CHAPTER - 1
Introduction
1.1 Description of Monsoon
1.1.1 Monsoon Systems
The precise definition of planetary-scale monsoon is a much
debated issue. "A monsoon is defined as a seasonal shift in wind
direction, being derived from the Arabic word "mausim", meaning
season." The Arabian Sea is generally considered to be in that area of
the world where the name "monsoon" was first used [Huschke, 1959] to
signify a seasonal wind regime in which surface winds blow persistently
from one general direction in summer and just as persistently from a
markedly different direction in winter. Monsoons can be viewed as
three-dimensional circulations associated with the global distribution of
land and sea. For example, Asian monsoon has often been compared
to a giant land-sea breeze. The seasonal variations of monsoon usually
can be seen as the change of amplitude of long waves. Generally it is
based on a wind speed at least 3 meters per second. In the place
heavily influenced by monsoon, seasonal wind reverses its direction
and causes a drastic change on precipitation and temperature. The
monsoon related phenomenon is the dominant feature of low-latitude
climates stretching from West Africa to the western Pacific Ocean
(figure 1.1).
I
Figure 1.1: Monsoon areas enclosed by solid line [Oliver and Fairbridge, 1987]
For hundreds of years, the monsoon has been observed by
sailors and was used for wind-driven ships. The systematic study of the
seasonal and interannual variations of the mean monsoon patterns
began in the late 19th century with the incentive to predict the Indian
rainfall after the great drought in 1877. Sir Gilbert Walker [1923, 1924
and 1928] figured out that the global climate possesses a coherent lowfrequency variability. As Webster [1987 and 1998] has noted in his
monsoon reviews, it has been recognized for hundreds of years that the
physics behind the annual monsoon cycle is the variation of incoming
solar radiation and the differential heating rate of the surface of land and
water. Sections of the earth's surface heat and cool at different rates
depending on their ability to absorb incoming solar radiation and the
magnitude of this incoming shortwave radiation. Heat capacity water is
much larger than soil and rock, which allows it to store energy more
efficiently than the land surface and therefore retains heat longer
duration. During summer the land heats more rapidly than the adjacent
ocean because of its smaller specific heat and shallow layer depth. In
the winter the land surface will cool much more quickly than the ocean
2
simply because there is little available heat in the subsurface that can
be made available to heat the surface on seasonal time scales [Webster
et al., 2003]. To maintain atmospheric energy balance, heat is generally
transferred from areas of surplus to deficit, and in the case of a landwater differential, this is accomplished through a phenomenon known
as the "land-sea breeze". On a larger scale when there is a land-water
contrast, such as a continent surrounded by oceans, heat build up on
land over summer time will warm air masses, lower their density,
increase their vertical volumes, and drive up the height of pressure
levels. Thus, in the same height of the upper troposphere, the pressure
is higher over continent than ocean. This build-up pressure gradient
between the land and ocean will drive air masses move from land to
ocean in the upper troposphere. As a result of this convection, the air
masses in the vertical column get much greater over the ocean than
land. Denser air associated with high pressure dominates ocean
surfaces. Surface air will blow from ocean to land resulting in a pressure
gradient and also construct the summer monsoon circulation.
Some common features in both summer and winter monsoons are:
a. Lower tropospheric winds flow from a high-pressure area into a
low-pressure area.
b. The low-pressure area is usually formed as a belt-shaped
zone, and is called the monsoon trough. Monsoon troughs are
typically associated with low-level convergence and cyclonic
vorticity, and cloudiness.
c. The wind speeds between the high-pressure area and the
monsoon trough are usually concentrated into a jet pattern. This
is referred to as a monsoon jet, monsoon surge, a low-level jet, or
a cross equatorial jet. Features associated with the jet typically
have local characteristics.
3
d. The monsoonal circulation has vertical structure. The rising air along
the monsoon trough diverges out of an upper-tropospheric anticyclone
along an upper-level jet-like flow.
1.1.2 Monsoon Indices
Monsoon indices have been developed to determine which areas
of the world actually experience monsoons. These indices typically
overestimate the monsoonal areas, including such regions as the
Russian Arctic and the Gulf of Alaska. Ramage [1971] and Hastenrath
[1994] listed four criteria which define a monsoon region as
(a) The prevailing wind direction shifts by at least 120 degrees
between January and July,
(b) The mean frequency of prevailing wind directions in January
and July exceeds 40%,
(c) The mean resultant winds in January and July exceed 3 m/s,
and
(d) Less than one cyclone-anticyclone alternation occurs every
two years in either January or July in a 5° latitude-longitude
rectangle.
Therefore, the monsoon region of the world as given by Ramage [1971]
is shown in figure 1.2.
4
Figure 1.2: Delineation of the world monsoon region. Hatched areas are
monsoonal areas according to surface wind criteria. Heavy line marks the
northern limits of region with low frequencies of surface cyclone-anticyclone
alternations in summer and winter [Ramage, 1971].
1.1.3 Indian-Asian Southwest Summer Monsoon
The Indian-Asian summer monsoon is the most widely studied.
