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importance is the large horizontal temperature gradient at the surface in midlatitudes.

The natural decomposition of the general circulation into those isentropic surfaces

that are everywhere above the tropopause (the ‘‘overworld’’), those that intersect the

tropopause (‘‘middleworld’’), and those that intersect the surface (‘‘underworld’’)

provides an alternative framework in which to view global-scale dynamical

processes in the extratropical atmosphere. The qualitative link between IPV and

the mass ﬁeld in (38) makes IPV an ideal tool with which to study planetary-scale

dynamical motions; examples of such motions include the str ucture of the mean

meridional mass transport in the troposphere, the interaction between thermal and

mechanical sources of IPV at Earth’s surface and their role in the overall general

circulation, and the exchange of air between the stratosphere and troposphere.

Coupled with the invertibility principle, this link makes IPV a powerful diagnostic

tool for studying the extratropical atmospheric circulation.

REFERENCES

Andrews, D. G., J. R. Holton, and C. B. Leovvy (1987). Middle Atmosphere Dynamics,

International Geophysics Series, Vol. 40, Academic, New York.

Gill, A. E. (1982). Atmosphere-Ocean Dynamics, International Geophysics Series, Vol. 30,

Academic, New York.

Held, I. M. (1982). Stationary and quasi-stationary eddies in the extratropical tropopshere:

Theory, In Large-Scale Dynamical Processes in the Atmosphere, B. J. Hoskins and

R. P. Pearce (Eds.), Academic, New York, pp. 127–168.

Held, I. M., and B. J. Hoskins (1985). Large-scale eddies and the general circulation of the

troposphere, Adv. Geophys. 28A, 3–31.

Holton, J. R. (1992). An Introduction to Dynamical Meteorology, 3rd ed., International

Geophysics Series, Vol. 48, Academic, New York.

Holton, J. R., P. H. Haynes, M. E. McIntyre, A. R. Douglass, R. B. Rood, and L. Pﬁster (1995).

Stratosphere–troposphere exchange, Rev. Geophys. 33, 403–439.

Hoskins, B. J. (1982). The mathematical theory of frontogenesis, Ann. Rev. Fluid Mech. 14,

131–151.

Hoskins, B. J., M. E. McIntyre, and A. W. Robertson (1985). On the use and signiﬁcance of

isentropic potential vorticity maps, Quart. J. Roy. Met. Soc. 111, 877–946.

Pedlosky, J. P. (1987). Geophysical Fluid Dynamics, 2nd ed. Springer, New York.

Pierrehumbert, R. T., and K. L. Swanson (1995). Baroclinic instability, Ann. Rev. Fluid Mech.

27, 419–167.

EXTRATROPICAL ATMOSPHERIC CIRCULATIONS

CHAPTER 5

DYNAMICS OF THE TROPICAL

ATMOSPHERE

GERALD MEEHL

1 THERMALLY DIRECT CIRCULATIONS AND MONSOONS

The heating of Earth’s surface that follows the annual cycle of solar forcing produces

fundamentally ageostrophic thermally direct circulations in the tropical atmosphere.

Hadley postulated that two thermally driven north–south cells, roughly straddling the

equator, must be necessary to transport heat and energy surpluses from the tropics

toward the poles. These meridional cells now bear his name. Subsequently, Sir

Gilbert Walker recognized that, in addition to the north–south Hadley cells, impor-

tant components of large-scale tropical circulations are associated with tremendous

east–west exchanges of mass. The term Walker circulation was ﬁrst used by Bjerknes

(1969) to describe these exchanges of mass in the equato rial Paciﬁc. Subsequently,

Krishnamurti (1971) showed that the Walker circulation in the equatorial Paciﬁc is

just a component of large-scale east–west circulations emanating from centers of

organized tropical convective activity associated with regional monsoon regimes.

