Potter T.D., Colman B.R. (co-chief editors). The handbook of weather, climate, and water: dynamics, climate physical meteorology, weather systems, and measurements

Подождите немного. Документ загружается.

than the climatological average. The tropospheric air was about 12



C colder than the

mean in Buffalo. Despite the fact that the most substantial synoptic-scale ascent

typically is located to the north and east of a developing surface low, owing to a

favorable combination of warm advection and cyclonic vorticity effects, extremely

large snowfall rates in excess of 5 cm=h were observed in the Buffalo area. Several

factors contributed to such an extreme snowfall rate. These include large-scale

ascent associated with a strong upper-level trough. Within this environment, several

crucial mesoscale conditions amplify the response to this advection. First, the cold

air traveling over the relatively warm waters will hydrostatically destabilize the air.

Second, the evaporation of water vapor from the lakes wi ll saturate the air, so that the

effective hydrostatic stability will be reduced even further.

Additional physical processes are respon sible for enhancing the snow. These

include the development of a surface trough (Fig. 12) that provides the larger-

scale ascent necessary to trigger moist convection in the presence of instability.

The existence of sensible heat transfer from the unfrozen lakes to the atmosphere

contributes to ascent. Additionally, the differential roughness between the lakes and

the surrounding land will creat e low-level convergence areas with accompanying

ascent. Often, as in the case of the December 17–19, 1985, event, mesoscale bands

of heavy snow will develop (Fig. 13), and a key forecast problem is to predict the

existence and movement of such band(s).

Freezing rain events can have devastating impacts on economic infrastructure.

The synoptic environment of freezi ng rain is characterize d typically by a surface

extratropical cyclone advecting warm, moist air above relat ively shallow, cold air

masses. Areas that experience persistent shallow, cold air are therefore prone to

freezing rain. Especially susceptible regions include larger valleys and basins.

Cortinas (2000) has documented the mean synoptic conditions for freezing rain

events in the Great Lakes region. Figure 14 illustrates that freezing rain occurs to the

northeast of a surface cyclone that advects warm, moist air poleward from either the

Gulf of Mexico or the Atlantic Ocean. Typically, the air to the northeast of the low

had been associated with a prior cold-air outbreak. The temperature stratiﬁcation

becomes very stable (Fig. 15), as the warm, moist air travels above the relatively cold

dense near-surface air. As Figure 15 shows, the lowest layer during a freezing rain

event is characterized by an inversion with near-surface air that is less than 0



C, with

a deep layer of air aloft that is greater than 0



C. The strong inversion helps to prevent

any turbulent mixing of the warm air aloft down to the surface. Furthermore, the

stratiﬁcation decreases markedly above the inversion, to the extent that nearly moist

adiabatic conditions in the free atmosphere may exist. The implication for precipita-

tion amounts is that substantial ascent, or even convection, may occur in these upper

weakly stratiﬁed layers. The lowest-layer air is typically from the east or northeast,

and this reinforces the inversion with cold-temperature advection from the down-

shear polar air mass. The wi nds at the top of the inversion are typically warm and

moist and blowing from the southwest. Therefore the vertical wind shear reinforces

the temperature inversion.

An extreme example of such an event occurred during January 5–9, 1998, when

an ice storm deposited greater than 100 mm of predominantly freezing rain in south-

556 WINTER WEATHER SYSTEMS

Figure 13 Precipitation echoes from the WSR-57 radar at Buffalo for the Lake Erie

snowstorm on 2330 UTC, December 17, 1985. Snowfall intensities are contoured for local

use, with linear attenuation, to depict the areas of heaviest snowfall more accurately. (From

Niziol, 1987; reprinted by permission of AMS).

Figure 14 Mean sea-level pressure (hPa, solid line) and 1000–666-hPa relative humidity

(%, dashed line) during freezing rain events over the central Great Lakes for the period

1976–1990. (From Cortinas, 2000; reprinted by permission of AMS).

