Download Large-Scale Tropical Circulations – Some General Aspects

Chapter 1
Large-Scale Tropical Circulations – Some
General Aspects
1.1 Introduction
Meteorology in general and tropical meteorology in particular have made tremendous strides during the last half a century or so after the introduction of computer
technology and space satellites. Undeniably, we now know much more about our
atmosphere and its behaviour than ever before. It is now well established that tropical circulation forms an integral part of the global general circulation and that there
is a continual exchange of heat, momentum and moisture between the tropics and
the rest of the atmosphere.
Yet, tropical circulation has some distinctive characteristics of its own which
need to be identified and studied independently in a more comprehensive manner.
At present, uncertainty prevails in several areas of interest and there are many
dark or twilight areas. In a seminal paper, Charney and Shukla (1981) had observed
that quite unlike the midlatitude weather systems the predictability of which was
limited by short period baroclinic wave activity, tropical weather systems which are
determined by circulations between long-period heat sources and sinks are more
predictable. Surface and boundary layer characteristics, such as ocean surface temperature, seasonal ground heating and cooling, have much longer lifespans and as
such amenable to prediction over longer periods of time. It is proposed to show
in the present text that tropical circulation systems are basically forced by boundary
layer heat sources and sinks, on different time and space scales, though wave activity
arising from flow instability plays a significant role in weather-forming processes.
However, several questions regarding tropical circulations remain unresolved.
For example, with a heat source centered at the equator and flanked by heat sinks in
both zonal and meridional directions, what sort of circulations and weather patterns
would evolve over the equatorial region? In this area, we haven’t gone far beyond
what we learnt from the excellent theoretical work of Matsuno (1966), Gill (1980,
1982) and others.
An uncertainty shrouds the existence or non-existence of double equatorial
troughs in some oceans and continents. For that matter, what is the origin of the
SW Pacific Convergence Zone (SPCZ), or the SW Atlantic Convergence Zone
(SACZ)? Are they to be regarded as equatorial, tropical, subtropical or extratropical convergence zones? Monsoon is probably one area where our knowledge
K. Saha, Tropical Circulation Systems and Monsoons,
C Springer-Verlag Berlin Heidelberg 2010
DOI 10.1007/978-3-642-03373-5_1, 3
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1 Large-Scale Tropical Circulations – Some General Aspects
and understanding lack precision. Firstly, its definition and then the manner of its
advance and retreat over different parts of the globe! In the past, monsoon has been
defined almost exclusively in terms of either rainfall or seasonal reversal of the prevailing surface wind. Little is known of the structure of monsoon and the manner of
its advance and retreat in different areas. Notwithstanding commendable progress
made in studies of monsoon (for a recent review of literature, see Webster et al.,
1998) over the globe in general, we yet do not have a clear idea of its structure and
how it advances or retreats over the Arabian Sea or the Bay of Bengal, for example,
during the northern summer. What brings up the monsoon current over the Arabian
Sea which remains relatively cold compared to its surrounding land areas during
summer? Then, again, what brings about the well-known intraseasonal oscillation
in monsoon rainfall and other meteorological parameters? In some parts of the tropics, the equatorial trough of low pressure which changes its orientation with season
shows strong correlation with the genesis of tropical cyclones/hurricanes/typhoons.
What is the connection between the two?
Another area where we lack precise information is how large mountain systems,
such as the Andes in South America, the Rockies in North America, the Mountains
of East Africa and the Great Himalaya complex with Tibetan Plateau in Asia affect
circulation systems over the respective continents.
Our knowledge of the effect of ocean currents and ocean surface temperatures on
equatorial circulation is also inadequate. Here we have not gone far beyond what we
learnt from the pioneering studies of Walker (1924), Bjerknes (1966, 1969) and others. There are many such areas where we need further and better information from
observations and their analysis and diagnosis, along with theoretical and dynamical
studies, than heretofore. In a lecture delivered during a seminar in 1979, Charney
had observed that ‘Data by themselves are not sufficient, nor is mind. We need a
combination of both’. A summary of this lecture was prepared by J. Shukla (1979)
from notes taken by him.
To-day, equipped with much better network of surface and upper-air observations
and improved methods of data analysis, we have much better opportunity to have
a fresh look into some of the above-mentioned outstanding problems of tropical
meteorology and apply our mind to them than ever before. Frankly, that is the main
objective of this book.
1.2 Tropical Circulation as Part of the Global General
Circulation – The Tradewinds
From time immemorial, mariners sailing over oceans in low latitudes for trading
purposes encountered a set of highly-steady seasonally-reversing winds which they
called the tradewinds and which they used to great advantage in moving from place
to place.
In late 17th century, Edmund Halley (1686) was the first to make a detailed study
of the tradewinds with the data then available and hypothesized that the observed
1.2
Tropical Circulation as Part of the Global General Circulation – The Tradewinds
5
winds at the surface were part of a direct thermally-driven circulation between a heat
source and a heat sink, which reversed its direction between winter and summer
But it was Hadley (1735) who offered an explanation for the cause of the
tradewinds as well as their observed reversal of direction on the basis of differential heating between the equator and the poles and the rotation of the earth. He
argued that a general equatorward drift of the tradewinds at low levels required a
compensating poleward drift at high levels in order to prevent an undue accumulation of mass near the equator. Further, a general westward drag by the tradewinds
near the earth’s surface at low latitudes due to the rotation of the earth required a
compensating eastward drag by the westerlies in high latitudes so as to prevent a
general slowing down of the earth.
It was found later that the general westward or eastward drift of the wind was
consistent with the principle of conservation of absolute angular momentum of the
earth. A parcel of air moving equatorward from high latitudes in order to conserve
the angular momentum of its original latitude would acquire an increasingly westward drift, while a poleward-moving parcel would acquire an increasingly eastward
drift. This was due to the fact that the earth’s surface moved faster at the equator
than at higher latitudes.
Hadley’s idealized single-cell circulation model held ground and went unchallenged for nearly a century and it was once thought that Hadley’s model was
representative of mean meridional circulation over all parts of the globe at all times
of the year. However, later observations called for a modification of Hadley’s idealized single-cell model. The new observations revealed the presence of a well-marked
high pressure belt over the subtropics and a low pressure belt further poleward near
60◦ latitude, which suggested a meridional pressure gradient and a poleward drift of
air, instead of an equatorward drift near surface, and a compensating equatorward
drift at some height, over the midlatitudes. Further, the westerly wind over the midlatitudes were found to be baroclinically unstable and characterized by large-scale
eddy motion.
Amongst the early attempts to modify Hadley’s original scheme were those of
Thomson (1857) and Ferrel (1859) who introduced a shallow indirect cell, characterized by a poleward flow near the surface and equatorward flow at some height,
over the midlatitudes, within the framework of the idealized single-cell Hadley
circulation model.
Further modifications to the meridional circulation model were made in the light
of later observations. That tropical circulation forms a distinct integral part of a
three-cell global general circulation model was first spelt out by Rossby (1947) and
others almost two centuries after Hadley (1735) had proposed his single-cell poleto-equator zonally-averaged annual-mean general circulation model.
The model proposed by Rossby (1947) and others, is in Fig. 1.1.
