garreaud and Wallace
1998; Liebmann et al. 1999; Marengo et al. 2002).
A series of works by Murakami and collaborators(Murakami 1987; Wang and Murakami 1987; Shresthaand Murakami 1988; Murakami 1988a,b; Murakamiand Sumathipala 1989) found low-level equatorward
wind surges on a variety of spatial and temporal scales
over the Indian Ocean and South Pacific. Their main
focus was placed on lower-frequency intraseasonal timescales, the so-called Madden–Julian oscillation (MJO;
Madden and Julian 1994) or a 30–60-day time scale,rather than higher-frequency time scales. In particular,
the former four works showed strong southerly surges
at 30–60-day time scales originating over the midlatitude Indian Ocean region that enhanced MJO-scaleconvection over south Asia. The latter two studies, Murakami (1988b) and Murakami and Sumathipala (1989),
also investigated higher-frequency fluctuations such as
the submonthly and synoptic time scales. These analyses showed that equatorward surges over the SH oceanscoincided with transient westerly bursts and enhanced
convection in equatorial regions. Notably, highfrequency equatorward surges will influence tropicalcirculation in purely oceanic regions. Unfortunately,
their discussions were limited for the local relationships
between the meridional surge and the westerly burst on
these shorter time scales in the Tropics.
From a viewpoint of large-scale tropical–extratropical connection over the Indian Ocean, some important
implications of extratropical forcing of tropical atmosphere were presented by also other authors. For example, Yasunari (1981a,b) suggested that the SH midlatitude westerly disturbance cause cold air penetration
into the tropical Indian Ocean region, which can play arole in initiating cloud activity associated with thelower-frequency (30–40 days) oscillation of the Indiansummer monsoon. On the other hand, Rodwell (1997)
proposed the mechanism for a midlatitude injectionthat acts as the triggering of a break in Indian monsoon
rainfall. His results, based on general circulation modelexperiments, showed that the injection of dry air produced by the SH midlatitude weather system occurs
through the low-level cross-equatorial jet along thewestern flank of the Mascarene high and leads to dry
conditions over India.
However, the structure and evolution of low-level
wind surges at submonthly time scales over the Indian

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Ocean and their influence on larger scales have notbeen as well defined as the structure, evolution, andinfluence of equatorward surges over east Asia and thewestern Pacific, and over the South Atlantic and SouthAmerica. The tropical–extratropical connection overthe Indian Ocean at submonthly time scales remains
poorly understood.
This study examines the structure and evolution of
low-level southerly surges and their influence on largescale circulations and convective activities over the Indian Ocean on submonthly time scales. A preferred
location for an equatorward low-level wind surge willbe identified for the SH winter, and a cause–effect relationship will be established between the surge, thetropical atmospheric conditions, and the midlatitude
systems in the location. During the SH winter, midlatitude disturbances at this high-frequency range favorably propagate along comparatively lower latitudes that
correspond to the region of the SH subtropical jet (e.g.,
Ambrizzi et al. 1995; Kiladis and Mo 1999). Thus, aclose connection between the midlatitude systems andtropical variability can be expected also over the Indian
Ocean. The purpose of the present study is to confirm
that sub–monthly scale extratropical wave forcing manifested as a meridional surge modulates the tropical atmosphere and to establish an associated tropical–extratropical connection over the Indian Ocean during theSH winter.
Data sources and the analysis procedures for this paper are described in section 2. Section 3 includes a statistical analysis that defines the basic characteristics ofthe sub–monthly scale surge activity at a low level in theatmosphere. In section 4, composite analysis results describe the space–time evolution of the large-scale fields
associated with the surge. Section 5 then shows evidence of the possible impacts of the surge on the tropical atmosphere. Surge characteristics and related aspects are discussed in section 6, followed by conclusions
in section 7.

2. Datasets and processing
Two datasets were used in this study to characterize
the large-scale aspects of atmospheric circulation andconvection associated with low-level wind surges. The
National Centers for Environmental Prediction (NCEP)–
Department of Energy (DOE) Atmospheric Model Intercomparison Project (AMIP-II) reanalysis (hereafterR2) dataset (Kanamitsu et al. 2002) includes zonal andmeridional wind components (u, ␷), temperature (T ),specific humidity (q), sea level pressure (SLP), geopotential height (Z ), and surface latent and sensible heat
fluxes (LHF and SHF). An interpolated outgoing long-JUNE 2005

