Influences of the Sierra Nevada on Intermountain Cold-Front Evolution
Gregory L. West
Department of Earth and Ocean Sciences, University of British Columbia and BC Hydro
Corporation, Vancouver, BC, Canada
and
W. James Steenburgh
Department of Atmospheric Sciences, University of Utah, Salt Lake City, UT
Submitted to Monthly Weather Review
December 10, 2010
Corresponding author address: Dr. W. James Steenburgh, Department of Atmospheric Sciences,
University of Utah, 135 South 1460 East Room 819, Salt Lake City, UT, 84112.
E-mail: jim.steenburgh@utah.edu
Abstract
Recent studies indicate that strong cold fronts develop frequently downstream of the
Sierra Nevada over the Intermountain West. To help ascertain why, this paper examines the
influence of the Sierra Nevada on the rapidly developing Intermountain cold front of 25 March
2006. Comparison of a Weather Research and Forecasting (WRF) model control simulation with
a simulation in which the height of the Sierra Nevada is restricted to 1500 m (roughly the
elevation of the valleys and basins of the Intermountain West) shows that the interaction of
southwesterly pre-frontal flow with the formidable southern "High Sierra" produces a leeward
orographic warm anomaly that enhances the cross-front temperature contrast. Several processes
generate this orographic warm anomaly including blocking of the penetration of low-level
Pacific air into the Intermountain West, airmass transformation, and indirect downstream effects
arising from airmass transformation. The latter includes increased pre-frontal sensible heating of
the boundary layer and diminished diabatic cooling from precipitation within the cloud and
precipitation shadow. In contrast, the post-frontal airmass experiences comparatively little
orographic modification as it moves across the relatively low northern Sierra Nevada. The direct
and indirect effects of the Sierra Nevada also contribute to a stronger frontal collapse, although
the role of indirect diabatic effects is likely maximized in this event since the front is moving
across the Intermountain West during the day. These results illustrate how the Sierra Nevada can
enhance cold-front development and contribute to the high frequency of strong cold-frontal
passages over the Intermountain West.
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1. Introduction
Frontal interactions with orography have been documented ever since Bjerknes and
Solberg (1922) described the first frontal-cyclone model. The topographically complex western
United States is one region where mountains play a major role in frontal evolution (e.g., Braun et
al. 1997; Colle et al. 1999; Steenburgh and Blazek 2001; Colle et al. 2002; Shafer et al. 2006;
Shafer and Steenburgh 2008; Steenburgh et al. 2009). As a cold front approaches the Pacific
coast, orographic blocking and friction produce enhanced prefrontal southerly flow, confluent
deformation, frontogenesis, and frontal deceleration (e.g., Braun et al. 1997; Doyle 1997; Colle
et al 1999; Yu and Smull 2000). Colle et al. (2002) illustrate these topographic effects with a
model sensitivity study in which coastal topography is removed, which results in a weaker, more
progressive cold front.
Further downstream, recent observational and modeling studies document several
manifestations of front-mountain interactions over the Intermountain West (see Fig. 1 for
geographic references). These interactions include orographic blocking and frontal retardation
windward of Sierra Nevada (Steenburgh and Blazek 2001; Shafer et al. 2006), discrete frontal
propagation across the Sierra-Cascade ranges and Intermountain West (Steenburgh et al. 2009),
and frontal distortions produced by basin-and-range and other topographic geometries
(Steenburgh and Blazek 2001; West and Steenburgh 2010). In addition, West and Steenburgh
(2010) describe how the Great Basin Confluence Zone (GBCZ), an area of confluent
deformation and convergence that extends downstream from the Sierra Nevada, can serve as a
locus for frontal development.
None of these studies have used numerical sensitivity experiments to examine the role of
the Sierra Nevada in Intermountain cold-front evolution. Oriented from north-northwest to
2
south-southeast, the Sierra Nevada occupy most of eastern California and join the southern
Cascade Mountains to the north (Fig. 1). The crest of the Cascade-Sierra ranges is relatively low
north of Lake Tahoe, but reaches altitudes of more than 3500 m in the southern "High Sierra"
south of Lake Tahoe (inset, Fig. 1).
Recent studies by Shafer and Steenburgh (2008),
Steenburgh et al. (2009), and West and Steenburgh (2010) suggest that flow interactions with the
Sierra Nevada, especially the High Sierra, contribute to the development of the GBCZ, frontal
