[Analysis of Anchorage Downslope Windstorms]{.underline}

John Papineau, NWS Anchorage

[1*. Introduction*]{.underline}:

Moderate-to-strong southeast winds are a common occurrence at summit
level (\~1200 m) in the western Chugach Mountains just to the east of
Anchorage (Figure 1) during the cooler months of the year, October
through April. On occasions when atmospheric conditions become favorable
a downslope windstorm forms and extends down the lee side of the Chugach
Mountains, "the "Hillside," and into east Anchorage. There have been a
number of these windstorms over the subsequent decades that have
produced considerable damage to structures in east Anchorage in addition
to toppling power poles and numerous trees. In addition these events are
also a serious threat to civil and commercial aviation throughout
northern Cook Inlet as areas of strong turbulence and low-level
windshear are common.

Southeast winds in the greater Anchorage area represent a continuum of
events; from damage producing downslope windstorms to what might be
referred to as 'generic' winds that are confined to higher elevations of
the Hillside and major valleys such as Turnagain Arm. Downslope
windstorms result from the formation of a mountain wave(s) over the
western Chugach Mountains as stable air is forced over the barrier. The
generation and evolution of these features has been elusive and
therefore this paper is an attempt to clarify the difference between
generic Hillside winds, many of which are weak downslope events, and the
far less frequent and stronger damaging downslope windstorms. The
ability to identify these features would lead to better forecasting of
these damaging events.

It should be noted that strong SE winds through Turnagain Arm, a
sea-level gap through the Chugach Mountains just south of Anchorage,
form in conjunction with both downslope windstorms and generic Hillside
winds. This northwest-to-southeast oriented 65 km long fjord is an
important terrain feature because it provides a near sea-level link
between Prince William Sound located on the northern Gulf of Alaska,
with northern Cook Inlet . Strong east-to-west pressure gradients
frequently form during the winter months along the Alaska coastline
bordering the Gulf of Alaska, and produce moderate-to-strong winds
through Turnagain Arm. Strong winds are a hazard to vehicular traffic
that traverse the fjord along the Seward Highway, the only road between
Anchorage and the communities located on the Kenai Peninsula. A study of
winds through Turnagain Arm will be addressed in a separate paper; hence
we will only note their characteristics as they directly relate to
Hillside winds.

There have been four damaging downslope wind events in Anchorage in the
past three decades: April 1, 1980; November 26-27, 1985; October 10,
1986; and December 1, 1992. The National Weather Service has christened
these events as the "Big Four" in order to distinguish them from more
frequently occurring moderate events. Hopkins (1994) summarized the
December 1, 1992 downslope event that produced an estimated six million
dollars of damage in east Anchorage. This event generated a one-second
gust of 49 ms^-1^ (95 kt) at the 730 m elevation, the highest value ever
recorded on the Hillside. In an effort to improve forecasting
techniques, Hopkins noted some common characteristics of Big Four
events: an extremely large Cordova (PACV) to Anchorage (PANC) mean
sea-level pressure (MSLP) gradient; a stationary ridge located over
western Canada with a surface low tracking northward across western
Alaska (\~165^o^W); presence of a critical layer (upper level winds are
zero or aligned along 220^o^) between mountain top level and the
mid-troposphere; and elevated inversions or stable layers. Table 1 lists
these events in addition to some important atmospheric parameters.

> Table 1: Significant Anchorage downslope wind events.

+--------------+-------+---------+-------+---------+-------+---------+
| Date & Hour  | PACV  | St      | Stab  | 850 mb  | 850   | C       |
|              | -PANC | ability | ility | winds   | mb    | ritical |
| (UTC)        |       |         |       |         | winds | layer   |
|              | Pre   | 700-850 | 50    | \       |       |         |
|              | ssure | mb      | 0-700 | [speed, | \[s   | Height  |
|              | diff  |         | mb    | dir\]   | peed, | (       |
|              |       | (^o^C   |       |         | dir\] | 210^o^) |
|              | (mb)  | km^-1^) | (^o^C | (       |       |         |
|              |       |         | km    | ms^-1^, | (ms   | \(m\)   |
|              |       |         | ^-1^) | deg)    | ^-1^, |         |
|              |       |         |       |         | deg)  |         |
+--------------+-------+---------+-------+---------+-------+---------+
| April 1,     | 12.3  | 6.9     | 5.9   | 16 185  | 19    | 6100    |
| 1980 00Z     |       |         |       |         | 195   |         |
|              | 16.2  | 2.3     | 7.8   | 19 115  |       | none    |
| 12Z          |       |         |       |         | 33    |         |
|              |       |         |       |         | 170   |         |
+--------------+-------+---------+-------+---------+-------+---------+
| Nov 26, 1985 | 11.3  | 6.0     | 3.3   | 17 135  | 8 180 | 4600    |
| 12Z          |       |         |       |         |       |         |
|              | 17.1  | 6.4     | 2.2   | 24 125  | 18    | 4600    |
| Nov 27, 00Z  |       |         |       |         | 145   | (       |
|              |       |         |       |         |       | 205^o^) |
+--------------+-------+---------+-------+---------+-------+---------+
| Oct 10, 1986 | 10.3  | 4.6     | 4.9   | 20 145  | 24    | 4000    |
| 00Z          |       |         |       |         | 190   |         |
+--------------+-------+---------+-------+---------+-------+---------+
| Dec 2, 1992  | 11.7  | 5.4     | 4.2   | 23 170  | 27    | 2900    |
| 00Z          |       |         |       |         | 210   |         |
|              | 12.0  | 6.2     | 5.1   | 27 160  |       | 4800    |
| 12Z          |       |         |       |         | 39    | (       |
|              |       |         |       |         | 195   | 205^o^) |
+--------------+-------+---------+-------+---------+-------+---------+
| Jan 7, 2001  | 7.9   | 3.0     | 4.5   | 12 105  | 31    | none    |
| 12Z          |       |         |       |         | 145   |         |
|              | 11.6  | 7.7     | 6.6   | 23 140  |       | none    |
| Jan 8 00Z    |       |         |       |         | 21    |         |
|              |       |         |       |         | 158   |         |
+--------------+-------+---------+-------+---------+-------+---------+
| Feb 5, 2003  | 12.6  | 3.1     | 5.8   | 27 130  | 37    | none    |
| 00Z          |       |         |       |         | 148   |         |
+--------------+-------+---------+-------+---------+-------+---------+

The ability to forecast the difference between a generic Hillside wind
event and a damaging downslope event remains a challenge. The main
objective is to understand the role played by

elevated inversions and critical wind layers since one or both are
always present during these events. Relying on output from WRF model
simulations of the Big Four and sparse observations of events after
since Hopkins paper was written, the role played by these two parameters
in the generation of strong downslope windstorms can hopefully be
clarified. This paper is organized as follows: in section 2 we give a
brief review of the literature of downslope windstorms. Observations of
the Big Four windstorms are given in Section 3. The output from WRF
model simulations of the Big Four are discussed in Section 4. In Section
5 we discuss the similarities and dissimilarities of the modeling
results. In section 6 the results of two additional model simulated
downslope wind events in order to compare with the Big Four data set. A
conceptual model of Anchorage downslope events is offered in section 7.
We make some final comments and suggestions for future study in section
8.

[2*. Literature Review*]{.underline}:

The study of flow over mountainous terrain and associated downslope
winds has garnered considerable attention from meteorologists since the
post-World War II era. The early work as typified by Scorer (1949, 1953)
and Long (1954) was based on analytical solutions to various forms of
the Navier-Stokes equations that govern fluid motion, in this instance,
stable air moving over irregular terrain. One of the main discoveries of
this early work was the influence that various vertical profiles of
temperature and wind (speed and direction) have on the resulting flow.
Discontinuities in either one of these parameters can create layers
within the atmosphere whereby surface generated mountain waves could
reflect and amplify. Under the proper conditions, gravity waves can
superimpose and form a large amplitude mountain wave to the lee of the
terrain, generating a downslope windstorm (Peltier & Clark 1979, Durran
& Klemp 1983, Clark *et al* 1994).