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wave radiation (OLR) dataset provided by the National
Oceanic and Atmospheric Administration (NOAA)/
Climate Diagnostics Center (CDC) was also used. Alldata records include daily averages for the 23-yr period(1979–2001) at global coverage on a 2.5° latitude–
longitude grid. In addition, streamfunction (␺), relative
vorticity (␨), and divergence (Div) were computed fromthe R2 winds using a spherical harmonic transform(SPHEREPACK 3.0; Adams and Swarztrauber 1999).
Potential vorticity (PV) on an isentropic surface wascalculated from relative vorticity and temperature inthree dimensions using the method described in Tomas
and Webster (1994).
Both filtered and unfiltered anomalous time series at
each grid point capture the space–time structure of thelarge-scale fields. Unfiltered daily perturbations of thevariables were computed by removing the 23-yr climatological seasonal cycle from the raw 365-day timeseries for each year. The climatological seasonal cycle
is defined as the sum of the 23-yr mean and the first
three harmonics of the daily climatological mean values
at each grid point. Submonthly (6–25-day) perturbations were computed by applying a Butterworth filter(Kaylor 1977) to the unfiltered perturbations. This
bandpass filter has a sharp cutoff and effectively reduces the effect of both higher-frequency (synoptic)
and lower-frequency intraseasonal (MJO or about 30–
60-day band) fluctuations. The focus of this study is theSH winter; analysis results are given for the SH winterseason, 1 June through 31 August [June–July–August(JJA)]. The reason is that significant extratropical–
tropical connections are expected associated with disturbance activity embedded in the SH midlatitude westerly flow during this season, as noted in section 1.

3. Surge activity over the Indian Ocean
a. Sub–monthly scale surge and transient eddy
activity
Figure 1 illustrates general features of variability at
submonthly time scales in the target region. Three variance statistics products are given for JJA. Daily 850hPa ␷ is a key measure of low-level surge activity. Figure 1a shows the geographical distribution of variancefor the 6–25-day filtered ␷ at 850 hPa and its percentage
of subseasonal total variance of unfiltered ␷. Local
maximum values of variance exceeding 4 m2 s⫺2 overlap the ratio exceeding 45% over the tropical eastern
Indian Ocean southwest of Sumatra. This high variability can be viewed as low-level meridional surge activity.
The box (17.5°–2.5°S, 87.5°–97.5°E) enclosing this local
maximum shows the key region (outlined in the figures)
used for making a surge index to construct compositesof various elements shown later. The surge index isdefined from the average of the 6–25-day filtered 850hPa ␷ time series over this key region. The variance andpercentage variance of the 6–25-day filtered OLR, as inVincent et al. (1998), are also presented in Fig. 1b to
confirm the sub–monthly scale convective activity. Comparatively high values, up to 700 W2 m⫺4, are located in
the surge key region, which suggests a possible relationship between tropical convection and surge activity
over the tropical eastern Indian Ocean. Disturbance
activity in the upper troposphere is commonly measured by the perturbation (eddy) kinetic energy (PKE)(e.g., Matthews and Kiladis 1999). Here, PKE is derived from the 6–25-day filtered eddy components of uand ␷, that is,1
PKE ⫽ 共u⬘2 ⫹ ␷⬘2兲,2

共1兲

where the overbar denotes the time mean, and theprime is the eddy component. The PKE is displayedwith the SH winter-mean u component (Fig. 1c). The
most striking feature is a band of maximum PKE exceeding 70 m2 s⫺2 that extends along the winter-mean
subtropical jet from the Indian Ocean through theSouth Pacific. Sub–monthly scale eddy activities aremost amplified near the subtropical jet core during theSH winter, which is consistent with Rossby wave activities with similar SH winter time scales investigated byAmbrizzi et al. (1995).

b. Surge index characteristics
Figure 2a shows the ensemble power spectrum of theunfiltered time series of 850-hPa ␷ averaged over thesurge key region for the 23 winters. Meridional surge
activity over the eastern Indian Ocean has a dominant
submonthly time scale. A fast Fourier transform (FFT)
method estimated the power spectra for 128 days (14May–18 September) centered on each individual winter(JJA, 92 days). A 99.9% significant peak at 14.4 days
and associated larger power in submonthly ranges areclearly evident. In the winter of 2000 (Fig. 2b), a significant peak and its associated prominent power occur
in the submonthly frequency range. Results in Fig. 2
support the existence of surges at submonthly scales
and justify applying a bandpass filter to the time series.
Figure 2c shows the 6–25-day filtered and unfilteredsurge indices during JJA 2000, an unfiltered time series
that corresponds to the power spectrum in Fig. 2b.
These two time series are very similar; the 6–25-day
filtered series is used as the southerly surge index tocomposite fields. Quasi-periodic, low-level wind

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FIG. 1. (a) Variance of 6–25-day filtered meridional wind at the 850-hPa level (contours) and its percentage of
total variance (shadings) for JJA (SH winter) 1979–2001. Contour interval is 2 m2 s⫺2. The surge key region is
enclosed by the grid box of 17.5°–2.5°S, 87.5°–97.5°E. (b) Same as (a) except for OLR. Contour interval is 100 W2
m⫺4. (c) Mean zonal wind (contours) and 6–25-day PKE (shadings) at the 200-hPa level during the same period.
Thin solid (dashed) contours indicate positive (negative) values. Thick solid contours are 0 values. Contour interval
is 5 m2 s⫺2.

maxima occur with approximately 10- to 20-day periodsin the time series. Positive extrema (denoted by inverted triangles) that exceed climatology by 1.5 standard deviations were selected as southerly surge peaks.
Further details of the compositing technique are included below.