During the winter, the temperature typically decreases poleward of the
equator. With the heating of the continents during summer, especially
the Tibetan plateau, the temperature gradient is reversed. The region of
highest temperature is now located near the southern base of the
Himalayas. The winter northeasterly flow reverses to southwesterly,
lasting from May to October. The surface wind chart for July is typical of
the monsoonal flow pattern. The south easterlies wind, trade wind, of
the South Indian Ocean turn southerly, then southwesterly north of the
equator. A speed maximum, often referred to as the Somalia jet, exists
in the Arabian Sea near the Gulf of Aden. The influence of the IndianAsian monsoon extends from eastern Africa well into the western North
Pacific near 150°E during August. The monsoon trough of the western
North Pacific is poorly depicted on a mean chart due to its migratory
nature. Aloft, a large anticyclone is situated over the Tibetan Plateau.
The easterlies to the south serve as the return branch of the surface
monsoonal flow. This results in large, persistent vertical shear, making
tropical cyclogenesis very rare during these months. During the period
5
1975-1990, there were an average of 0.5 Tropical Cyclones (TCs) in
June, 0.0 in July, 0.1 in August, and 0.2 in September. Of all of the
phenomena associated with the Indian-Asian monsoon, the onset is
probably the most important and interesting. The interest lies in the
sudden arrival of the onset, which can typically be determined using
rainfall as the indicator. The onset over Southeast Asia typically occurs
during May and there after gradually the onset takes place over Indian
region. The actual date of the onset varies throughout the region.
1.1.4 Southeast Asian Winter Monsoon
The Southeast Asian winter monsoon owes its existence in large
part to the Tibetan plateau. The east-west orientation of the Himalayas
blocks the synoptic scale exchanges of cold polar air with warm tropical
air. The only avenue of exchange is east of the Himalayas, over
Southeast Asia. Consequently, cold air from the Siberian anticyclone
flows southward across eastern China, over the South China Sea, and
into Southeast Asia. Numerical modeling studies show that if the
Himalayan Mountains were "removed", the Siberian anticyclone would
no longer be a semi- permanent feature. Rather, the region would
resemble North America with transitory cyclones and anticyclones. The
onset of the northeast monsoon is far more distinct than the summer
monsoon. Cold air begins to penetrate North Vietnam during late
August and September. By the end of October, the entire Indochina
Peninsula is covered by the winter north easterlies. The monsoon
typically lasts through March. During the northeast monsoon season,
the South China Sea region is characterized by strong vertical shear.
The low-level monsoon north easterlies are overlaid with strong
westerlies from India. This environment makes tropical cyclogenesis
rare. More importantly, TCs moving into this region are typically
sheared, with the low-level circulation steered to the southwest. As with
the summer southwest monsoon, the winter monsoon, while persistent,
6
experiences surges and lulls. Surges occur when an upper-level trough
develops over northern China and moves off the east coast 24 hours
later. Lulls are associated with surface bubble highs moving off the east
coast of China, creating easterly or southeasterly flow over the South
China Sea.
1.1.5 Northwest Australian Summer Monsoon
Like its Indian counterpart, the Australian summer monsoon is
largely driven by the strong heat lows which form over Northern
Australia. However, since Australia lacks a mountain range like the
Tibetan Plateau, the monsoonal flow is neither as strong nor as steady
as the Indian monsoon. The onset of the monsoonal flow typically
occurs during January. As the Southeast Asian winter monsoon
strengthens, its flow penetrates deeper and deeper into the tropical
latitudes. By January, the flow crosses the equator near Indonesia and
turns westerly. These equatorial westerlies flow across New Guinea, the
Solomon Islands, and Northern Australia. The season lasts until April
when easterly trades return covering the entire area south of about
10°S. As with the Indian monsoon, a semi-permanent anticyclone is
located over the surface heat lows. Upper-level south easterlies to the
north of this anticyclone cross the equator, providing the return flow of
the Hadley circulation. This reversal in flow (upper-level south easterlies
over low-level north westerlies) results in a large vertical shear similar to
the Indian summer monsoon. However, in this case, the shear region is
only located equator ward of 10°S. Thus, unlike the Indian monsoon
season, tropical cyclogenesis still occurs during the Australian monsoon
season, but is confined to south of 10°S.
1.1.6 African Monsoon
The magnitude and the thickness (less than 6 km) of the air layers
of monsoonal circulations over Africa are smaller and shallower than
7
that over Asia. In the West Africa, a large continental area north of the
equator to about 15°N, there is a difference between the two lower
tropospheric monsoon winds: northeasterly trade winds in January and
southerlies to south westerlies in July. The air masses which the two
monsoon winds bring are different. The northeasterly trade winds
prevail to an elevation of about 3000 m and bring dry, stable, and often
dusty air, these winds are called the "harmattan". The southwesterly
monsoon winds are warm and humid. In the East Africa, the continent
stretches on both sides of the equator, these two monsoon winds differ
only in direction and the air masses which they bring are similar. In
January the Inter Tropical Convergence Zone (ITCZ) is located at about
15°S and most of East Africa is under the influence of northeasterly
winds, which become north westerlies south of the equator. In July the
ITCZ is situated at about 15°N and most of East Africa is under the
influence of southeasterly and southerly.