The majority of work on this subject in the 1970s and 1980s was
accomplished by the use of various two-dimensional numerical models that
contained idealistic terrain. By the mid-1980s our understanding of
downslope windstorms had evolved to the level whereby three production
mechanisms were proposed (Durran 1990). One mechanism is wave breaking
resulting from vertically propagating mountain waves that reflect from
environmental critical layers. Wave breaking in turn creates it own
critical layer, *i.e.-*turbulent zone, that is typically lower then the
environmental critical layer (Clark and Peltier 1977, 1984). A second
mechanism is wave reflection off of vertical discontinuities in
stability (Lilly and Klemp 1979, Durran 1986). Hydraulic amplification
is the third production mechanism for downslope windstorms (Smith 1985,
Smith and Sun 1987). It should be noted that these mechanisms do not
necessarily work independent of each other. Wave reflection or wave
breaking may induce hydraulic amplification (Durran 1990). In other
words, there may be multiple atmospheric forcings leading to the
development of a downslope windstorm. However, attempts to relate
theoretical considerations such as the vertical profiles of the Scorer
parameter, Brunt-Vaisala frequency, and non-dimensional mountain height,
to real world complex terrain with limited observations, remains a
challenge.

Due to constraints on computer speed it was not until the 1990s that
three-dimensional weather modeling was able to incorporate realistic
terrain in high resolution simulations. One of the seminal
three-dimensional modeling efforts was that of Colle and Mass (1998a,b)
in their investigation of northern Cascade downslope and gap windstorms.
Using MM5 they found that windstorms of various intensities were
produced by a broad spectrum of lower and mid-tropospheric wind and
temperature profiles. Mid-tropospheric critical layers and elevated
stable layers were fundamental to the generation of strong lee-side
winds; nevertheless they found that no single combination of atmospheric
parameters was responsible for the strongest windstorms.

The Mesoscale Alpine Program (MAP) was an multinational field research
program conducted over a number of seasons throughout various locales in
the Alps. One of these studies (Jiang and Dole 2004) unmasked the
synergic role of various terrain scales; gravity waves generated by the
flow of air over shorter wavelength terrain help condition the
mid-troposphere to be able to amplify gravity waves produced by the
larger-scale terrain. These two scales can interfere constructively and
produce large amplitude lee-side waves. More recent work (Smith *et al*
2007) has focused on the role which boundary layer process, latent heat
effects and upstream blocking play in downslope windstorms. The
resolution of three-dimensional modeling is now down below 1-km as
represented in the work of Agustsson & Olafsson (2007).

The net result of these observational and modeling studies is that
downslope windstorms develop under a variety of tropospheric conditions,
which in turn will produce a continuum of responses. For example, in the
Front Range of Colorado downslope windstorms tend to form when the
mid-troposphere contains an inversion (or stable layer) but lacks a
critical layer. In other locations, the formation of downslope winds
seems to depend on the presence of a critical layer (Coleman and
Dierking 1992, Colle and Mass 1998b). In other words more than one set
of tropospheric parameters will most likely generate a lee side wind
response at any given location. However there is a subset of parameters
that for a given initial set of conditions and terrain will produce the
strongest response (Durran 1986). The challenge for meteorologists and
those responsible for forecasting downslope wind events is to determine
that parameter space for their particular area of responsibility.

[3*. The Big Four*]{.underline}:

As noted in the *introduction* there have been four damaging downslope
wind events in east Anchorage in the collective memory of the
forecasting community. In this section we will explore the similarities
and dissimilarities of these events based on the limited surface
observations and radiosonde flights from the airport (PANC) in west
Anchorage. Figure 2 shows temperature profiles observed at or near the
period of maximum winds for the four events. Notice the elevated
inversions in April 1980 and November 1985 cases, although this feature
is lacking in the remaining two cases. The low-level inversion for
December 1992 is not representative of conditions over the mountains or
upstream of the barrier, note however that there is a relatively stable
layer from 775 mb to 625 mb at this time. Figure 3 shows the radiosonde
winds observed on the same flights. The most notable characteristic is
the veering of the wind with height. Between 700 mb and 600 mb, the
winds are parallel or nearly parallel to the long-axis of the ridgeline
of the western Chugach Mountains (\~205^o^). This condition would
produce an environmental critical layer (Coleman and Dierking 1992,
Colle and Mass 1998b). The April 1980 and Nov 1985 cases lack these
southwest winds. Although there are some subtle differences one might
suspect that there are additional controlling factors to these high wind
events. It should be noted that it is uncommon to have a strong or even
moderate downslope wind event where the winds are southeast throughout
the troposphere.

Figure 4 depicts the mean sea-level pressure (MSLP) differences between
Anchorage and Cordova, the latter being a small community located on the
Gulf of Alaska some 250 km southeast of Anchorage. This pressure
difference, as noted by Hopkins (1994), has historically been one of the
fundamental parameters monitored by forecasters in an attempt to
forecast the onset and decline of downslope winds. Maximum pressure
differences range from 17 mb to 23 mb. It is interesting to note that
the April 1980 case had the 'weakest' pressure difference despite the
fact that to this day this event is considered the most destructive, not
only of the Big Four, but of any other event as well. Although this
pressure difference provides a qualitative check, it does not produce a
reliable measure of the strength of either the summit level wind speeds
or those in the lee-side windstorm.

It should be noted that during the cooler months, air moving from the
Gulf of Alaska over the western Chugach Mountains is significantly
warmer than the ambient air over northern Cook Inlet. In other words,
despite the winter timeframe during which these events occur, these are
not cold advection (boras) cases, but rather chinooks or fohns. The one
exception was the November 1985 case in which cold air from Canada moved
across the Gulf of Alaska and then over the western Chugach.
Nevertheless, despite the modest cold advection aloft, surface
temperatures in Anchorage increased once the SE surface winds commenced.

In order for a downslope wind to develop in Anchorage, two general
conditions must be satisfied. First, upstream vertical profiles of
pressure, winds, and temperature must be conducive to the acceleration
of SE winds over the final ridgeline of the western Chugach Mountains.
Second, the temperature and wind profiles within the boundary-layer over
Anchorage have to be conducive to the descent of winds from aloft. The
fact that strong downslope wind events in Anchorage are infrequent
reflects the difficulty of both conditions being met simultaneously. The
first requirement is met in various degrees a number of times each
winter. It is important to consider the fact that the climate regime in
northern Cook Inlet is significantly colder and drier then 80 km to the
southeast over the Gulf of Alaska where the downslope winds originate.

There have been many situations favoring strong winds over the mountains
and the potential for a downslope event, yet cold deep or relatively
deep air over northern Cook Inlet has prevented these strong winds from
reaching the surface. Cold air, in the form of a nocturnal inversion or
isothermal layer, acts as 'effective terrain' (Lee *et al* 1989) with
respect to the flow of air from the mountains to the southeast. Once
strong downslope winds have been generated over the lee of the Chugach
Mountains, they will attempt to erode the upper zone of the stable
boundary layer until the winds reach the surface. Therefore the erosion
of a dome of cold air over Anchorage is a function of the strength of
the SE winds, as well as the mean sea-level pressure (MSLP )gradient in
northern Cook Inlet and the Susitna Valley. It is common to have light
(\<5 ms^-1^) northwest-to-northeast surface winds over Anchorage, with
moderate to strong southeast winds over the Hillside, in conjunction
with a Talkeetna (PATK) minus Anchorage MSLP gradient of +2 to +4 mb. A
pressure gradient oriented down Cook Inlet helps to reinforce cold air
over Anchorage, as it frequently does when a surface low in western
Alaska is south of 60^o^N (Hopkins 1994). Once the low tracks further
north, the pressure gradient weakens or reverses, making it more
difficult for the cold air to remain in place. For example, at 12Z
November 26, 1985 there was a 100 mb deep layer of cold stable air over
Anchorage, despite increasing SE winds over the Hillside and an 11 mb
pressure difference between Cordova and Anchorage. It was not until 6
hours later that the surface winds over west Anchorage (PANC) abruptly
backed to the southeast while the sustained speeds jumped from 5 ms^-1^
to 14 ms^-1^ with gusts to 19 ms^-1^. The descent of warmer air that
originates from the Gulf of Alaska can produce some substantial
temperature increases over Anchorage in a matter of hours, once the cold
dome has been eroded. It is also common for surface air temperatures to
warm at a given site before the onset of moderate-to-strong winds at
that same location. Typical temperature increases of 10^o^ to 15^o^ C
over the course of three to six hours in west Anchorage is not unusual,
while temperatures at the base of the mountains to the east can rise
that much in one or two hours.

[4*. Model Configuration and Limitations*]{.underline}*:*

Modeling downslope windstorms is important in the effort to gain a more
in depth understanding because of the scarcity of available
observations, both at the surface and within the atmosphere, especially
for the Big Four events. This is true particularly over the Chugach
Mountains and upstream over the Gulf of Alaska. Model output also allows
us to explore the dynamical nature of these events that could not be
ascertained by any other means.