4. Space–time evolution of southerly surges andlarge-scale fields
a. Composite method
In this section, compositing techniques are used toestablish relationships between southerly surges andJUNE 2005

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FIG. 2. (a) The 23-winter ensemble power spectrum of 850-hPa meridional wind averaged over the surge key
region. Thin solid curve is the 99.9% confidence level. Dashed curve is the red noise spectrum. (b) Same as (a)except for 2000 winter case. (c) The 6–25-day filtered (thick solid line) and unfiltered (thick dashed line) 850-hPa
meridional wind time series over the surge key region (surge index) for JJA 2000; 23-winter climatological 1.5 std
dev (thin dashed lines), surge peaks (reversed triangles), and antisurge peaks (triangles) are also shown.

large-scale atmospheric conditions. Southerly surge
peaks, as displayed in Fig. 2c, are referred to as day-0
phases; daily lag composites for various elements forday ⫺5 through day ⫹5 are constructed. A total of 62
surge events were identified during the 23 SH winterseasons, and space–time composites were defined byboth filtered and unfiltered anomalies. A Student’s tstatistic procedure (two-tailed t test) for the two-sample
case determined the local statistical significance of thecomposites. The procedure is almost identical to that in
Murakami (1987) and requires composite mean differences between paired reversal phases at each timestage. To do this in the present study, a total of 58
negative surge index peaks (as shown by triangles inFig. 2c) were detected, and the composite differences
between the positive and negative surges were tested
for statistical significance. Additionally, critical values
for area significance were determined from Monte
Carlo experiments to confirm spatial significance of thecomposite maps (cf. Murakami 1987). Then, the spatial
significance of the composite patterns was tested using
these critical values.

b. Horizontal structure of southerly surge
Figure 3 shows the horizontal flow structure of the
low-level surge at 850 hPa. Left panels (Figs. 3a–c)
show composites of total wind vectors and the meridional wind component (␷: shading and contour) for days
⫺3, 0, and ⫹3. At day 0, southeasterly–southerly flow
penetrates from the SH subtropics into the southern
Bay of Bengal over the eastern Indian Ocean. This
surge flow converges with Asian monsoon westerly
flow that extends from the Arabian Sea to the Philippines. The meridional surge flowing into the NH is notpresent 3 days before and after the peak surge.
Right panels (Figs. 3d–f) show similar composites
with unfiltered ␷ anomalies. Unfiltered anomalies were

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FIG. 3. (a)–(c) Composites of 850-hPa total wind vectors and meridional wind component (shadings and dashedcontours) at day ⫺3, day 0, and day ⫹3 based on the surge index. Positive (negative) values are shadings (contours),
and 0 values are omitted. Contour interval is 2 m⫺1. (d)–(f) Composites of unfiltered meridional wind anomalies
at the same timing. Solid (dashed) contours are positive (negative) values. Contour interval is 0.5 m s⫺1. Positive(negative) anomalies statistically significant at the 99.9% level are darkly (lightly) shaded.

computed by subtracting the seasonal cycle as noted in
section 2. An east–west-oriented wave pattern of ␷
anomalies and a clear penetration of a positive ␷ (southerly) anomaly into the southern Bay of Bengal is evident on day 0. The positive ␷ anomaly propagates eastward across the Indian Ocean from day ⫺4 through day⫹2 (not shown) and decays by day ⫹3.
Composites of the 6–25-day filtered ␷ anomalies (notshown) show patterns similar to those of the unfilteredanomalies. The eastward-propagating ␷ anomalies contribute an oscillating off–on–off structure to the crossequatorial flow that connects the SH midlatitude westerly flow to the NH monsoon westerly flow in the lower
troposphere. These quasi-periodic surges over the eastern Indian Ocean might play an important role in meridional airmass exchange between the two hemi-

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spheres, as is shown in the case of the east Asian wintermonsoon surge by Carrera and Gyakum (2003).

c. Midlatitude waves and tropical convection
The eastward propagation of low-level ␷ anomalies(Figs. 3d–f) suggests that midlatitude waves induce
equatorward surges over the eastern Indian Ocean. In
addition, the low-level surges can alter the tropical atmospheric condition, which may lead the occurrence oftropical convection. The evolution in space and time forthe broader domain of the large-scale circulation andthe convection associated with surges are presented
here to capture the cause and effect of southerly surges.
Figure 4 illustrates the relationships between thesoutherly surges, low-level circulation, and convective
activity using composite maps of the 6–25-day filtered
850-hPa wind vectors, streamfunction, and OLR
anomalies every 2 days from day ⫺2 to day ⫹4. Composite wind vectors were tested statistically, and only
those vectors for which the u or ␷ component exceeded
the 95% confidence level were plotted. In addition,
OLR values less than ⫺5 W m⫺2 were shaded.
From day ⫺2 through day ⫹2, a wave with cyclonic
and anticyclonic circulation anomalies propagatesalong midlatitudes from the southwestern Indian
Ocean into the South Pacific. This wave appears to
produce the southerly surge flow as it amplifies. At day⫺2 (Fig. 4a), a wave train extends from south of Africa
to south of Australia. Southwesterly and southerly wind
anomalies, which are an early surge stage, appear in the
key region; these wind anomalies originate from themidlatitude wave train.
At day 0 (Fig. 4b), the surge peaks when strengthened southerly winds cross the equator and reach thesouthern Bay of Bengal as part of a meridionally elongated trough–ridge couplet over the Indian Ocean. The
midlatitude wave train that includes this trough–ridge
couplet is well organized at day 0. Two days after the
peak (day ⫹2; Fig. 4c), the anomalous surge flow maintains its strength, but the direction shifts from southerly
to southeasterly in association with the eastward traveling midlatitude trough–ridge pair. At the same time, anegative OLR anomaly begins to spread over the eastern equatorial Indian Ocean west of Indonesia.
By day ⫹4 (Fig. 4d), the wave train and the southerly
surge have decayed, but the negative OLR anomaly
reaches a maximum over the key region. This OLR
signal reflects enhanced convection that is likely triggered by the surge. The location of the OLR anomaly is
consistent with that of the local maximum of OLR variance in the surge key region (Fig. 1b). The strengthened
southerly surge could create favorable atmosphericconditions for local convective development in the key