4 PRECIPITATION 557

ern Quebec and upstate New York. Approximately 44 fatalities and damages in

excess of $4 billion occurred (NOAA, 2000). The synoptic conditions at 1200

UTC, January, 8, 1998 (Fig. 16), were typical for the event. The freezing rain

occurred in a strong geostrophic deformation zone that extended northeast of a

surface low from upstate New York into New England. Surface temperatures

during the event ranged from 5to1



C in extremely large 1000–500-hPa thick-

nesses of 552 dam or greater. The anomalously strong Bermuda anticyclone assisted

Figure 15 Median values of dry-bulb temperature (



C, dark solid line) and dewpoint

temperature (



C, dark dashed line) from a distribution of freezing rain soundings at Flint,

Michigan. Data are plotted on a skew T–log p diagram, with pressure (vertical axis) given in

hPa and winds shown by full barb ¼5m=s and half barb ¼2.5 m=s. (From Cortinas, 2000;

reprinted by permission of AMS).

558

WINTER WEATHER SYSTEMS

the tropospheric warming of the affected regions with northward transports of warm

and moist air from the Caribbean Sea. Surface winds advected cold air from the

northeast along the Saint Lawrence River Valley. These surface winds occurred

along the southern ﬂank of a bitterly cold 1044-hPa surface anticyclone centered

in the Canadian Northwest Territories. While the instantaneous synoptic-scale

features were unusually strong for a freezing rain event, the factor that produced

such an extreme event was the persistence for 5 days of this freezing rain–producing

pattern.

5 SUMMARY

We have discussed several meteorological phenomena associated with winter

weather. These events, associated with extreme cold, wind, and precipitation, have

the potential to produce substantial economic impact (NOAA, 2000). We have

described these events in terms of the larger-scale meteorological patterns that are

conducive to their existence. Additionally, we have seen how these events are

produced by the modulation of larger-scale atmospheric patterns by interesting

topographic features, such as mountains, lakes, and oceans. Although we have

focused our attention on North American phenomena, similar events exist in other

Figure 16 As for Fig. 1, except for 1200 UTC, January 8, 1998.

5 SUMMARY 559

regions of the world, where similar interactions exist between the atmospheric ﬂows

and topography.

REFERENCES

Cortinas, J., Jr. (2000). A climatology of freezing rain in the Great Lakes region of North

America, Mon. Wea. Rev. 128, 3574–3588.

Eichenlaub, V. L. (1979). Weather and Climate of the Great Lakes Region, Notre Dame, IN,

University of Notre Dame Press.

Kalnay, E., M. Kanamitsu, R. Kistler, W. Collins, D. Deaven, L. Gandin, M. Iredell, S. Saha,

G. White, J. Woollen, Y. Zhu, A. Leetmaa, B. Reynolds, M. Chelliah, W. Ebisuzaki,

W. Higgins, J. Janowiak, K. C. Mo, C. Ropelewski, J. Wang, R. Jenne, and D. Joseph

(1996). The NCEP=NCAR 40-year reanalysis project, Bull. Am. Meteor. Soc. 77, 437–471.

Klemp, J. B., and D. K. Lilly (1975). The dynamics of wave-induced downslope winds,

J. Atmos. Sci. 32, 320–339.

Kocin, P. J., P. N. Schumacher, R. F. Morales, Jr., and L. W. Uccellini (1995). Overview of the

12–14 March 1993 superstorm, Bull. Am. Meteor. Soc. 76, 165–182.

Lackmann, G. M., and J. R. Gyakum (1999). Heavy cold-season precipitation in the north-

western United States: Synoptic climatology and an analysis of the ﬂood of 17–18 January

1986, Wea. Forecasting 14, 687–700.

NOAA (2000). Billion Dollar US. Weather Disasters, 1980–2000 (available on the World Wide

Web at http:==www.ncdc.noaa.gov=ol=reports=billionz.html).

Niziol, T. A. (1987). Operational forecasting of lake effect snowfall in western and central

New York, Wea. Forecasting 2, 310–321.

Niziol, T. A., W. R. Snyder, and J. S. Waldstreicher (1995). Winter weather forecasting

throughout the Eastern United States. Part IV: Lake effect snow, Wea. Forecasting 10,

61–77.

Quiroz, R. S. (1984). The climate of the 1983–84 winter—a season of strong blocking and

severe cold in North America, Mon. Wea. Rev. 112, 1894–1912.

