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of high cloud with a sharp western edge. This cloud band represents the so-called

warm conveyor belt, within which air from relatively warm latitudes is carried

northward and upward through the storm roughly parallel to the cold front before

curving anticyclonically beyond the warm front (Fig. 11). Precipitation beneath this

conveyor belt may be convective, stratiform, or nonexistent. If the cyclone is sufﬁ-

ciently strong, the conveyor belt cloud band may curve cyclonically around the

approximate position of the low center (as in Fig. 7b), or a separate cloud mass

known as the cold conveyor belt may be present (as in Fig. 8b over the Dakotas and

Fig. 10b over the Great Lakes and upper Midwest). This conveyor belt consists of air

originating north of the warm front and ascending ahead of and on the poleward side

of the cyclone before spreading northward or wrapping around the cyclone center. A

third upper-tropospheric airstream, consi sting of air descending behind the cyclone

and curving cyclonically to the nor theast, is known as the dry airstream and is

indicated in Figure 11.

These cloud structures may be used to diagnose the location and intensity of the

extratropical cyclone center, but they are most directly related to the horizontal and

vertical wind ﬁelds in the upper troposphere and the distribution of precipitation.

Because surface pressure features tend to be strongly modiﬁed by orography, the

structure of the surface cyclone can be inferred reliably from cloud structures only

over water. Notice the superﬁcial similarity of Figures 7 b ,8b, and 10b despite the

wildly different low-level wind and pressure ﬁelds.

Statistical Climatology

The climatological distribution of extratropical cyclones reﬂects the inﬂuence of

orography (Fig. 12). Cyclones are most common over the oceans between latitudes

of 30



N and 60



N. In contrast, extratropical cyclones are rare over high orography.

Figure 11 Infrared satellite images showing the location of the warm conveyor belt (dotted

streamlines) and the dry airstream (solid streamline). ‘‘LSW’’ marks the leftmost streamline in

the warm conveyor belt, which is known as the ‘‘limiting streamline’’ and typically is

manifested as a sharp cloud boundary in satellite imagery. (From Carlson, 1991).

526

LARGE-SCALE ATMOSPHERIC SYSTEMS

Given the general tendency of cyclones to move toward the east or northeast, one can

infer from Figure 12 the general distribution of cyclogenesis (formation and inten-

siﬁcation of extratropical cyclones) and cyclolysis (weakening and dissipation of

extratropical cyclones). Cyclones tend to form downstream of major mountain

barriers such as the Rocky Mountains and the Alps, as well as over the midlatitude

oceans. Semipermanent large-scale cyclones may be found in the extreme North

Atlantic (the Icelandic low) and North Paciﬁc (the Aleutian low). These lows are

periodically reinvigorated as strong extratropical cyclones migrate northwestward

across the jet and merge with them. At higher latitudes, small-scale intense extra-

tropical cyclones known as polar lows sometimes form near the sea ice boundary.

These polar lows are partly driven by deep convection as air masses are destabilized

by the underlying warmer ocean surface, and share some of the characteristics of

hurricanes.

The midlatitude extratropical cyclones that become particularly intense typically

do so by deepening rapidly. Such explosively deepening cyclones, known a s bombs,

Figure 12 Distribution of total number of cyclones passing through 5 5 latitude–longitude

boxes during 20 Januarys, 1958–1977. (From Whittaker and Horn, 1984).

6 EXTRATROPICAL CYCLONES 527

develop almost exclusively over the oceans, preferentially near the eastern edge of

continents (Fig. 13). Conditions favoring explosive deepening include a strong

upper-level disturbance, a strong low-level horizontal temperature gradient, and a

source of warm, moist air. The Kuroshio current (in the Paciﬁc) and the Gulf Stream

(in the Atlantic) help provide the latter two conditions by inﬂuencing the distribution

of heat and moisture in the overlying atmosphere.

Vertical Structure and Dynamics of Extratropical Cyclones

For development, extratropical cyclones require vertical shear. This means that

cyclones will typically develop beneath an upper-tropospheric jet stream. Since

balanced vertical shear implies a horizontal temperature gradient, cyclones therefore

develop withi n a large-scale temperature gradient or along a frontal zone. Discussing

the vertical structure of extratropical cyclones requires the introduction of the terms

upshear, meaning in the direction opposite the vertical shear vector (typically, the

direction opposite the upper-tropospheric jet stream), and downshear, meaning in the

same direction as the vertical shear vector (typically, downwind relative to the upper

troposphere). These concepts are illustrated in Figure 14.

Figure 13 Distribution of the locations at which rapidly deepening extratropical cyclones

undergo their most rapid intensiﬁcation, expressed as total number within each

5 5 latitude–longitude box (normalized by latitude) for the period 1976–1982. (From

Roebber, 1984).

528

LARGE-SCALE ATMOSPHERIC SYSTEMS

An intensifying extratropical cyclone, as deﬁned by the height and pressure

distribution, typically tilts upshear with height. This means that the trough in the

upper troposphere will be located upshear of the surface cyclone. The temperat ure

distribution is tilted in the opposite sense, so that the warmest temperatures at the

surface are just downshear of the surface cyclone position and the warmest tempera-

tures in the upper troposphere are well downshear of the surface cyclone position.