Simulation of the Big Four events was conducted with the Weather
Research & Forecasting (WRF) Environmental Modeling System configured
with an outer 4-km and a nested 1-km grid with one-way nesting from the
outer to the inner grid. Each simulation was initialized with the North
American Regional Reanalysis (NARR) data set using 6 hour boundary
files. Terrain data consisted of the USGS's 30 arc-second data set.
There were 31 vertical grid levels ranging from 20 m to 15,800 m (100
mb), with the upper 5 km of the atmosphere set to dampen vertical
velocities and hence limit unwanted wave reflection off the top of the
model domain. The microphysics routine was based on a five class scheme.
The planetary boundary layer was parameterized using the Yonsei
University scheme, while surface fluxes are based on Monin-Obukhov
similarity theory. In order to make comparisons from one simulation to
another, the grid configuration and model physics were unchanged from
one simulation to the next. The duration of these simulations ranged
from 36 to 48 hours depending on the nature of the event.

Model output was overall realistic; the main bias existed in the MSLP
fields. The modeled MSLP fields tended to be 5 to 10 mb too low compared
to the scattered observations. The MSLP gradient between Prince William
Sound and northern Cook Inlet was realistic as the pressure bias
appeared to be aerially uniform; the exception was the December 1, 1992
event in which we have observations from Whittier. For this case the
model produced a maximum Whittier-to-Anchorage pressure gradient on the
order of 10 mb during the time of peak winds, while the observations
indicated a 16 mb gradient. We are not sure why this discrepancy
occurred, although the geopotential heights for this case were in line
with the Anchorage radiosonde measurements. Unfortunately we do not have
any observations from Whittier for the older events.

Model generated 700 mb heights over Anchorage tend to range from +10 m
to -30 m when compared to the radiosonde flights. Near surface air
temperatures in the model for the April 1980 and December 1992 events
were very close to the observations taken in Anchorage. The October 1986
event had a considerable warm bias in the first half of the simulation
but then came into alignment with the observations during the second
half. The November 1985 simulation for unknown reasons displayed a large
warm bias throughout the length of the run. The near surface air
temperatures in winter in this region are highly dependent on
windspeeds. As noted above, even modest mixing can cause temperatures to
rise 10^o^ C in an hour or two; hence it is not surprising that modeled
temperatures deviated from observations at times.

Despite these limitations it is believed that model output can be useful
in interpreting the evolution of downslope windstorms, although some
caution has to be exercised. In the following subsections a brief
description of the model results for each of the Big Four events will be
presented. This section will be concluded by noting some significant
similarities and dissimilarities of the four cases.

**December 1992**

This event had the strongest observed Cordova-to-Anchorage MSLP
gradient, on the order of

23 mb. The model produced a peak Whittier-to-Anchorage MSLP gradient of
only about 10 mb compared to the observed 16 mb difference. However, the
850 mb height gradient from northern Prince William Sound to northern
Cook Inlet is 73 m, the same as for the October 1986 event and

9 m higher than in April 1980. We are not sure why the model MSLP
gradient was so weak, it could be related to the orientation of the
isobars. Northwest-to-southeast oriented isobars in the model may not
generate as high of pressure along the northern Gulf of Alaska as occurs
in nature with this orientation.

The peak Hillside winds generated by the model were on the order of 41
ms^-1^, the timing corresponding well with the qualitative observations
(peak gust of 51 ms^-1^). Mountain wave activity was confined to levels
below 650 mb. Prior to the maximum Hillside wind (06Z/2), the model
atmosphere formed a modest mountain wave over the lee slope with smaller
amplitude trapped lee waves downstream. At 06Z/2, the time of maximum
winds, there was a single mountain wave (Figure 5a & 6a) that had little
tilt from the surface 800 mb and then upstream tilt from 800 mb to 650
mb indicating upward propagating wave energy (Durran 1990). The vertical
profile of winds (Figure 7a & 8a) and temperature indicated that a
critical layer (SW flow) was positioned around 700 mb, however there
were no stable layers (lapse rate \~ 7.0^o^ C/km). Additionally, between
700 mb and 500 mb there was modest (8 ms^-1^) forward wind shear, speeds
increasing with height. As will be seen shortly, this event shares many
of the same characteristics as the November 1985 and October 1986
events; what is notably absent in this particular case is a
mid-tropospheric stable layer. This case resembles Smith's (1985)
hydraulic amplification conceptual model in which the air over the
western Chugach Mountains accelerates down the Hillside in the absence
of high or even moderate amplitude mountain waves aloft. This type of
acceleration is called "shooting flow" in hydraulics.

**October 1986**

This wind event differed from the other three in that it was embedded in
a larger scale disturbance that produced copious amounts of rain and
subsequent flooding across a large part of Southcentral Alaska. The
model generated a substantial lee-side windstorm, however like the
October 1985 modeling case, the simulated surface sea-level pressure
field was from 4-7 mb too low as were the lower and mid-tropospheric
geopotential heights. The model heights came into close agreement with
PANC sounding observations during the second half of the simulation.

This event had a larger southerly wind component in the flow over the
western Chugach Mountains than the November 1985 case. A lee-side
hydraulic jump developed from 12Z/9-05Z/10 as the upstream dividing
streamline, the level at which turbulent air is separated from
relatively non-turbulent air (Smith 1985),rose and windspeeds at the
base of the Hillside increased (Figures 5C-8C). This fits well with the
qualitative observations from the Hillside which assign the strongest
winds from 01Z-06Z/10. The ascending limb of the jump moved downstream
and weakened between 12Z-18Z/10; the shooting flow over the Hillside
weakened during this period as well.

As the downslope windstorm developed, the amplitude of the low-level
mountain wave over the Hillside amplified and extended down the slope.
However, the model indicated minimal increase in wave energy above the
650 mb level. The approximate form of the Scorer parameter in the
equation *S=\[N^2^/U^2^\],* where *N* is the Brunt-Vaisala frequency and
*U* is the upstream mean windspeed at barrier height (Smith *et al*
2002), , steadily decreased with height. This was to be expected as the
winds increased with height; their values were taken upstream of the
Chugach Mountains over northwestern Prince William Sound. However, there
were no sharp discontinuities in the Scorer Parameter that indicated the
presence of a reflective layer, or trapped lee waves downstream of the
Chugach. An analysis of the Brunt-Vaisala frequency, *N=\[g/Θ dΘ/dZ*\],
over the western Chugach indicated that in the 800-600 mb layer the
stability was increasing as the mountain wave descended the lee-slope.
In addition, an environmental critical layer located between 750 mb and
700 mb was present throughout the development period. Flow was \~ 10-15
ms^-1^ at 220^o^. The vertical velocity field indicated that narrow
bands of ascending and descending air within the waves were confined to
the troposphere below the 700 mb level (Figure 9).

In order to investigate the role of latent heating (Durran & Klmep 1983,
Smith *et al* 2007), this case was rerun without the model microphysics.
The results indicated that upstream model atmosphere below 600 mb was
indeed 3-6^o^ C cooler in this simulation when compared to the one with
full microphysics. However, tropospheric temperatures over the western
slopes of the Chugach Mountains were warmer in the absence of
microphysics. Several temperature profiles over the Chugach and further
upstream over northern Prince William Sound indicated that without
microphysics, a pronounced isothermal layer existed between 800-700 mb
that was absent with the full microphysics simulation. In addition,
during the period of maximum precipitation, the wind speeds over the
Hillside were some 15-25% stronger in the no-microphysics case as were
the amplitudes of the gravity waves. The conclusion is that the release
of latent heat over the barrier can have a significant impact on the
stability in the 850-600 mb layer and hence alter the resulting flow
over the lee-slope. Whether the release of latent heat increases or
decreases the winds over the Hillside and Anchorage might however depend
on the amount of heating and at what levels it is released (Doyle &
Smith 2003).

**November 1985**

This event differed from the other Big Four events in that cooler air
from western Canada via the northern Gulf of Alaska was advected across
the Chugach Mountains into northern Cook Inlet. From the early stages
(09Z/26) a lee-side mountain wave formed, centered at an elevation of

2500 m. There was a strong inversion over Anchorage stretching from Cook
Inlet to the base of the Hillside. During the next five hours this
mountain wave deepened and worked its way down the upper part of the
Hillside, coincident with a minimal 2 ms^-1^ increase in upstream wind
speeds. By 17Z/26 the pattern of isentropes resembled shooting flow
similar in nature to the two previous cases (Figures 5D-8D).