intensification over Nevada, and the increasing frequency of strong cold-frontal passages as one
moves eastward across the Intermountain West.
Here we examine how the Sierra Nevada contribute to the development and strength of
the Intermountain cold front of 25 Mar 2006, which propagated discretely across the northern
Sierra Nevada and southern Cascade Mountains (see Steenburgh et al. 2009 for a detailed
analysis), intensified rapidly over Nevada (summarized in Fig. 2), and produced the sixth largest
temperature change in the 25-year cold-front climatology of Shafer and Steenburgh (2008).
Specifically, we compare a full-terrain control simulation by the Weather Research and
Forecasting (WRF) model with a simulation in which the height of the Sierra Nevada is
restricted to 1500 m, roughly the elevation of the valleys and basins of the Intermountain West.
Analysis of the two simulations shows that the interaction of pre-frontal southwesterly flow with
the southern High Sierra produces a leeward orographic warm anomaly that enhances the crossfront temperature contrast and contributes to a stronger frontal collapse.
Discrete frontal
propagation occurs, however, in both simulations and thus appears not to be a consequence of
troughing downstream of the Sierra Nevada as hypothesized by Steenburgh et al. (2009).
3
2. Data and Methods
All simulations of the 25 Mar 2006 cold front use the Advanced Research core of the
Weather Research and Forecasting (WRF) model version 2.2.1 (Skamarock et al. 2005). The full
terrain control simulation (FULLTER), is identical to the control (CTL) run described by
Steenburgh et al. (2009), and features a 36-km outer domain, 12-km inner nest (the only domain
presented in this paper), 34 half-η levels in the vertical, and unaltered topography (Fig. 3a).
Physics packages include the Rapid Radiative Transfer Model (RRTM) longwave radiation
scheme (Mlawer et al. 1997), the Dudhia shortwave radiation scheme (Dudhia 1989), the Noah
land surface model (Chen and Dudhia 2001), the Mellor–Yamada–Janjić planetary boundary
layer parameterization (Mellor and Yamada 1982; Janjić 2002), the new Kain–Fritsch cumulus
parameterization [a modified version of the scheme described by Kain and Fritsch (1990, 1993)],
and Thompson et al. (2004, 2006) cloud microphysics parameterization. North American
Mesoscale Model (NAM) analyses provide the cold-start atmospheric and land surface (e.g., soil
temperature, soil moisture, snow cover) initial conditions at 0000 UTC 25 March 2006 and
lateral boundary conditions through the integration period, with some modifications to the lowertropospheric analysis in complex terrain and the land-surface initialization as described in
Steenburgh et al. (2009).
The NOSIERRA simulation is identical to FULLTER except that we restrict the height of
the Sierra Nevada and southern Cascades of California to an elevation of 1500 m, the
approximate mean elevation of the valleys and basins of the Intermountain West (Fig. 3b). The
resulting terrain height differences are relatively small over the northern Sierra Nevada and
southern Cascades, but much larger south of Lake Tahoe along the High Sierra. The atmosphere
in the volume previously occupied by topography derives from North American Mesoscale
4
(NAM) model initial analyses or, at levels below the NAM model surface, assumes a dryadiabatic lapse rate and uses winds from the lowest above ground model level. Given the small
volume of topography removed and the 15-h integration time before incipient frontal
development, results are likely insensitive to these prescribed initial conditions.
Fake-dry
(FKDRY) simulations based on FULLTER or NOSIERRA topography do not include diabatic
heating and cooling associated with cloud and precipitation processes. They do, however, allow
simulated clouds and precipitation to interact with other model physics packages, such as the
radiation scheme.
For figure clarity, all horizontal contour and color-fill analyses of geopotential height,
potential temperature, potential-temperature gradient magnitude, frontogenesis, and potentialtemperature difference between simulations are smoothed using a 7-point cowbell spectral filter
(Barnes et al. 1996). This enables a clearer presentation without eliminating the mesoscale
terrain signal. The 850-hPa geopotential height analysis is based on hydrostatic extrapolation
where that pressure level is below the model terrain. Analyses may differ slightly from those in
Steenburgh et al. (2009) due to the minor correction of the spatial differencing algorithm.
We use several diagnostic quantities to examine the mechanisms responsible for frontal
development.
As in Steenburgh et al. (2009), surface frontogenesis is defined following
Petterssen (1936) and Miller (1948) as
F
d
 ,
dt 
(1)
where

