2914

JOURNAL OF CLIMATEVOLUME 25

Seasonal and Intraseasonal Modulation of Tropical Cyclogenesis Environment overthe Bay of Bengal during the Extended Summer Monsoon
WATARU YANASE AND MASAKI SATOH
Atmosphere and Ocean Research Institute, The University of Tokyo, Kashiwa, JapanHIROSHI TANIGUCHIInternational Pacific Research Center, University of Hawaii at Manoa, Honolulu, Hawaii

HATSUKI FUJINAMI
Hydrospheric Atmospheric Research Center, Nagoya University, Nagoya, Japan(Manuscript received 12 April 2011, in final form 16 September 2011)
ABSTRACTThe environmental field of tropical cyclogenesis over the Bay of Bengal is analyzed for the extended
summer monsoon season (approximately May–November) using best-track and reanalysis data. Genesis
potential index (GPI) is used to assess four possible environmental factors responsible for tropical cyclogenesis: lower-tropospheric absolute vorticity, vertical shear, potential intensity, and midtropospheric relative
humidity. The climatological cyclogenesis is active within high GPI in the premonsoon (;May) and postmonsoon seasons (approximately October–November), which is attributed to weak vertical shear. The
genesis of intense tropical cyclone is suppressed within the low GPI in the mature monsoon (approximately
June–September), which is due to the strong vertical shear. In addition to the climatological seasonal transition, the authors’ composite analysis based on tropical cyclogenesis identified a high GPI signal moving
northward with a periodicity of approximately 30–40 days, which is associated with boreal summer intraseasonal oscillation (BSISO). In a composite analysis based on the BSISO phase, the active cyclogenesis
occurs in the high GPI phase of BSISO. It is revealed that the high GPI of BSISO is attributed to high relative
humidity and large absolute vorticity. Furthermore, in the mature monsoon season, when the vertical shear is
climatologically strong, tropical cyclogenesis particularly favors the phase of BSISO that reduces vertical
shear effectively. Thus, the combination of seasonal and intraseasonal effects is important for the tropical
cyclogenesis, rather than the independent effects.

1. Introduction
The Bay of Bengal (BoB) in the north Indian Ocean(NIO) is one of the regions where tropical cyclones(TCs) form frequently (Fig. 1). Although the TC frequency over the BoB is less than those over the westernand eastern North Pacific, TCs over the BoB have affected human society in the surrounding countries. For
example, the Bhola cyclone caused a catastrophic death
toll of 500 000 in November 1970, and tropical Cyclone
Nargis took more than 130 000 victims in April–May
2008 (Webster 2008; Fritz et al. 2009).
Corresponding author address: Wataru Yanase, Atmosphereand Ocean Research Institute, The University of Tokyo, 5-1-5
Kashiwanoha, Kashiwa, Chiba 277-8564, Japan.
E-mail: yanase@aori.u-tokyo.ac.jp
DOI: 10.1175/JCLI-D-11-00208.1
Ó 2012 American Meteorological Society

A unique characteristic of the BoB is that TCs withmaximum wind speeds exceeding 34 kt or ;17.5 m s21[referred to as tropical storms (TSs)] form in two separate seasons (black circles in Fig. 2) under the influenceof the Asian summer monsoon (ASM): most of the TSs
form around the premonsoon (;May) and around thepostmonsoon (approximately October–November),whereas few TSs form during the mature monsoon (approximately June–September). Such a bimodal TS distribution is not observed over the other regions, where TS
formation is active around summer (black circles inFig. 13) with a unimodal distribution, except over the
Arabian Sea in the NIO. On the other hand, TCs withwind speeds less than 34 kt throughout lifetimes [tropicaldepressions (TDs)] still form during the mature monsoon
season over the BoB, which are called low pressure15 APRIL 2012YANASE ET AL.

2915

FIG. 1. Distribution of the number of tropical cyclogenesis (contour interval is five events per
58 3 58) in 27 yr (1982–2008).

systems (LPSs; Goswami et al. 2003; Krishnamurthyand Ajayamohan 2010). These seasonal characteristicsof tropical cyclogenesis (including TS and TD) over the
BoB could be related to environmental conditions in
the ASM region. The Bay of Bengal is a suitable region
to examine the local environmental field responsible
for tropical cyclogenesis, because the surrounding landon the northern, eastern, and western margins of the BoB
obstructs disturbances moving from the other oceans to
the BoB (Fig. 1).
Previous studies have proposed several environmental factors responsible for tropical cyclogenesis (Gray
1975; Emanuel and Nolan 2004; Camargo et al. 2007).
Whereas the sea surface temperature (SST) is a wellknown important factor, the activity of tropical cyclogenesis also seems to be controlled by other dynamical
and thermodynamical factors including moist static stability and planetary vorticity. Emanuel and Nolan (2004)
developed the genesis potential index (GPI) based on thefollowing four factors: large lower-tropospheric absolute
vorticity, weak vertical shear, high potential intensity,and high midtropospheric humidity. Weak vertical shear
is necessary for tropical cyclogenesis because the shear
breaks a warm core structure of TC (a ventilation effect).
The potential intensity is a thermodynamical factor considering the effects of SST and moist static instability (seesection 2). Camargo et al. (2007) demonstrated that theseasonal transition of GPI is in good agreement with thatof tropical cyclogenesis in the major oceans including thebimodal TS season over the NIO. Another interesting
characteristic of tropical cyclogenesis over the BoB or
NIO is related to the boreal summer intraseasonal oscillation (BSISO). Using 30 yr of data, Kikuchi and Wang(2010) showed that the BSISO modulates the topical
cyclogenesis activity over the NIO, and that the GPI ishigh during the active cyclogenesis phase of the BSISO.
Yanase et al. (2010) proposed that the superposition ofseasonal transition and BSISO caused the high GPI
during the genesis of Cyclone Nargis in the premonsoon
of 2008. More statistical analysis is necessary to elucidate
the relation between the seasonal transition and shortertime-scale modulation that is responsible for tropical
cyclogenesis over the BoB.
In the present study, we assess the statistical relation
between the tropical cyclogenesis and environmental
field over the BoB by addressing the following issues: 1)
Are there any GPI signals on seasonal and shorter timescales in a composite analysis based on cyclogenesis
events? 2) What is the relation between the seasonaltransition and the shorter time-scale modulation for the
cyclogenesis? 3) Which factor contributes to the modulation of GPI? These characteristics over the BoB arecompared with those over the other oceans including the
Arabian Sea and the North Pacific. The remainder of themanuscript is organized as follows. Section 2 describes
the methodology employed in the data analysis. Section 3