1.1.7 American Monsoon
The southwest region of North America is very arid, under the
general influence of a subtropical ridge of high pressure associated with
the thermal contrast between land and adjacent ocean. The North
American Monsoon (NAM) system develops in early July [Higgins and
Mo, 1997]; the prevailing winds over the Gulf of California undergo a
seasonal reversal, from northerly in winter to southerly in summer,
bringing a pronounced increase in rainfall over the southwest USA and
ending the late spring wet period in the Great Plains [Bordoni et ai,
2004]. The projection of smaller warming over the Pacific Ocean than
over the continent, and amplification and northward displacement of the
subtropical anticyclone, is likely to induce a decrease in annual
precipitation in the south-western USA and northern Mexico. It affects
Mexico along the Sierra Madre Occidental as well as Arizona, New
Mexico, Nevada, Utah, Colorado, West Texas, and California. It pushes
8
as far west as the Peninsular Ranges and Transverse Ranges of
southern California but rarely reaches the coastal strip. The North
American monsoon is known to many as the summer, Southwest,
Mexican or Arizona monsoon. It is also sometimes called the Desert
Monsoon as a large part of the affected area is desert. The monsoon
extends into the southwest United States as it matures in mid July when
an area of high pressure, called the monsoon or subtropical ridge,
develops in the upper atmosphere over the four corners region, creating
an easterly to southeasterly wind flow aloft. This wind flow pattern
directs moisture originating in the Gulf of Mexico, Gulf of California and
the tropical Pacific by way of northern Mexico into the region. As much
as 70% of rainfall in the region occurs during the summer monsoon.
Many desert plants are adapted to take advantage of this brief wet
season. Monsoons play a vital role in managing wildfire threat by
providing moisture at higher elevations and feeding desert streams.
1.2 Role of Tropical Ocean-Atmosphere interactions
One example of the way in which ocean-atmosphere interaction
introduces an Interannual Variability (IAV) of the Monsoon Annual Cycle
(MAC) is the mutual interaction between the El Nino Southern
Oscillation (ENSO) and the monsoon. This interaction is primarily
through the change in the equatorial Walker circulation influencing the
regional Hadley circulation associated with the Asian monsoon
[Goswami, 1998; Lau and Nath, 2000; Webster et al., 1998]. Since the
pioneering work of Sir Gilbert Walker [Walker, 1924], the influence of
the ENSO on the Asian Summer Monsoon (ASM) has been recognised
[Rasmusson and Carpenter, 1983; Shukla, 1987; Sikka, 1980]. As the
strong heat source associated with the ASM could influence the
atmospheric circulation in a significant way, it has also been recognized
that it could modify the surface stresses over the central and western
Pacific and influence the strength and evolution of the ENSO [Chung
9
and Nigam, 1999; Kirtman and Shukla, 2000; Yasunari, 1990]. These
independent studies of ENSO influence on the ASM and ASM influence
on the ENSO indicate that the ENSO and the ASM are not independent
phenomena but part of a coupled ocean-atmosphere oscillation.
Several recent studies with coupled ocean-atmosphere
GCMs
[Loschnigg et al., 2003; Wu and Kirtman, 2004; Yu et al., 2003]
investigate IAV of the ASM due to air-sea interaction involving ASM
and the ENSO. These studies indicate that the observed biennial
tendency of the ASM may be a result of such air-sea coupling. Based
on analysis of simulations of their coupled model, Wu and Kirtman
[2004] proposed a plausible mechanism through which a Tropical
Biennial Oscillation (TBO) may be generated. In the monsoon ENSO
connection, the impact of Indian monsoon on ENSO is most prominent
in JJAS when ENSO is developing. The effect of ENSO on the
monsoon transition occurs during DJF and MAM. A schematic diagram
as shown in figure 1.3 is used to summarize the mechanism. A strong
ASM during June-August (JJA) can enhance surface easterlies in the
central equatorial Pacific, induce an eastward-propagating upwelling
Kelvin wave, and give rise to negative SST anomalies in the eastern
Pacific that amplify through air-sea interactions. Colder SST in the
eastern Pacific is also associated with warmer SST in the western
Pacific. A strong ASM also cools the Indian Ocean through enhanced
evaporation and upwelling. Associated intensification of the Walker
circulation leads to divergence of moisture supply in the western Indian
Ocean. Reduced moisture supply at low levels together with upper-level
subsidence leads to a weaker ASM during the next summer.
10
DJF(2)i
Warm Indian Ocean
1......
1 /* Reduced **.
1 '--EyaporatiQD--'
JJA(l)
Weak"' Cold Western Pacific /Weak"', Warm Central Pacific
V.WQ.''
!.__
/'
Enhanced ~\
^''Air-Sea**
-interaction'-•.Jjyaporation-''
Anomalous Surface
Westerly
Weak Indian Monson
Warm Central Pacific
...1...
Reduced Moisture, ^
*j
s
A
DJF(l)
Cold Indian Ocean
,
~ >
"^Strong*, Warm Western Pacmt •Strong* - Cold"Central Pacific
\WC/'
.....I
I......
,*" Enhanced \
H
- -Ey aporatign- -'
JJA(O)
f
Anomalous Surface
Easterly
Strong Indian Monson
Indian Ocean
•' Air-Sea'*
-interaction'
,'"' Reduced **\
" - -EyapjqratiQD- -'
%
Cold Central Pacific
Pacific Ocean
•+H-
Figure 1.3: Schematic diagram showing the processes of monsoon-ENSO
interaction in the biennial oscillation during the strong monsoon year to the
weak monsoon year. Arrows in the figure denote the interactive processes.