Windspeeds in the core of the lee-slope mountain wave increased by 6
ms^-1^ in conjunction with a 2.6 mb increase in the Prince William Sound
to Anchorage MSLP gradient. It is possible that wave amplification is
due to wave reflection off of a critical layer located around 650 mb.
There was little vertical tilt in the mountain waves located below 800
mb with noticeable upstream tilt in the low amplitude waves located
between 800-600 mb. At the beginning of the simulation, 02Z/26, there
was a shallow inversion centered around 750 mb. This layer rose to about
675 by 15Z/26 as cold air advection occurred in the 850-700 mb layer.
The inversion turned into a \~150 mb thick isothermal layer towards the
end of the simulation.

The downslope SE winds reached the surface in west Anchorage around
15Z/26 despite the continued north-south MSLP gradient down Cook Inlet.
The observed wind speeds at PANC correlated fairly well with the
simulated surface winds. After a period of moderate SE winds the model
and the observations indicated a return to northwest winds as the
downslope event began to weaken and the northern Cook Inlet MSLP
gradient was able to reassert itself. The modeled NW winds were however
several hours ahead of the observations.

The core of maximum winds, which extended from the top of the Hillside
down over the base of the mountains to west Anchorage, was on the order
of 50 ms^-1^. The elevation of this core averaged 500-1000 m AGL, \~975
to 925 mb. Beneath this core lay SE surface winds which were about half
the strength of those in the core. There were only a few scattered
observed wind values; a weather spotter located on the mid-Hillside
reported peak gusts of 35-38 ms^-1^ from 18Z/26 through 23Z/26. The bulk
of the damage reports were from east Anchorage where peak gusts were
estimated in the vicinity of 50 ms^-1^.

The lee-side mountain wave reached its maximum amplitude around 0Z/27 at
which time there was an area of wave breaking as evident by a decrease
in potential temperatures in the upper reaches (700-650 mb) of the wave.
By 04Z/27, the flow pattern indicated the development of a hydraulic
jump that initially formed over Cook Inlet and then propagated to the
east over Anchorage continuing on towards the Hillside. The 4-km grid
with its downstream boundary further to the west also indicated the
presence of a hydraulic jump; however it is considerably weaker than the
feature portrayed on the 1-km grid. By 10Z/27 the ascending limb of the
jump was over east Anchorage but the amplitude had decreased
significantly form the previous several hours. At the conclusion of the
simulation, (18Z/27) the jump resembled the elevated lee-side wave that
developed at the beginning of the simulation.

The simulation of this mountain wave (figure 10) also visually resembled
the hydraulic amplification process noted by Smith (1985) and Smith and
Sun (1987), with a dividing streamline that formed over the lee of the
Chugach Mountains. The height of the dividing streamline, determined as
the lowest isotherm over the lee-slope that descends at least 500 m,
varied from a height of 2400 m at the start of the simulation, to 3400 m
during the period of peak winds and then descended to 2700 m as the
event tapered off. We suggest that the initial wave amplification
process appears related to the reflection of mountain waves off of the
critical layer, close to the height of the dividing streamline. In
addition, it is evident that wave breaking, limited as it is in this
simulation, does occur once the mountain wave responsible for the
shooting flow reaches large amplitude. Essentially all of the downslope
forcing mechanisms noted in the introduction are present at some point
in this simulation. The role of the critical layer and elevated
inversion are important, but this one run does not allow us to determine
the importance of each separately.

**April 1980**

Strong SE winds were observed along the Hillside beginning around 04Z/1
with the strongest winds occurring between 09Z and 14Z. The highest
recorded gust was \~47 ms^-1,^ but estimates in the NWS storm report
reach as high as 60 ms^-1^. Most of the structural damage from this
storm occurred in the Muldoon area at the base of the Hillside, rather
than on the Hillside itself. The highest model generated winds in the
mountain wave over the Hillside at \~500 m above the terrain, were on
the order of 52 ms^-1^ at 07Z/1.

Recall that the observed MSLP gradient between Cordova and Anchorage was
the weakest of the four events discussed in this paper: \~16 mb. The
model generated a maximum Whittier-to-Anchorage MSLP gradient of 8 mb,
the weakest pressure gradient of the four simulated events. The maximum
windspeeds over the Hillside however were comparable to the November
1985 case and considerably stronger then the other two events (Figures
5B-8B).

From 0Z/1 to 6Z/1 as the winds over the Hillside increased, not only did
the peak-to-trough amplitude (PTA) of mountain waves increase, but the
energy propagated to higher levels in the model atmosphere, \~400 mb. As
the wave amplified, (Figure 11) the trough centered at 725 mb deepened
prior to any amplification that occurred at the surface. By 06Z/1 the
model had generated two elevated hydraulic jumps over Anchorage located
in the 900-600 mb layer with a PTA around 1100 m and minimal tilt. These
wave patterns were transient, however, as the stability in the middle
troposphere decreased concurrent with an increase above 400 mb. These
plots show that two of the phase lines are nearly vertical below 700 mb
where the stability is high. The trough positioned near 61.1 tilts
downstream and then upstream once above 750 mb. Above this level the
phase lines tilt upstream for the longer wavelength feature, but remain
nearly vertical for the shorter wavelength features. Wave tilt is a
function of the tropospheric stability; an upstream tilting wave
typically indicates a decrease in stability (Durran 1990). Mountain
waves generated by shorter wavelength terrain features tend to be
confined to regions of the atmosphere where the stability is high, as
they cannot propagate through regions of low stability.

The main difference between the April 1980 windstorm and the other three
members of the Big Four is that the model output indicated a higher
amplitude mountain wave in the mid-troposphere, with a significant
amount of mountain wave energy transmitted above 600 mb. Recall that
there was no critical layer in this event; the winds were predominately
southerly throughout the middle and upper troposphere. However, a strong
inversion formed early on April 1, mainly over the southern part of the
domain as a result of shallow warm advection, which may have played a
major role in the development of the mountain wave. The increase in
stability allowed for the amplification of mountain waves, especially
those of shorter wavelengths. High amplitude mountain waves also pose a
greater threat to aviation than the flows resulting from the hydraulic
amplification process.

Although we only have qualitative observations of windspeeds on the
Hillside or east Anchorage, we do know that this event produced the most
physical damage of any of the four events. It is possible that stronger
gusts were produced in this event because of the presence of a high
amplitude mountain wave located in the 800-600 mb layer, as compared to
the shooting flow generated in the other three events. We speculate that
shooting flow produces higher sustained [near surface]{.underline}
windspeeds over the lee-slope while a high amplitude mountain wave
produces higher gusts via shear instability. Although these gusts are
generated some distance above the surface they are advected towards the
ground by rotors and Kelvin-Helmholtz (K-H) waves (Peltier & Scinocca
1990).

[5*. Discussion*]{.underline}*:*

Since mountain waves play a prominent role in downslope wind
development, a comparison of wave structure resolved on the 4-km and
1-km girds was conducted. In general, mountain waves generated on the
1-km grid show considerably more fine scale structure when compared to
the same features on the 4-km grid. The 4-km grid produces a mountain
waves with peak-to-trough amplitude and lee-side winds comparable to
that of the 1-km grid. At times the winds generated by the mountain wave
on the 4-km grid are higher then on the 1-km grid. However, hydraulic
jumps and lee waves, whether stationary or transient, are
[not]{.underline} resolved on the 4-km grid as well as on the 1-km grid.
In addition, the more vertical isotherms on the latter grid, which are
in large part a product of shorter wavelength terrain that is resolved
on the 1-km grid, will lead to more rapid wave breaking (Figure 12). It
is also evident that shorter wavelength mountain waves propagate higher
into the model troposphere than longer wavelength features. Jiang &
Doyle (2004) found in 2-D simulations that wave breaking generated by
the shorter wavelength terrain features help pre-condition the model
atmosphere to longer wavelength waves by generating elevated turbulent
areas, wave induced critical layers.

Table 2 shows an inter-simulation comparison of the MSLP, 850 mb, 700 mb
height differences across the eastern Chugach from northern Prince
William Sound to northern Cook Inlet, at the time that the model
generated the maximum lee-side wind. It is interesting to note the
considerable height and MSLP differences (Whittier-to-Anchorage) between
the November 27 and the April 1 cases, despite a similar maximum wind
speed. We believe that these differences reflect the two different
mechanisms that generated these two events. As noted above, the November
27 event was similar to Smith's (1985) hydraulic amplification in which
the flow over the western Chugach Mountains transitions from subcritical
to supercritical flow.