2916

JOURNAL OF CLIMATEVOLUME 25

FIG. 2. Time–latitude diagram of climatological variables averaged over the BoB (808–1008E): (a) SST (8C), (b)OLR (W m22), (c) zonal wind at 850 hPa (m s21), and (d) MTG (1026 K m21). Black, red, green, and blue curves atthe bottom of the panels show the averages in 08–258N, 58–258N, 58–208N, and 108–208N, respectively (the meridional
ranges are shown by the lines with the same color at the top left of the panels). Black (gray) circles indicate the timeand latitude of TS (TD) genesis.

presents the climatological relation between the seasonaltransition of environmental field and tropical cyclogenesis.
This section confirms the results of previous studies. Section 4 provides a composite analysis based on the tropical
cyclogenesis to identify environmental modulations thathave shorter time scales than the seasonal transition. Because the dominant shorter time-scale modulation was
revealed to be the BSISO, section 5 analyzed the cyclogenesis and GPI based on the phases of the BSISO. We
extended the result of Kikuchi and Wang (2010) to identify the main environmental factors that modulate the GPI
within the BSISO. Section 6 discusses the cyclogenesis
over the other oceans and the role of BSISO. Finally,
section 7 summarizes the characteristics of the environmental field responsible for the cyclogenesis over the BoB.

2. Methodology
a. Data
The large-scale environmental field analyzed in thisstudy comprises daily data on the atmosphere and ocean
between 1982 and 2008. Atmospheric analysis data are

the Japanese 25-yr Reanalysis (JRA-25) for the period
between 1979 and 2004 (Onogi et al. 2007) and the Japan
Meteorological Agency (JMA) Climate Data Assimilation System (JCDAS) for the period from 2005, which
are provided by the JMA and the Central ResearchInstituted of Electric Power Industry (CRIEPI); because JCDAS is based on the same system as that employed by JRA-25, the two products, when combined,
provide a homogeneous dataset for the period from 1979
to 2008. The horizontal resolutions of the products are
1.258 grid intervals. The original 6-hourly dataset ofJRA-25/JCDAS is averaged every four time steps tocreate a daily dataset. Convective activity is represented
by outgoing longwave radiation (OLR) of the National
Oceanic and Atmospheric Administration (NOAA).
The analysis data of SST is the product of NOAA Optimum Interpolation from late 1981, which uses in situ
and satellite data (Reynolds et al. 2002). The original
weekly dataset of SST, with a horizontal resolution of 18,
is linearly interpolated in time and space to create a daily
dataset on the same grid system as that used in the atmospheric dataset.

15 APRIL 2012

2917

YANASE ET AL.

TABLE 1. Monthly cyclogenesiss number for TDs and TSs over the
BoB between 1982 and 2008.
Month
TD
TS

FIG. 3. Wind speeds of tropical cyclones over the BoB relative tothe cyclogenesis (0 h). Dashed line indicates the wind speed of;12.9 m s21 (25 kt).

In any analysis of the environmental field of topical
cyclogenesis, it is important that the temporal and spatial scales of environmental modulation are larger thanthose of TC itself because the TC influences the atmospheric and oceanic fields transiently. To avoid thisproblem, TC-scale fields are removed by the spatial filterof Kurihara et al. (1993); the parameter of the spatial
filter is set to remove the wavelength shorter than 8.758
in this study. In addition, a 15-day running mean is used
to filter out the TC-scale signal in the time direction.
Herein, this spatially and temporally filtered data is referred to as an environmental field.
Information regarding the locations and intensities ofindividual TCs for the period 1982–2008 is derived froma best-track dataset obtained from the International
Best Track Archive for Climate Stewardship (IBTrACS;
Knapp et al. 2010). Here, a tropical cyclogenesis is defined as the time when the surface wind speed first exceeded the threshold of 12.9 m s21 (25 kt) in the lifetime.
This definition includes TSs and intense TDs. Intense
TDs are analyzed to obtain supplementary samples (graycircles in Fig. 2) in the mature monsoon season when few
TSs are observed over the Bay of Bengal. Figure 3 showstime series of surface wind speed of TCs over the BoB
analyzed in the present study, and Table 1 summarizes
the monthly number of TCs observed over the Bay of
Bengal between 1982 and 2008.

b. Genesis potential index
The environmental field responsible for tropical cyclogenesis is assessed by the GPI proposed by Emanueland Nolan (2004), which is defined based on four environmental factors:
GPI 5 j105 hj3/2 (1 1 0:1Vshear )22

3  3
H
,
50
70

V

pot

1
2
2

2
0
1

3
0
2

4
0
4

5
4
13

6
9
5

7
6
1

8
13
0

9
8
3

10
16
17

11
5
23

12
5
8

where h is the absolute vorticity at 850 hPa (s21), Vshear
is the magnitude of the vertical wind shear between 850
and 200 hPa (m s21), Vpot is the potential intensity (m s21),
and H is the relative humidity at 600 hPa (%). The potential intensity Vpot considers SST and vertical profiles oftemperature and specific humidity in the troposphere,
and is defined as
2
Vpot
5

Ts Ck
(CAPE* 2 CAPEb ),T 0 CD

where TS is SST, T0 is the mean outflow temperature at
the level of neutral buoyancy, Ck is the exchange coefficient for enthalpy, CD is the drag coefficient, CAPE*
is the convective available potential energy (CAPE) foran air parcel at the radius of maximum winds, and
CAPEb is the CAPE for an air parcel lifted from thelowest data grid (1000 hPa) in the ambient atmosphere(Bister and Emanuel 2002; Camargo et al. 2007).