Here, A and C denote anticyclone and cyclone, respectively; WC refers to
Walker circulation fWu and Kirtman, 2004].
A weak ASM induces opposite effect and can lead to a stronger
monsoon next year. This indicates that ocean-atmosphere interaction
could generate IAV of the ASM via generation of TBO signal. This twoway interaction between the monsoon and ENSO led to floating of the
"monsoon year" concept of climate year in the Tropics. However, the
recent analysis of Ailikun and Yasunari [2001] indicates that the biennial
transition takes place within the summer monsoon season. They also
found that the variability of the Asian monsoon in the early summer
(June) is associated with the anomalous state of the ENSO in the
previous winter while that of the mid-late summer (July-September) is
associated with the anomalous state of the ENSO in the following
winter. They find that the all-India monsoon rainfall in September is well
correlated with that in the following June indicates a continuity of the
climate year of the coupled ocean-atmosphere system. The fact that
all-India rainfall of June is uncorrelated with that of July-September
li
(JAS) rainfall indicates that the biennial transition takes place between
June and July. Based on these findings, they proposed a slightly
modified conceptual model monsoon year [Yasunari, 1991; Yasunari
and Seki, 1992] for the coupled monsoon/atmosphere-ocean system as
shown in figure 1.4.
eastern Pacific
eastern Pacific
^
Weak Monsoon Year
^*)
: weak monsoon
i
^>
•4
: strong monsoon
Strong Monsoon Year
•
BO: biennial oscillation
Figure 1.4: Schematic diagram of two successive monsoon years for the
coupled ENSO-monsoon system [from Ailikun and Yasunari, 2001].
This schematic diagram depicts sequence of two successive monsoon
years, a year starting from the mid-late northern summer (JAS) and
persists till early next summer (June). In the figure, a sequence of a
"weak" monsoon year is shown followed by a "strong" monsoon year.
The east-west Walker circulation is weakened associated with a weak
monsoon and leads to colder SST in the western Pacific and warmer
SST in the eastern Pacific in the following seasons (an El Nino
condition). The colder SST in the western Pacific is crucial in
maintaining anomalous ENSO state from winter until next June actively
leading to a weaker Asian monsoon during early summer. The right side
of the figure 1.4 depicts the processes in a strong monsoon year.
12
The strong convective activity during the mid-late summer leads
to a cold condition in the eastern Pacific and a warm condition in the
western Pacific during the following winter (a La Nina condition). The La
Nina signal in the western Pacific carried by the ocean until the next
June maintains a strong Asian monsoon in June. Thus, there is
evidence that air-sea interaction plays an important role in the observed
TBO. However, it is not clear that air-sea coupling is essential for
existence of the observed TBO. It may be noted that most Coupled
GCMs (CGCMs) that simulate TBO have a significant systematic bias in
simulating the climatological mean Annual Cycle (AC). Systematic
biases in simulating the climatological mean may play a significant role
in simulating a TBO in many of these CGCMs. Hence, it is not well
settled that air-sea coupling is essential for the existence of the TBO. It
is possible that a TBO may be triggered by atmospheric internal
dynamics but amplified through air-sea coupling. In addition to the
ENSO-related
ocean-atmosphere
interaction,
local
warm-ocean
atmosphere interaction over the Indian Ocean (10) and Western North
Pacific (WNP) can also give rise to IAV of the Monsoon Annual Cycle
(MAC).
The recently discovered Indian Ocean Dipole (IOD) mode [Saji et
al., 1999; Webster et al., 1999] is a good example of manifestation of
such air-sea interaction. This mode is not an equatorially confined
zonal mode. The SST dipole is coupled with the south 10 anticyclonic
anomalies. In the presence of the summer monsoon background flow,
the ocean to the east of the anticyclone near Sumatra cools as a result
of coastal upwelling, evaporation, and entrainment. Reduction of
convection associated with the cooling excites westward-propagating
descending Rossby waves and reinforces the anticyclone [Li et al.,
2003; Wang et al., 2003]. However, the role of the IO dipole mode on
IAV of the south Asian summer monsoon is unclear at this moment.
Similarly warm ocean-atmosphere interaction involving the WNP
13
anticyclone leads to IAV of the East Asian Monsoon (EAM) [Wang etal.,
2003].
We provide some empirical evidence that the observed MAC is
modulated by the coupled ocean-atmosphere interaction. The anomaly
of all-India monthly rainfall [Parthasarathy et a/., 1994] is obtained as
departure of observed monthly means from a climatological monthly
annual cycle. Anomaly in annual cycle (AC) is defined as the difference
between the JJA mean anomaly minus the following DJF mean
anomaly. The correlation between the two is -0.61 for the 130 years
period. ENSO is a set of specific interacting parts of a single global
system of coupled ocean-atmosphere climate fluctuations that come
about as a consequence of oceanic and atmospheric circulation. The
irregularity of ENSO makes predicting it of high interest, as it is
demonstrably connected to seasonal, even yearly, regional climatic
effects on large areas. ENSO is the most prominent known source of
inter-annual variability in weather and climate around the world (about 3
to 8 years), though not all areas are affected. ENSO has signatures in
the Pacific, Atlantic and Indian Oceans. El Nino changes the distribution
of rainfall, causing floods in some areas and drought in others.