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region. The surge-related change in the tropical atmosphere is discussed in section 5.
Another notable feature in the composite fields is thewestward propagation of low-level westerly wind
anomalies north of the equator. Between days ⫺2 and⫹4, significant westerly winds propagate westward
from the South China Sea to the Arabian Sea between
5° and 15°N. The westerly wind signal is on the southern flank of an anomalous cyclonic circulation that is
associated with negative OLR anomalies. This westward-moving disturbance may be similar to those discussed by Fukutomi and Yasunari (1999, 2002). In addition, the confluence of the cross-equatorial southerly
surge flow and the monsoon westerly flow at day 0 in
the anomaly map (Fig. 4b) coincides with the flow inthe total wind map (Fig. 3b). Development of low-level
westerly winds over the Bay of Bengal is coincident
with cross-equatorial southerly flow. This is similar to
the relationships between the east Asian cold surge andsubsequent development of transient westerly windbursts over the western Pacific (e.g., Murakami and
Sumathipala 1989; Kiladis et al. 1994; Meehl et al. 1996;
Compo et al. 1999).
The behavior of the upper-level circulation associated with the evolving low-level surge was investigatedto clarify the connection between the surge and midlatitude wave development. Figure 5 shows composite
anomalies of 6–25-day filtered 200-hPa wind vectors,
streamfunction, and OLR, similar to those in Fig. 4.
However, OLR anomalies are mapped only for regionswith 99.9% statistical significance, and negative (positive) anomalies are darkly (lightly) shaded.
The composite sequence shows a wave train thatstretches from the southern Indian Ocean to the South
Pacific, with eastward propagating wave anomalies andwave packets. The dominant zonal wavenumbers are
5–7. As the wave packet moves eastward across Australia during the surge period (day ⫺2 through day ⫹2),
the wave anomalies grow downstream, suggesting
Rossby wave energy propagation to the east. An arc of
anticyclonic and cyclonic circulation centers curving toward the equator is nearly superposed on the JJA mean
subtropical jet over the eastern Indian Ocean, Australia, and the South Pacific (Fig. 1c). At the surge peak(day 0; Fig. 5b), an amplified trough–ridge couplet overthe eastern Indian Ocean shifts westward about 10°
relative to corresponding features at 850 hPa; this structure is consistent with a baroclinic midlatitude wave.
The low-level southerly surge could therefore be related to baroclinic development of the eastward-propagating wave in the entrance region of the SH subtropical westerly jet core over the eastern Indian Ocean.
This is analogous to the east Asian cold surge that oc-

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FIG. 4. Composites of 6–25-day filtered OLR, 850-hPa wind vectors and streamfunction anomalies from day ⫺2 through day ⫹4 with 2-day interval based on the surge
index. OLR anomalies less than ⫺5 W m⫺2 are shaded. Solid (dashed) contours are
positive (negative) streamfunction values. Zero contours are omitted. Contour interval
is 5.0 ⫻ 105 m2 s⫺1. Only 95% statistically significant vectors are plotted.

curs in the entrance region of the Asian–Pacific subtropical jet during the NH winter as illustrated with aconceptual model by Slingo (1998). At the mature stageof convection in the key region (day ⫹4), the wave train
weakens, but it still maintains the arc-like structure withnew anticyclonic circulation growth at its leading edge
over the South Pacific.

Characteristics of the upper-level waves in the present study agree with those in Ambrizzi et al. (1995).
The motion of the anomalous troughs and ridges of the
wave train (Fig. 5) is controlled by a waveguide along
the subtropical and subpolar westerly jets for the SH
winter. The waves follow an interesting route in thepresent case. The wave pattern moves northeastward

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FIG. 5. Same as Fig. 4 except for 200-hPa wind fields and 99.9% statistically significant
OLR anomalies. Contour interval is 1.0 ⫻ 106 m2 s⫺1. Negative (positive) OLR anomalies
statistically significant at the 99.9% level are darkly (lightly) shaded.

into the Tropics on the upstream side over the Indian
Ocean and the Australian continent; it then moves
southeastward into midlatitudes on the downstreamside over the South Pacific. The turning point for thewaves is over the Coral Sea east of Australia (near
20°S, 165°E). Overall, the behavior of waves in thecomposite sequence is consistent with waves trapped
inside the SH winter waveguide and reflected at a criti-

cal line at the boundary between the subtropical westerlies and tropical easterlies near 10°N (Fig. 1c).
Significant negative OLR anomalies that are associated with the leading edge of the southwest–northeastoriented upper-level wave train develop in the subtropical South Pacific. From day ⫺2 through day ⫹2,
OLR anomalies ahead of a trough east of Australia
expand eastward as the wave train propagates down-

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stream. This OLR signal represents convective activity
at submonthly scales in the South Pacific convergencezone (SPCZ) region in conjunction with the midlatitude
wave activity (e.g., Schrage and Vincent 1996).