For example, applications of the Hadley and Walker circulation concepts have

linked the monsoon regimes of Asia and northern Australia to weather and climate in

other parts of the tropics and subtropics. Figure 1 shows one such interpretation of

the east–west Walker circulation connecting the Asia–Australia region to the eastern

Paciﬁc (Webster et al., 1998). Another branch of east–west circulation, termed the

transverse monsoon here, connects those monsoon regions to eastern Africa. The

north–south Hadley components in Figure 1 are termed lateral monsoon and involve

large-scale mass overturning from the respective summer monsoon regimes to

Handbook of Weather, Climate, and Water: Dynamics, Climate, Physical Meteorology, Weather Systems,

and Measurements, Edited by Thomas D. Potter and Bradley R. Colman.

ISBN 0-471-21490-6 # 2003 John Wiley & Sons, Inc.

the winter hemisphere subtropics, thus depicting the classic Hadley cell type of

circulation.

The latitude where the sun is most directly overhead in its seasonal march is

where the surface is most strongly heated in the tropics. The air directly above that

warm surface rises and draws nearby air in to take its place. The warm air rises, cools

and condenses, and forms convective clouds. The low-level convergence at that

latitude forms a zonally oriented line of convection called the inter tropical conver-

gence zone (ITCZ). The arrangement of land and ocean in the tropics dictates that

the ITCZ does not simply follow the sun’s motion from northern to southern hemi-

sphere with the seasonal cycle. Off-equatorial land masses in the tropics tend to heat

more than the adjacent oceans, thus giving rise to seasonal regional convective

systems called monsoons. Though the word monsoon has a number of interpreta-

Figure 1 Schematic of the (a) south Asian and (b) Australian monsoon divergent wind

circulations. Three major components identiﬁed are the transverse monsoon (Walker-type

east–west circulation west of the monsoons), lateral monsoon (Hadley-type north–south

circulations), and Walker circulation (east–west circulation east of the monsoons) (Webster

et al., 1998).

DYNAMICS OF THE TROPICAL ATMOSPHERE

tions, monsoon regimes exist wherever there is a regional organization of convection

associated with an off-equatorial land mass and an equatorward ocean moisture

source. This particular distribution of land and ocean disrupts the normal zonal

orientation of the ITCZ and contributes to the east–west Walker-type circulations

(e.g., Fig. 1). These regional monsoon regimes are recognized to exist in tropical

regions of south Asia, northern Australia, southwestern North America, west Africa,

and, in some deﬁnitions, South America.

All these monsoon regimes share certain characteristics. With the approach of

local summer, intensifying solar radiation heats the land surface forming a heat low.

Since the land surface has heated faster than the adjacent ocean, the heat low begins

to draw in moist low-level air from the ocean. This moist warm air rises and forms

convective clouds and rainfall. The latent heating from the convective systems

powers a positive feedback, with even stronger low-level moist inﬂow from the

ocean intensifying convection and precipitation over land. An upper level high

forms overhead with outﬂow in all directions from the regional convective center

associated with the monsoon circulation. This feeds into the ther mally direct mass

circulations mentioned above and connects the monsoon regimes with other areas in

the tropics and subtropics. The massive amounts of latent heating associated with the

monsoon regimes provide a major energy source for the general circulation of the

global atmosphere.

2 EL NIN

O–SOUTHERN OSC ILLATION

It was the efforts of Walker and others to try to understand the facto rs that inﬂuence

the Indian monsoon in particular that led to the discovery of the east–west

circulations evidenced by large-scale exchanges of mass between the Indian and

Paciﬁc Oceans described above. This phenomenon became known as the Southern

Oscillation. These investigators associated the ﬂuctuations of mass between the

Indian and Paciﬁc regions with temperature and precipitation variations over much

of the globe. Bjerknes (1969) later linked sea surface temperature (SST) anomalies

in the equatorial Paciﬁc (which came to be known as El Nin

o and La Nin

a), and

changes in the east–west atmospheric circulation near the equator, to the Southern

Oscillation. The entire coupled ocean–atmosphere phenomenon is called El Nin˜o–

Southern Oscillation (ENSO).