3. Climatological seasonal transition
The seasonal transition seems to be the most dominant environmental modulation responsible for tropical
cyclogenesis. Figure 2 shows time–latitude diagrams ofbasic variables averaged over the BoB (808–1008E). The
SST over the Bay of Bengal shows a bimodal annual
distribution with the peaks in the pre- and postmonsoon
seasons (shadings and curves in Fig. 2a), which is uniqueto the BoB and the Arabian Sea in the NIO. Although
the bimodal annual distribution of SST is qualitatively
similar to that of TS genesis, the SST alone cannot explain the following characteristics: the TS genesis is mostactive in the postmonsoon, whereas the SST is highest in
the premonsoon; and the SST exceeding 28.58C in themature monsoon season is enough high for TS genesis.
The most active convection over the BoB (the lowest
OLR in Fig. 2b) occurs in the mature monsoon season,
when the TS genesis over the BoB is reduced. Note thatthe numbers of tropical cyclogenesis over the North Pacific and the North Atlantic are increased during summer
characterized by the active convection (not shown). Overthe BoB, a monsoon flow represented by the zonal windchanges the sign in the pre- and postmonsoon seasons(Fig. 2c). Apparently, most of the TS genesis occurred
during these transitions of the monsoon flow. Meridional

2918

JOURNAL OF CLIMATEVOLUME 25

FIG. 4. As in Fig. 2, but for (a) GPI, (b) absolute
vorticity at 850 hPa (105 s21 ), (c) vertical shear between 850 and 200 hPa (m s 21 /650 hPa), (d) potential intensity (m s21 ), and (e) relative humidity at
600 hPa (%).

temperature gradient (MTG) in the upper troposphere(200–500 hPa) is another index of the ASM (Mao and Wu
2007). In general, the MTG is negative in the Northern
Hemisphere because of the meridional difference of solar
radiation. However, the MTG during the mature ASM is
uniquely positive, which is attributed to the heating of the
Asian continent to the north of the BoB. Thus, the MTG
changes the sign in the pre- and postmonsoon seasons(Fig. 2d), when most of the TS formations are observed.
To understand why the pre- and postmonsoon seasonsare favorable for TS genesis, GPI and its four factors areanalyzed (Fig. 4). GPI shows bimodal annual distribution(Fig. 4a) with the peaks in the pre- and postmonsoon,

which is in good agreement with the TS activity overthe BoB. This is consistent with the result of Camargo
et al. (2007). The GPI is moderate during the mature
monsoon, when only TDs form. In addition, GPI qualitatively explains the meridional distribution of TS genesis, including a northward shift in the premonsoon anda southward shift in the postmonsoon. Note that TS
genesis in the Southern Hemisphere also occurs during
the high GPI seasons from October to May.
The four GPI factors give further information on theenvironment responsible for the bimodal TS genesis.
The bimodal GPI distribution is dominantly attributed
to weak vertical shear in the pre- and postmonsoon

15 APRIL 2012YANASE ET AL.

(shadings and curves in Fig. 4c). In the mature monsoon
season, strong vertical shear seems to prevent TCs fromdeveloping to TS intensity by the ventilation effect. Basedon the thermal wind balance, the weak vertical shear isassociated with the small MTG in the pre- and postmonsoon over the BoB (Fig. 2d). The potential intensity
also shows a weak bimodal distribution (Fig. 4d), whichreflects the SST distribution (Fig. 2a). The seasonaltransition of absolute vorticity is relatively weak (Fig. 4b)
because of the large planetary vorticity. Midtroposphericrelative humidity is highest during the mature monsoon(Fig. 4e), which is related to the active convection asshown in Fig. 2b. In summary, the bimodal annual distribution of climatological GPI is attributed to the weakvertical shear dominantly and to the high potential intensity secondarily.

4. TC-based composite analysis
The seasonal transition of climatological GPI in section 3 is in good agreement with the tropical cyclogenesis: TS genesis occurs within the high GPI environment
in the pre- and postmonsoon seasons, whereas only TD
forms within the moderate GPI environment in themature monsoon. If the large-scale environment modulates on a time scale shorter than the seasonal transition, it is of interest to examine whether GPI can explain
the cyclogenesis even on such a short time scale statistically. To extract signals of environmental modulation,
we attempt a composite analysis based on the cyclogenesis events (TC-based composite analysis).
As the climatological seasonal transition is the large
environmental modulation, the original field (total field)
is separated into climatological and anomalous ones inadvance of the composite analysis: the climatologicalfield is the average of 27 yr, whereas the anomalous fieldis the difference between the total and the climatologicalones. Then the TC-based composite is performed for the
total, climatological, and anomalous fields. As the cyclogenesis activity is different between seasons (Fig. 4), thecomposite field is assessed for three seasons: April–May(AM; premonsoon over the BoB), June–September(JJAS; mature monsoon) and October–November (ON;postmonsoon). To obtain the time evolution of environmental modulation, the composite is performed for theperiod from 30 days before the cyclogenesis (230 relativeday) to 30 days after the cyclogenesis (30 relative day).
It should be noted again that the analysis is performed
using the environmental fields (spatially and temporally filtered data).
Figure 5 shows time–latitude diagrams of the TCbased composite GPI. Here, the relative day 0 is thetime when the cyclogenesis events occur, and the short