1.3 Variability of the Southwest Monsoon
1.3.1 Intraseasonal Variability
Some of the studies (Krishnamurti and Bhalme, 1976; Sikka and
Gadgil, 1980; Goswami, 1994 and recently Kripalani ef a/., 2004) have
shown that the precipitation distribution over India varies considerably
from day to day, while over major parts of the country rains occur in
spells under the influence of favorable circulation conditions. Hence,
total seasonal monsoon strength may be related to the spells of active
and break periods. This intermittent behavior of rain is related to a
hierarchy of quasi-periods like 3-7 days, 10-20 days and 30-60 days.
Three to seven days periodicity is associated with oscillation of the
14
monsoon trough, while 10-20 days periodicity is associated with the
synoptic scale convective systems generated over a warm Bay of
Bengal propagating inland and contributing substantial precipitation.
Active spells of monsoon are characterized by a sequence of cluster of
such systems, whereas no disturbances occur during the break periods.
This study clearly shows that the seasonal monsoon strength may
depend on the frequency and duration of the spells of active and break
periods associated with these intra-seasonal oscillations. With the
availability of the satellite data (e.g. NOAA OLR, TMI SST, QuikSCAT
surface winds etc.) and reanalysis products (NCEP/NCAR and ERA),
better description of spatio-temporal characteristics of monsoon
Intraseasonal Oscillations (ISOs) have evolved over the past decade or
so. Such observations have revealed that the active and break of south
Asia monsoon (SAM) or the wet and dry spells over the Indian
continent, are manifestation of repeated northward propagation of the
Tropical Convergence Zone (TCZ) from the equatorial position to the
continental position [Sikka and Gadgil, 1980; Yasunari, 1979] and
results from superposition of a 10-20 day and a 30-60 day oscillations.
Both the 10-20 day oscillation and the 30-60 day oscillation contribute to
the total Intraseasonal Variability (ISV) in the SAM region. The 30-60
day mode is characterized by a northward propagation while the 10-20
day mode is characterized by a westward propagation. It indicates that
relative frequency of occurrence of active and break phases could
influence the seasonal mean and contribute to the Interannual
Variability (IAV) of the SAM. Major advances have been made during
the past two decades in understanding the temporal scale selection and
northward propagation of the 30-60 day mode and temporal scale
selection and westward phase propagation of the 10-20 day mode (see
Goswami [2004a]; Wang [2004] for detail). Until recently no clear
physical mechanism for the selection of period, wavelength and
westward phase propagation of the quasi-biweekly mode was known.
15
With the availability of high resolution reliable SST from satellite on daily
time scale and time series data from some moored buoys in the Bay of
Bengal [Sengupta and Ravichandran, 2001] it became clear that the
ISV of SST over the north Indian Ocean (IO) has large amplitude and
large spatial scale similar to that of the atmospheric ISV [Sengupta et
al., 2001]. Coherent northward propagation of intraseasonal SST,
surface wind speed, net heat flux at the surface and OLR (or
precipitation) are found [Sengupta et al., 2001]. Coupled modeling
studies [Fu et al., 2003; Zheng et al., 2004; Rajendran et al., 2004]
demonstrate that air-sea interaction is required to explain the observed
space-time spectra of summer ISO in SST and precipitation. The
monsoon ISOs are a crucial building block of the ASM. Through multiscale interactions with synoptic activity on one hand and the seasonal
cycle on the other, they determine not only the probability of occurrence
of daily precipitation but also the IAV of the seasonal mean. Hence,
conditions for cyclogenesis are much more favorable during an active
phase compared to a break phase. Like MJO, do the monsoons ISOs
too modulate the synoptic activity in the region during northern
summer? Using genesis and track data for Low Pressure Systems
(LPS) for 40 years (1954-1993), Goswami et al. [2003] show that
genesis of an LPS is nearly 3.5 times more favorable during an active
condition (147 events corresponding to normalized index > +1)
compared to a break condition (47 events corresponding to normalized
index < -1) of the monsoon ISO. They also show that the LPS are
spatially strongly clustered to be along the monsoon trough region
during an active condition. Since the day to day fluctuation of
precipitation is essentially governed by these synoptic activities, the ISO
phase modulates the probability of occurrence of daily precipitation.
Thus, the ISOs also have the potential to produce IAV of the seasonal
mean precipitation. The amplitude of the ISV (e.g. the coefficient of
variability) is much larger than the amplitude of IAV of the SAM. This
16
fact (the intraseasonal signal being strong) and the fact that the
monsoon ISOs are associated with quasi-periodic oscillations, indicate
potential predictability of the ISO phases beyond the medium range
weather prediction. Estimates made by Goswami and Xavier [2003];
Waliser et al. [2003] show potential predictability of the break phase of
monsoon ISO to be about three weeks while that of the active phase
being smaller. Earlier, Lo and Hendon [2000]; Mo [2001]; Jones et al.
[2004] demonstrated usefulness of empirical techniques in making
useful prediction of the ISO phases.