*Table 2: Simulation comparison between northern Prince William Sound
and northern Cook Inlet. Maximum winds are in elevated wind core above
the Hillside.*

+---------+-----+------------+------------+------------+------------+
| Date    | T   | 850 mb ht  | 700 mb ht  | MSLP       | Maximum    |
|         | ime |            |            |            |            |
|         |     | Difference | Difference | Difference | Wind       |
|         |     | (m)        | (m)        | (mb)       | (ms^-1^)   |
+---------+-----+------------+------------+------------+------------+
| Oct 10, | 06Z | 73         | 53         | 9.5        | 41         |
| 1986    |     |            |            |            |            |
+---------+-----+------------+------------+------------+------------+
| Dec 2,  | 06Z | 73         | 53         | 9.0        | 41         |
| 1992    |     |            |            |            |            |
+---------+-----+------------+------------+------------+------------+
| Nov 27, | 06Z | 92         | 36         | 12.1       | 52         |
| 1985    |     |            |            |            |            |
+---------+-----+------------+------------+------------+------------+
| Apr 1,  | 07Z | 64         | 43         | 6.2        | 52         |
| 1980    |     |            |            |            |            |
+---------+-----+------------+------------+------------+------------+

In the April 1 event a high amplitude mountain wave formed which we
suspect generated localized areas of very strong turbulence due to wave
breaking, this resulted in damaging wind gusts reaching the ground
(Belusic & Klaic 2004). It is also possible that shear layers within a
high amplitude wave produce gusts in the 50-60 ms^-1^ range that are
periodically transported to the surface by turbulent eddies as discussed
in Brasseur (2001). It is these gusts that are responsible for the
considerable structural damage sustained in east Anchorage during the
April 1980 windstorm.

Using model output a simplified Scorer Parameter, *ℓ=(N^2^/U^2^)^1/2^*,
(Smith *et al* 2002) was calculated at 50 mb intervals for each of the
four simulations. Various authors have shown that a sharp discontinuity
(subject to interpretation) in the Scorer Parameter indicates a
'boundary' from which wave energy can be reflected back towards the
surface. The discontinuity in ℓ is a result of a either a significant
discontinuity in either the stability *(N)* or wind speed *(U),* as
would be found in a layer of speed shear. We selected the period during
which the winds were increasing in order to estimate if any
discontinuities in the Scorer Parameter were present. The analysis
revealed mixed results: the November 1985 event had a sharp
discontinuity centered around 600 mb, the October 1986 event had a weak
increase from 950-850 mb then a steady decrease from 750-500 mb. The
December 1992 profile indicated a weak but steady decrease from 800-500
mb. The April 1980 case displayed the lowest Scorer Parameter. It is
noteworthy that in the November 1985 event the discontinuity in the
Scorer Parameter corresponds to the same level as that of the critical
layer. This is not too surprising since the critical level corresponds
to southwest flow which often advects warmer air for at least some
initial period, as in an occluded front.

In a study of downslope windstorms over the west slopes of the Cascades,
Colle and Mass (1998) found that layers of wind shear could also be
important to the evolution and strength of the event. In order to see if
there are any strong shear layers, an examination of the upstream wind
profile of NW Prince William Sound near the time of maximum winds was
conducted. All four events contained forward shear (speeds increasing
with height) layers in the mid-troposphere, ranging from a low of 4.3
ms^-1^ km^-1^ in December 1992, to a high of 7.2 ms^-1^ km^-1^ in April
1980. Only the November 1985 and April 1980 cases had layers containing
reverse shear (speeds decreasing with height) in the lower troposphere
\~ 5.6 ms^-1^ km^-1^ in 900-700 mb. It is interesting to note that these
are the two events in which the model generated the strongest Hillside
winds. It is difficult to ascertain the importance of a reverse shear
layer since in the November 1985 and April 1980 events it occurred in
conjunction with other lower tropospheric features such as critical
layers and elevated stable layers. Colle and Mass (1998) found in their
3-D simulations that when their model was initialized with
mid-tropospheric reverse shear, the resulting lee-slope mountain waves
were stronger than in cases which were initialized with forward shear.

A NW-SE transect of 850 mb height tendencies from northern Cook Inlet
across the western Chugach Mountains to Prince William Sound indicates
that during the amplification phase of mountain wave development, the
largest decrease in geopotential height occurred over the lee-slope.
This area also had the minimum 850 mb height of the entire transect. We
suspect that these developing mountain wave(s) produce a localized
perturbation pressure (Smith 1985, Durran 1986) which acts to further
accelerate the winds across the top of the lee ridge into the developing
wave. This acts in the same direction as the meso-ά or synoptic-scale
pressure gradient that exists across the mountain barrier. The net
result is a local acceleration *(U ∂U/∂X)* of the winds that allows the
air to transition from subcritical *(Fr\<1)* to supercritical flow
*(Fr\>1)*. Using model estimates for each term of the definition of the
Froude number *\[-Fr^2^=(U ∂U/∂X)/(g ∂D/∂x)\]* (Durran 1990 eq. 4.35)
indicates that if Fr is set equal to one, using a realistic depth
gradient *(∂D/∂X)* over the lee slope of 0.05, and setting *U=*25 ms^-1^
results in a *∂U/∂X* on the order of 0.02 s^-1^ or \~20 ms^-1^ per km.
This type of speed gradient can only occur on the boundary of a mountain
wave.

For each of the Big Four events the value of the upstream , summit level
and Hillside winds was computed and then compared. The maximum speeds
over the Hillside (typically 500 m above the slope) are roughly twice
the speed of the upstream flow at the same elevation as it transitions
from the Gulf of Alaska to the Chugach Mountains. In addition, there is
considerable hysteresis in the data (not shown). By way of example, for
a given 850 mb windspeed over the Gulf of Alaska, windspeeds over the
Hillside may be 200% of the upstream speed while the mountain wave is
developing (Hillside winds are increasing). However, when the mountain
wave is dissipating (Hillside winds are decreasing), for the same 850 mb
level windspeed over the gulf, the hillside wind may only be on the
order of 100%-150%. This suggests that there is stronger dynamic forcing
when a mountain wave is forming compared to the period when it is
dissipating, possibly an isallobaric (pressure tendency) contribution-
this is a topic that warrants further consideration.

Model output clearly indicates the interaction between downslope winds,
Turnagain Arm winds and northerly surface winds generated by the
Talkeetna-to-Anchorage MSLP gradient. The net result is the generation
of boundary layer eddies over Anchorage. During periods when Hillside
winds remain aloft, surface winds out of Turnagain Arm and some other
valleys can bend around the exit regions and form eddies (Figure 13).
This corresponds with the observations from Merrill Field in Anchorage
which on April 1, 1980 have wind directions veering and backing hourly.
The observations at Anchorage International Airport are frequently
biased by outflow from Turnagain Arm, and hence do not necessarily
reflect that southeast winds from the downslope component stretch
east-to-west across Anchorage.

Another interesting aspect of the low-level flow in the model atmosphere
is the inhomogeneity of the near surface wind pattern to the lee of the
western Chugach that exists at certain times. Outflow from Turnagain Arm
and Eagle River Valley clearly generate perturbations in the level flow
below 850 mb. However, at times even in areas that are not punctuated by
a valley or irregular terrain, windspeeds and potential temperatures can
vary considerably along the western lee-slope of the Chugach. In
general, the largest inhomogeneities can be found during periods of the
strongest mountain wave activity because lee-side mountain waves are
often transient in nature and a product of both atmospheric parameters
and terrain forcing. Hydraulic flows, in contrast, tend to produce more
uniform flow along the western slopes of the Chugach as mountain wave
activity is limited.

[6*. Additional Downslope Wind Events*]{.underline}:

Two additional moderate downslope windstorms were modeled in order to
compare and contrast with the four damaging storms discussed above. The
January 7-9, 2001 and February 6-8, 2003 events were selected because
they displayed some interesting characteristics.