d




 u
v
 η ,
dt t
xη
yη
η
5
(2)
  i


j
,
x η
y η
(3)
.
the subscript η denotes differentiation along the terrain following lowest η level, and η is the ηcoordinate vertical velocity. Following Miller (1948), Eq. (1) may be written as
FFW FT FD,
(4)
where
FW  
1    u  u     v  v  


,


 
 η  x  x x y y  y  x x y y 
FT  
FD  
   η  η  
,

 
   x x y y 
(6)
   d   d 


,
 x x d t y y d t
(7)
1
 η
1
η
(5)
and the subscript η has been dropped for convenience. FW, FT, and FD are the frontogenesis
components produced by horizontal deformation and divergence (hereafter kinematic
frontogenesis), tilting, and horizontal gradients in diabatic heating and cooling (hereafter
diabatic frontogenesis), respectively. Although FT is non-zero due to the presence of a surfacebased stable layer in the morning and a super-adiabatic layer in the afternoon, it does not appear
to contribute significantly to frontal development and is not presented. FD contains two
components
FD  FM  FBL ,
(8)
where FM is the diabatic frontogenesis produced by the WRF-model cloud microphysics and
6
cumulus parameterizations (hereafter moist frontogenesis) and FBL is the diabatic frontogenesis
produced by the boundary layer and radiation parameterizations (hereafter boundary layer
frontogenesis). We calculate FM and FBL from heating rates obtained directly from the WRFmodel parameterizations.
3. Results
Steenburgh et al. (2009) provide a thorough analysis of the 25 Mar 2006 case, including
validation of FULLTER (their CTL). Here we concentrate on the influence of the Sierra Nevada
by comparing FULLTER and NOSIERRA.
a. Antecedent conditions
At 1500 UTC 25 March 2006 a weakening occluded front moves inland across the
northern California coast, its position virtually identical in FULLTER and NOSIERRA (cf. Figs.
4a,b). Ahead of the occluded front, confluent south–southwesterly large-scale flow develops over
northwest Nevada, initiating Intermountain cold-front development. In FULLTER, the Sierra
Nevada disrupt the confluent flow, producing two weak troughs (dashed lines) and an airstream
boundary (dotted line) that separates south–southeasterly flow over southern and central Nevada
from southwesterly flow over northwest Nevada (Fig. 4c). In contrast, NOSIERRA produces a
single trough and wind shift (Fig. 4d). FULLTER also generates more spatial variability in
potential temperature over northwest Nevada (cf. Figs. 4c,d), although differences between the
two simulations are less than 2 K (Fig. 5a).
More substantive differences are found further south where the Sierra Nevada are higher
and the crest-height difference between the two simulations is larger (cf. Figs. 3a,b).
7
In
particular, FULLTER produces stronger windward ridging and lee troughing across the High
Sierra (cf. Figs 4a,b), southeasterly–southerly rather than southerly–southwesterly flow
downstream of the High Sierra (cf. Figs. 4c,d), and a broad region of higher (2–5 K) potential
temperature over central Nevada that we refer to as the orographic warm anomaly (Fig. 5a). The
aforementioned airstream boundary (dotted line) lies near the northern edge of the southeasterly–
southerly flow and the orographic warm anomaly, which develops early in the simulation and
strengthens and spreads outward across the Intermountain West from the lee of the High Sierra.
Fifteen-hour (0000–1500 UTC) three-dimensional trajectories illustrate how flow
interaction with the Sierra Nevada contributes to the development of the orographic warm
anomaly at 1500 UTC (Fig. 6). Trajectory calculations follow Petterssen (1956, p. 27) and
Seaman (1987), use three-dimensional WRF-model horizontal winds and vertical velocities at
15-min intervals, and are constrained to remain on the lowest-η level if they approach the model
surface. A dense trajectory grid encompassing the Sierra Nevada was examined, but for clarity