Rossby’s three-cell meridional circulation model shows a direct circulation cell
over the tropical belt with rising motion over the equatorial region and sinking
motion over the subtropical belt, an indirect circulation cell over the midlatitudes
and a direct circulation cell over the polar latitudes, with a polar front located at a
latitude of about 60◦ . The Rossby model has, by and large, stood the test of time
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1 Large-Scale Tropical Circulations – Some General Aspects
Fig. 1.1 Schematic of a three-cell meridional circulation model proposed by Rossby (1947) and
others
and found general acceptance by the scientific community to this day. In this model,
the classical Hadley circulation has been shown to have its rightful place over the
tropical belt only, as against the original model with a pole-to-equator circulation.
In the Rossby model, the tradewinds blowing towards the equator pick up
increasing amounts of heat and moisture from the warm ocean surface and deliver
them in rising currents over the equatorial belt to form cloud and precipitation and
then turn around in the upper troposphere to blow as anti-trades towards the subtropical belt where they subside to low levels to blow as low-level trades or tradewinds
again. This constitutes the direct kinetic energy-producing tradewind circulation cell
of low latitudes shown in the Rossby model.
1.3 Poleward Boundary of the Tropical Circulation
In the past, there appears to have been some controversy and divergence of
opinion regarding the definition of tropics, particularly its poleward boundary.
Climatologists divide climates on the earth’s surface into three distinct latitudinal
belts, calling that which lies within the first 30◦ parallels of latitude, the tropics or
the hot climate zone, as distinct from the other 30-degree belts which lie poleward
with temperate and frigid climates. Defined in this way, the tropics is the hottest
part of the earth’s surface with the sun always shining overhead somewhere over the
region. Some people even put the poleward boundary of the tropics at the solstices,
that is at 231/2◦ of latitude north and south of the equator.
However, meteorologists, by and large, are aware that the poleward boundary of
the tropical circulation is not determined by any fixed latitude, since it meanders
about its annual mean location considerably with longitude and season. To them,
the ridge of the global subtropical high pressure belt at surface, which divides the
easterly tradewinds from the midlatitude westerly winds, would appear to be a more
reasonable and acceptable choice.
That this last is the correct view is also suggested by a three-dimensional examination of the structure of the tropical atmosphere vis-à-vis that of the higher
1.4
Heat Sources and Sinks
7
latitudes. Besides surface temperatures, the vertical distribution of temperature over
the two belts is such as to create a fairly sharp thermal and circulation divide
between the tropics and the midlatitudes, which creates a strong horizontal temperature gradient across a narrow boundary zone to drive a westerly wind maximum
known as the subtropical jet (STJ) along the poleward boundary of the tropics.
One of the earliest to find the existence of this jet was Palmen (1951) who studied
the wintertime mean meridional circulation on the earth’s surface. Palmen’s finding
has been substantiated by several other studies (e.g., Krishnamurti, 1961; Wallace
and Hobbs, 1977; Galvin, 2007).
According to observations, the boundary moves equatorward of its annual-mean
latitude during the winter and poleward during the summer in a hemisphere. So, the
tropics of our concern in the present text will be the latitudinal belt between the
subtropical belts of the two hemispheres, bounded in each hemisphere by the ridge
of the subtropical high pressure at surface and the subtropical jet in the upper air.
The tropical circulation, however, is not isolated. In some parts of the globe,
a tropical circulation system may extend to high extratropical latitudes, or an
extratropical circulation system extend to very low tropical latitudes, and there is
continual interaction and exchange of heat, mass, and momentum between the two
circulations.
1.4 Heat Sources and Sinks
Before proceeding further, we need to define what we mean by heat sources and
sinks and state how we can identify them in the atmosphere.
1.4.1 Definition of Heat Sources/Sinks
At any given time of the year, when the atmosphere over a place is continually
heated (cooled) relative to its surrounding, the differential heating creates a heat
source (sink) over the place. We define a heat source or a heat sink by the following
criteria:
dH/dt = 0,
∇ 2 dH/dt < 0,
Heat source
(1.4.1)
dH/dt = 0,
∇ 2 dH/dt > 0,
Heat sink
(1.4.2)
Where H denotes the steady-state heat content of the air at a given place at time t,
and ∇ is a Del operator. Here, H = ρcp T, where ρ is density, cp is specific heat at
constant pressure p, and T is temperature determined by local heat balance.
Thus, to maintain a steady-state, a heat source must give up its excess heat to the
environment, and a heat sink must make up its deficit heat by acquiring heat from
the environment.
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1 Large-Scale Tropical Circulations – Some General Aspects
1.4.2 Diabatic/Adiabatic Heat Sources/Sinks
1.4.2.1 Diabatic Heating/Cooling
A diabatic heating or cooling process is one in which a working sample of the
atmosphere is free to exchange heat with its environment. There are three important processes of diabatic heating or cooling in the atmosphere in the tropics. These
are:
(i) Absorption or emission of short- and long-wave radiation;
(ii) Latent heat released by condensation of water vapour, or lost through evaporation of water, and mixing; and
(iii) Sensible and latent heat gained from, or lost to, a boundary surface through
turbulence and convection.
1.4.2.2 Adiabatic Heating/Cooling
In an adiabatic process, the air sample has to work in a closed system and is
barred from sharing its heat with the environment. This means that any change that
occurs in its temperature is due to its own expansion or compression. In this process, air with upward velocity cools by expansion to a lower pressure and that with
downward velocity warms up by compression to a higher pressure.
1.4.2.3 Temperature Change in the Atmosphere Due to Condensation Heating
Both diabatic and adiabatic processes are important in producing temperature
change at a point in the atmosphere. The most important diabatic heat source in
a tropical disturbance is the release of latent heat of condensation of water vapour in
the atmosphere which amounts to nearly 2.5 × 106 J kg–1 and this value increases
with temperature.
The temperature change produced by this process may be computed by using the
approximate relation (Anthes, 1982)
(∂T/∂t)cond = (L/cp )C,
(1.4.3)
where C is the local condensation rate (mass of water vapour condensed per unit
mass of air per unit time) and L is the latent heat of condensation of water vapour.
The local condensation rate in a saturated updraft, for example in a mature
thunderstorm, may be computed by using the approximate relation
C̃ ∼ −ω∂qs /∂p
(1.4.4)
where ∂qs /∂p represents the quantity of water released between two pressure surfaces assuming a saturated adiabatic lapse rate of temperature. Anthes (loc. cit.)
finds that for a saturated adiabat through, say, 24◦ C, this term would yield about
1.5
Some Physical and Dynamical Constraints and Conservation Laws
9
5 g kg–1 of water in the 200 mb-layer between 800 and 600 mb. For an updraft
velocity of 1000 mb/h, the value of C is 25 g kg–1 h–1 . Substitution of this value of
C in (1.4.3) yields a condensation heating rate of about 1500 K (day)–1 . Of course,
such an enormous heating rate is never observed in the atmosphere. This is because
the diabatic heating is almost totally compensated by adiabatic expansion and cooling, so that actual change of temperature at any level is only a small imbalance
between diabatic heating and adiabatic cooling. It is only near the lower boundary
where vertical motions are constrained to be small that diabatic heating produces
large changes of temperature.
1.4.2.4 Identification of Heat Sources and Sinks
Heat sources and sinks are created in the atmosphere by both diabatic and adiabatic
processes and may be identified generally by their effects on the distribution of
temperature and pressure at surface and aloft.