[a*. January 7-8, 2001*]{.underline}

The January 2001 windstorm produced observed peak gusts on the order of
28-33 ms^-1^ on the Hillside and 13 ms^-1^ in east Anchorage. At the
1200 m high Nike site located northeast of Anchorage, a gust of 60
ms^-1^ was observed at 18Z/7. A strong Talkeetna-to-Anchorage MSLP
gradient and the northerly surface winds it generated on January 7^th^
prevented the well developed mountain wave from reaching the surface
across most of Anchorage. On January 8^th^ however, as the
Talkeetna-to-Anchorage MSLP gradient weakened, winds exiting Turnagain
Arm (gap winds) fanned out over the city.

The North American Regional Reanalysis indicates that the MSLP gradient
along the northern Gulf of Alaska for this event was not as strong as
any of the Big Four. The Whittier-to-Anchorage MLSP gradient was \~7 mb
at its maximum. During their respective periods of peak mountain wave
development, both the April 1980 and January 2001 cases had speeds on
the order of 50 ms^-1^ in mountain wave core, and similar tropospheric
structures: a stable layer from 750-600 mb and strong mountain wave
activity in mid-troposphere. One notable difference between the two
events was that in the January 2001 case, the storm system that produced
the area of high pressure along the Gulf of Alaska coast was
considerably deeper then any of the Big Four, and it was positioned
closer to Cook Inlet. The question of interest is why were the observed
winds in east Anchorage in the April 1980 event a factor of two or three
stronger than in the January 2001 event, despite this similar
tropospheric structure? Was it solely due to the MSLP gradient in Cook
Inlet and the Susitna Valley, or were other factors contributing as
well? It is these types of atmospheric subtleties that make it difficult
to differentiate between a moderate and severe event.\
Analysis of the peak lee-slope winds originating from the mountain wave
indicates that in April 1980 the core of the strongest winds was closer,
that is lower in elevation, to the Hillside despite similar
Talkeetna-to-Anchorage MSLP gradients (+2 to +3 mb). The shear zone also
appears to be stronger in April 1980 case and hence wind gusts were
stronger as well. The primary difference between these two events is
that the cold dome is stronger and remains intact over Anchorage in the
January 2001 event longer when compared to the April 1980 case. In
addition, only when the Talkeetna-to-Anchorage MSLP gradient approaches
zero as it did early on January 8^th^, does it allow downslope winds to
reach the surface in Anchorage. However by this time the mountain wave
has weakened dramatically. The net result is that the winds which
reached east Anchorage early on January 8^th^ were considerably weaker
than they would have been if they had been able to reach the surface
hours earlier. In other words the peak mountain wave winds over and to
the lee of the Hillside occurred some six hours prior to any surface
observation from PANC indicating SE winds. In this case, the low
pressure center of \~960 mb tracked between Kodiak Island and Bristol
Bay along \~160^o^W, producing a strong north-south MSLP gradient down
Susitna Valley and Cook Inlet that could not be eroded by the strong SE
winds above the boundary layer. As a rule-of-thumb with all other
atmospheric parameters equal, a mid-winter cold dome over Anchorage is
more difficult to erode than one occurring during autumn or spring
events.

b*.* *[February 3-5,2003]{.underline}*

This moderate event generated observed peak gusts on the Hillside of
30-35 ms^-1^ and 15-

20 ms^-1^ in east Anchorage. This event was modeled and compared it to
the Big Four because it displayed a variety of lower and middle
tropospheric temperature and wind structures. For example, at 12Z/3 a
strong inversion was centered at 625 mb with a critical layer at 600 mb.
Model output displays considerable mountain wave activity extending
through 500 mb (Figure 14). At 00Z/4 the critical layer has now
descended to 775 mb and the inversion has dissipated. A comparison of
isentropes at 12Z/3 with 00Z/4 indicates that mountain wave amplitude
has diminished by 00Z/4, especially at the shorter wavelengths, despite
an increase in the cross-barrier MSLP gradient. The cross-section of
potential temperatures at 00Z/4 resembles hydraulic flow similar to the
three of the four Big Four cases. Model winds over the upper Hillside
are slightly stronger (22-25 ms^-1^) at 00Z/4, but they extend further
to the base of the Hillside at 00Z/4 when compared to their mid-Hillside
position at 13Z/3. This is probably due to the stronger cross-barrier
MSLP gradient and a weakening of the MSLP gradient in northern Cook
Inlet at 00Z/4.

The incoming flow however continues to evolve so that by 12Z/4 a
critical layer is located in the vicinity of 700 mb, but by 00Z/5 it is
well into the upper troposphere. Model winds did not back to the SE as
much as the PANC sounding indicates. In addition, there is an increase
in stability in the 800-650 mb layer at 00Z/5. The difference in wave
structure and lower tropospheric winds is pronounced. The lower right
panel of Figure 14 shows that mountain wave amplitude is significantly
larger in the 700-500 mb layer at 00Z/5 compared to earlier times. We
believe that the evolution of this flow and the resulting wave structure
confirms the idea that a critical layer in the lower troposphere limits
wave activity, but has the potential to generate strong hydraulic flow.
In contrast an elevated stable layer or inversion with either a weak or
absent critical layer, supports high amplitude mountain waves in the
middle and/or upper troposphere. This event displays both types of flow
regimes: hydraulic amplification at 00Z/4 and high amplitude mountain
waves at 00Z/5, with transitional of hybrid flow at other times.

[c*. Gusts*]{.underline}:

The National Weather Service captured a limited number of 1-second wind
observations from the primary Hillside weather station (GAHA2) during
the winter of 2007-2008. The data was extracted during periods of
moderately strong southeast winds in order to gain an understanding of
the frequency and duration of the highest gusts. In other words, do the
highest gusts occur as a single one second burst or do they occur as a
cluster over several seconds?

Using 10 hours of collected data from three events, the gust ratio, the
ratio of the 1-minute average and the highest 1-sec winds were
determined. Using the 10 strongest gusts during each hour and the one
minute average we found that the gust ratio ranges from 1.9 to 2.5; a
good average is 2.3. Therefore, a 1-minute wind of 20 ms^-1^ will
generate a peak gust on the order of 46 ms^-1^. Wavelet analysis
preformed on one hour long samples reveals a host of resonant
frequencies for this data. The most common ranges from one to nine
minutes, although at times frequencies of 30 to 60 seconds are common as
well.

Figure 15 shows a five minute window of winds which were observed on
January 22, 2008 just after 08Z. The high variability is readily
apparent. The lower chart in Figure 15 shows a one minute slice from the
upper chart. Note that around 30 seconds there is a sharp increase from
7 ms^-1^ to

29 ms^-1^ to 35 ms^-1^ in a span of two seconds. Also note there are
occasional gust clusters, one in the 30-35 second window where five
consecutive values exceed 25 ms^-1^. For this particular set of data
these gust clusters occur about once out of every four or five single
gusts. The high variability of these winds over the Hillside indicates
that there is considerable turbulence in the atmosphere. It appears that
at this time the flow can be classified as shooting regime, since there
is southwest flow at 700 mb (environmental critical layer), and weak
stability in the lower half of the troposphere. The ultimate goal is to
determine if high amplitude regimes produce the same high frequency
gusts as shooting flow regimes. This of course is going to require a
much larger data then we have currently archived.

[7*. Conceptual Model of Downslope Winds*]{.underline}*:*

Once a low amplitude mountain wave forms over the crest of the lee-slope
around the 850 mb level, the decrease in pressure (or geopotential
heights) in the trough of the wave enhances the cross-barrier flow above
and beyond that of the synoptic pressure gradient (Smith 1985). This is
due to the fact that in the descending limb of a mountain wave the air
is warmer than its surroundings due to adiabatic warming. As a result
the column of air underneath this portion of the wave is at a lower
pressure (geopotential height) then its surroundings. Once this begins
to occur, air that is subcritical may transition, if the added
accelerating is sufficient, to supercritical flow.

At this point we propose that there are two mechanisms responsible for
generation of downslope winds in the Anchorage area. The first scenario
requires a critical layer in the vicinity of 750 mb, that is roughly
1000-1500 m above the terrain. If the increase in windspeeds across the
barrier is sufficient, the flow can transition from a subcritical regime
to that of a supercritical one where the flow continues to accelerate
down the lee-slope. The exact nature of the transition from subcritical
to supercritical is elusive. However the critical layer seems to
restrict mountain wave activity which may allow the flow directly
upstream of the lee-slope to accelerate more rapidly than if wave energy
were not trapped at these lower levels. Equally important is the fact
that the depth of this accelerated layer cannot increase because the
environmental critical layer or any subsequent lower wave induced
critical layer is restricting the depth of the air, at least over the
lee-slope. As illustrated in Durran (1990), the net result is that the
pressure gradient term (change in depth of the fluid), cannot compensate
for the nonlinear acceleration and the air accelerates as it descends
down the lee-slope. Once the cross-barrier pressure gradient weakens,
the shooting flow diminishes over a number of hours and the flow
transitions back to a subcritical regime where the highest windspeeds
are found over the barrier itself and not to the lee. Changes in
upstream stability due to warm or cold air advection and/or the release
of latent heat can generate a variety of responses that would modify
this simple model.

