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downgradie nt ageostrophic wind in the entrance region of a jet streak, and a corre-

sponding upgradient ageostrophic wind in the exit region.

For an idealized straight jet streak (Fig. 18), this ageostrophic wind conﬁguration

produces characteristic patterns of convergence and divergence. Divergence is found

in what is known as the right entrance region and the left exit region, with conver-

gence in the left entrance region and the right exit region. The pattern is reversed in

the Souther n Hemisphere. And, since jet streaks are located near tropopause level

and the overlying stratosphere inhibits vertical motion, the divergence and conver-

gence imply rising and sinking motion beneath the jet streak in the middle tropo-

sphere. Consequently, clouds and precipitation are favored to the left or poleward

side of an approaching jet streak and to the right or equatorward side of a receding

jet streak.

An along-stream intensiﬁcation of the height gradient, such as is found at the

entrance region of a jet streak, implies conﬂuence of the balanced geostrophic wind.

Figure 18 Idealized four-quadrant jet streak. ( a) Patterns of divergence and convergence

(DIV and CONV) and associated transverse ageostrophic winds (ar rows) associated with a

straight jet streak. (b) Vertical sections along A–A

and B–B

from (a) showing the vertical

circulations in the entrance and exit regions. Also plotted in (b) are the jet position (J) and two

representative isentropes (dotted lines). (From Kocin and Uccellini, 1990).

536

LARGE-SCALE ATMOSPHERIC SYSTEMS

Beneath the jet streak is a horizontal temperature gradient, very often in the form of

an upper-level front. The conﬂuence implies frontogenesis, and the ageostrophic

ﬂow at jet streak level constitutes the upper branch of the direct circulation asso-

ciated with frontogenesis.

Short and Long Waves

Just as straight, concentrated gradients of potential vorticity in the upper troposphere

are associated with jets, waves in the potential vorticity gradient are associated with

upper-tropospheric troughs and ridges. Indeed, the concentrated gradie nts serve as a

medium along which the waves, known as Rossby waves on the synoptic scale and

Rossby–Haurwitz waves on the planetary scale, propagate.

The propagation mechanism is fairly straightforward and can be applied also to

the upper and lower waves in a baroclinic extratropical cyclone. Imagine that a

straight jet has been perturbed into a periodic wave pattern (Fig. 19). Each equato r-

ward excursion of the potential vorticity contours will be associated with a cyclonic

circulation, by the principles outlined in the ﬁrst part of this chapter, and each

poleward excursion will be associated with an anticyclonic circulation. In a frame

of reference moving with the large-scale westerlies, these circulations imply an

alternatin g pattern of northerlies and southerlies, which act to redistribute the poten-

tial vorticity so as to cause the potential vorticity waves to propagate westward. The

speed of propagation is proportional to the strength of the potential vorticity gradient

and to the wavelength. As a general rule, the propagation speed of short waves

(wavelengths less than 5000 km or so) tends to be much less than the speed of

the westerlies in which they are embedded, so short waves (also known as mobile

Figure 19 Schematic Rossby wave train. Top panel: undisturbed upper-level jet showing

winds and PV variations. Bottom panel: Rossby waves along upper-level jet with associated

north–south component of winds. The pluses and minuses represent cyclonic and anticyclonic

potential vorticity anomalies, respectively.

7 UPPER-TROPOSPHERIC JETS AND TROUGHS 537

troughs for this reason) tend to move in the same direction and slower speed as the

upper-level winds in which they are embedded. Midlatitude long waves may move

slowly, be stationary, or even move westward against the large-scale ﬂow. Stationary

ridges that persist for one or more weeks are known as blocks or blocking highs

because extratropical cyclones follow the jet and mig rate poleward around the ridge,

and are blocked from entering the affected area.