At the lower boundary of the atmosphere, a diabatic heat source (sink) may be
identified with a ‘heat low’ (‘cold high’). In the atmosphere above, a diabatic heat
source is created by condensation of water vapour and a diabatic heat sink by the
reverse process of evaporation of drops, or by radiative cooling.
In the atmosphere, in the absence of precipitation, convection generally leads
to adiabatic cooling and subsidence to adiabatic warming. The processes lead to a
‘cold low’ and ‘warm high’ respectively. This means that a ‘cold low’ is produced
when warm air rises and expands adiabatically by rising to a lower pressure, and a
‘warm high’ when a sample of cold air is compressed and heated adiabatically by
subsidence to a higher pressure, i.e., to lower elevations.
1.5 Some Physical and Dynamical Constraints
and Conservation Laws
Circulations forced by heat sources and sinks are, however, required to comply with
certain physical and dynamical constraints and conservation laws. Some of these are
stated below.
1.5.1 Direct and Indirect Circulations
When a heat source over a place gets continually heated up, it tends to maintain
itself as a source by transferring the excess heat to a heat sink in the neighborhood.
Likewise, when a heat sink continually gets colder, it tends to maintain itself as
a heat sink by transferring its excess cold to a neighboring heat source. The heat
exchanges are assumed to take place via a direct kinetic energy-producing vertical
circulation in which warm air rises and cold air sinks. But due to adiabatic processes
in the atmosphere, the rising air cools and the sinking air warms up. So, the transfer
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process requires an independent source of energy to provide for the kinetic energy
of the direct circulation. This source must be an indirect circulation which would
generate sufficient available potential energy to provide for the kinetic energy of the
direct circulation.
1.5.2 Energy Transformations
Since the total energy of the circulations in a closed frictionless domain must remain
constant, it follows that the required available potential energy (both zonal and eddy,
as defined by Eqs. (1.5.1) and (1.5.2) below), must be generated in one part of a
system to provide for the kinetic energy of the other part, in accordance with the
following relationships (after Krishnamurti and Surgi, 1987):
Generation of zonal available potential energy (Pz )
γ [ < T > ][ < Q > ]dm
(1.5.1)
G(Pz ) =
m
Generation of eddy available potential energy (Pe )
γ [ < T Q > + < T >∗ < Q >∗ ]dm
G(Pe ) =
(1.5.2)
m
Conversion from zonal available potential energy (Pz ) to zonal kinetic energy
(Kz )
< Pz .Kz >= − [ < ω > ] [ < α > ] dm
(1.5.3)
m
Conversion from eddy available potential energy (Pe ) to eddy kinetic energy (Ke )
[ < ω α > + < ω >∗ < α >∗ ]dm
< Pe .Ke >= −
m
where γ is a stability parameter, and the symbols used with meanings are:
ω vertical p-velocity, where p is pressure,
T temperature,
Q total diabatic heating per unit mass (m) of air,
α specific volume,
Deviation from zonal mean,
Deviation from horizontal area average,
m integration over mass, m,
<> zonal average,
[] time average, and ∗ deviation from time average.
(1.5.4)
1.5
Some Physical and Dynamical Constraints and Conservation Laws
11
1.5.3 Energy Transfer Process – Carnot’s Cycle
In physics, Carnot’s heat engine works in a closed system and the process of heat
transfer between a source and a sink is in a cycle which is strictly reversible. In
the open atmosphere, the principle of reversibility clearly does not hold, since the
working substance which is a part of the environment is exposed to the environment
with which it can exchange heat. Yet, the general principle of the Carnot’s heat
engine has been found to be useful in interpreting the atmospheric heat transfer
processes.
Let us see how it works. In the first stage, the working substance which is in
contact with a heat source which in the atmospheric case is latent heat released by
condensation draws a certain quantity of heat from the source by rising and expanding isothermally to a somewhat lower pressure; in the second stage, it expands
adiabatically by rising to a much lower pressure and temperature; in the third stage,
it descends isothermally to a somewhat higher pressure and delivers a certain quantity of heat to the heat sink with which the working substance is now placed in
contact, from where an adiabatic compression to a higher pressure and temperature
during the fourth stage will restore the working substance to its original pressure
and temperature.
Thus, the working of the Carnot’s cycle in the atmosphere represents a heat transfer process in which cold air is raised adiabatically to become colder and warm
air lowered adiabatically to become warmer, a process which generates available
potential energy by an indirect circulation.
1.5.4 Conditional Instability and Convection
The tropical atmosphere most often remains in a state of conditional instability in the
sense that it is unstable (∂θ e /∂z < 0) in the lower troposphere below about 700 hPa,
but stable (∂θ e /∂z > 0) above, where θ e is the equivalent potential temperature of
the air which may be defined as ‘the temperature attained by a parcel of air which
is raised adiabatically from its existing level to a level where all its moisture is
condensed out and the latent heat of condensation added to it, and then brought
down dry-adiabatically to a standard level, usually 1000 mb.’
However, it does not follow from this stability condition that it will automatically
lead to convective overturning. In the tropics, moisture is usually confined to a shallow boundary layer in contact with the earth’s surface. So, convective overturning is
only possible when this low level moisture is lifted to higher levels by synoptic-scale
convergence. It is not surprising, therefore, that most of the tropical disturbances
such as depressions and cyclones develop over oceans where low level convergence
is able to lift the boundary layer moisture to higher levels for large-scale release
of the latent heat of condensation which will potentially lead to explosive growth
of the disturbance. Such developments may also occur over land areas where there
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is a copious supply of moisture from neighboring oceans and strong lapse rate of
temperature may develop during afternoon-evening hours. In all cases, cyclogenesis
must lift the low-level moisture to higher levels for the latent instability energy of
the atmosphere to be released.
1.5.5 Cellular Structure – Shallow and Deep Convection
It is well-known from laboratory experiments as well as theoretical and observational studies that when the atmosphere is thermodynamically and/or hydrodynamically unstable, it breaks down into convective cells of different dimensions
depending upon the depth of the layer involved and the amount of moisture
present.
These convective cells may appear in the sky in the form of clouds of different
horizontal and vertical extents. For example, when instability is present in a shallow
layer above the earth’s surface and there is a limited supply of moisture, we may
see only small-scale fair weather cumulus type of clouds. On the other hand, when
instability involves a deeper layer of the atmosphere and there is deep moisture convergence at low levels with divergence above, clouds may appear in different layers
and some of the cloud cells may grow to great heights. The cellular structure of the
atmosphere then becomes quite clearly visible. However, when more than one layer
is involved, the cells must arrange themselves vertically so as to secure sufficient
vertical compensation to ensure a small pressure change that is actually observed at
the earth’s surface. In a developing system, however, there is continual rearrangement of the cells which may result in a large pressure change. In the tropics, the
cellular structure of atmospheric circulation is ubiquitous.
However, the actual structure of the cells in any case depends upon the configuration and dimensions of heat sources and sinks and their orientation. It must also
satisfy the requirements of energy conservation and transformations and physical
and dynamical constraints discussed earlier in this section.
1.5.6 Coriolis Control-Variation with Latitude
The relationship between pressure and wind is largely controlled by the Coriolis
force, apart from friction and gravity. In higher latitudes where Coriolis control is
strong, pressure rapidly adjusts to the wind field, but in the tropics where the Coriolis
control is weak, the wind tends to adjust to the pressure field. So, in the absence of
friction or any other force, when there is a pressure gradient across the equator where
the Coriolis control vanishes, there can be a direct cross-equatorial airflow down the
pressure gradient. However, away from the equator, the wind becomes increasingly
quasi-geostrophic.