1.3.2 Interannual Variability
The interannual variability of the south Asian monsoon (SAM) is
rather modest with the interannual standard deviation being about 10%
of the seasonal mean. However, larger excess or deficits of all India
rainfall are associated with large spatial scale covering most of the
country [Shukla, 1987]. Extremes in monsoon rainfall leads to
devastating floods and droughts [Shukla, 1987; Mooley and Shukla,
1987; Webster et al., 1998; Sikka, 1999] leading to enormous economic
loss and human misery. Therefore, understanding of the physical
processes responsible for the observed IAV of SAM is crucial for
advancing the capability for predicting the IAV. One notable connection
with the IAV of the SAM is that with the ENSO. There is a tendency for
the El Nino's to be associated with droughts and La Nina's to be
associated with above normal conditions over India. While a connection
between the SAM and the ENSO exits, it is not very strong. It is worth
noting here that many droughts and floods of the SAM occur without El
Nino or La Nina. Since the pioneering work of Sir Gilbert Walker
[Walker, 1924], this influence of the ENSO on the SAM has been
investigated [Sikka, 1980; Rasmusson and Carpenter, 1983; Shukla,
1987]. The strong heat source associated with the ASM could indeed
influence the atmospheric circulation in a significant way and could
17
modify the surface stresses over the central and western Pacific and
influence the strength and evolution of the ENSO [Yasunari, 1990;
Chung and Nigam, 1999; Kirtman and Shukla, 2000]. These
independent studies of ENSO influence on the ASM and ASM influence
on the ENSO, made it clear that the ENSO and the ASM are not
independent phenomena but part of a coupled ocean-atmosphere
oscillation. Associated intensification of the Walker circulation leads to
divergence of moisture supply in the western Indian Ocean. Reduced
moisture supply at low levels together with upper level subsidence leads
to a weaker ASM during the next summer. A weak ASM induces
opposite affects and can lead to a stronger monsoon next year. The
SAM is not driven by the land-sea surface temperature gradient but by
the tropospheric temperature gradient. In fact Liu and Yanai [2001] find
a significant positive correlation between March-April-May (MAM) upper
tropospheric temperature over western Europe and All India monsoon
rainfall (AIR). In addition to the ENSO related ocean-atmosphere
interaction, local warm-ocean atmosphere interaction over the Indian
Ocean (IO) and western north Pacific can also give rise to IAV of the
Monsoon Annual Cycle (MAC). Recently discovered Indian Ocean
Dipole Mode (IODM, Saji et al. [1999]; Webster et al. [1999]) is a good
example of manifestation of such air-sea interaction. This mode is not
an equatorially confined zonal mode. The SST dipole is coupled with
the south IO anticyclonic anomalies. In the presence of the summer
monsoon background flow, the ocean to the east of the anticyclone near
Sumatra cools due to coastal upwelling, evaporation and entrainment.
Reduction of convection associated with the cooling excites westward
propagating descending Rossby waves and reinforces the anticyclone
[Li et al., 2003; Wang et al., 2003]. This air-sea interaction also
contributes to a quasi-biennial signal of the monsoon [Loschnigg et al.,
2003; Li et al., 2003]. Similar warm ocean-atmosphere interaction
18
involving the Western North Pacific (WNP) anticyclone leads to IAV of
the EAM [Wang etal., 2003].
1.4 Winds and Currents over the Equatorial Indian Ocean
Surface Winds
The seasonal reversal of the surface wind fields over the tropical Indian
Ocean is far more dramatic than in the other regions of the low
latitudes. Over the large part of the Indian Ocean the surface wind
forcing completely reverses between the boreal winter and boreal
summer monsoons (figure 1.5). In January, the strong Siberian high
produces flow off the Asian continent and across the equator towards
the Inter Tropical Convergence Zone (ITCZ). Therefore, winds are north
easterlies over most parts of the Indian Ocean north of the equator.
While between equator and 15°S northwesterly winds prevail in this
month. The southeasterly winds exist throughout the year in the region
south of 15°S. Similar type of conditions sustained in the entire winter
monsoon season (December to February). During March to May, winds
are weak over the Indian Ocean north of 10°S. The onset can be an
abrupt change from weak pre monsoonal winds into the fully developed
southwest monsoon in early to mid-June [Schott and McCreary, 2001].
At lower levels of the atmosphere, during the summer season owing to
the presence of the land - ocean contrasts and orographic barriers, the
seasonal wind reversal, the associated south to north migration of the
ITCZ is highly modified. The southwest monsoon winds become very
strong in July and August with strong south westerlies in the north of the
equator and south easterlies south of the equator. By October
southwest monsoon winds get weaken and light winds prevail over the
north Indian Ocean. Similar conditions prevail in the Month of November
(figure 1.5). In the equatorial Indian Ocean unique wind forcing pattern
occurs, which is unlike the pattern in other equatorial oceans, it involves
19
the occurrence of semiannual eastward winds over the equator during
April to June and October to November [Schott and McCreary, 2001].
a) JAN
V^tfrffl' . < • • • •
. ^ „ . • « » . . . . • • • • • • -
d) JULY
^>
- ».*£
r,
•
*
- ^*
MC
^T
c ) MAY
0 NOV
Figure 1.5: Bi-monthly surface (10 m) wind pattern over the Indian Ocean from
NCEPNCAR climatology (1950-2006).