5. Impacts of the southerly surge on the tropicalatmospherea. Cold and dry air advection from midlatitudes
into the Tropics
Basic circulation structures and convective signaturesassociated with low-level southerly surges have been
identified. Figure 6 shows a sequence of 925-mb temperature and specific humidity anomaly composites to
elucidate changes in the tropical lower-troposphereconditions during surges.
Composites of temperature anomalies (Figs. 6a–d)
reveal northeastward movement of significant cold air
anomalies from midlatitudes into the Tropics during
the surge period (from day ⫺2 through day ⫹2). Lowlevel circulation structures (Fig. 4) suggest that amplification of a band of cold air (negative anomalies) coincides with the enhancement of the southerly andsoutheasterly surge flow along the eastern flank of ananomalous anticyclone. As the midlatitude wave propagates eastward, the band of cold air is pushed toward
the Tropics. At the time of the surge peak (day 0; Fig.
6b), the band of cold air is over the eastern Indian
Ocean and prevails over the surge region. By day ⫹2, anorthwest–southeast-oriented cold front is establishedover the tropical and subtropical eastern Indian Ocean.
After day ⫹4, the northeastward movement terminates,
and the cold anomaly weakens. Composite maps suggest that the coincident development of the cold frontand southerlies characterizes an equatorward cold
surge.
Specific humidity fields (Figs. 6e–h) show the formation of dry (negative) anomalies by the surge peak andthe development of moist (positive) anomalies afterthat in the southerly surge region. By day 0, drying has
approached the key region and a dry front has formed
along the southerly and southeasterly surge (Fig. 4)
over the tropical and subtropical eastern Indian Ocean.
Here, a series of composite maps including day ⫺1phase (not shown) reveals that a weak preexisting dry
anomaly in the key region mixes into the dry front. At
the same time, a moist anomaly expands southwardfrom north of the key region. From day ⫹2 through day⫹4, a northwest–southeast elongated band of moist air
anomalies prevails as the dry front erodes in the Tropics. This widespread moistening might produce convectively unstable conditions in the tropical eastern Indian
Ocean region. This increase in low-level moisture may

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result from enhanced sea surface evaporation driven byincreased surface winds in the southerly surge.
Maps of low-level temperature and humidity anomalies clearly depict the equatorward advection of coldand dry air in the southerly surge that originates in
midlatitudes. Transport of midlatitude air into the
Tropics during the surge is further confirmed by the
behavior of PV anomalies on an isentropic surface using Lagrangian interpretation of airmass movement.
Figure 7 shows the development of PV anomalies on
the 315-K potential temperature surface from day ⫺1through day ⫹2. This isentropic surface is in the middle
to upper troposphere (500–200 hPa) at midlatitudes
poleward of 30°S and in the midtroposphere (700–500hPa) at low latitudes equatorward of 30°S during JJA(not shown).
The pattern of PV anomalies in midlatitudes is characterized by a well-defined wave train structure as in
the circulation fields of Figs. 4 and 5. Recall that negative (positive) PV anomalies have cyclonic (anticyclonic) vorticity in the SH. The map sequence showsan anomalous PV trough extending toward the Tropics
as the midlatitude wave propagates downstream. At
day 0, a tongue of negative PV anomaly extends equatorward along the northeastern flank of the positive
PV anomaly. At day ⫹1, a streamer of negative PV
anomaly oriented northwest-southeast stretches farther
northwestward. On day ⫹2, a PV anomaly of less than
⫺0.03 PVU becomes cut off from the midlatitude wave
train; this isolated cyclonic PV anomaly subsequently
drifts westward around 20°S. Features in the PV evolution corroborate evidence of a midlatitude airmass
intrusion into the lower latitudes.

b. Convectively unstable conditions in the Tropics
Surge-induced evaporation from the sea surface may
be an important moisture source for the lower troposphere during the period between the surge peak andthe convective peak. Composite lead–lag relationships
of a variety of components will be compared to elucidate the roles of surge flow in preconditioning the convective environment. Figure 8 shows space–time structures of sea surface LHF and SHF anomalies presented
in a similar way to the low-level temperature and moisture fields in Fig. 6. The evolution of LHF anomalies(Figs. 8a–d) shows that a positive LHF anomaly(evaporation) occurs coincident with the southerly
surge core (Figs. 3 and 4) over the tropical and subtropical eastern Indian Ocean. The LHF in the surge
region maximizes from day 0 through day ⫹2, contributing to the local moisture accumulation preceding theconvection (Figs. 6e–h). The SHF anomalies (Figs. 8e–