Rossby waves are dispersive, and the group velocity is oriented opposite the

phase velocity, so wave energy tends to propagate downstream at a speed faster

than the jet stream winds. This phenomenon is illustrated schematically in

Figure 20, which shows an isolated upper-tropospheric trough. The cyclonic

winds associated with the trough act to move the trough westward, but at the

same time they generate a ridge to the east of the trough. As time goes on, that

ridge would then generate a trough farther downstream, and so on. Meanwhile, the

original trough would weaken, unless it is involved in cyclogenesis or has some

other energy source to allow it to maintain its strength. The process by which one

wave triggers another wave downstream is known as downstream development. An

example of downstream development in the atmosphere is shown in Figure 21.

Long waves may be formed by interaction between the jet and the underlying

large-scale topography, by interaction between the jet and the embedded short waves,

and by large-scale latent heating produced by convection. The latter mechanism is

the primary means by which equatorial oceanic or atmospheric conditions affect the

midlatitude and polar circulations. Equatorial tropospheric heating by convection

and the associated upper-tropospheric divergence alters the potential vorticity pattern

and initiates a wave train whose energy propagates poleward and eastward. Ideally,

the waves follow a great circle path and ultimately return to the equator, but the exact

path of the waves is inﬂuenced by the local wind and potential vorticity character-

istics of the medium within which they are embedded. The wave train persists as

long as the equatorial forcing persists.

Figure 20 Schematic example of energy propagation associated with an isolated upper-level

trough on a jet such as the one depicted in Fig. 19. As the pattern evolves, successive troughs

and ridges appear downstream (to the east), while the original wave weakens. While individual

crests and troughs propagate upstream, the wave packet as a whole propagates downstream.

538

LARGE-SCALE ATMOSPHERIC SYSTEMS

8 WATER VAPOR IMAGERY

The primary satellite tool for detecting upper-tropospheric weather features is the

water vapor image. This image is based on a near-infrared wavelength for which

water vapor is a strong absorber and emitter, so that the radiation reaching the

satellite typically comes from water vapor in the middle to upper troposphere.

Where the upper troposphere is particularly dry, the radiation reaching the satellite

originates from the lower troposphere, from water vapor at a higher temperature. The

variation in intensity of radiation caused by the differing temperatures of the water

vapor emitting the radiation is used to infer the distribution of moisture (or lack

thereof) in the upper and middle troposphere.

A water vapor image can be used to pinpoint the locations of troughs and jets.

Often, the jet stream is marked by a strong upper-level moisture gradient, with low

Figure 21 Example of downstream development in the atmosphere, Feb. 25–28, 1997.

Contours represent 250-mb height at 120-m intervals. First panel: relatively straight jet across

North Atlantic. Second panel: trough develops over west Atlantic associated with surface

cyclogenesis (not shown); ridge begins forming in east Atlantic. Third panel: trough begins

weakening over central Atlantic; ridge moves eastward and reaches maximum intensity over

Great Britain; new trough forming over central Europe. Fourth panel: trough nearly gone over

central Atlantic, ridge weakening over northern Europe, trough over central Europe attains

maximum intensity.

8 WATER VAPOR IMAGERY 539

moisture on the cyclonic side of the jet and higher moisture on the anticyclonic side

of the jet. This pattern is partly related to the conﬂuence of different air masses as air

enters the jet, and partly related to the lower tropopause on the cyclonic side of the

jet or in troughs, with comparatively low water vapor content in the stratosphere

compared to the troposphere. However, similar features can be produced by ordinary

advective processes and vertical motion, so loops of images or maps of upper-level

height and wind must be consul ted to properly interpret features appearing in water

vapor images.

Figure 22 shows the water vapor satellite image for the time of the strong Atlantic

jet shown in Figure 16b. The water vapor image features a band of moisture crossing

Cuba and becoming brighter as it passes alongside the East Coast and moves south

of Newfoundland. The sharp left edge of this band is associated with the axis of

maximum winds at jet level. By contrast, the polar jet crossing into the Atlantic from

Canada does no t appear as a distinct feature in this water vapor image; as a general

rule, southwesterly jets are easier to identify in water vapor imagery than north-

westerly jets. The two upper-level vortices over Florida and off the coast of Spain are

also prominent in the water vapor image. The dark area east of Newfoundland, north

of the subtropical jet, indicates the position of an upper-level trough and is caused by

the extremely low tropopause there. The trough appears to be initiating cyclogenesis,

judging from the developing water vapor pattern just to its east (compare Fig. 22 to

Fig. 10b). The much larger dark areas over the eastern Caribbe an are associated

Figure 22 Water vapor image of strong Atlantic jet, 1115 UTC Dec. 15, 1997. Compare

with Figs. 16b and 17.