1.5
Some Physical and Dynamical Constraints and Conservation Laws
13
1.5.7 Conservation Laws
In steady state, tropical circulation is required to satisfy certain conservation laws.
These include:
(a) The Law of conservation of heat energy, and
(b) The Law of conservation of potential vorticity.
(a) Conservation of Heat Energy: As mentioned earlier, condensation heating
is one of the powerful diabatic heat sources in the atmosphere but its effect is
almost totally offset by adiabatic processes. The balance between the two processes
is governed by the thermodynamic energy equation in the form
∂T/∂t = −(u∂T/∂x + v∂T/∂y) + σ ω + (1/cp )δQ/dt
(1.5.5)
Where T is temperature, u, v are the components of the wind along the coordinate
axes x, y respectively along the pressure surface p, ω is the vertical p-velocity, cp
is the specific heat at constant pressure, δQ/dt is the rate of diabatic heating, and σ
is the static stability parameter given by the relation, σ (= κT/p–∂T/∂p), where κ =
R/cp , and R is Gas constant. The balance is reached when the left-hand side vanishes.
In the tropics, the third term on the right hand side of (1.5.5) balances largely with
the second term, while in the midlatitudes, the first term becomes important as well.
(b) Conservation of Potential Vorticity: The principle of conservation of potential
vorticity stated in the form (1.5.6)
(ζ +f )/(δθ/δp) = Constant
(1.5.6)
(Where ζ is relative vorticity, f is Coriolis parameter, θ is potential temperature,
and p is pressure) is a powerful constraint on tropical circulation though the equation
was derived on the assumptions that the atmosphere was frictionless, adiabatic, and
barotropic.
Obviously, the assumptions are not quite realistic, since frictional forces are
always at work in the earth’s boundary layer, and adiabatic processes may be more
or less compensated by diabatic processes, and baroclinicity cannot be ruled out
from the tropical atmosphere.
Yet, the principle of conservation of potential vorticity is useful in interpreting
the distribution of vorticity in airflows negotiating large mountains. For example,
during northern winter, the cold NE-ly winds with negative relative vorticity over
the Tibetan Plateau, on descending to the Plains of Northern India, develops a narrow zone of positive relative vorticity before resuming the normal negative relative
vorticity over Central India. This change appears to be in conformity with the principle of conservation of potential vorticity. The principle appears to explain similar
changes in relative vorticity in several other mountainous regions of the globe.
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1.6 Equatorial Circulations
In the annual mean, the equatorial region of the earth receives the maximum solar
radiation and may act as a perennial heat source. However, owing to distribution of
continents and oceans and differential heating between them, heat sources and sinks
are created along the equator. Heat sources and sinks are also created over equatorial
oceans by powerful warm and cold ocean currents, especially between their western
and eastern parts.
1.6.1 Circulation with Heat Sources and Sinks Placed Alternately
Along the Equator – Walker Circulations
A general problem of this type was first addressed by Matsuno (1966) theoretically
by using a one layer homogeneous divergent barotropic model and later by Gill
(1982) and others. In the first part of his treatment, Matsuno studied the different
types of wave motions like Rossby and inertio-gravity oscillations that are excited in
the equatorial atmosphere when he placed mass sources and sinks alternately along
the equator, but in the second part he addresses the problem of forced stationary
circulation and obtained the following interesting results:
(i) In latitudes somewhat away from the equator, the surface tends to be raised
where mass is added, and lowered where mass is extracted;
(ii) In the immediate vicinity of the equator, however, the deviations of the surface
elevation is less than that in the higher latitudes in magnitude. Consequently,
high and low pressure cells tend to be divided into two parts, one on each side
of the equator;
(iii) Strong zonal flow is created along the equator when mass flows from source
to sink. The zonal flow along the equator is intensified by convergence of flow
from higher latitude circulations and weakened by divergence of equatorial
flow to higher latitudes;
(iv) In the higher latitude region, the velocity fields are in geostrophic balance.
When he applied the same boundary conditions to the two-level model of the
atmosphere in which heat sources and sinks are placed alternately along the equator, the differential heating produces low pressure and high pressure respectively.
The wind blows geostrophically in the high latitude region. The convergence or
divergence of such winds along the equator induces vertical motions. These vertical motions counteract to the imposed heat sources and sinks, and their effects are
strongest along the equator. Consequently, the warm air associated with the heat
source is split into two parts, by the cold air belt located at the equator. Similarly,
the cold air associated with the heat sink is split into two parts by the warm air belt
located at the equator.
Matsuno’s results in the case of the forced stationary circulation are depicted in
Fig. 1.2(a, b, c).
In Fig. 1.2, mass sources and sinks are shown in terms of high (H) and low (L)
pressure systems, along with the induced vertical circulations in the lower panel
1.7
Meridional Circulation with Heat Source at the Equator and Heat Sinks
15
Fig. 1.2 Forced stationary
horizontal (b) and vertical (c)
circulations caused by
imposition of mass sources
(+) and mass sinks (–)
alternately along the equator
(a). In (b), pressure replaces
mass; H for source, L for
sink, CV denotes
convergence, DV divergence
(adapted from Matsuno,
1966)
(c) which has come to be known as the Walker Circulation. The arrows show the
directions of air motion.
It is easy to see from Fig. 1.2(c) that the regions of low-level convergence and
upper-level divergence are those of dense clouding and heavy precipitation (as
shown by hatching below the bottom line in the lower panel) and where they diverge
at low levels and converge aloft are relatively cloud-free areas with little or no precipitation. According to observations, heavy clouding and precipitation occur along
the equator in some preferred regions, such as equatorial Eastern Indian Ocean,
Western Pacific Ocean, the Amazon basin of South America, and some parts of
equatorial Africa.
1.7 Meridional Circulation with Heat Source at the Equator
and Heat Sinks in Higher Latitudes – The Hadley
Circulations
The classical Hadley circulation cell visualizes a zonally-symmetric annual-mean
meridional circulation between a heat source with its cyclonic circulation centered
at the equator and a heat sink characterized by anticyclonic circulation centered at
16
1 Large-Scale Tropical Circulations – Some General Aspects
the ridge of the subtropical high pressure with rising motion at the equator and subsidence over the subtropical belt. Tradewinds diverging from the high pressure belt
were assumed to travel equatorward, converge at the equator and rise in penetrative
convection producing cloud and rain and thereby releasing latent heat of condensation of water vapour carried by the tradewinds. At the equator, the rising currents
are assumed to diverge in the upper troposphere and flow poleward as anti-trades to
give up heat and subside over the tropical belt in order to flow back again towards
the equator as tradewinds in some kind of a meridional-vertical circulation. It was
this mean meridional circulation which was assumed to transfer sensible and latent
heat from the equator poleward.
However, when applied to the real atmosphere, the idealized Hadley circulation
cell model as described in the preceding para, faces several difficulties. Observation
shows that the equatorial tropopause with a temperature of about –80◦ C at a height
of about 16 km above sea level is the coldest place in the tropical atmosphere and as
such a transfer of heat from the equator poleward in the upper troposphere against
a temperature gradient is not possible, as envisaged in the classical model. The
inadequacy of the classical single-cell model in this regard was first pointed out
by Fletcher (1945) who during a flight across the equatorial eastern Indian Ocean
during the Second World War found two well-organized zonally-oriented cloud
bands, one on each side of the equator, with little or no cloud in between over the
equator.