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1647

FIG. 6. (a)–(d) Composites of 925-hPa air temperature anomalies from day ⫺2 through day ⫹4 with 2-day interval based on the surge
index. Solid (dashed) contours are positive (negative) values. Zero contours are omitted. Contour interval is 0.3 K. Negative (positive)
anomalies statistically significant at the 99.9% level are darkly (lightly) shaded. (e)–(h) As in (a)–(d) but for 925-hPa specific humidity
anomalies. Contour interval is 0.2 g kg⫺1.

h) mimic the behavior of the LHF anomalies, but theirmagnitude is considerably smaller; thus, the SHF
anomalies contribute less to the net surface heat budget. Strengthened surge winds, increased near-surface
humidity deficit and air–sea temperature differences

could enhance fluxes into the atmosphere during thesurge period.
Finally, daily lag composites in the surge region wereused to investigate the temporal phase relationship between the surge index (850-hPa ␷), OLR, moist static

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FIG. 7. Composites of isentropic PV (IPV) anomalies on the 315-K potential temperature surface from day ⫺1 through day ⫹2 with
1-day interval. Solid (dashed) contours are negative (positive) values. Zero contours are omitted. Contour interval is 0.03 PVU (10⫺6
K s2 kg⫺1). Negative (positive) anomalies statistically significant at the 99.9% level are darkly (lightly) shaded.

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FUKUTOMI AND YASUNARI

1649

FIG. 8. (a)–(d) Composites of LHF anomalies from day ⫺2 through day ⫹4 with 2-day interval based on the surge
index. Solid (dashed) contours are positive (negative) values. Zero contours are omitted. Contour interval is 20 Wm⫺2. Positive (negative) anomalies statistically significant at the 99.9% level are lightly (darkly) shaded. (e)–(h) Asin (a)–(d) but for SHF anomalies. Contour interval is 5 W m⫺2.

energy, instability index, low-level (850 hPa) divergence, and SLP. Moist static energy h is defined as h ⫽CpT ⫹ gZ ⫹ Lq, where Cp is the specific heat capacityof air, g is the gravitational acceleration, and L is thelatent heat of condensation. The instability index is represented here as the difference in h between 925 and700 hPa [h(925 hPa) ⫺ h (700 hPa)]. Kemball-Cook andWeare (2001) showed that this index is a useful mea-

sure of atmospheric stability. A positive index value
represents convectively unstable atmospheric conditions. Figure 9 shows composite anomalies of the above
variables averaged over a grid box (15°S–0°, 85°–
100°E). This grid box well enclosing the tropical convective signal after the surge peak (Figs. 4 and 5) wasused to construct composite time series of all variables
except the surge index. Surge index and OLR time se-

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MONTHLY WEATHER REVIEW

FIG. 9. (a) Composite evolution of 6–25-day filtered 850-hPa meridional wind (solid) and OLR (dashed) anomalies. OLR anomalies
are averaged over the box 15°S–0°, 85°–100°E. (b) Composite evolution of LHF (solid) and SHF (dashed) anomalies over the samebox with OLR anomalies. (c) As in (b) but for moist static energy
anomalies at the 925- (thin solid) and 700-hPa (dashed) levels, andinstability index (thick solid) anomalies. (d) As in (b) but for 850-hPa
divergence (solid) and SLP (dashed) anomalies.

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ries were bandpass filtered; all other variables were unfiltered except for removal of the annual cycle. It isnoted here that if we use the same box with the surge
key region for constructing composites of above variables, the temporal relationships among those basicallydo not change.
Composites of the surge index and the OLR anomaly(Fig. 9a) show the timing relationship between thesoutherly surge and subsequent convection. An OLR
minimum, which can be regarded as a convective peak,occurs 3 to 4 days after the southerly surge peak (day0). During the mature stage of the surge (day ⫺1 to day⫹1), the LHF anomaly increases and reaches a positive
peak at day ⫹1 (Fig. 9b). The surge forces sea surface
evaporation.
The SHF shows similar variability, but its contribution is small. Instability index changes (Fig. 9c) areroughly in phase with those of the surge index and the
LHF, but with a 1-day lag. After a sharp increase fromday ⫺1 to day 0, instability is sustained from day 0
through day ⫹2, with a peak at day ⫹1. This local build
up of instability followed by a surge is dominated by a
rapid drop in h(700 hPa) and a gradual increase inh(925 hPa). During the mature stage of the surge, therapid decrease of h(700 hPa) is linked to the mid–lower-tropospheric cold and dry air advection associated with the surge; 700-hPa T and q anomalies become
negative at this time (not shown). A concurrent increase of h(925 hPa) (day 0 to day ⫹2) is associatedwith the positive 925-hPa q anomaly shown in Figs.
6e–h. Initial moistening in the near-surface layer isdriven by surge-linked evaporation and resulting
transport of moisture into the convective region. Lowlevel convergence is not dominant during this period(Fig. 9d).
The 850-hPa divergence and SLP anomaly (Fig. 9d)
imply strong convergence from day ⫹2 through day ⫹4,by which time instability is decreasing (Fig. 9c). However, the 925-hPa q continues to increase, reaching amaximum when the low-level convergence peaks. Thus,
low-level convergence at this stage can increase nearsurface q. The peak in convergence, however, occurs 2days after the peak in instability, suggesting that lowlevel convergence is not responsible for the buildup of
instability following the surge, but is linked directly tothe subsequent convective signal captured by the OLR.
Overall, the phase relationships among these variables
highlight the crucial role played by the surge in creating
the low-level instability prior to convection over thetropical eastern Indian Ocean. Low-level instability induced by the surge and the accompanying air–sea interaction can set the stage for subsequent tropical convection.