ENSO and its connections to the large-scale east–west circulation can best be

illustrated by correlations of sea level pressure between one node, representing the

Indian–western Paciﬁc region (in this case Darwin on the north coast of Australia),

with all other points (Fig. 2). Positive values cover most of the south Asian and

Australian regions, with negative values east of the dateline in the Paciﬁc. Also note

positive values over northeastern South America and the Atlantic. Thus, when

convection and precipitation in the south Asian or Australian monsoons tends to

be below normal, there is high pressure in those regions with a corresponding

decrease of pressure and enhanced rainfall over the tropical Paciﬁc, with higher

pressure and lower rainfall over northeastern South America and tropical Atlantic.

2ELNIN

O–SOUTHERN OSCILLATION 59

Since the atmosphere and ocean are dynamically coupled in the tropics, ENSO

and its associated ﬂuctuations of atmospheric mass is associated with SST anomalies

as a consequence of the mostly ageostrophic surface winds involved with sea level

pressure anomalies such as those implied in Figure 2. This dynamic coupling is

reﬂected with the association between El Nin

o and La Nin

a events and the Southern

Oscillation noted above. For example, anomalously low pressure and enhanced

precipitation over the tropical Paciﬁc are usually associated with anomalously

warm SSTs there. In the extreme, this is known as an El Nin

o event. In the opposite

extreme, anomalously cold SSTs east of the dateline have come to be known as a

La Nin

a event. These are often associated with low pressure and strong Asian–

Australian mons oons west of the dateline, with high pressure and suppressed rainfall

east of the dateline. The conditions that produce El Nin

o and La Nin

a events, with a

frequency of roughly 3 to 7 years, involve dynamically coupled interactions of ocean

and atmosphere over a large area of the tropical Indian and Paciﬁc Oceans (Rasmus-

son and Carpenter, 1982). Figure 3 shows low-level wind, precipitation, and SST

anomalies for the northern summer season during the 1997–1998 El Nin

o event.

Westerly surface wind anomalies occur to the west of the largest positive SST

anomalies in the central and eastern equatorial Paciﬁc. These wind anomalies

reduce upwelling of cool water and contribute to the positive SST anomalies. The

warmer water is associated with greater evaporation and increased precipitation. As a

result of the anomalous Walker circulation, there is suppressed precipitation over

much of Southeast Asia.

El Nin

o and La Nin

a events are extremes in the system that tends to oscillate

every other year to produce the tropospheric biennial oscillation, or TBO (Meehl,

1997). Thus there are other years, in addition to the Paciﬁc SST extremes, that have

similar coupled processes but lower amplitude anomalies also involving the Asian–

Australian monsoons and tropical Indian and Paciﬁc SST anomalies (Fig. 4).

The TBO encompasses all the ENSO processes of dynamically coupled air–sea

interaction in the tropical Indian and Paciﬁc Oceans, large-scale east–west cir-

Figure 2 Distribution of the zero-lag correlation of surface pressure variations with the

Darwin surface pressure and all other points for monthly data for the period 1890–1993

(Webster et al., 1998).

DYNAMICS OF THE TROPICAL ATMOSPHERE

culations in the atmosphere, the Asian–Australian monsoons, and coupled SST

anomalies in Indian and Paciﬁc involving internal ocean wave dynamics.

Figure 4 illustrates these coupled interactions for a sequence from a weak

Australian monsoon during anoma lously warm conditions in the tropical Paciﬁc

that, in the extreme, is an El Nin

o event, through to the onset of anomalously

cold SSTs in the Paciﬁc (in the extreme, a La Nin

a event) associated with a

strong Indian monsoon, and on to a strong Australian monsoon the following south-

ern summer. In the December–January–February (DJF) season prior to a strong

Indian monsoon, there is a relatively weak Australian monsoon (Fig. 4a), with

warm SST anoma lies to the west in the Indian Ocean, to the east in the central

and eastern Paciﬁc, and relatively cool SSTs north of Australia. Relatively strong

convection over the western Indian Ocean and eastern African areas, along with

strong convection over the central equatorial Paciﬁc, are associated with an atmos-

pheric Rossby wave response over Asia such that there is an anomalous ridge of

positive 500 hPa heights and warm land temperatures. Anomalous easterly winds in

the western Paciﬁc induce upwelling Kelvin waves in the ocean that propagate

eastward. These begin to raise the thermocline in the central and eastern Paciﬁc

over the course of the next few months, as indicated in Figure 4b. Westerly anom-