The largest variations in pressure and temperature are located in the upper tropo-

sphere and at the surface but are substantial at all levels in between. This pressure

and temperature distribution is entirely consistent with thermal wind balance, so this

particular pressure distribution implies this particular temperature distribution and

vice versa. Winds, being in approximate geostrophic balance, vary in tandem with

pressure.

The vertical motion is an integral part of the extratropical cyclone structure

because without vertical motion cyclones would not intensify. Cyclone intensiﬁca-

tion can be thought of a s an increase in the circulation about the cyclone center and,

for large-scale frictionless ﬂow, the fractional rate of increase of total circulation

(including the circulation associated with Earth’s rotation, which is essentially

constant) is proportional to the horizontal convergence of the wind. Mass conserva-

tion implies upward motion in the middle troposphere wherever there is low-level

convergence (excluding exotic effects associated with a sloping lower boundary), so

vertical motion is not just a consequence of cyclogenesis; it is a necessary compo -

nent of cyclogenesis. Since the stratosphere inhibits vertical motion, upper-

tropospheric convergence (and intensiﬁcation of the upper-tropospheric part of the

cyclone) implies midtropospheric downward motion. In a typical intensifying

cyclone, there is midtropospheric upward motion (with the associated clouds and

precipitation) ahead of and over the surface cyclone, and downward motion beneath

and behind the upper-tropospheric troug h (Fig. 15).

Figure 14 Schematic showing the deﬁnition of upshear and downshear for a simple, typical

vertical distribution of horizontal wind. The low-pressure system in the ﬁgure tilts upshear

with height.

6 EXTRATROPICAL CYCLONES 529

To diagnose and predict extratropical cyclogenesis, meteorologists use a relation-

ship between the structure of the large-scale balanced winds and temperatures and

the vertical motion known as the omega equation. The omega equation is a three-

dimensional analog to the Sawyer–Eliassen equation and is based on the same

Figure 15 Distribution of upward and downward motion in an idealized extratropical

cyclone. Depicted are a perspective view (top) and two-dimensional view (bottom) of

height contours at 1000 mb (dashed) and 600 mb (thin solid), surface fronts, and ascending

and descending airstreams (wide arrows) with their projections onto the 1000-mb or 600-mb

surfaces. (From Palmen and Newton, 1969).

530

LARGE-SCALE ATMOSPHERIC SYSTEMS

dynamical principles of continuous adjustment to preserve nearly balanced ﬂow. The

forcing, the wind and temperature patterns that imply upward motion, is rather

complicated when expressed mathematically, and several versions of the equation,

some involving simplifying assumptions, are in use. The most common qualitative

application roughly equates upward motion with the sum of two terms, one propor-

tional to the vertical derivative of vorticity advection and the other proportional to

the temperature advection. The ﬁrst term tends to dominate in the upper troposphere

and implies upward motion ahead of an upper-level trough and downward motion

behind it. The second term tends to dominate in the lower troposphere and implies

upward motion ahead of the surface cyclone (where warm advection is typically

found) and downward motion behind it. Between the upper-level trough and the

surface cyclone, it will be noted that the two terms are of opposite sign, leading to

difﬁculty in inferring the vertical motion ﬁeld there using this method. With the

increasing use of computers, many of these simplifying assumptions are being

discarded in favor of explicit mathematical calculation of the vertical motion forcing,

but the qualitative interpretation is still essential for relating the vertical motion ﬁeld

to the large-scale features in the lower and upper troposphere.

In contrast to the simultaneous treatment of winds, temperatures, and vertical

motion, cyclogenesis can also be understood from a dynamical point of view in

terms of potential vorticity. Indeed, the basic theoretical paradigm for cyclogenesis,

known as baroclinic instability, requires a potential vorticity gradient aloft opposite

in orientation to the surface potential temperature gradient as a necessary condition

for instability. Cyclogenesis itself, by any mechanism, requires an increase in the

integrated perturbation potential vorticity and=or surface perturbation potent ial

temperature in the vicinity of the surface cyclone. Baroclinic instability accom-

plishes this by having two waves exist simultaneously, one on the surface potential

temperature gradient and the other on the upper-tropospheric potential vorticity

gradient; if the two waves are of large enough horizontal scale and the upper one

is upshear of the lower one, the circulation associated with each wave causes the

other wave to intensify. Other mechanisms for growth that do not involve instability

include rearrangement (speciﬁcally, compaction) of potential vorticity by horizontal

shear, and a decrease in the vertical tilt between the upper and lower perturbations.

Using the electrostatics analogy, the instability mechanism involves increasing the

electric charge, while the other two mechanisms involve rearranging the electric

charge into a compact area. The potential vorticity approach and the omega equation

approach are mutually consistent and complementary.