To overcome the difficulty of explaining the observation with the classical Hadley
circulation model, Fletcher proposed a revised model with a small-scale meridionalvertical circulation cell interposed between the equator and the Hadley cell. The
equatorial cell was supposed to have subsidence at the equator and rising motion
where the Hadley cell circulation would converge into the equatorial circulation at
some distance away from the equator. We call this convergence zone, the Tropical
Convergence Zone (TCZ). The Fletcher model was supported by Rossby (1947) and
several others. But it soon became apparent that the Fletcher model which visualized
a single equatorial vertical circulation cell may also have problems in explaining
some observed phenomena in the atmosphere. One of these concerns the formation
of different types of clouds and release of condensation heating in different layers
of the atmosphere.
In small-scale cumulus-type clouds, maximum heating is likely to be small and
confined mostly to a lower level. In large cumulonimbus-type of clouds, maximum
heat is released by condensation in the upper troposphere. In the vertical distribution
of diabatic heating in the mean tropical atmosphere, the level of maximum heating
appears at about 400 mb (Holton, 1979). This fact appears to suggest that more than
one layer of convective clouds may be involved in releasing latent heat of condensation in the atmosphere and that the upper layer may, in fact, contribute a greater
amount of diabatic heat to the warming of the equatorial troposphere than the lower
layer. In the present text, we, therefore, suggest a further revision to the Classical
Hadley cell model, by dividing the single equatorial cell into two cells, a Lower
Equatorial (LE) and an Upper Equatorial (UE), as shown in Fig. 1.3.
1.8
Seasonal Migration of the Equatorial Heat Source
17
Fig. 1.3 Schematic showing the suggested revised model of the Hadley circulation with two
equatorial cells (UE, Upper-Equatorial) and (LE, Lower Equatorial) and the location of the TCZ;
Symbols DV means Divergence, CV Convergence. Arrow shows the direction of air motion
1.8 Seasonal Migration of the Equatorial Heat Source
1.8.1 Origin of Monsoon
A northward movement of the equatorial heat source from the equator is accompanied by change in the structure of the associated Hadley circulations described
in the preceding section. The SE’ly tradewinds, which were earlier confined to the
southern hemisphere, cross the equator directly and turn into SW/W’ly tradewinds.
So the movement of the equatorial heat source brings about a latitudinal expansion
of the belt of the W’ly tradewinds to the south of its trough with influx of cool,
moist airmass from across the equator. On converging into the circulation around
the heat source on its equatorward side at low levels, the convergence produces the
well-known Intertropical Convergence Zone (ITCZ) where the converging winds
rise in penetrative convection, precipitate, and diverge in the upper troposphere. A
branch of the diverging currents returns equatorward as NE-ly antitrades to sink
over the region of the heat sink where on sinking it joins the low-level diverging
current which converges into the ITCZ in order to complete a vertical circulation
which has come to be known as the Monsoon circulation, while the other branch of
the diverging upper currents moving poleward sinks over higher latitudes to form
the upper branch of the Hadley circulation. However, within the framework of the
Hadley circulation of the northern hemisphere, there is a secondary vertical circulation associated with the TCZ where the converging low-level E/NE tradewinds rise
in subdued convection, often producing clouding and light rain. The structure and
properties of the monsoon circulation appear in clear perspective when we consider
a general case of horizontal circulation around the equatorial heat source, interacting with the tradewind circulations at low levels, as shown schematically in Fig. 1.4,
for summer in: (a) Northern Hemisphere, and (b) Southern Hemisphere.
18
1 Large-Scale Tropical Circulations – Some General Aspects
Fig. 1.4 Schematic showing
the structure of the horizontal
circulation around the
equatorial heat source
(EQ.TR.) and the locations of
the ITCZ and the TCZ where
the tradewinds (thick bold
lines with arrows) of the two
hemispheres converge: (a)
Northern Hemisphere (N.H.),
(b) Southern hemisphere
(S.H.)
1.8.2 The Wave Structure
The circulations depicted in Fig. 1.4 reveal the following general features:
1. The equatorial heat source as shown has two troughs of low pressure,
one zonally-oriented (marked EQ.TR.) and the other meridionally-oriented
(unmarked); they separate out four heat sinks, two on each side;
2. The winds diverging from each heat sink appear to converge into the circulation
around the heat source, forming two distinct convergence zones with a zone of
divergence in between on each side of the EQ.TR: these are the tradewinds; (a)
the cold NE trades on the poleward side representing the Hadley circulation, and
the cold SW trades on the equatorward side representing the Monsoon circulation
in the Northern Hemisphere (N.H.); (b) the corresponding flows in the Southern
Hemisphere (S.H.) are the cold NW Monsoon current on the equatorward side,
and the cold SE Monsoon current on the poleward side;
3. The tradewind circulation on either side of the EQ.TR appears to have the structure of a wave characterized by two convergence zones separated by a zone of
divergence; a wave associated with the TCZ on the poleward side, and a wave
associated with the ITCZ on the equatorward side of the EQ.TR;
1.8
Seasonal Migration of the Equatorial Heat Source
19
4. Over most parts of the globe, the average zonal wavelength of the monsoon wave
appears to be about 2000–2500 km;
5. In the waves, it is the convergence zones where penetrative convection and precipitation occur, with relatively clear, or less cloudy, conditions in the divergence
zone in between;
6. The Monsoon and the Hadley circulations appear to co-exist with the equatorial
heat source circulation at all times and in all monsoon climes. They may interact
with traveling waves in forming tropical disturbances.
1.8.3 Forcing for the Seasonal Movement of the Equatorial
Heat Source
It is well-known that in the summer hemisphere, given the same solar radiation,
temperature rises and pressure falls over both land and ocean, but the changes occur
much more rapidly and with greater amplitude over land than over neighboring
ocean due to much lower heat capacity of the land compared to that of the ocean.
Conversely, for the same reason, in the winter hemisphere, temperature drops and
pressure rises much faster and with greater amplitude over the land than over the
neighboring ocean. An example of this difference in pressure tendency between
land and ocean is presented in Fig. 1.5, which shows the latitudinal distribution of
mean sea level pressure along longitude 70◦ E over the Arabian Sea and adjoining
Western India during January (DJF) and July (JJA).
Fig. 1.5 Latitudinal distribution of mean sea level pressure along 70◦ E over the Arabian Sea sector
during January (DJF) and July (JJA)
20
1 Large-Scale Tropical Circulations – Some General Aspects
So, it is the gradient of pressure tendency (also called isallobaric gradient)
between the land and the ocean that appears to force the equatorial trough of low
pressure over the ocean to move towards the heat source over the neighboring land
in the summer hemisphere with a velocity given approximately by the well-known
kinematic relation (Petterssen, 1956)
C = −∇(∂p/∂t)/∇ 2 p
(1.8.1)
where C is the velocity vector, p is pressure, t is time, ∇ is Vector Del operator,
(∂p/∂t) is pressure tendency, the denominator is the curvature of the pressure field
in the trough, and the isallobaric gradient is in a direction at right angle to the axis
of the trough of the equatorial heat source.