Figure 3 (a) Precipitation and surface wind anomalies for the 1997 June–July–August–

September seasonal average during the 1997–1998 El Nin

o event; dark gray shading indicates

precipitation anomalies greater than 0.5 mm=day, light gray less than 0.5 mm=day, scaling

arrow below part (a) indicates surface wind anomalies of 2 m=s; (b)asin(a) except for SST

anomalies with dark gray greater than 0.5



C, light gray less than 0.5



2ELNIN

O–SOUTHERN OSCILLATION 61

62 DYNAMICS OF THE TROPICAL ATMOSPHERE

alous surface winds in the far wester n Indian Ocean set off downwelling equatorial

Kelvin waves in the ocean and contribute to a deepening of the thermocline in the

eastern equatorial Indian Ocean in March–April–May (MAM, Fig. 4b).

In the MAM season prior to a strong Indian monsoon (Fig. 4b), the upwelling

Kelvin waves from the easterly anomalous winds in the western Paciﬁc during DJF

continuing into MAM act to raise the thermocline in the central and eastern Paciﬁc.

During June–July–August–September (JJAS) (Fig. 4c), convection and precipita-

tion over the Indian monsoon region is strong. Anomalous winds over the equatorial

Indian ocean remain westerly, and the shallow thermocline in the west is evidenced

by anomalously cool SSTs appearing in the western equatorial Indian Ocean. There

are strong trades, and a shallow thermocline and cool SSTs in the central and easter n

equatorial Paciﬁc, thus completing the SST transition there.

The September–October–November (SON) season after the strong Indian

monsoon (Fig. 4d ) has anomalously strong convection traversing with the seasonal

cycle to the southeast, encountering warm SSTs set up by the dynamical ocean

response to the westerly winds in the equatorial Indian Ocean in JJAS with a

dipole of SST anomalies in the Indian Ocean.

Finally, in Figure 4e in the DJF following a strong Indian monsoon, there is a

strong Australian monsoon, and strong easterlies in the equatorial Paciﬁc as part of

the strengthened Walker circulation associated with cool SSTs there. Anomalous

westerlies in the far western equatorial Paciﬁc associated with the strong Australian

monsoon convection start to set off downwelling equatorial oceanic Kelvin waves,

which begin to deepen the thermocline to the east and set up the next transition to

warm SSTs in the central and eastern equatorial Paciﬁc the following MAM–JJAS.

3 SUBSEASONAL TROPICAL ATMOSPHERIC WAVES

The westerly winds in the far western equato rial Paciﬁc involved with transitions of

SST anomalies in the Paciﬁc have been associated with a subseasonal phenomenon

called the Madden–Julian Oscillation or MJO (Madden and Julian, 1972). The MJO

is a domi nant mode of subseasonal tropical convection. It is associated with an

eastward-moving envelope of convection that progresses around the global tropics

with a period of about 30 to 70 days and is most evident in clouds and convection in

Figure 4 Schematic representation of coupled processes for (a) the DJF season after a weak

Indian monsoon with a weak Australian monsoon and anomalously warm SSTs in the tropical

Paciﬁc, (b) MAM, (c) a strong Indian monsoon associated with anomalously cold SSTs in the

Paciﬁc, (d ) SON and (e) DJF with a strong Australian monsoon. Note that conditions in DJF

preceding the strong monsoon in (a) are opposite to those in the DJF following the strong

monsoon in (e) indicative of the TBO in the coupled system. The TBO encompasses El Nin

and La Nin

a events along with other years with similar but lower amplitude anomalies. Thus

the coupled processes depicted here apply to TBO and ENSO (Meehl and Arblaster, 2002).