7 UPPER-TROPOSPHERIC JETS AND TROUGHS

A meteorologist looks at the surface weather map to see the current weather but

looks at an upper-tropospheric map to understand the current weather. The up per-

tropospheric map, generally the 500-mb constant-pressure surface or above, shows

the meteorologist the upper-level featu res associated with the current vertical motion

7 UPPER-TROPOSPHERIC JETS AND TROUGHS 531

ﬁeld, the current motion of the upper-level features and related surface features, and

the potential for intensiﬁcation of surface cyclones and anticyclones. By examining

the large-scale height ﬁe ld at or near jet stream level, the meteorologist can infer the

likely distribution of surface temperatures and the likely path of any storms that

might develop.

Jet Streams and Jet Streaks

The jet stream is often described as a band of strong winds in the upper troposphere

(8 to 12 km above sea level) encircling the globe. Typically, however, the location of

the jet stream is not so simple. There may be two or more jet streams at a given

longitude, or the Northern Hemisphere jet stream may originate in the subtropical

Atlantic, cross southern Asia, pass over the Paciﬁc Ocean at about 35



N and the

Atlantic at 45



N, and eventually weaken over northern Asia. Three examples of the

wintertime Northern Hemisphere jet stream and one of the wintertime Southern

Hemisphere jet stream are shown in Figure 16.

The ﬁrst example (Fig. 16a) possesses many features common to the Northern

Hemisphere wintertime circulation. The strongest jet is over the Paciﬁc Ocean,

although it is unusually far south in this example. Consistent with this departure

from normal is an unusual northward displacement of the jet over North America;

long-term departures from the average circulation tend to have wavelengths of about

8000 km. Over the Atlantic Ocean and over Asia there are at least two jets. The

southern jet is called the subtropical jet and the northern jet is called the polar jet.

The second example (Fig. 16b) is unusual because the jet over the Atlantic is

stronger than the jet over the Paciﬁc. This unusually strong Atlantic jet corresponds

to a location where the polar and subtropical jet streams appear to merge. Over the

extreme northeastern Atlantic and northern Europe is a northward displacement of

the jet stream known as a block, because weather disturbanc es are carried northward

around the block rather than from west to east as would be normal. This particular

block is known as an omega block because of the resemblance of the jet stream

pattern to the Greek letter O.

The third example (Fig. 16c) occurred during an El Nin

o year and is also quite

unusual. It features a subtropical jet that is essentially continuous around the globe,

with a weaker polar jet that is nearly continuous except for its merger with the

subtropical jet over the Paciﬁc. Some resemblance may be found to the wintertime

jet pattern over the Souther n Hemisphere (Fig. 16d), which also features two nearly

continuous jets. As a general rule, because the topographic and sea surface tem pera-

ture variations in the Southern Hemisphere are comparatively small, the planetary-

scale waves in the jet stream tend to be weaker there. Also (not shown), the jet

streams actually tend to be stronger during the summer in the Southern Hemisphere

than during the winter because the pole-to-subtropics temperature contr ast increases

during the summer season. In the Northern Hemisphere, temperatures are much

more uniform during the summer and the upper-tropospher ic jet speeds are corre-

spondingly weaker.

532 LARGE-SCALE ATMOSPHERIC SYSTEMS

Figure 16 Wind speed (shaded, 10 m=s interval, beginning at 20 m=s) and height (solid,

120-m interval) at 250 mb for four wintertime hemispheres. (a) 0000 UTC Jan. 1, 1997. (b)

1200 UTC Dec. 15, 1997. (c) 0000 UTC Feb. 1, 1998. (d) 0000 UTC July 1, 1997.

7 UPPER-TROPOSPHERIC JETS AND TROUGHS 533

Figure 16 (continued )

534 LARGE-SCALE ATMOSPHERIC SYSTEMS

In a cross section through a jet stream (Fig. 17), the close relationship between the

horizontal derivative of tem perature and the vertical derivative of wind speed is

apparent. Beneath the jet is a strong horizontal temperature gradient, and the wind

speed increases with height; above the jet the opposite is true. At jet steam level, the

vertical derivative of wind speed is (by deﬁnition) zero and, to the extent that the

airﬂow is balanced, the horizontal derivative of temperature is zero too. The height of

the tropopause (marked by the transition between a rapid decrease of temperature

with height and nearly uniform temperatures with height) also varies across the jet

stream, being low on the poleward side and high on the equatorward side. Since

potential vorticity is much higher in the stratosphere than the troposphere, a strong

potential vorticity gradient would also be found in the vicinity of the jet.

A jet streak is essentially a local wind maximum within a jet stream. Being of

smaller scale than a jet stream, air parcels frequently undergo rapid accelerations as

they enter and exit a jet streak, implying strong ageostrophi c winds and the like-

lihood of patterns of divergence and convergence. For example, as an air parcel

enters a jet streak, it experiences a stronger pressure (or height) gradient than before,

and accelerates to its left in the direction of lower heights. This downgradient-

directed ageostrophic wind then implies a Coriolis force directed in the original

direction of motion, causing the air parcel to speed up and ultimately approach

balance with the height gradient. In terms of the ageostrophic wind, one expects a

Figure 17 Vertical section taken through the strong Atlantic jet shown in Fig. 16c along



W. South is to the left. Solid contours show the component of wind into the page (m=s) and

dashed contours show temperature (



C).

7 UPPER-TROPOSPHERIC JETS AND TROUGHS 535