In a study of the onset, advance and withdrawal of summer monsoon over the
Indian Subcontinent, Saha and Saha (1980) tested the validity of this hypothesis
qualitatively by applying it to the case of advance and withdrawal of summer monsoon along a narrow longitudinal belt parallel to the West Coast of India (about
73◦ E). Since pressure and height are related hydrostatically, the tendencies of mean
monthly height of the 850 mb surface at eight stations along the meridian were
calculated using the mean height data available from Ramage and Raman (1972).
Results are shown in Fig. 1.6.
Fig. 1.6 Values of 850 mb mean monthly height tendency (gpm/month) in different months at
Minicoy (MNC), Bangalore (BNG), Bombay (BOM), Ahmedabad (AHM), Jodhpur (JDP), Delhi
(DLH), Lahore (LHR), and Peshawar (PWR). Height data are taken from Ramage and Raman
(1972). Arrows show the direction of the height tendency gradient during periods of advance and
withdrawal of monsoon
1.8
Seasonal Migration of the Equatorial Heat Source
21
In regions where monsoon is interhemispheric, the forcing for cross-equatorial
flow appears to be provided by the isallobaric gradient between the heat source on
one side and the heat sink on the other. The crossing appears to open a floodgate
for cold, humid airmass of the winter hemisphere to rush into the summer hemisphere to start the process of advance of summer monsoon in that hemisphere. A
reversal in the direction of the isallobaric gradient is what causes the retreat of
the monsoon. The isallobaric gradient is clearly northward (indicated by a long
arrow in Fig. 1.6) during the period April–June, and southward during the period
September–November.
1.8.4 Intraseasonal Oscillation of Monsoon
The advance and retreat of monsoon as a result of the movement of the equatorial heat source and its associated heat sinks would produce an intraseasonal
oscillation in usual meteorological fields of the region. Figure 1.7 illustrates the
rationale behind occurrence of intraseasonal oscillation on this account in the field
of atmospheric pressure.
This intraseasonal oscillation appears to result from the superposition of the
monsoon perturbation upon the seasonal oscillation of pressure over the region.
An example of intraseasonal oscillation in mean monthly maximum and minimum
temperatures at Delhi (India) is shown in Fig. 1.8.
Fig. 1.7 Illustrating the
formation of intraseasonal
oscillation in atmospheric
pressure (Ṕ denotes
perturbation pressure) during
advance and retreat of
monsoon: Symbols used:
H – High, L – Low, C – Cold,
W – Warm. Other symbols
have their usual meanings
22
1 Large-Scale Tropical Circulations – Some General Aspects
Fig. 1.8 Distribution of
mean monthly maximum
(continuous line) and
minimum (dashed line)
temperatures at Delhi
Intraseasonal oscillation is observed in other meteorological parameters as well;
for example, in temperature, humidity, rainfall, etc. It is well-known that people
living in monsoon regions experience prolonged rainy spells twice during the year,
once during advance and a second time during retreat of monsoon. Also, as for temperature, it is not always cool and humid during the monsoon season. During certain
periods within the season, temperature rises and weather becomes unbearably hot
and humid, at least twice during the season, once during advance and a second time
during retreat. For example, at New Delhi, with the arrival of monsoon towards the
end of June, the daytime air temperature may rapidly drop from about 46 to 40◦ C
or even lower, but it does not remain at that reduced value for long, since it rises
again at least twice, the first in August when the monsoon wave moves up to the
Western Himalayas, and a second time in October during the retreat of the trough.
An intraseasonal oscillation of this kind occurs in other airmass properties as well,
such as sunshine hours, air quality, physical comfort level, etc.
Similar intraseasonal oscillations are believed to be occurring in monsoons in
other regions as well.
1.9 Co-existence of Monsoon, Hadley and Walker
Circulations – Inclined Troughs
Over most parts of the tropics, the equatorial troughs of low pressure are seldom
zonally or meridionally oriented. They are inclined to latitudes or longitudes. There
may be several reasons for the observed inclination, but the main reason appears
to be the general inclination of the coastline. In summer, when a powerful heat
low develops over the land, its sphere of influence extends over the neighboring
ocean and tends to draw an oceanic trough of low pressure which may come within
its sphere of influence. So, the oceanic trough gets inclined to the heated land. In
winter, when a cold high pressure develops over the land, the oceanic trough will
move away from the cold land. Thus, in general, the trough will be inclined towards
the nearest land during summer and away from it during winter.
1.9
Co-existence of Monsoon, Hadley and Walker Circulations – Inclined Troughs
23
Inclined troughs are found over all the global oceans, on both sides of the equator, during both summer and winter. The areas which are particularly noted for
occurrence of inclined troughs are the following:
1.
2.
3.
4.
The western and the Eastern parts of the Pacific and the Atlantic Oceans;
Eastern Arabian Sea and the Bay of Bengal;
A wide area of the southwestern Indian Ocean close to the coast of Madagascar;
A wide area of the southeastern Indian Ocean close to the Indonesian Islands and
the coast of northwestern Australia;
5. A long stretch of the southwestern Pacific Ocean, extending southeastward from
New Guinea across the Coral Sea.
Fig. 1.9 Streamlines showing the circulations around an NW–SE oriented inclined trough during
northern summer: (a) a plan view; (b) vertical circulation in a zonal section through the center of
the trough
24
1 Large-Scale Tropical Circulations – Some General Aspects
There is re-organization of the associated circulation cells when a traveling wave
interacts with a quasi-stationary inclined equatorial trough. The TCZ is then activated by additional convergence at low levels and divergence at high levels which
favor development.
Figure 1.9 shows schematically (a) a plan view of a NW–SE oriented inclined
trough, and (b) the likely zonal-vertical circulation across the trough in the N.H.
An inclined trough is more than of academic interest, for it appears to mark out
the regions noted for development of tropical cyclones. But, what is the connection
between the trough and the cyclone and how is development favored by an inclined
trough? This aspect of the question will be looked into further in the next chapter after we have introduced the easterly and the westerly waves which initiate the
development.
1.10 Definition of Monsoon
In the past, the word ‘monsoon’ has been defined exclusively in terms of either seasonal rainfall (Rao, 1976), or reversal of the direction of the prevailing surface wind
between summer and winter (Ramage, 1971). For people wholly dependent upon
rainfall for water resources, food production, etc., especially in developing countries of Southeast Asia, Africa and South America, a year with subnormal rainfall
means to them poor monsoon and that with above-normal rainfall good monsoon.
To these people, the word ‘monsoon’ is almost synonymous with rainfall.
The Arabs first noted the seasonal reversal of the surface wind direction while
sailing over the Arabian Sea more than a thousand years ago and coined the word
‘Mawsim’ for the phenomenon. This definition has also continued to this day.
Ramage (1971) defines monsoon by the following criteria:
(1) The prevailing wind direction shifts by at least 120◦ between January and
July;
(2) The average frequency of prevailing wind directions in January and July
exceeds 40%;
(3) The mean resultant winds in at least one of the months exceed 3 m s–1 ; and
(4) Less than one cyclone-anticyclone alternation occurs every two years in either
month in a 5◦ latitude-longitude rectangle.
Definitions on similar lines, seeking to identify monsoon by either rainfall, or
the seasonal reversal of the prevailing surface wind, have been advanced by several workers. None of the other characteristic features of the monsoon are ever
mentioned in these definitions.