540

LARGE-SCALE ATMOSPHERIC SYSTEMS

with dry, subsiding air within the subtropics and are of no particular dynamical

signiﬁcance.

REFERENCES

Carlson, T. N. (1991). Mid-Latitude Weather Systems, London, HarperCollins.

Kocin, P. J., and L. W. Uccellini (1990). Snowstorms Along the Northeastern Coast of the

United States: 1955 to 1985, Meteor. Monogr., 22, No. 44, Boston, American Meteor-

ological Society, 280pp.

Nese, J. M., and L. M. Grenci, (1998). A World of Weather: Fundamentals of Meteorology, IA,

Kendall=Hunt.

Roebber, P. J. (1984). Statistical analysis and updated climatology of explosive cyclones, Mon.

Wea. Rev., 112, 1577–1489.

Shapiro, M. A. (1991). Frontogenesis and geostrophically forced secondary circulations in the

vicinity of jet stream— frontal zone systems, J. Atmos. Sci., 38, 954–973.

Shapiro, M. A., and D. Keyser (1990). Fronts, jet streams, and the tropopause, Extratropical

Cyclones: The Erik Palmen Memorial Volume (C. Newton and E. Holopainen, eds.),

Boston, American Meterological Society, 167–191.

Thorpe, A. J. (1986). Synoptic scale disturbances with circular symmetry, Mon. Wea. Rev., 114,

1384–1389.

Whittaker, L. M., and L. H. Horn (1984). Northern hemisphere extratropical cyclone activity

for four mid-season months, J. Clim, 4, 297–310.

REFERENCES 541

CHAPTER 27

WINTER WEATHER SYSTEMS

JOHN GYAKUM

1 INTRODUCTION

Winter weather in the extratropical latitudes is deﬁned by the individual and collec-

tive contributions of cold air, wind, and precipitation. Each of these meteorological

phenomena is determined by a combination of synoptic-scale weather systems and

topography.

Economic impacts of winter weather can be substantial. Among the 48 weather

disasters that cost at least $1 billion in the United States during the period

1980–2000, 11 were winter weather events (NOAA, 2000). These included 4 ﬂooding=

heavy-rain events, three coastal cyclones, two freezes in Florida, and two ice storms.

This chapter focuses on the dynamics of winter weather systems and how

topographic features affect these systems.

The weather maps used in this chapter show ﬁelds of sea level pressure (SLP),

500–1000 hPa thickness, and their anomalies from a climatological state. Eac h ﬁeld

is taken from the National Centers for Environmental Prediction (NCEP) global

reanalysis of meteorological data (Kalnay et al., 1996). The thickness ﬁelds display

the 500–1000 hPa layer mean virtual temperature. An incremental change of 60 m of

500–1000 hPa thickness represents an approximate temperature change of 3



C. The

anomalies of this quantity are also displayed on these maps. We deﬁne an anomaly

as the difference between the actual 500–1000 hPa thickness and the appropriate

monthly climatological mean. This mean value is computed for the period

1963–1995. Therefore, a thickness anomaly ﬁeld quantiﬁes how much colder or

warmer a particular region is, compared with its expected value for the season.

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.

543

2 COLD AIR

Winter is characterized by relatively cold temperatures in extratropical latitudes.

These colder temperatures are a manifestation of equatorward displacements of

the zonal jet stream and cold-surface anticyclones. One especially interest ing

component of winter weather is the production of cold-air outbreaks. These

outbreaks are associated with anomalously cold temperatures, where an anomaly

is deﬁned as the difference between the actual temperature and the long-term

climatological mean.

