Coastal mountains form a barrier to the wind field that may affect
both the downstream and upstream evolution of the flow. The problem is
characterized by two free parameters, the Froude number Fr, defined by
U/(Nh
) and the Rossby number Ro, defined by U/(fl
), where
U is the speed of the air stream, h
is the height of the barrier,
f is the Coriolis parameter, l
is the half width of the barrier,
N is the Brunt-Vaisala frequency and is equal to (g/
/
z)
, g is gravity, and q
is the mean potential temperature (the temperature of a parcel of
air moved dry adiabatically to a pressure of 1000 mb). Generally
blocking of the air flow occurs when Fr is < 1 [ Smith, 1989],
which for a typical value of N of 10
s
can occur with
elevations as low as 100 m. The influence of the earth's rotation
on the deceleration of the upstream flow is considered through Ro.
Deceleration is insignificant when Ro < 1. In steep topography it
has been shown that the deceleration zone will grow upstream to a
width defined by the Rossby radius of deformation l
, which is
equal to Nh
/f [ Pierrehumbert and Wyman, 1985; Chen
and Smith, 1987]. Steep topography is defined by the non-dimensional
slope (h
/l
)(N/f), being greater than 1. This may also be
written as Ro/Fr. In the coastal region the mountains often represent
a wall such that l
is typically greater than l
, where Ro >
1, and the flow is not expected to be geostrophic (i.e., the flow will
not remain perpendicular to the pressure gradient). The offshore
influence of mountain coasts is given by the Rossby radius of
deformation, which typically varies from 10 to 100 km.
Localized regions of high or low pressure generated in the coastal zone can become trapped and propagate along the coastline. Specific phenomena of particular interest are the energetic coastally-trapped disturbances that exist for between two and six days, and have length scales of 1000 km in the alongshore direction and 100-300 km across-shore. These features generally cause significant changes in the local weather in the coastal zone, for example, replacing clear skies with clouds and fog, and causing intensification and reversals of the wind field as the system moves along the coast. The existence of this class of feature has been well documented in California [e.g., Beardsley et al., 1987; Dorman, 1985, 1987; Mass and Albright, 1987; Winant et al. 1988; Zemba and Friehe, 1987; Eddington et al., 1992], the Pacific Northwest [e.g., Hermann et al., 1990], and Alaska [e.g., Overland and Bond, 1993; Macklin et al., 1993; Bond and Macklin, 1993].
When uniform onshore flow, characterized by a low Froude number, encounters a coastal barrier the steady-state response is a pressure ridge, a phenomenon referred to as damming [ Overland and Bond, 1993]. The topographically induced pressure fields produce along-ridge pressure gradients that can result in barrier jets [e.g., Doyle and Warner, 1993c]. The best examples of the phenomenon are found along the east coast of the U.S. associated with cold air damming between a coastal front and the Appalachian Mountains [e.g., Xu, 1990; Doyle and Warner, 1993c]. A similar structure can occur when the incident flow is not uniform, such as in the vicinity of a storm [ Mass and Ferber, 1990; Overland and Bond, 1993].
Overland and Bond [1993] investigated an unforecast wind
storm in the vicinity of Yuakutat, Alaska, which illustrates the
importance of ageostrophic dynamics (the acceleration of
the flow in response to an imposed pressure gradient) in a
coastal zone with steep terrain. They observed a pressure
perturbation that propagated northward along the coast as a
surge or a pulse when the geostrophic flow was directed onshore
(the wind determined by the alongshore pressure gradient).
Strong surface winds, 20-30 m s
, coincided with the
leading edge of the surge. In two other less dramatic storms
they observed damming and downslope flow at the coast. An
ageostrophic momentum balance in the alongshore direction occurs
when the coastal mountains are hydrodynamically steep, and trapped
mesoscale features propagate along the shore due to Kelvin waves
(waves that have a significant amplitude only within a distance
l
of a barrier) or density currents. These results indicate
that, to correctly forecast winds in the coastal zone, numerical
weather prediction models must resolve terrain slopes to determine
the correct flow dynamics.
Interaction between regional or mesoscale pressure gradients and
topography lead to complex wind and turbulence fields over the coastal
ocean [ Macklin et al., 1988; Bond and Macklin, 1993].
Strong ageostrophic winds occur in topographically restricted channels
when synoptic scale disturbances (features on a scale of 1000 km
that may last several days) produce an along-channel pressure gradient,
resulting in large accelerations of the wind downstream of
these topographic features [ Lackmann and Overland, 1989;
Macklin et al., 1990; Bond and Macklin, 1993]. These
phenomena are generally referred to as gap winds [
Overland and Walter, 1981; Bond and Macklin, 1993].
The criteria for the development of these wind systems are similar
to those required for coastal trapping. Low Froude number flows
are constrained to flow along rather than over barriers. Moderate
Froude number flows have more of a tendency to flow over barriers
inducing gravity waves over the mountains and associated extreme
downslope winds. Bond and Macklin [1993] investigated the effects
of upstream orography of the Alaskan peninsula on the low-level flow
in the coastal region using observations from NOAA P-3 aircraft flights
in the vicinity of Wide Bay. The topography is characterized by a low
sill, about 300 m high and 80 km wide, flanked by 900 m terrain to the
southwest and 1500 m terrain to the northeast. When Fr >> 1 the
boundary layer winds were approximately 30 m s
downstream
of the gap and moderate terrain. The cross-terrain component of
the wind, above the boundary layer, was 24 m s
upstream of
the barrier and up to 45 m s
approximately 70 km downstream
of the barrier. When Fr << 1 large wind speeds were
isolated to the region downstream from the gap at Wide Bay.
Over the coastal ocean, when the boundary layer inversion is
below the height of the terrain, the flow is constrained to follow
the coastline [e.g., Samelson, 1992]. Irregularities in the
coastline control the depth of the boundary layer and hence the
velocity of the wind. In the
vicinity of capes, expansion fans occur, the boundary layer
height decreases and the wind speed
increases downstream [ Winant et al., 1988]. Further
changes in the coastline cause the flow
to decelerate and a hydraulic jump may occur. Eddington et
al. [1992] used a nonlinear
numerical model of a well-mixed marine layer to study
topographically forced mesoscale
variability off coastal California near Point Conception.
Northwesterly flow is blocked by the
headland on the upwind side, which causes the marine layer height
to rise. On the downwind
side, the northwesterly flow removes mass from the region and the
marine layer height decreases.
The perturbation in the height of the marine layer creates a
pressure gradient force that is
responsible for a wind speed maximum. When the large-scale
forcing is sharply reduced a
coastally-trapped Kelvin wave forms to the north of the headland
and a marine layer eddy
develops to the south [ Eddington et al., 1992; Wilczak
et al., 1991].
These topographically constrained or modified flows have a
significant effect on the exchange of
heat, moisture, momentum and trace constituents between the ocean
and atmosphere. The
horizontal gradients in the wind velocity will modify the wind
stress curl affecting upper ocean
mixing [ Enriquez and Friehe, 1995], and the temperature
gradients, and will feedback to the
scalar fluxes. A drop in SST of 1--3 
C offshore to
approximately 150 km (the Rossby
radius), and a rise in SST of similar magnitude within
approximately 15 km of the shore were
observed by Hermann et al. [1990] in the coastal zone of
the Pacific Northwest. Since they
are topographically forced these flows are frequently stationary
allowing persistent modification
of particular areas of the coastal ocean.