2919

horizontal bars indicate the latitude of cyclogenesis. In
the premonsoon (the left column in Fig. 5), a high GPI
signal in the total field shifts northward from 58 to 208N
during the period between 220 and 10 days (Fig. 5a). This
northward-moving signal is attributed to both the climatological and anomalous modulations. In the climatological field (Fig. 5d), a high GPI signal slowly
shifted northward during the period from 230 to 30 days,
which is the seasonal transition observed in Fig. 4a. In theanomalous field (Fig. 5g), a positive GPI signal moves
northward quickly from 220 to 10 days, which is enoughintense to modify the climatological field. Most of thecyclogenesis in the premonsoon occurs within the positive
anomaly of the northward-moving GPI signal. The positive GPI anomaly is followed by a negative one, and thenby a weak positive one again, showing a periodicity of
approximately 30–40 days. For the mature monsoon(the middle column in Fig. 5), a high GPI signal in the
climatological field is stationary around approximately
158–208N. In the anomalous field, on the other hand,a high GPI signal moves fast northward as observed in
the premonsoon season. In the postmonsoon season (theright column in Fig. 5), a high GPI signal in the climatological field slowly shifts southward, which is the seasonaltransition shown in Fig. 4a. In the anomalous field, a high
GPI signal still moves northward. Thus, the northwardmoving anomaly with the intraseasonal time scale (approximately 30–40 days) is observed in all the analyzed
seasons, although it is a little obscure in the postmonsoon
season.
The northward-moving signal in the anomalous field issimply observed in the time–latitude diagram of zonal
wind at 850-hPa pressure level (Fig. 6). In the climatological field, the westerly flow over the BoB increases
in the premonsoon, remains almost static in the mature
monsoon, and then decreases in the postmonsoon(Figs. 6d–f). The anomalous fields of zonal wind show
fast northward-moving signals in all the seasons again(Figs. 6g–i): a westerly anomaly to the south of the cyclogenesis and an easterly one to the north. In the longitude–
time diagrams of the composite field (not shown) thesignal almost remain static, and shows the zonal scale of408 or more. Because the signal has the large horizontal
scale and long time scale, it is not merely the influenceof TC itself.
To elucidate the environmental modulation responsible
for the high GPI, the four GPI factors are assessed using
the TC-based composite analysis. Figure 7 shows time–
latitude diagrams of the anomalous fields of the four
factors in the three seasons. The absolute vorticity(Figs. 7a–c) shows a positive signal quickly moving
northward in all the seasons. As planetary vorticity does
not change in time, the northward-moving signal is

2920

JOURNAL OF CLIMATEVOLUME 25

FIG. 5. Time–latitude diagram of GPI over the BoB in the TC-based composite analysis. Diagrams for months of (left) AM, (middle)
JJAS, and (right) ON. Diagrams for (top) total, (middle) climatological, and (bottom) anomalous fields. Dotted areas correspond todifferences in the anomaly at p , 0.1 using Student’s t test. Short horizontal bars at day 0 indicate the latitudes of individual cyclogenesis.

caused by the change in relative vorticity, which is large
on the northern (southern) side of the westerly (easterly)
axes in Figs. 6g–i. Most of the cyclogeneses occur aroundthe positive anomaly of absolute vorticity. In detail, thecyclogeneses in the mature monsoon prefer the northernside of the axis of the positive anomaly—that is, the cyclogeneses in the mature monsoon occur farther norththan those in the pre- and postmonsoon relative to thenorthward-moving signal, which is discussed in section 6.
The vertical shear (Figs. 7d–f) also shows the northwardmoving signal particularly in the pre- and mature monsoon seasons. The cyclogenesis in the mature monsoon
prefers the region of weak vertical shear (Fig. 7e),whereas the cyclogenesis in the pre- and postmonsoon
is observed both in positive and negative anomalies(Figs. 7d,f). The potential intensity (Figs. 7g–i) does notshow the oscillation of approximately 30–40 days,

although some longer-period signal is observed in the
premonsoon (Fig. 7g). Similar to the absolute vorticity,
the cyclogenesis occurs within the positive anomaly of
relative humidity (Figs. 7j–l), and tends to occur on thenorthern side of the axis of positive anomaly in the mature monsoon.
In summary, the tropical cyclogenesis in the pre-,mature-, and postmonsoon seasons occurs within thenorthward-moving signal of high GPI anomaly, which
had a zonal scale more than 408 and the periodicity of
approximately 30–40 days. The high GPI is related to
the large absolute vorticity and high relative humidity
in all the seasons, whereas it is also related to the weakvertical shear in the mature monsoon. The cyclogenesis
in the mature monsoon occurs farther north than thosein the pre- and postmonsoon relative to the axis ofnorthward-moving signals. The northward-moving signal

15 APRIL 2012YANASE ET AL.

2921

FIG. 6. As in Fig. 5 but for zonal wind at 850 hPa.

of high GPI is more significant in the pre- and mature
monsoon seasons than in the postmonsoon season. The
TC-based composite is also useful to assess whether
the environmental modulation is the cause or effect of
the tropical cyclogenesis; given that the signal of GPI
appears 20 days prior to the cyclogenesis, reaching its
maximum before the genesis, the signals are not merely
the response of the cyclogenesis. As the signal is accompanied by the modulation of zonal flow and relative humidity, it seems to be associated with the BSISO (Jiang
et al. 2004). In the next section, the relation between thecyclogenesis and the BSISO is analyzed.