In the second scenario the primary requirement is the existence of a
relatively deep stable layer or inversion in the mid-troposphere. This
layer allows for the generation of moderate to high amplitude mountain
waves throughout the middle and at times upper troposphere. These waves
will exhibit wave breaking characteristics throughout much of the period
of high winds. If a critical layer exists at all, it is located in the
upper troposphere above 500 mb. Events of this nature will tend to
produce higher gusts due to shear induced instabilities (K-H waves) in
the highly turbulent air aloft, compared to shooting flow regimes where
wave breaking appears to be more limited. Of the Big Four events
discussed in this study it is noteworthy that the April 1980 event
produced the most structural damage in east Anchorage despite the fact
that it had the weakest observed cross-barrier MSLP gradient. We believe
that the high amplitude mountain waves generated in the absence of a
pronounced critical layer allowed for the development of very strong
gusts which propagated down to the surface.

The question then arises: for any given cross-barrier pressure gradient,
will an event with large amplitude mountain waves consistently produce
stronger surface winds when compared to a hydraulic amplification type
event? Limited wave breaking in a particular hydraulic amplification
scenario could generate surface gusts comparable to those generated in
the high amplitude mountain wave scenario, although we are not able to
determine from model output, we believe that gustiness is more common in
the high amplitude mountain wave regimes. An additional complicating
factor is the presence of hydraulic jumps to the lee of the Chugach
Mountains. For example, a jump located directly over west Anchorage
tends to limit strong winds from reaching the surface, with the
exception of the Hillside and possibly east Anchorage. The model
simulations presented in this paper generate transient jumps of varying
amplitudes, no doubt produced by the variation in the height of the
critical layer, changes in lower tropospheric stability, and the
presence or lack thereof a cold dome over northern Cook Inlet.

Smith *et al* (2002) and Qiang *et al* (2005) have shown that the upper
boundaries of low-level stagnate or reverse flows can be important
reflectors of mountain waves that have propagated downward, i.e.
reflected down from a stable or critical layer. An important unanswered
question is whether a critical layer located in the vicinity of 600 mb
to 500 mb (well above our 750 mb criteria), will also generate a high
amplitude mountain wave(s). In their idealized model simulations, Colle
& Mass (1998) found that for the cold air advection cases over the
northern Cascades, a reduction in stability despite the presence of a
critical layer reduced wave amplitude and windspeeds over the
lee-slopes. To see if this holds true for Anchorage windstorms as well,
similar idealized simulations would have to be conducted using the
Anchorage terrain configuration and flow characteristics. Another
important factor with regard to the extent of downslope winds is the
magnitude of the northern Cook Inlet and Susitna Valley MSLP gradient.
With cold air in place over northern Cook Inlet prior to the
development, a modest +2 to +4 mb Talkeetna-to-Anchorage MSLP is often
enough to support the cold air dome and restrict the strongest winds to
the Hillside or a layer over the cold dome at \~950 to 900 mb. The
scenario is further complicated by the presence of moderate to strong
Turnagain Arm winds during downslope wind events. The specific nature of
Turnagain Arm winds will be addressed in a separate paper. However, they
frequently play an important role during downslope events in that they
can move over western and central Anchorage converging with the
downslope winds. In fact the vast majority of cases of SE winds at PANC
during a downslope scenario are probably winds exiting from Turnagain
Arm. It is rare to have downslope SE winds extending from the Hillside
across all of Anchorage, as evidenced by the few cases of strong SE
observed winds at Merrill Field or Elmendorf Air Force Base.

[8*. Conclusion*]{.underline}:

Essentially at least two mechanisms are responsible for the generation
of moderate to strong downslope windstorms in Anchorage. Hydraulic
amplification produces lee-side shooting flow when a critical layer
exists around the 750 mb level. High amplitude mountain waves are
generated in the absence of lower tropospheric critical layers but are
associated with elevated stable layers. Model output clearly shows that
the area of strongest downslope winds extends from above the lee-slope
of the Chugach to the base of the Hillside. At times this core of strong
winds extends NW across Anchorage. Whether these winds remain aloft
around 950 mb or extend to the surface is a function of the MSLP
gradient in northern Cook Inlet and the Susitna Valley and the presence
of hydraulic jumps.

The difference between a moderate and strong downslope wind event in
Anchorage is a function of several factors. Foremost is the presence of
a cold dome over Anchorage. There are cases in which the winds at crest
level are very strong with a developing downslope event, but the cold
dome prevents SE winds from reaching the surface and restricts further
amplification of the mountain wave. Another set of important factors is
the height of the critical layer. This is difficult to quantify until a
series of idealized simulations can be preformed. In addition the
height, strength and duration of elevated inversions and stable layers
play a role as well. Forecast model output can be used to monitor
changes in these parameters. The fact that severe downslope windstorms
in Anchorage are uncommon while moderate events are fairly common
suggests that the differences between the two are subtle, at least in
our ability to detect or forecast them.

As alluded to above, in order to further knowledge of these events and
answer critical questions, several idealized model simulations isolating
the importance of critical layers compared to stable layers should be
run. Specifically, is a transition seen from shooting flow to high
amplitude mountain waves as a critical layer is raised from the lower
into the middle troposphere? Is there a given stability that is more
conducive to high amplitude wave formation then others? Also, the role
of small-scale valleys located on the crest of the Hillside bears
further consideration; what is the nature of winds through these
features during periods of downslope events? Another worthy line of
inquiry would be to analyze the formation and propagation of hydraulic
jumps. In addition, are hydraulic jumps more common in high amplitude
mountain wave regimes then in shooting flow regimes? Finally, it would
be beneficial to continue to observe 1-second winds on the Hillside to
determine if there is any difference in gust strength and frequency
during events produced by high amplitude breaking mountain waves versus
those produced during shooting flow.

***[References]{.underline}*:**

Agustsson, H., H. Olasfsson. 2007: Simulating a severe windstorm in
complex terrain.

Met. Zeits. **16** 111-122

Belusic, D., M. Pasaric., M. Orlic. 2004: Quasi-periodic bora gusts
related to the structure of the troposphere. *Q. J. R. Meteorol. Soc*.
**130** 1103-1121.

Belusic, D., Z.B. Klaic. 2004: Estimation of bora wind gusts using a
limited area model. *Tellus*. **56A** 296-307

Brasseru, O., 2001: Development and application of a physical approach
to estimating wind gusts. *Mon Wea Rev*. **129** 5-25.

Clark. T.L., W.D. Hall., R.M. Banta. 1994: Two and Three-dimensional
simulations of the 9 January 1989 severe Boulder windstorm: comparison
with observations. *Jr. Atmos. Sci*. **51.** 2317-2343

Clark. T.L.., W.R. Peltier. 1977: On the evolution and stability of
finite-amplitude mountain waves. *Jr. Atmos. Sci*. **34**. 1715-1730.

Clark. T.L.., W.R. Peltier. 1984: Critical level reflection and the
resonant growth of nonlinear mountain waves. *Jr. Atmos. Sci*. **41**.
3122-3134

Colle. B.A., C.F. Mass. 1998a: Windstorms along the western side of the
Washington Cascade Mountains: A high-resolution observational and
modeling study of 12 February 1995 event.

*Mon. Wea. Rev*. **126** 28- 52

Colle. B.A., C.F. Mass. 1998b: Windstorms along the western side of the
Washington Cascade Mountains: Part II: Characteristics of past events
and three-dimensional idealized simulations.

*Mon. Wea. Rev*. **126** 53-71

Colman, B.R., C.F. Dierking. 1992: The Taku wind of Southeast Alaska:
its identification and prediction. *Wea. & Forecasting.* **7**. 49-64

Doyle, J.D., M.A. Shapiro. 2000: A multi-scale simulation of an extreme
downslope windstorm over complex topography. *Meteorol. Atmos. Phys*.
**74**. 83-101.

Doyle, J.D., D.R. Durran. 2002: The dynamics of mountain-wave-induced
rotors. *Jr. Atmos Sci*. **59**, 186-201.