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FUKUTOMI AND YASUNARI

6. Discussion
Low-level meridional wind variance at submonthly
time scales (Fig. 1a) shows that the tropical eastern
Indian Ocean is a preferred region for equatorward
surges. Factors that allow surge activity in this geographical location are important, but little is known
about the cause of this geographical dependence. Major
meridional surge activities typically occur near continental-scale mountain ranges such as the Tibetan Plateau of Asia, the Rockies of North America, and theAndes of South America (Garreaud 2001). Past studies
have suggested that the prominent topography plays acrucial role in intensifying cold surges and determining
the geographically fixed locations of surge activities
downstream of the topography. Several studies recognized that dynamic interaction between midlatitude
waves developing eastward and the large-scale topography drives low-level equatorward flow (e.g., Murakami and Sumi 1981; Hsu 1987; Murakami and Ho
1981; Garreaud 2000).
In contrast, oceanic surge activity in the present casedoes not depend on upstream large-scale topography;
the basic atmospheric flow over the eastern Indian
Ocean can produce surge activity despite the lack of
topography. One essential feature that controls thispreferred geographical location is the mean subtropicaljet stream. As shown in sections 3 and 4, surge activity
is north of the subtropical jet entrance region over theeastern Indian Ocean off the west coast of Australia in
the SH winter. Present composite analyses suggest thatthe subtropical jet entrance region favors strong baroclinic development of Rossby waves that are carried
eastward by the subtropical jet. Therefore, amplification of the wave within the jet entrance environment
may cause the low-level surge, even in the absence of
upstream topography.
We have emphasized that the presence of the wave
activity along the SH subtropical westerly jet is responsible for the low-level meridional surge activity in theeastern Indian Ocean. However, the results presented
here are derived based on only the SH winter (JJA)
season. The seasonal dependence of the surge activity
related to the location and strength of the subtropicaljet was not examined, and that is open to discussion. As
shown in Fig. 1 of Kiladis and Weickmann (1997), comparatively weaker subtropical jet exists also in the fall
(March–May) and spring (September–November) seasons. Therefore, it seems reasonable to suppose that thesimilar type of tropical–extratropical connection is observed in this region during these seasons. In the summer (December–February) season, that may not be expected because of a lack of the subtropical jet. In ad-

1651

dition, the primary wave propagation path in the SH
midlatitude for the SH summer season (Hsu and Lin
1992) is different from that for the SH winter season(Ambrizzi et al. 1995). It calls for further investigationof the seasonal dependence of the surge activity forthese other seasons.
As moist instability is released, low-level convergence increases until the convective maximum isreached in and around the surge key region. Increasing
low-level convergence can further concentrate moistureand facilitate ascent and convective development. The
mechanism that forces low-level convergence following
the surge peak is still ambiguous. A trough feature eastof the southeasterly surge flow in the SH equatorial
Tropics (Fig. 4c; positive streamfunction values) may
contribute to force initial low-level convergence. The
development and westward extension of an apparent
anomalous cyclone after the surge peak was confirmed
in composites of SLP and 850-hPa z fields (not shown).
This anomalous cyclone is likely an equatoriallytrapped wave signature in response to the surge. Oncethe initial low-level convergence induces ascending motion in convectively unstable conditions, convective
processes in the low-level cyclonic environment may
further promote subsequent growth of low-level convergence. One of the possibilities for causing the initialconvergence after the surge peak may be boundary
layer friction associated with the strengthened surge
flow. Boundary layer frictional effect on the near-surface surge winds may result in the convergence in theboundary layer. Upward motion induced by the boundary layer frictional convergence further reinforces convergence in the lower free troposphere, and that may becapable of producing convection in the moist unstablecondition. Diagnostic analysis of boundary layer flow
applying the model of boundary layer motion (e.g.,
Wang and Li 1994) to our case using observational
near-surface winds is needed for confirmation of thisprocess.
Another important aspect of the surge is the transport of midlatitude dry air into the Tropics. The behavior of a cyclonic PV anomaly in the surge region (Fig. 7)
implies an intrusion of dry air from the midlatitude
midtroposphere into the tropical lower troposphere
over the eastern Indian Ocean. One of the preferred
locations for equatorward dry air intrusion events is theeastern Pacific westerly duct region during the NH winter. A strong equatorward extension of PV accompanied by Rossby wave propagation from the midlatitudes into the Tropics facilitates dry-air intrusion in thisregion (e.g., Kiladis and Weickmann 1992; Tomas andWebster 1994; Kiladis 1998; Waugh and Polvani 2000;
Waugh and Funatsu 2003). Other studies have detected