5. BSISO-based composite analysis
The TC-based composite analysis in section 4 clarified
that the most dominant signal in the anomaly field appears to be relevant to the BSISO. This section focuseson the environmental modulation responsible for tropical
cyclogenesis in terms of the phase of the BSISO. Basedon the previous studies (Yasunari 1981; Murakami et al.
1984; Goswami and Mohan 2001), the BSISO is definedby 30–60-day filtered zonal wind at 850 hPa. Here, thecyclogenesis from May to October is assessed, given thatthe amplitude of the BSISO mode is large during thisperiod (Kikuchi and Wang 2010). Note that the positive
MTG over the BoB during this period (the curves inFig. 2d) corresponds to the ASM season. To minimize
the influence of the TC itself on the BSISO signal, thezonal wind is averaged in the large area (108–208N and
608–1108E; the rectangle in Fig. 9), which is determinedbased on the structure of the BSISO shown in the previous studies (Murakami and Nakazawa 1985; Goswami
and Mohan 2001). Figure 8a shows the time series of the
BSISO index in 2008. Each minimum–maximum–
minimum cycle is divided into eight phases in the following

2922

JOURNAL OF CLIMATEVOLUME 25

FIG. 7. As in Figs. 5g,h,i but for the anomalous fields of (a)–(c) absolute vorticity at 850 hPa (1025 s21), (d)–(f) vertical shear[m s21(650 hPa)21], (g)–(i) potential intensity (m s21), and (j)–(l) relative humidity at 600 hPa (%).

manner (Fig. 8b): phases 1–4 (5–8) are defined by separating the period between the first minimum and the
maximum (between the maximum to the next minimum)
evenly. If the amplitude of the signal (the half of thedifference between a maximum and a minimum) is lessthan one standard deviation, the period is not analyzed

because of the weak BSISO. Composite of the environmental fields (spatially and temporally filtered data as
described in section 2) and count of cyclogenesis are
performed for each phase of the BSISO.
Figure 9 shows composite patterns of zonal wind at
850 hPa and OLR for phases 1, 3, 5, and 7. In phase 1,

15 APRIL 2012YANASE ET AL.

2923

FIG. 8. Time series of 30–60-day bandpath-filtered zonal wind at 850 hPa averaged in the region (108–208N,
608–1108E). (a) Time series in 2008. (b) Definition of eight phases of the BSISO.

a westerly anomaly is observed near the equator over the
Indian Ocean, and active convection (negative OLR)
occurs to the north of the westerly axis. From phases 1 to
5, the westerly axis and active convection shift northwardfrom the equator to the BoB. In phase 7, the westerly axis
almost disappears over the Indian Ocean, which is theend of a BSISO cycle.
Figure 10a shows the number of tropical cyclogeneses
for each phase. The cyclogenesis is significantly modulated by the phase of the BSISO: the number of cyclogeneses in phases 3 and 4 are 20 and 23, respectively,
whereas those in phases 1 and 8 are only 2. Phases 3 and
4 are characterized by active convection over the BoB(Fig. 9b). This is consistent with the result of Kikuchi
and Wang (2010) using an extended empirical orthogonal function (EEOF) for the BSISO definition. As thenumbers of analyzed days are not much different between the phases (approximately 392–467 days; Fig. 10b),
the difference in cyclogenesis numbers between thephases is attributed to the BSISO dynamics. For the weak
BSISO period, the number of cyclogenesis is 21 for 1638
analyzed days. This frequency (0.013 events per day) is
nearly the same as those for phases 5–7.
The phase–latitude diagrams of GPI and its four factors in the BSISO-based composite analysis are shown in
Fig. 11 together with the months of the cyclogenesis(digits in the figure). The total GPI is largest from phases
2 to 4 between 158 and 208N, where the cyclogenesis isactive (Fig. 11a). The anomalous field of GPI shows apositive signal moving northward from phases 1 to 4
(Fig. 11a). The cyclogenesis is active around the positive
region of GPI anomaly, although the timing seems to bedelayed by ;1 phase. This lag may indicate the time
during which a TC grows gradually. To assess the individual contributions of the four GPI factors quantitatively, the GPI is recalculated using the climatologicalfields (1982–2008) of three out of the four factors, butusing the unmodified daily fields of the target factor (cf.
Camargo et al. 2009). Figures 11c–f show the contribution of the four factors to the GPI anomaly. The relative
humidity dominantly contributes to the high GPI. The
absolute vorticity changes in a similar manner to therelative humidity, although the contribution is relatively
small. The GPI anomaly caused by vertical shear isaccompanied more by the cyclogenesis from June to
September than by those in May and October (Fig. 11d),
which is consistent with the result of TC-based composite
analysis (Fig. 7e).
As the sign of GPI anomaly in a BSISO phase alsochanges in the meridional direction, the analysis is more
effective if the BSISO phase is adjusted based on thelatitude. Therefore, all the cyclogeneses are moved to
the latitude of 208N (the dashed line in Fig. 11) along theline parallel to the northward movement of the BSISO
signal (the solid line in Fig. 11). Figure 12 shows the cyclogenesis number recalculated in the latitude-adjusted
BSISO phases. The cyclogenesis distribution with thepeak around phase 4 in the adjusted phase is more symmetric than that in the original BSISO phase (Fig. 10a).
The ratio of cyclogenesis from June to September (graybars) is high between phases 2 and 4, whereas those in
May and October (black and white bars, respectively) arehigh between phases 4 and 7. This indicates that the cyclogenesis in the mature monsoon occur in an earlier
phase (or farther north relative to the BSISO phase) thanthose in the pre- and postmonsoon. Thus, the cyclogenesis in the mature monsoon prefers the BSISO phase that
reduces the vertical shear effectively.