560

WINTER WEATHER SYSTEMS

CHAPTER 28

TERRAIN-FORCED MESOSCALE

CIRCULATIONS

JOHN HOREL

1 INTRODUCTION

The characteristics of the Earth’s surface affect climate and weather on all spatial

scales (Barry, 1992). However, many weather phenomena that are inﬂuenced by

surface inhomogeneities in elevation, moisture, temperature, snow cover, vegetation,

or roughness are organized on the mesoscale, which spans the range from 2 to

200 km. According to Pielke (1984), weather systems on the mesoscale can be

divided into two general categories: those that are forced primarily by instabil ities

in traveling large-scale disturbances (e.g., squall lines or mesoscale convective

complexes) and those that are forced by surface inhomogeneities (e.g., mountain=

valley circulations, sea breezes, or urban circulations).

The impact of mesoscale variations in the underlying land surface is evident in an

estimate of the average annual precipitation over the United States (Fig. 1). As a

general rule, precipitation increases locally as terrain height increases, for example

over the Appalachians and the mountain ranges of the western United States.

A further general rule is that precipitation is higher near the coasts, where moisture

is more abundant, than further inland.

This section emphasizes mesoscale weather phenomena that are strongly modu-

lated by the characteristics of the underlying surface. After discussion of thermally

driven and mechanically driven ﬂows, the impact of terrain upon precipitation

processes will be presented. The inﬂuence of other variations in surface properties

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.

561

upon mesoscale circulations is followed by discussion of the predictability of

terrain-forced mesoscale circulations.

2 IMPACT OF TERRAIN UPON WIND

Thermally Driven Flows

Mesoscale wind circulations are produced frequently by temperature contrasts that

develop as a result of terrain variations. As noted by Pielke and Segal (1986) , the

general nature of diurnal sea and land breezes was well understood in ancient times:

‘‘southward goes the wind, then turns to the north; it turns and turns again’’ (Eccle-

siastes 1:6). Sea (or lake) and land breezes are driven by horizontal tem perature

contrasts that develop between water bodies and adjacent land surfaces (Whiteman,

2000). As shown in Figure 2 for the Great Salt Lake (the largest body of water in the

continental United States to the west of the Great Lakes), differences in air tempera-

ture develop over the water and land surfaces as a result of the higher heat capacity

of the water relative to that of the surrounding land. In the case of the Great Salt

Lake during summer, the air over the lake is roughly 1



C warmer than the air over

the surrounding land at night and 4



C cooler during the day. These temperature

Figure 1 (see color insert) Average annual precipitation in inches, as determined from the

Parameter-elevation Regression on Independent Slopes (PRISM) model by Chris Daly, based

on 1961–1990 normals from National Oceanic and Atmospheric Administration (NOAA)

cooperative stations and Natural Resources Conservation Service (NRCS) SNOwpack

TELemetry (SNOTEL) sites. Modeling sponsored by USDA-NRCS Water and Climate

Center, Portland, Oregon. Available from George Taylor, Oregon State Climatologist,

Oregon Climate Service. See ftp site for color image.

562

TERRAIN-FORCED MESOSCALE CIRCULATIONS

differences drive the air from over the water to over the land during the day (lake

breeze) and drive the air from over the land to over the water at night (land breeze).

The resulting convergence of the surface winds at night over the Great Salt Lake and

divergence during the day are evident in Figure 2.

Mountain–valley circulations arise as a result of differential heating between the

ground in regions of complex terrain and the free atmosphere at the same elevation.

There are two broad categories for mountain–valley circulations: slope ﬂows and

along-valley winds (Whiteman, 1990, 2000). Typically, slope ﬂows are driven by

horizontal temperature contrasts between the air over the valley sidewalls and the air

over the center of the valley. Since a larger diurnal temperature variation occurs near

the ground, the higher terrain above the valley serves as a heat source during the day

and a heat sink at night. These temperature differences between the slope and free air

over the valley drive cold, dense air ﬂowing down the slope at night and warm, light

air surging up the slope during the day. Upslope ﬂows tend to be deeper than

the shallow nocturnal drainage ﬂows as a result of strong turbulent mixing during

the day.