It was Edmund Halley (1686) of England who appears to have been the first to
conceive monsoon as a seasonally-reversing tradewind circulation when he wrote
in his celebrated paper in the Philosophical Transactions of the Royal Society of
London, the following:
1.11
Global and Regional Distribution of Monsoons
25
But as the cool and dense air, by reason of its greater gravity, presses on the hot and rarefied,
it is demonstrable, that this latter must ascend in a continued stream, as fast as it rarefies,
and that being ascended, it must disperse itself to preserve the equilibrium; that is, by a
contrary current, the upper air must move from those parts where the heat is greatest: so by
a kind of circulation, the north-east tradewind below will be attended with a south-westerly
above, and the south-easterly with a north-west wind above. And that this is more than a
mere conjecture, the almost instantaneous change of the wind to the opposite point, which
is frequently found in passing the limits of the trade winds, seems to assure us; but that
which above all confirms this hypothesis, is the phenomenon of monsoons, by this means
most easily solved, and without it hardly explicable.
Our analysis of monsoon as a divergent tradewind circulation which converges
into the seasonally-migrating equatorial heat source, producing a wave on the equatorward side of its trough prompts us to offer the following holistic, comprehensive,
and yet simple, alternative definition of tropical monsoon:
“Monsoon is a large-scale perturbation of the tradewind circulation associated with the seasonal movement of the equatorial heat source, which converges into the circulation around
the heat source on its equatorward side at low levels, producing a wave with two zones of
precipitation and a zone of clearance in between, and a host of other characteristic changes
in airmass properties, during its advance and retreat.”
The foregoing definition brings within its fold not only the observed rainfall and
reversal of wind direction, but also a host of other characteristic changes in prevailing meteorological conditions, such as changes in air temperature, cloudiness, etc.
It also produces an intraseasonal oscillation in meteorological parameters during
advance and retreat because of its wave structure.
It is believed that the definition given here can be used as a dependable criterion
for identifying and tracking monsoon over any part of the globe at any time of
the year. Our definition implies co-existence of monsoon with Hadley and, at some
locations, Walker-type east-west circulations.
However, in offering the above definition of monsoon, we should not lose sight
of some other regions of the globe, especially the extratropics, where such thermal
contrast may arise between a large continent and a neighboring ocean and force a
seasonal movement of the oceanic trough of low pressure between land and ocean.
Of course, in such regions, the prevailing winds and eastward-propagating baroclinic wave disturbances may often interfere with the weak monsoon circulation
that may develop and make its identification difficult.
1.11 Global and Regional Distribution of Monsoons
1.11.1 Tropical Monsoons
Since tropical monsoon constitutes a perturbation of the tradewind circulation that
converges into the ITCZ, the seasonal movement of the ITCZ offers a practical
means to identify the leading edge of the monsoon over a region at any time of the
year (e.g., Riehl, 1954; Saha, 1978; Walisser and Gautier, 1993; Saha et al., 1998).
26
1 Large-Scale Tropical Circulations – Some General Aspects
In other words, the region swept out by the ITCZ in the course of its movement
between summer and winter constitutes the domain of the tropical monsoon in that
region. A definition of a monsoon region on similar lines based on the seasonal
movement of the ITCZ was also advanced by Asnani (1993).
In general, tropical monsoons are interhemispheric in character and usually follow the seasonal movement of heat lows over the land from one hemisphere to the
other. Three continental sectors stand out in this respect, viz., Asia, Africa, and the
Americas. In the Asia sector, the movement is between Australia and Asia; in the
African sector, between Northern and Southern Africa; and in the American sector
between South America, and Central and North America.
The monsoon regions in these sectors are as follows:
Asia–Australia sector
Region I Indian Sub-continent and adjoining SE Asia
Region II Eastern Asia including China, Korea and Japan
Region III Maritime Continent including Indonesia and Philippines
Region IV Australia
Africa sector
Region V North and South Africa
American sector
Region VI South America
Region VII Central America and adjoining States of Southwestern North
America
In general, the domains of oceanic monsoons are much narrower, and confined
to latitudes closer to the equator. However, they are much wider over the western
and the eastern parts of the Atlantic and the Pacific Oceans where they tend to join
up with the continental monsoons along inclined troughs.
1.11.2 Extratropical Monsoons
Monsoonal-type atmospheric circulations between continents and oceans are also
observed in some extratropical regions of the globe where seasonal contrasts in
temperatures between land and ocean and differential heating between them bring
about seasonal movements of the circulation systems and associated rainbelts across
the coastline and to higher latitudes. They are also observed over some extratropical
mountain plateau regions where the circulations and the associated rainbelts are
affected by seasonal reversals of temperature and pressure between the Plateau and
the neighboring Lowlands or oceans. However, it is often difficult to see the full
impact of such monsoons in the face of interference from strong planetary wind
systems and frequent movement of baroclinic wave disturbances across the region.
The concept of a monsoon in extratropical latitudes is not new, though it has been
opposed by some meteorologists (e.g., Ramage, 1971). It was advocated by several
1.11
Global and Regional Distribution of Monsoons
27
workers (e.g., Alisov, 1954; Khromov, 1957 and others), though the steadiness of
the prevailing surface wind, which was cited to be one of the main characteristics of
a monsoon circulation, was found to be somewhat low on account of the interference
from disturbances (Klein, 1957).
It is proposed to show in Part 3 of this text that the idea of an extratropical
monsoon is not just a fantasy, but a stark reality. It is driven by differential heating between a heat source and a heat sink in much the same way as in a tropical
monsoon, but in a different environment. A tropical monsoon results from the seasonal movement of an oceanic heat source called the equatorial heat source from
one hemisphere to the other and its movement is limited to the tropical belt. For
this monsoon, the subtropical high pressure cells with their anticyclonic circulations
and diverging tradewinds serve as heat sinks. On the other hand, for an extratropical
monsoon, the same subtropical high pressure cells would act as heat sinks but heat
sources would lie over the extratropical belt. The exact mechanism of this process
will be elaborated in Chaps. 5 and 11 with a few case studies.
Figure 1.10 shows the approximate domains of tropical and extratropical monsoons over different parts of the globe including the oceans.
Fig. 1.10 Global distribution of principal monsoon regions (stippled areas) bounded by the
January and July locations of the ITCZ: January (full line); July (dashed line). Regions marked
I to VII are areas of tropical continental monsoons; VIII and IX are domains of Extratropical
monsoon; SPCZ denotes Southwest Pacific Convergence Zone, and SACZ the Southwest Atlantic
Convergence Zone
1.11.3 Zonal and Meridional Anomalies
As shown in Fig. 1.10, large zonal and meridional anomalies exist in the distribution
of monsoons over the globe. These anomalies arise from the following:
28
1 Large-Scale Tropical Circulations – Some General Aspects
(a) Geographical distribution of continents and oceans;
(b) Influence of ocean currents – oceanic monsoons; and
(c) Effects of large-scale orographic barriers.
(a) Geographical Distribution of Continents and Oceans: Whatever may be the
geological origin of continents, their locations in a certain alignment on the earth’s
surface, appears to be very important for the seasonal movement of monsoons. This
happens in three major continental sectors, viz., Asia-Australia, North and South
Africa, and North and South America. A quick glance at the world map shows that in
each sector, the landmasses are oriented more or less in a general NW–SE direction,
and monsoons in general follow this direction in their interhemispheric movement
following the seasonal movement of the sun.