6. Discussion
a. Comparison with other oceans
The previous sections described the seasonal andintraseasonal characteristics of the environmental field

2924

JOURNAL OF CLIMATE

FIG. 9. Composite of zonal wind anomaly at 850 hPa (shading)and negative OLR anomaly (white contours with contour intervalof 2 W m22) for phases (a) 1, (b) 3, (c) 5, and (d) 7 of the BSISO.
The black rectangle shows the region in which the zonal wind isaveraged for making the BSISO index.

responsible for tropical cyclogenesis over the BoB. To
understand whether such characteristics are unique tothe BoB, cyclogenesis over the following oceans are examined: the Arabian Sea in the NIO (558–758E), thewestern North Pacific (1258–1458E), the eastern North
Pacific (958–1158W), and the North Atlantic (308–508W).
The above regions of 208 in longitude cover the mostactive cyclogenesis region (Fig. 1).
To compare the seasonal transitions, time–latitude
diagrams of the climatological GPI field for the four

VOLUME 25

regions are shown in Fig. 13 (see also Fig. 4a for the
BoB). The tropical cyclogenesis and high GPI over the
Arabian Sea show bimodal annual distribution (Fig. 13a),which is similar to those over the BoB, whereas thoseover the other regions are characterized by unimodal
distribution from boreal summer to autumn (Figs. 13b–d).
Therefore, the bimodal distribution of cyclogenesis andhigh GPI is unique to the NIO including the BoB and theArabian Sea, which is consistent with the previous studies(Camargo et al. 2007; Kikuchi and Wang 2010; Evan and
Camargo 2011). The high GPI in the pre- and postmonsoon over the Arabian Sea is associated with the
weak vertical shear and high potential intensity (notshown), which is also the same characteristics as the
BoB. The other region shows weak vertical shear andhigh potential intensity in summer (not shown). The
strong vertical shear over the NIO in summer is related
to the large positive MTG in the ASM region throughthe thermal wind balance. Therefore, the heating of the
Asian continent to the north of the NIO during thesummer is responsible for the bimodal annual distribution of TSs over the NIO.
The intraseasonal characteristics over the NIO arealso different from those over the other oceans. Figure 14
shows the time–latitude diagram of the anomalous
GPI field for the four oceans in the TC-based composite analysis (see also the bottom row in Fig. 5 for the
BoB). Here, the tropical cyclogenesis used in the composite analysis over the western and eastern North
Pacific and North Atlantic is limited to the cases between 108 and 208N, as the composite analysis would be
obscure if cyclogenesis locations are diverse. In the
JJAS season, the anomalous GPI field over the Arabian
Sea shows a northward-moving signal, whereas thoseover the other oceans are weak or nearly stationary.
The northward-moving signal over the Arabian Sea isassociated with the BSISO in the ASM because thezonal scale of the signal extends over the entire NIO(see Fig. 9). In the AM season, the northward signal isalso observed over the western and eastern North
Pacific as well as the Arabian Sea. Because the signal
over the western North Pacific is also accompanied bysignificant eastward movement in the longitude–timediagram (not shown), it seems to be related to the MJO(Madden and Julian 1972). On the other hand, thesignal over the eastern North Pacific is almost stationary in the zonal direction (not shown), whichshould be studied in a future work. Thus, the BoB andthe Arabian Sea in the NIO have similar dynamics fortropical cyclogenesis both on the seasonal and intraseasonal time scales, although more samples over the
Arabian Sea may be necessary to make a more reliableconclusion.

15 APRIL 2012YANASE ET AL.

2925

FIG. 10. Numbers of the events in each phase of the BSISO. (a) Tropical cyclogenesis. (b) Total observed day.

b. Dynamical relation between tropical cyclogenesis
and BSISO
The TC-based composite analysis showed that thenorthward-moving signal of high GPI is attributed to thelarge relative vorticity, high relative humidity, and weakvertical shear (Figs. 5 and 7). These characteristics are
consistent with the structure and dynamics of the BSISO(Jiang et al. 2004): the lower-tropospheric vorticity
causes active convection through the convergence in the
Ekman layer, whereas the convection intensifies the lowertropospheric vorticity through the stretching effect ofvortex. This results in the good correlation between thelower-tropospheric vorticity and midtropospheric humidity. The structure of the vertical shear in the BSISO
results from the superposition of the BSISO flow andclimatological one: because the lower-troposphericpositive vorticity in the BSISO is accompanied by theeasterly (westerly) to the north (south), the vertical shearof the zonal flow is westerly (easterly) to the north(south). Given the climatological easterly vertical shear
over the Bay of Bengal during the ASM season, themagnitude of the total vertical shear of zonal flow is weak
(strong) to the north (south) of the lower-troposphericpositive vorticity. The modulation of vertical shear by the
BSISO works more effectively in the mature monsoon(Fig. 11d), presumably because the very strong vertical
shear in the climatological field plays a key role to inhibit
the cyclogenesis (Fig. 4c). It is also speculated that the
BSISO modulation may not give sufficient time for TDs
in the mature monsoon to develop to the TS intensity.
The present study has demonstrated that the active
tropical cyclogenesis occurs within the high GPI causedby the BSISO. However, other dynamics might explain
the relation between the cyclogenesis and the BSISO.
Krishnamurti et al. (1981) proposed that barotropic

instability might promote the cyclogenesis over the
Arabian Sea during the onset of ASM. They analyzed
the meridional gradient of the absolute vorticity (MGAV)
using the zonally averaged westerly wind at 850 hPa, as
the necessary condition for the existence of barotropic
instability is MGAV vanishing somewhere on the meridional plane. They demonstrated that the zero line of
MGAV was observed during the period of cyclogenesis
in June 1979. To examine the barotropic instability over
the Arabian Sea and the BoB, TC-based composite of
MGAV is analyzed for both oceans. Figure 15 shows thetotal field of the MGAV over the Arabian Sea and theBoB. Here, any spatial filter is not used, because the zeroline of MGAV has very narrow structure in the meridional direction. Over the Arabian Sea, the MGAV in the
JJAS season shows the zero line around 208N (Fig. 15b),
where most of the cyclogenesis occurs. This is consistent
with the result of Krishnamurti et al. (1981). On the other
hand, MGAV over the BoB shows no zero line in all theseasons (Figs. 15d–f). In detailed analysis for individual
cases, the cyclogenesis around the ASM onset over the
Arabian Sea occurs near the zero line of MGAV in mostcases (not shown), whereas those over the BoB showed
weaker relation with the zero line of MGAV. Therefore,
the barotropic instability seems to work more dominantly
over the Arabian Sea than over the BoB, although more
careful analysis is necessary to obtain a clear conclusion.
For individual tropical cyclogenesis, finite-amplitude
disturbances (seeds) may be necessary (Rotunno and
Emanuel 1987), although the GPI can explain the statistical probability of cyclogenesis. As the detailed dynamics of individual cyclogenesis are not the target in
the present study, the following studies only suggest
several candidates for the cyclogenesis seeds over the
BoB. Kikuchi and Wang (2010) proposed that a westerly
wind burst in the BSISO can provide a synoptic-scale