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MONTHLY WEATHER REVIEW

the dry air intrusion and its potential effect on the tropical atmosphere over the Tropical Ocean Global Atmosphere Coupled Ocean–Atmosphere Experiment(TOGA COARE) domain of the western Pacific (e.g.,
Numaguti et al. 1995; Yoneyama and Fujitani 1995;
Yoneyama and Parsons 1999; Parsons et al. 2000; Yoneyama 2003). Furthermore, Yatagai and Sumi (1998)
noted that both the Pacific and Indian Oceans werelocations for frequent dry air intrusions. However, few
studies have addressed the role of dry air intrusions
over the Indian Ocean.
We have been concerned with the possible impact of
the low-level surge on the subsequent development of
unstable atmospheric condition for convection in theeastern Indian Ocean. One good example that is comparable with the present result can be a case of the dry
air intrusion event that occurred over the TOGACOARE domain in mid-November 1992. This event istreated in the TOGA COARE case studies (e.g., Numaguti et al. 1995; Ding et al. 1995; Yoneyama and Parsons 1999; Parsons et al. 2000). In this case, a dry air
mass that originated in the SH subtropics moved into
the equatorial western Pacific accompanied bystrengthened low-level southerly flow. Based on thisevent, Parsons et al. (2000) proposed a conceptualmodel of a dry intrusion and the following developmentof convection. This model is constructed from consideration of atmospheric boundary layer dynamics andair–sea interaction. The brief description is as follows.
An injection of dry air into the tropical middle andlower troposphere suppresses widespread convection
because of decrease of buoyancy in boundary layer.
Then boundary layer entrainment promotes downward
transport of dry air and momentum, which leads to anincrease of sea surface latent flux into the boundary
layer. In turn, moistening of the boundary layer via thesea surface latent flux occurs. Furthermore, moisteningof the lower troposphere aids in recharge of convectiveenergy. The time for recovery from the extreme dry
condition to the moist condition and the next convective organization is shorter than a week for this case(Fig. 20 of Parsons et al. 2000). It is possible that thissequence of local processes is applicable also to dry air
intrusions over the eastern Indian Ocean. As evidenced
in the previous section, a recovery of low-level moist
unstable condition occurs within a week after the peakof the low-level surge and associated drying. Nevertheless, it would be difficult to confirm existence of thesame boundary layer process from the present results.
Because our composite results do not resolve air motion in the boundary layer, that is a limitation of thisstudy. To confirm whether this conceptual model is applicable or not, dry air intrusions associated with low-

VOLUME 133

level surges over the eastern Indian Ocean should be
investigated in detail with case studies based on satellite
and in situ observational data in the future.

7. Conclusions
This study investigated the large-scale characteristics
of low-level southerly surges on submonthly time scales
over the eastern Indian Ocean during the SH winter.
The tropical and subtropical eastern Indian Ocean is apreferred region of cold surge activity at this high-frequency time scale. Composite analyses highlight thetropical–extratropical links associated with surges andsurge impacts on the tropical atmosphere and give rise
to discussions of surge-related atmospheric phenomenon.
Low-level southerly surges over the eastern Indian
Ocean are caused by strong baroclinic development of
midlatitude Rossby waves in the subtropical jet entrance region. The strengthened low-level winds constitute a cross-equatorial flow that stretches from the subtropical eastern Indian Ocean into the southern Bay of
Bengal; this flow is confluent with the monsoon westerly flow. During the surge, a midlatitude wave train
that is the origin of the surge extends from the southern
Indian Ocean into the South Pacific. Midlatitude waves
show a clear signature of downstream propagation anddispersion along the subtropical jet, suggesting thatthey are trapped within the SH winter waveguide.
Surges that accompany wave passage force advection ofcold and dry air from midlatitudes into the Tropics. The
resulting development of a cold and dry front in thesoutherly surge is prominent during the surge period.
The low-level southerly surge associated with themidlatitude wave passage is an important source oftropical atmospheric variability over the eastern Indian
Ocean. The subsequent organization of tropical convection 2–4 days after the surge peak is an importantsurge impact. The low-level surge forces advection ofcold and dry air and subsequent near-surface moistening owing to the enhanced sea surface evaporation.
This increases the instability in the surge region. Consequently, surge-induced moist instability acts as a trigger for the local flare up of subsequent convection.
The southerly surge on submonthly time scales is an
important bridge that links midlatitudes and the Tropics over the Indian Ocean. Potential vorticity analyses
suggest that the surge is manifested as an intrusion of
dry air from the midlatitude midtroposphere into thetropical lower troposphere over the eastern Indian
Ocean. Further studies exploring physical processes inthe dry-air intrusion would facilitate an understandingof tropical–extratropical connections in this region.

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FUKUTOMI AND YASUNARI

Acknowledgments. The authors would like to thankthree anonymous reviewers for valuable comments that
helped to improve the quality of this paper. Both the
NCEP–DOE AMIP II reanalysis and the NOAA OLR
dataset were obtained from the NOAA/CDC (ftp://ftp.
cdc.noaa.gov). The Generic Mapping Tools (GMT)
graphics package outlined by Wessel and Smith (1995)
was used for drawing all the figures.
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