3 SUBSEASONAL TROPICAL ATMOSPHERIC WAVES 63

the Indian and Paciﬁc sectors. Particularly active MJO events, as they move across

the western Paciﬁc, are associated with bursts of westerly anomalous winds that

force eastward propagating Kelvin waves in the ocean. These oceanic Kelvin waves

affect the depth of the thermocline to the east and can contribute to a transition of

SST anomalies in the eastern Paciﬁc associated with the onset of some El Nin

events (McPha den, 1999). For example, Kessler et al. (1995) show how successive

MJO westerly wind events in the western Paciﬁc can lead to the forcing of oceanic

Kelvin waves and successive or stepwise advection of warm SSTs from the western

Paciﬁc toward the east to contribute to El Nin

o onset.

The MJO is just one of a number of convectively coupled tropi cal waves

(i.e., atmospheric wave dynamics coupled to convection such that their signatures

are seen in analysis of tropical clouds) that exist in the equatorial atmosphere.

Subseasonal tropical convection is organized on larger scales by certain modes

that are unique to the equatorial region. An entire class of equatorially trapped

modes organizes atmospheric motions and convection at low latitudes. The theory

for these ‘‘shallow water’’ modes was developed by Matsuno (1966).

Formally, equatorial wave theory begins with a separation of the primitive equa-

tions, linearized about a basic state with no vertical shear, governing small motions

in a three-dimensional stratiﬁed atmosphere on an equatorial Beta plane, into the

‘‘vertical structure’’ equation and ‘‘shallow-water’’ equations. Four parameters char-

acterize the equatorial wave modes that are the zonally (and vertically) propagating

equatorially trapped solutions of the shallow-water equations. These include meri-

dional mode number n, frequency v, ‘‘equivalent depth’’ h of the ‘‘shallow’’ layer of

ﬂuid, and zonal wavenumber s. The internal gravity wave speed c is related to the

equivalent depth as

c ¼

ﬃﬃﬃﬃﬃ

ð1Þ

and links the vertical structure equation and shallow-water equati ons as a separation

constant. The equatorial Rossby radius R, is related to the vertical wavelength of free

waves and to the meridional scaling by

ﬃﬃﬃﬃﬃ



1=2

ð2Þ

where b is the latitudinal gradient of the Coriolis parameter.

The theoretical dispersion relationship will fully characterize the wave, given the

meridional mode number and wave type, provided two out of h, v, and s are speci-

ﬁed. It is assumed that tropical waves associated with convective heating are internal

modes with wavelike vertical structures, and the resulting solution of the shallow-

water equations are either antisymmetric or symmetric about the equator. Convection

is usually related to the divergence or temperature ﬁeld, and modes of even meri-

dional mode number n are antisymmetric, and odd n are symmetric. Figure 5 shows

examples of circulation and divergence associated with some of these symmetric

64 DYNAMICS OF THE TROPICAL ATMOSPHERE

modes, the n ¼0 mixed Rossby gravity, n ¼1 equatorial Rossby, n ¼1 K, and

n ¼1 westward inertiogravity waves.

A space–time analysis of satellite-observed outgoing longwave radiation (OLR),

a proxy for tropical convection, is shown in Figure 6. By separating the OLR into its

antisymmetric and symmetric components about the equator, as well as removing

a red background spect rum, the spectral peaks of these convectively coupled equa-

torial waves are shown in Figure 6a for the antisymmetric components and Figure 6b

for the symmetric (Wheeler and Kiladis, 1999). The modes occur for a wide range of

space and time scales, from synoptic to planetary, and submonthly to intraseasonal.

Waves depicted in Figure 5 appear in Figure 6 as Kelvin waves, equatorial Rossby

Figure 5 Depiction of various atmospheric tropical waves from shallow water theory:

(a) n ¼0 mixed Rossby gravity, (b) n ¼1 equatorial Rossby, (c) n ¼1 Kelvin wave, and

(d ) n ¼1 westward inertiogravity wave. Hatching indicates convergence, shading is diver-

gence.

3 SUBSEASONAL TROPICAL ATMOSPHERIC WAVES 65