2926

JOURNAL OF CLIMATEVOLUME 25

FIG. 11. Phase–latitude diagram obtained by the BSISO-based composite analysis: (a) GPI (total), (b) GPI(anomaly), (c) contribution of absolute vorticity to GPI (anomaly), (d) contribution of vertical shear to GPI(anomaly), (e) contribution of potential intensity to GPI (anomaly), and (f) contribution of relative humidity to
GPI (anomaly). The digits represent the months of the individual cyclogenesis (October is indicated by ‘‘0’’ forbrevity). Solid line indicates the meridional phase speed of 4.48 per phase. Dashed line indicates the latitude of 208N
on which cyclogenesis is projected parallel to the solid line in the latitude-adjusted phase analysis in Fig. 12.

disturbance on its northern side based on the composite
analysis. Yokoi and Takayabu (2010) showed that a cold
surge in the eastern coastal area of the Eurasian continent in April triggered the genesis of Cyclone Nargis in
2008. Frank and Roundy (2006) showed that the tropical
cyclogenesis over the NIO seems to be related to sometropical waves. On the other hand, instabilities of theenvironmental field could develop a finite-amplitude

disturbance from an infinitesimal noise. Barotropic instability mentioned above is one of the candidates ofsuch environmental instability. Sufficiently high GPI may
also be an environmental instability in which cyclogenesis
occurs spontaneously without a preexisting disturbance(Emanuel and Nolan 2004; Nolan et al. 2007). Further
case studies are necessary to obtain a comprehensive
understanding about the cyclogenesis seeds.

15 APRIL 2012YANASE ET AL.

FIG. 12. Numbers of the tropical cyclogenesis in the latitudeadjusted phase of the BSISO. Black, gray, and white bars indicate
the numbers in May, June–September, and October, respectively.

7. Summary
The environmental field responsible for tropical cyclogenesis over the BoB is examined using the TC besttrack data and atmosphere–ocean analysis data. The GPI

2927

index and its four factors (Emanuel and Nolan 2004) areapplied to the seasonal and intraseasonal environmental
modulations. In the climatological field, the bimodal annual
distributions of cyclogenesis and high GPI (Figs. 4a and13a) are unique to the Bay of Bengal and the Arabian Sea
in the north Indian Ocean, which is consistent with previous
works (Camargo et al. 2007; Kikuchi and Wang 2010; Evan
and Camargo 2011). This is attributed to the weak vertical
shear during the premonsoon (;May) and postmonsoon(approximately October–November) over the NIO.
The TC-based composite analysis clarified that anorthward-moving signal of high GPI with approximately
30–40-day periodicity is the dominant environmental
modulation responsible for the cyclogenesis. The structure of the signal is consistent with the BSISO, which isalso unique to the NIO in summer season. Thus, the TCbased composite analysis is useful to identify the dominant environmental modulation of time scales shorter
than the seasonal transition. It is interesting to investigate
the relation between cyclogenesis and environmental
modulation over the other oceans, focusing on the MJO(Camargo et al. 2009) and other intraseasonal dynamics.
The BSISO-based composite analysis showed that thecyclogenesis is active within the high GPI phase of the

FIG. 13. As in Fig. 4a, but for (a) the Arabian Sea, (b) the western North Pacific, (c) the eastern North Pacific, and(d) the North Atlantic.

2928

JOURNAL OF CLIMATEVOLUME 25

FIG. 14. Time–latitude diagram of the anomalous GPI field over the BoB in the TC-based composite analysis. Diagrams for months of(left) AM, (middle) JJAS, and (right) ON. (a)–(c) The Arabian Sea, (d)–(f) the western North Pacific, (g)–(i) the eastern North Pacific,and ( j)–(l) the North Atlantic. The dots and short horizontal bars are as described in Fig. 5.

BSISO, which is consistent with Kikuchi and Wang(2010). We have extended this result to identify the main
environmental factors that modulate the GPI within the
BSISO. It is revealed that the high GPI phase is attributed to the high relative humidity and large absolute
vorticity within the convective phase of the BSISO in all

the analyzed seasons. The dominant influence of therelative humidity on the high GPI observed in the BSISO
is similar to the characteristics of the MJO described by
Camargo et al. (2009). In the mature monsoon, when thevertical shear is climatologically strong, the BSISO also
increases GPI by reducing the vertical shear. These15 APRIL 2012YANASE ET AL.

2929

FIG. 15. Time–latitude diagram of the total MGAV field (10211 m21 s21) in the TC-based composite. Diagrams for months of (left) AM,(middle) JJAS, and (right) ON. (a)–(c) The Arabian Sea and (d)–(f) the BoB. The short horizontal bars are as described in Fig. 5.

results propose that the combination of seasonal andintraseasonal effects is important for the tropical cyclogenesis, rather than the independent effects.
Acknowledgments. We acknowledge helpful comments by Dr. Tetsuo Nakazawa, Dr. Yukari Takayabu,
Dr. Satoru Yokoi, Dr. Amato T. Evan, Dr. Suzana
J. Camargo, and three anonymous reviewers. The besttrack data was obtained from the IBTrACS website. The
atmosphere and ocean datasets made use of JRA-25/
JCDAS reanalysis, NOAA OI-SST reanalysis, and NOAA
OLR data. Potential intensity was calculated using the
FORTRAN program from the website of Dr. Kerry
Emanuel. This research was supported by Core Research
for Evolutional Science and Technology, Japan Scienceand Technology Agency (CREST, JST), and Grants-inAid for Scientific Research 20740266 and 23740349.
REFERENCES
Bister, M., and K. A. Emanuel, 2002: Low frequency variability oftropical cyclone potential intensity 1. Interannual to interdecadal variability. J. Geophys. Res., 107, 4801, doi:10.1029/
2001JD000776.
Camargo, S. J., K. A. Emanuel, and A. H. Sobel, 2007: Use ofa genesis potential index to diagnose ENSO effects on tropical
cyclone genesis. J. Climate, 20, 4819–4834.
——, M. C. Wheeler, and A. H. Sobel, 2009: Diagnosis of the MJO
modulation of tropical cyclogenesis using an empirical index.
J. Atmos. Sci., 66, 3061–3074.