–1

–3

0.0 6.0

12.0 18.0 0.0

0.0 6.0 12.0 18.0 0.0

1.0e–04

5.0e–05

0.0e–00

–5.0e–05

–1.0e–04

Figure 2 Top panel: Diurnal variation in the difference in air temperature (



C) over the land

surrounding the Great Salt Lake, Utah, versus the air over the lake during August 1999. The

air over the lake is warmer than the air over the land during the night. Bottom panel: Diurnal

variation in surface wind divergence (s

1

) over the Great Salt Lake. The air converges over the

lake at night and diverges away from the lake during the day (a similar diurnal variation in

surface wind divergence is evident over Lake Ontario; Chen, 1977).

2 IMPACT OF TERRAIN UPON WIND 563

Along-valley winds develop as a result of variations in the strength of slope ﬂows

within the valley or differences in temperature between the air in the valley and that

found over the adjacent lowlands. As an example of along-valley circulations,

Figure 3 shows the evolution of wind and temperature at a station located roughly

midway along the length of the Big Cottonwood Canyon in the Wasatch Mountains

of northern Utah. During this 3-day period, a warming trend was underway and the

skies remained clear; further, the station location within this narrow, steep canyon

was shaded during the morning and late afternoon. The diurnal temperature swing of

roughly 10



C drove a reversal in the wind from downvalley at night (from the east)

to upvalley during the day (from the west). The sudden onsets of the wind reversals

were particu larly striking during these 3 days.

Dynamically Driven Flows

As summarized by Whiteman (2000), the degree to which a hill or mountain affects

the air ﬂow depends upon the characteristics of the terrain feature (e.g., height,

width, roughness, orientation relative to ﬂow direction) and the upstream speed

and stability of the air. Since the speed and stability of the ﬂow can vary signiﬁcantly

with height, the impact of the terrain upon the ﬂow can change from one

atmospheric layer to another (Fig. 4).

Figure 3 Diurnal mountain–valley circulation in Big Cottonwood canyon in the Wasatch

Mountains of northern Utah. Top panel: temperature (



C) from midnight local standard time

(LST) October 15 to midnight October 18, 2000. Bottom panel: wind speed (solid line in m=s)

and wind direction (circles) for the same period. The weather station is located where the

canyon is oriented east–west.

564

TERRAIN-FORCED MESOSCALE CIRCULATIONS

Whether an obstacle signiﬁcantly obstructs the ﬂow depends upon whether or

not the approaching air has enough kinetic energy to lift the air over it. The non-

dimensional Froude number (deﬁned as Fr ¼u=Nh, where u is the speed of the

upstream ﬂow, N is the Brunt–Vaisala frequency, a measure of stability, and h is

the height of the terrain) provides a conceptual framework to assess the impact of

terrain height and ﬂow speed and stability. If the ﬂow approaching a relatively low

obstacle has strong winds and weak stability (Froude number greater than 1), then

there is sufﬁcient kinetic energy to cross the obstacle. On the other hand, if the ﬂow

approaches a relatively high barrier with weak winds and strong stability (Froude

number less than 1), then there is not enough kinetic energy to force the air over the

obstacle and the ﬂow is either channeled through gaps in the terrain or forced to

travel laterally around the obstacle. For example, Figure 5 shows fog spilling through

a gap in the coastal mountains of northern California as a result of strong stability at

crest level. The ﬂow is more likely to be blocked and required to travel around

obstacles if the barrier has a relatively short lateral extent or the ﬂow impinging upon

the mountain range is very shallow.

As stable air ﬂows across a mountain barrier (with Froude number less than 1),

mountain gravity waves are often created over or in the lee of the barrier (Durran,

1990). The structure of the mountain waves exhibited at any particular time depends

upon the charateristics of the barrier (e.g., height, width, mul tiple ridges) as well as

the stability, orientation of the ﬂow relative to the ridge crest, and vertical change of

wind with height. If sufﬁcient moisture is present, the uppermost portion of the wave

Figure 4 (see color insert) Cumulus clouds, indicative of shallow instability, above the

Wasatch Mountains in northern Utah are capped by lenticular wave clouds, indicative of stable

air ﬂowing across the mountain barrier (photo by J. Horel). See ftp site for color image.

2 IMPACT OF TERRAIN UPON WIND 565