(b) Influence of Ocean Currents – Oceanic Monsoons: Over all the world oceans,
the subtropical anticyclonic gyres and the coastal ocean currents exercise large control on the distribution of ocean surface temperature. Their seasonal shift gives rise
to oceanic monsoons and in this process ocean currents play a very major role.
For example, in the Pacific Ocean, the seasonal location of the ITCZ is affected
by the following major ocean currents: (1) The poleward-flowing warm Kuroshio
Current in the northwest; (2) the equatorward-flowing cold California Current in the
northeast; and (3) the equatorward-flowing cold Peruvian or Humbolt Current in the
southeast. On account of the cross-equatorial flow of the Peruvian Current, the equatorial eastern Pacific Ocean remains cold almost throughout the year. Similarly, in
the Atlantic Ocean, the corresponding ocean currents are: (1) the poleward-flowing
warm Gulfstream in the northwest; (2) the equatorward-flowing cold Canaries current in the northeast; and (3) the equatorward-flowing cold Benguela current in the
southeast. The equatorward-flowing cold Benguela Current keeps the temperature of
the equatorial eastern Atlantic Ocean low almost throughout the year. On account of
these cross-equatorial flows of cold ocean currents, the ITCZ seldom appears south
of the equator over these oceans.
The situation over the Indian Ocean is somewhat different. Here, in the western
Indian Ocean, because of the dominant influence of the cold Somali current during
northern summer, a zonal anomaly of surface temperature is maintained between the
western and the eastern parts of the equatorial ocean almost throughout the year. In
the Southern Indian Ocean, the equator-flowing cold West Australian current keeps
the surface temperature of the eastern part relatively cold, while in the western part
the poleward-flowing warm Agulhas current maintains a warm ocean surface. These
movements of the ocean currents are reflected in the equatorial distribution of ocean
surface temperature shown in Fig. 1.11.
In some of the oceans where equatorial belt remains cold, the ITCZ is displaced
from the equator to the relatively warmer parts of the ocean or to the nearest warm
continent. For example, the SW Pacific Convergence Zone which runs from equatorial New Guinea eastsoutheastward across the Coral Sea area may be looked upon as
the displaced ITCZ of the southern Pacific Ocean. Similarly, in the western part of
the equatorial Indian Ocean which remains relatively cold throughout the year, the
ITCZ is displaced to about 15◦ S during northern winter and to about 25◦ N during
1.12
Co-existence of Monsoon with Desert Circulation
29
Fig. 1.11 Equatorial distribution of mean ocean surface temperature during February (continuous
line), and August (dashed line) (after Defant, 1961)
northern summer. Similar displacements may also be observed over other parts of
the tropical oceans and continents.
(c) Effects of Mountain Barriers: Large mountain systems, such as the Himalayas
of Asia, the Mountains of East Africa, the Rockies of North America, the Andes of
South America, and the Great Dividing Range of Australia, all appear to affect the
circulations in their respective areas by relocating the ITCZ and other troughs and
ridges of pressure through their mechanical and thermal effects on circulation. In
accordance with the principle of conservation of potential vorticity, the blocking of
a W’ly flow without horizontal shear approaching a north-south oriented mountain
range will produce a trough of low pressure in the run-up to the base of the mountain,
then a high pressure as it climbs up to the top, and then a low pressure as it descends
on the leeside, till it resumes its zonal flow. In the case of an E’ly flow approaching
such a mountain range, there is a difference. The flow senses the presence of the
mountain barrier from a long distance and accordingly adjusts the pressure in such
a way that will enable it to get over the top and emerge as an E’ly current on the
leeside. This it does by first producing a low pressure trough and then a high pressure
ridge of such intensity as will enable it to cross the mountain and emerge on the other
side as an E’ly current.
The ITCZ or the TCZ, as the case may be, is relocated whenever the airflow
around the equatorial trough has to negotiate these high mountain barriers.
1.12 Co-existence of Monsoon with Desert Circulation
It is by no means an accident that the world’s principal monsoons co-exist with
subtropical deserts on their poleward sides. In the northern hemisphere, the great
central Asian desert, the great Saharan desert in Africa, the Mojave-Sonoan desert
30
1 Large-Scale Tropical Circulations – Some General Aspects
in Central America and adjoining southwestern North America, all lie on the poleward sides of monsoons in the respective regions. Similar co-existence of monsoons
with deserts on their poleward sides is also observed in the southern hemisphere in
all the continents; for example, the Great Gibson and Western Deserts in Australia,
the Namib-Kalahari desert in Southern Africa and the Patagonia desert in South
America all lie on the poleward sides of their respective monsoons. So, a coexistence of monsoons and desert circulations appears to be a global phenomenon
which has drawn the attention of meteorologists for several decades.
Ramage (1966) observed a seesaw type inverse relationship between surface
pressure anomaly in the ‘heat low’ over the Thar Desert in Pakistan and monsoon
rainfall anomaly over the neighboring Arabian Sea off Bombay. He observed that
a fall (rise) of surface pressure in Thar Desert over Pakistan was correlated with
increase (decrease) of rainfall over the neighboring Arabian Sea.
In a theoretical study of the dynamics of deserts and recurrent drought in the
Sahel at the southern periphery of the Saharan desert, Charney (1975) observed that
to maintain the radiative equilibrium of the atmosphere over a surface with high
albedo, such as a sandy soil surface, against radiative heat loss, air must descend
and that it is this descent or adiabatic subsidence and warming of the air that leads
to continued dryness and maintenance of the desert. He also recognized the contribution of the descending branch of the Hadley circulation to the desertification
process but expressed the view that the impact of the radiative heat loss was greater
than subsidence warming.
Fig. 1.12 Mean vertical circulations (resultant streamlines) along the Greenwich meridian at 12
GMT during August, showing the monsoon and desert circulations STF marks the location of the
subtropical front between the desert and the Mediterranean Sea circulations (after Saha and Saha,
2001b)
1.12
Co-existence of Monsoon with Desert Circulation
31
Blake et al. (1983) who carried out a detailed observational study of heat balance
of the atmosphere over a part of the Saudi Arabian desert (Rub-al-Khali) in May
1979 found that there was strong subsidence in the middle and upper troposphere
which often descended to very low levels at night. Strong convection over the desert
surface during daytime was limited to the lower troposphere only and there was
strong outflux of sensible heat in the lower and middle troposphere from the desert
to the monsoon over the adjacent Arabian Sea.
Saha and Saha (2001b), using NCEP reanalysis data, computed the mean vertical
circulation along the Greenwich meridian over the western part of the Saharan desert
at 12 GMT during August and showed how the monsoon and the Hadley circulations
co-exist with the desert circulations and are linked to one another. Their results are
presented in Fig. 1.12.
Figure 1.12 shows that in the meridional circulation, while the monsoon winds
converge into the ITCZ along about 12◦ N at low levels producing strong penetrative
convection which diverges in the upper troposphere both equatorward as well as
poleward, there is strong sinking motion in the Hadley circulation over the desert.
In fact, the sinking motion over the desert area is so strong that it prevents the
thermally-direct low-level desert convection from rising beyond 700 mb. Along the
northern boundary of the desert heat low which extends beyond latitude 30◦ N, cool,
moist winds diverging from a high pressure area over the neighboring Mediterranean
Sea converge into the heat low circulation, forming a subtropical front (STF), the
presence of which was first reported by Soliman (1958). In the present text, we have
designated the STF as the TCZ