Surface Currents
The circulation of the Indian Ocean north of 20°S is characterized
by seasonally reversing currents in response to monsoon wind forcing
[McPhaden, 1982]. A schematic representation of climatological
20
circulation in the Indian Ocean is shown in figure 1.6 [adopted from
Shankar et ai, 2002]. These currents are derived from ship drift
climatology [Cutler and Swallow, 1984] and from drifters [Molinari era/.,
1990; Shenoi era/., 1999]. In the Indian Ocean, the region south of 10°S
is not subjected to the seasonal reversal of winds and circulation pattern
[Swallow et ai, 1988]. The westward flowing South Equatorial Current
(SEC), driven by the southeast trades, exists within the latitudinal range
12°S - 25°S during all seasons. During winter monsoon the atmospheric
and oceanic circulation patterns are similar to those found in the Pacific
and Atlantic. Winter Monsoon Current (WMC) originating southeast of
Sri Lanka (figure 1.6) flows westward centered along 7°N. Along the
west coast of India, a part of WMC circulate around Lakshadweep High
(LH) and feeds into the West Indian Coastal Current (WICC). The
westward flowing branch of WMC turns southward along the Somali
Coast and forms the Somali Current (SC). The SC and the East African
Coastal Current (EACC) confront south of equator and flows eastward
as South Equatorial Counter Current (SECC) at 2-4°S. East India
Coastal Current (EICC) flows towards the south in the Bay of Bengal.
Equatorial Current (EC) flows east to west at the equator. During the
summer monsoon the intense winds in turn drive the ocean currents
eastward north of the equator. The eastward flowing low latitude
Southwest Monsoon Current (SMC) reaches south of Sri Lanka in May.
By mid-June a part of SMC turns northward into Bay [Vinayachandran
era/., 1999], flowing around the cyclonic Sri Lankan Dome (SD) as East
Indian Coastal Current (EICC). During the summer monsoon season,
the SEC and EACC supply the northward flowing SC. A series of gyres
are seen along the east African coast with the establishment of summer
monsoon, as Southern Gyre (south of 4°N) and Great Whirl (4°N-10°N).
Most of the supply to the SMC appears to be from SC and a part of its
source water feeds from southward flowing West Indian Coastal Current
(WICC).
21
Schematic of circulation in the Indian Ocean
30
N
•
:
»' \
'
\ Jamwrv \
20 N ': \
<
' -' •
(
\
< \
Arjhun
V
-'
' ' ••••-. of
^ t w ^ \ V
10 N
:\J«il>
V - ^
•A
:\ i
20 N • \ \
< \
10
- - ^
r
lilJui
^ >
"J?"'
N-
<->
//li,v,i
SMI
•.'!.••'! A
.^
<
>10 S
f
/
^
W\
t
•«..v»>^'-.;
Sri U n k i V f r ? * ^
-
_
Sit <
: t(«v'
sir
•T
40 E
;
SMi
U
•
>
An^f^VV
50 E
60 E
70 E
.-,,
80 E
90 E
100 E
Figure 1.6: Schematic representation of the circulation in the Indian Ocean
during January (winter monsoon) and July (summer monsoon). The
abbreviations are as follows: SC, Somali Current; EC, Equatorial Current;
SMC, Summer Monsoon Current; WMC, Winter Monsoon Current; EICC, East
India Coastal Current; WICC, West India Coastal Current; SECC, South
Equatorial Counter Current, EACC, East African Coastal Current; SEC, South
Equatorial Current; LH, Lakshadweep high; LL, Lakshadweep low; and GW,
Great Whirl [Shankar et al., 2002].
The transition
between the
monsoons
(April-May
and
October-
November) is marked in the equatorial Indian Ocean (5fiS-5QN) by
westerly winds and strong eastward jets as first identified by Wyrtki
[1973]. Reppin et al. [1999] observed that these equatorial jets or Wyrtki
fall jets are stronger (> 1.2 ms"1) than spring jets (1 ms"1).
22
1.5 Prediction methods of southwest monsoon
India Meteorological Department (IMD) has been issuing Long
Range Forecast (LRF) of the southwest monsoon rainfall since 1886.
The extensive and pioneering work of Gilbert Walker [1923 and 1924],
led to the development of the first objective models based on statistical
correlations
between
monsoon
rainfall
and
antecedent
global
atmosphere, land and ocean parameters. Since then, IMD's operational
LRF system has undergone changes in its approach and scope from
time to time. In the last 50 years many reviews on the LRF of Indian
Southwest Monsoon Rainfall (ISMR) are available in the literature e.g.
Normand [1953]; Jagannathan [1960]; Thapliyal and Kulshreshtha
[1992]; Hastenrath [1995]; Krishna Kumar etal. [1995]; Rajeevan [2001]
and Gadgil et al. [2005]. In a very recent study, Gadgil et al. [2005]
addressed the major problems of the statistical and dynamical methods
for LRF of monsoon rainfall in view of the recent forecast failures in
2002 and 2004. Their analysis revealed that IMD's operational forecast
skill based on statistical methods has not improved over seven decades
despite continued changes in the operational models. For the LRF of
the ISMR, three main approaches are used. The first is the statistical
method, which uses the historical relationship between the ISMR and
global atmosphere-ocean parameters [Walker, 1914 and 1923;
Thapliyal, 1982; Gowariker et al., 1989 and 1991; Navone and
Ceccatto, 1995; Singh and Pai, 1996; Guhathakurta et al., 1999;
Rajeevan et al., 2000, 2004 and 2005; Delsole and Shukla, 2002; Sahai
et al., 2003; Pai and Rajeevan, 2006]. The second approach is the
empirical method based on a time series analysis. This method uses
only the time series of past rainfall data [Goswami and Srividya, 1996;
Iyengar and Raghukanth, 2004; Kishtawal et al., 2003] and do not use
any predictors. The third approach is based on the dynamical method,
which uses general circulation models of the atmosphere and oceans to
simulate the summer monsoon circulation and associated rainfall. In
23
spite of its inherent problems, at present, statistical models perform
better than the dynamical models in the seasonal forecasting of ISMR.