For a cold-air outbreak to have a large impact, cold air must travel rapidly

equatorward from its source region in the higher latitudes of continents. The rapid

displacement is essential for a strong cold anomaly to occur, because cold air will

warm at lower latitudes owing to the warmer ground or ocean surface and to stronger

solar insolation. A favorable upper-tropospheric pattern includes an anomalously

strong wave in the jet stream with the ridge and trough being ampliﬁed relative to

the mean climatological state.

A severe cold-air outbreak occurred during December 1983 with damage to the

Florida citrus crop exceeding $2 billion. The structure of the highly ampliﬁed 1000–

500 hPa thickness ﬁeld (Fig. 1) included a strong ridge in Alaska with an anomaly

Figure 1 Sea-level pressure (light solid, interval of 8 hPa), 1000–500-hPa thickness

(light dashed, interval of 6 dam), and 1000–500-hPa thickness anomaly (heavy solid=dashed

dashed for positive=negative with interval of 120 m) for 0000 Universal Time Coordinated

(UTC), December 26, 1983.

544

WINTER WEATHER SYSTEMS

exceeding þ24 dam and a cold anomaly (44 dam) in the eastern middle Atlantic

states. The structure shown by Figure 1 is an extreme case of a tropospheric ﬁeld that

persisted for much of December 1983 (e.g., Fig. 8b of Quiroz, 1984).

The qualitative structure of Figure 1 illustrates the necessary and sufﬁcient condi-

tions for a signiﬁcant cold outbreak. The tropospheric ridging in the thic kness ﬁeld

develops as a result of strong surface cyclogenesis in the North Paciﬁc, as poleward

ﬂow to its east advects warm air into Alaska. This strong ridging in western North

America provides anomalously strong equatorward steering ﬂow for cold-surface

anticyclones. A cold-surface ridge axis extends southeastward from Alaska to the

Mexican border east of the Rockies, which provides a natural barrier from the

moderating inﬂuence of the North Paciﬁc. The combination of the surface anti-

cyclones and a deep cyclone centered south of Greenland provides a northerly

low-level geostrophic ﬂow extending from the Arctic to the Caribbean Sea.

Topographic features, such as the Rockies, are often important to the evolution

of the cold-air outbreak. A relatively warm sea surface temperature (SST ) in the

vicinity of the tropospheric thickness ridge would enhance its amplitude. A snow

cover over the eastern part of North America acts to amplify the thickness trough.

As in North America, cold-air outbreaks in other regions of the globe are cha-

racterized by an upper-level pattern that favors a cold continental surface anticyclone

traveling into the affected region. Signiﬁcant cold-air outbreaks in western Europe,

for example, are characterized by southwestward-traveling anticyclones from the

continental regions of Russia.

3 WIND

High winds during the winter are often associated with blizzard conditions. The

ofﬁcial deﬁnition of a blizzard includes large amounts of falling or blowing snow,

with wind speeds greater than 56 km=h and visibilities less than 0.4 km for a least

3 h. Blizzards are substantial threats to North American east coastal regions because

intense surface cyclones and their accompanying strong horizontal pressure gradi-

ents occur preferentially in these regions. Coastal residents are therefore especially

vulnerable to the effects of these deep surface lows, which may include hurricane-

force winds and storm surges comparable to those associated with land-falling

hurricanes. Other locations in Nor th America that experience blizzard conditions

include the high-plains region extending east of the Rocky Mountains. Blizzard

conditions are more likely to occur in these areas in association with high winds

and blowing (with small amounts of falling snow) over relatively smooth terrain.

North American east coast cyclones develop in response to a middle-u pper tropo-

spheric mobile trough that typically travels along a northwesterly upper-level ﬂow.

The initial development of the coastal surface low may be in a preexisting zone of

coastal frontogenesis that separates warm, moist marine air from cold, dry air that is

dammed east of the Appalachians. The interaction of this lower-tropospheric cyclo-

nic potential vorticity anomaly with its upper-level mobile counterpart deﬁnes the

cyclogenesis.

3 WIND 545