Emanuel, K. A., and D. S. Nolan, 2004: Tropical cyclone activity
and global climate. Proc. 26th Conf. on Hurricanes andTropical Meteorology, Miami, FL, Amer. Meteor. Soc.,
240–241.
Evan, A. T., and S. J. Camargo, 2011: A climatology of Arabian
Sea cyclonic storms. J. Climate, 24, 140–158.
Frank, W. M., and P. E. Roundy, 2006: The role of tropicalwaves in tropical cyclogenesis. Mon. Wea. Rev., 134, 2397–
2417.
Fritz, H. M., C. D. Blount, S. Thwin, M. K. Thu, and N. Chan, 2009:
Cyclone Nargis storm surge in Myanmar. Nat. Geosci., 2, 448–
449, doi:10.1038/ngeo558.
Goswami, B. N., and R. S. A. Mohan, 2001: Intraseasonal oscillations and interannual variability of Indian summer monsoon.
J. Climate, 14, 1180–1198.
——, R. S. Ajayamohan, P. K. Xaviers, and D. Sengupta, 2003:
Clustering of synoptic activity by Indian summer monsoon
intraseasonal oscillations. Geophys. Res. Lett., 30, 1431, doi:
10.1029/2002GL016734.
Gray, W. M., 1975: Tropical cyclone genesis. Colorado State
University, Department of Atmospheric Science Paper 234,
121 pp.
Jiang, X., T. Li, and B. Wang, 2004: Structures and mechanisms ofthe northward propagating boreal summer intraseasonal oscillation. J. Climate, 17, 1022–1039.
Kikuchi, K., and B. Wang, 2010: Formation of Nargis (2008) andtropical cyclones in the northern Indian Ocean associated with
with tropical intraseasonal oscillation. J. Meteor. Soc. Japan,
88, 475–496.
Knapp, K. R., M. C. Kruk, D. H. Levinson, H. J. Diamond, andC. J. Neumann, 2010: The International Best Track Archive
for Climate Stewardship (IBTrACS) unifying tropical cyclone
data. Bull. Amer. Meteor. Soc., 91, 363–376.

2930

JOURNAL OF CLIMATE

Krishnamurthy, V., and R. S. Ajayamohan, 2010: Composite
structure of monsoon low pressure systems and its relation to
Indian rainfall. J. Climate, 23, 4285–4305.
Krishnamurti, T. N., P. Ardanuy, Y. Ramanathan, and R. Pasch,
1981: On the onset vortex of the summer monsoon. Mon. Wea.
Rev., 109, 344–363.
Kurihara, Y., M. A. Bender, and R. J. Ross, 1993: An initialization
scheme of hurricane models by vortex specification. Mon.
Wea. Rev., 121, 2030–2045.
Madden, R. A., and P. R. Julian, 1972: Description of global circulation cells in the tropics with a 40–45 day period. J. Atmos.
Sci., 29, 1109–1123.
Mao, J., and G. Wu, 2007: Interannual variability in the onset of thesummer monsoon over the eastern Bay of Bengal. Theor.
Appl. Climatol., 89, 155–170.
Murakami, T., and T. Nakazawa, 1985: Tropical 45 day oscillations
during the 1979 Northern Hemisphere summer. J. Atmos. Sci.,
42, 1107–1122.
——, ——, and J. He, 1984: On the 40–50 day oscillations during
the 1979 Northern Hemisphere summer. Part I: Phase propagation. J. Meteor. Soc. Japan, 62, 440–468.
Nolan, D. S., E. D. Rappin, and K. A. Emanuel, 2007: Tropical
cyclogenesis sensitivity to environmental parameters in

VOLUME 25

radiative–convective equilibrium. Quart. J. Roy. Meteor.
Soc., 133, 2085–2107.
Onogi, K. J., and Coauthors, 2007: The JRA-25 Reanalysis.
J. Meteor. Soc. Japan, 85, 369–432.
Reynolds, R. W., N. A. Rayner, T. M. Smith, D. C. Stokes, andW. Wang, 2002: An improved in situ and satellite SST analysis
for climate. J. Climate, 15, 1609–1625.
Rotunno, R., and K. A. Emanuel, 1987: An air–sea interaction
theory for tropical cyclones. Part II: Evolutionary study using
a nonhydrostatic axisymmetric numerical model. J. Atmos.
Sci., 44, 542–561.
Webster, P. J., 2008: Myanmar’s deadly daffodil. Nat. Geosci., 1,
488–490.
Yanase, W., H. Taniguchi, and M. Satoh, 2010: The genesis oftropical cyclone Nargis (2008): Environmental modulationand numerical predictability. J. Meteor. Soc. Japan, 88, 497–
519.
Yasunari, T., 1981: Structure of an Indian summer monsoon system with around 40-day period. J. Meteor. Soc. Japan, 59,
336–354.
Yokoi, S., and Y. Takayabu, 2010: Environmental and external
factors of tropical cyclone Nargis in April 2008 over the Bay of
Bengal. J. Meteor. Soc. Japan, 88, 425–435.