The dynamical models have not shown the required skill to accurately
simulate the salient features of the mean monsoon and its interannual
variability [Latif et al., 1994; Gadgil and Sajani, 1998; Krishnamurti et
al., 2000; Kang et al., 2002; Gadgil et al., 2005; Krishna Kumar et al.,
2005; Wang et al., 2005]. During the period of 1988-2002, IMD's
operational forecasts were based on the
16-parameter
power
regression and parametric models [Gowariker et al., 1989 and 1991].
The forecasts issued during this period were qualitatively correct.
However, the mean forecast error during this period was more than the
mean error of the forecasts based on climatology alone. This model
failed to predict the severe drought of 2002. Following the failure of
forecast in 2002, a critical evaluation of the 16-parameter power
regression and parametric models was made and in 2003, two new
models (8 and 10 parameter models) were introduced for the
operational work. Further a two-stage forecasting strategy was also
adopted with the provision for a forecast update by end of June/first
week of July [Rajeevan et al., 2004]. According to this new strategy,
IMD's operational forecasts for the seasonal ISMR for the country as
whole are issued in two stages. The first stage forecast is issued in mid
April and an update or second stage forecast is issued by the end of
June. While the 2003 and 2005 operational forecasts for the southwest
monsoon rainfall based on these new models were accurate, the
forecast for the 2004 monsoon was false. In spite of all available
literatures about Southwest Monsoon (SWM), understanding and
prediction using different meteorological parameters like wind, SST,
OLR, rainfall, Wind Shear, etc. do not give complete understanding for
the small / large temporal scale variability. Generally prediction of SWM
considers total seasonal rainfall prediction using different methods
(Synoptic, Dynamic and Statistical). But in the present study the main
24
intension is to understand and predict the onset date of SWM at Kerala
coast and to determine the impact of tropical Indian Ocean sea surface
temperature (SST) over southwest monsoon variability.
1.6 Motivation of the study
The Asian summer monsoon, manifested in all its glory and fury
over the Indian subcontinent, is the largest seasonal abnormality of the
global climate system. During the monsoon, the equatorial region is
colder than the regions to the north. The summer monsoon rains that
result are critical for food production, water supply, and the economic
well-being of the Asian society. There is thus great interest in predicting
the waxing and waning of the Asian monsoon. What are the prospects
for predicting monsoon rainfall over India and the surrounding regions?
Why has the accuracy (or "skill") of monsoon forecasts been so low?
What are the projected impacts of global warming on the Asian summer
monsoon? Monsoon forecasting has a long history in India. After the
subcontinent had experienced a devastating drought and famine in
1877, the British Government asked the recently established India
Meteorological Department (IMD) to forecast monsoon rainfall. It is
important to understand the mechanisms responsible for the monsoon
and its variability at intraseasonal, interannual and decadal time scales.
In fact, a simple tool for use of common man is required for quantitative
seasonal rainfall prediction of Indian monsoon. But, there is no clear
consensus on the mechanisms of monsoon interannual variability, and
the influence of the Indian Ocean on monsoon variability remains an
open question [Slingo et al., 2004]. The limited understanding has been
a big barrier to predict monsoon accurately. Hence, predictions are
uncertain and often fail.
25
1.7 Objectives of the Thesis
The main objective of this thesis is to understand and predict the
southwest
monsoon variability
using oceanic
and
atmospheric
parameters over Tropical Ocean. Several observations such as OLR
(NOAA), GPCP daily rainfall, monthly all India rainfall data from IITM
website, historical SST (HadlSST) data from Hadley centre, reanalysis
daily wind (NCEP/NCAR) datasets and TMI (Tropical
Rainfall
Measuring Mission (TRMM) Microwave Imager) SST are used in this
study. In addition a daily 1° x 1° gridded rainfall data product from IMD
has been also used. Detailed information about the observational
datasets and method of derived parameters are provided in the
Chapter-2. Intraseasonal, interannual and interdecadal variation in
tropospheric wind shear and tropical convection over southwest
monsoon region west of 80°E are studied using wind, rainfall and OLR
observations and are discussed in Chapter-3. Chapter-4 describes the
interannual and intraseasonal variability of southwest monsoon in terms
of drought and flood years, active and break spells, temporal scale of
southwest monsoon, daily and seasonal rainfall variability and number
of synoptic systems using SST and rainfall data. In Chapter-5, two new
methods are suggested for prediction of onset date over Kerala coast
using SST and OLR data. Chapter-6 addresses about the sea surface
temperature over equatorial Indian Ocean and its effect on Indian
summer monsoon performance and northern limit of southwest
monsoon. Finally, in Chapter-7 important results of the research work
are summarized.
26