Doyle, J.D., R.B. Smith. 2003: Mountain Waves over the Hohe Tauren:
influence of upstream diabatic effects. *Q. J. R. Meteorol. Soc*.
**129** (588) 799-823.

Durran, D.R., J.B. Klemp. 1983: A compressible model for the simulation
of moist mountain waves. *Mon. Wea. Rev*. **111.** 2341-2361.

Durran, D.R. 1986: Another look at downslope windstorms. Part I: The
development of analogs to supercritical flow in an infinitely deep,
continuously stratified fluid. *Jr. Atmos. Sci*. **43**. 2527-2543.

Durran. D.R., J.B. Klemp. 1987: Another Look at downslope winds. Part
II: Nonlinear amplification beneath wave-overturning layers. *Jr. Atmos.
Sci*. **44**. 3402-3412.

Durran, D.R. 1990: Mountain waves and downslope winds. Pp. 59-81 in
*Atmospheric processes over complex terrain*. Ed. W. Blumen. American
Meteorological Society, Boston, MA

Hopkins, R.H. 1994: Anchorage windstorm of 1 December 1992. *Wea &
Forecasting.* 9. 469-478

Jiang, Q., J.D. Doyle., R.B. Smith. 2005: Blocking, descent and gravity
waves: observations and modeling of a MAP northerly fohn event. *Q. J.
R. Meteorol. Soc*. **131**. 675-701.

Jiang, Q., R.B. Smith. 2003: Gravity wave breaking in two-layer
hydrostatic flow. *Jr. Atmos Sci.* **60**. 1159-1172.

Jiang, Q., J.D. Doyle. 2004: Gravity wave breaking over the central
Alps: Role of complex terrain.

*Jr. Atmos. Sci*. **61**. 2249-2266.

Lee. T.J., R.A. Pielke., R.C. Kessler., J. Weaver. 1989: Influence of
cold pools downstream of mountain barriers on downslope winds and
flushing. *Mon. Wea. Rev.* **117**. 2041-2058

Lilly, D.K.., J.B. Klemp. 1979: The effects of terrain shape on
nonlinear hydrostatic mountain waves. *Jr. Fluid Mech.* **95**. 241-261

Mass. C.F., M.D. Albright. 1985: A severe windstorm in the lee of the
Cascade Mountains of Washington State. *Mon. Wea. Rev*. **113**. 1261-
1281.

Peltier, W.R., T.L. Clark. 1979: The evolution and stability of
finite-amplitude mountain waves. Part II: Surface wave drag and severe
downslope windstorms. Jr. Atmos. Sci. **36**. 1498-1529.

Peltier. W.R., J.F. Scinocca. 1990: The origin of sever downslope
windstorm pulsations. *Jr. Atmos. Sci*. **47**. 2853-2870.

Rognvaldsson, O., J.W. Bao., H. Agustsson., H. Olafsson. 2007: Downslope
windstorm in Iceland-

WRF/MM5 model comparison-I.

Smith, R.B. 1985: On severe downslope winds. *Jr. Atmos Sci.* **42**.
2597-2603

Smith. R.B., J. Sun. 1987: Generalized hydraulic solutions pertaining to
severe downslope winds.

*Jr. Atmos. Sci*. **44**. 2934-2939.

Smith. R.B. 1990: Why can't stably stratified air rise over high ground?
*Atmospheric Processes over Complex Terrain*. Ed- W. Blumen. A.M.S.
vol.**23**. no.45. Appendix to Ch.5. 105-107.

Smith, R.B., S. Skubis., J.D. Doyle., A.S. Broad., C.Kiemle., H.
Volkert. 2002: Mountain waves over Mount Blanc: Influence of a stagnant
boundary layer. *Jr. Atmos. Sci*. **59**. 2073-2092.

Smith, R.B., J.D. Doyle, Q. Jiang., S.A. Smith. 2007: Alpine gravity
waves: lessons from MAP regarding mountain wave generation and breaking.
*Q. J. R. Meteorol. Soc.* **133**. 917-936.

![](media/image1.jpeg){width="5.996527777777778in"
height="4.631944444444445in"}

*Figure 1: Southcentral Alaska topography.*

![](media/image2.jpeg){width="5.993055555555555in" height="6.05in"}

*Figure 2: Anchorage soundings during or close to period of maximum
Hillside*

*winds for the Big Four events.*

![](media/image3.jpeg){width="5.997916666666667in"
height="5.647916666666666in"}

*Figure 3: Anchorage raob winds during or near the period of peak
Hillside*

*winds for the Big Four events.*

![](media/image4.wmf){width="5.997222222222222in"
height="3.036111111111111in"}

*Figure 4: MSLP pressure difference (mb) between Cordova (PACV) and*

*Anchorage (PANC) for the Big Four wind events. X-axis interval is one*

*hour intervals with relative scaling.*

![](media/image5.jpeg){width="5.995833333333334in"
height="5.252083333333333in"}

*Figure 5: North-South plots of potential temperature (ci= 1^o^ C,
labels= 4^o^ C)*

*through 149.5^o^W. This cuts across the mid-Hillside from 60.1^o^N to
61.2^o^N.*

*Notice the mountain wave structure in the April 1, 1980 event. These
plots*

*were captured near the time of maximum model winds.*

![](media/image6.jpeg){width="5.998611111111111in"
height="5.002777777777778in"}

*Figure 6: East-West plots of potential temperature (ci= 1^o^ C, labels=
4^o^ C)*

*through 61.15^o^N. This cuts across the Hillside and through west
Anchorage.*

*Notice the mountain wave structure in the April 1, 1980 event.*

![](media/image7.jpeg){width="5.997916666666667in" height="5.29375in"}

*Figure 7: North-South plots of vector winds (ci= 1m/s, labels= 2 m/s)*

*through 149.5^o^W. This cuts across the mid-Hillside from 60.1^o^N to
61.2^o^N.*

![](media/image8.jpeg){width="5.996527777777778in"
height="4.948611111111111in"}

*Figure 8: East-West plots of vector winds (ci= 1 m/s, labels= 2 m/s)
through 61.15^o^N. This cuts across the Hillside and through west
Anchorage.*

> ![](media/image9.jpeg){width="4.125in" height="6.356944444444444in"}

*Figure 9: East-West cross-section of vertical velocity*

*(ci=10 Pa/s) through the Hillside and west Anchorage for*

*Oct 9 & 10, 1986. ( 10 Pa/s \~1 m/s)*

![](media/image10.jpeg){width="5.9944444444444445in"
height="4.561805555555556in"}

*Figure 10: North-South cross-section of potential temperatures (ci=
1^o^ C, labels= 4^o^ C) through 149.8^o^W (east Anchorage) on November
27, 1985 at 00Z. White line at 600 mb represents approximately of
environmental critical layer.*

![](media/image11.jpeg){width="5.999305555555556in"
height="4.779861111111111in"}

*Figure 11: North-South cross-section through 149.8^o^W (east Anchorage)
for*

*April 1, 1980 at 05Z (ci= 1^o^ C, labels= 4^o^ C). Black lines
represent estimation*

*of mountain wave phase lines.*

![](media/image12.jpeg){width="5.194444444444445in"
height="7.740277777777778in"}

*Figure 12: North-South cross-section 149.8^o^W (east Anchorage) on*

A)  *4-km grid and B) 1-km grid for April 1, 1980 at 10Z (ci= 1^o^ C,*

> *labels= 4^o^ C).*

![](media/image13.jpeg){width="5.996527777777778in"
height="4.496527777777778in"}

*Figure 13: Plot of streamlines on April 1,1980 at 09Z showing some eddy
structures (E). Anchorage International Airport is labeled with orange
dot. Coastline is outlined in white.*

![](media/image14.jpeg){width="5.998611111111111in"
height="4.950694444444444in"}

*Figure 14: East-West cross-sections through 61.15^o^N for Feb 4 & 5,
2003 (across the Hillside and through west Anchorage). Note changes in
mountain wave structure from one panel to the next as a result of
changes in the stability and height of the environmental critical layer.
Features labeled with a HJ are hydraulic jumps.*

> ![](media/image15.jpeg){width="5.998611111111111in"
> height="6.197916666666667in"}

*Figure 15: Upper chart is 1-sec wind data (m/s) from GAHA2 on January
22, 2008 at 08Z. Red line is 60-sec centered running mean, green line
indicates period expanded in the bottom chart.*