we present selected forward trajectories that begin at 850 hPa in a line parallel to and upstream
of the Sierra Nevada (Group X, brown) and selected backwards trajectories that end at 850 hPa
(or the lowest-η level if the model terrain rises above 850 hPa) along two lines parallel to and
downstream of the Sierra Nevada (Group Y and Z, green and light green, respectively).
The
FULLTER and NOSIERRA forward (backward) trajectories have the same Earth-relative
beginning (ending) location.
The Group X trajectories exhibit diffluence in both simulations, but in FULLTER split
more strongly and circumscribe the High Sierra to the south and the north (cf. Figs. 6a,b). The
pronounced influence of the High Sierra is illustrated well by trajectory 5, which is blocked and
deflected southward in FULLTER, but penetrates directly into the Intermountain West in
8
NOSIERRA. The FULLTER Group Y and Z trajectories either curve around the southern
periphery of the Sierra Nevada and then move northward along the barrier, or traverse the Sierra
Nevada and subside to their lee (Fig. 6a). In contrast, the NOSIERRA Group Y and Z trajectories
penetrate directly into the Intermountain West (Fig. 6b).
Selected Group Y and Z trajectories (labeled 1–4 in Fig. 6) illustrate how the differing
origins of the FULLTER and NOSIERRA trajectories contribute to the orographic warm pool
development. FULLTER trajectories 1 and 2 begin near 700 hPa and subside abruptly in the lee
of the Sierra Nevada at ~1000 and ~0600 UTC, respectively (Fig. 7a). In NOSIERRA, however,
trajectories 1 and 2 begin below 900 hPa, ascend over the lower windward slopes of the Sierra
Nevada, and penetrate into the Intermountain West without experiencing subsidence. Since
FULLTER and NOSIERRA trajectories 1 and 2 experience similar net cooling, the greater initial
altitude (and hence potential temperature) of FULLTER trajectories 1 and 2 accounts for most of
the difference in potential temperature in the northern portion of the orographic warm pool at
1500 UTC.
In the southern portion of the orographic warm pool, FULLTER trajectories 3 and 4
originate south of the Sierra Nevada over the Mohave Desert, whereas NOSIERRA trajectories 3
and 4 originate over the Central Valley of California (cf. Figs. 6a,b).
This difference in
geographic origin (and hence potential temperature) explains more than half of the final potential
temperature difference between FULLTER and NOSIERA (Fig. 7b).
Comparison of FULLTER and NOSIERRA cross sections taken across the High Sierra
and central Nevada further elucidates the structure and development of the orographic warm
anomaly at 1500 UTC (Fig. 8, see line-segment XY in Figs. 4c,d and 5 for cross-section
location). The orographic warm pool is deepest near the Sierra Nevada, but extends several
9
hundred km downstream below ~650 hPa (Fig. 8c).
Consistent with the above trajectory
analysis, the High Sierra produce a high-amplitude mountain wave with strong lee side
subsidence in FULLTER, whereas cooler Pacific air penetrates directly into the Intermountain
West in NOSIERRA (cf. Figs. 8a,b; compare the fate of the 292–296 K air). As a result, a
pronounced low-level cloud and precipitation shadow extends downstream of the High Sierra in
FULLTER, whereas scattered precipitation and related diabatic effects, including localized lowlevel diabatic cooling, lie over the Intermountain West in NOSIERRA (cf. Figs. 4a,b and Figs.
8a,b).
These cloud and precipitation differences represent a secondary contributor to the
orographic warm pool development and explain some of the variability in potential temperature
change along FULLTER and NOSIERRA trajectories 3 and 4 (Fig. 7b).
For example,
NOSIERRA trajectory 4 experiences a 4K decrease in potential temperature from 1300–1500
UTC as it moves into an area of diabatic cooling beneath the southernmost area of precipitation
in southern Nevada.
In addition, the air penetrating across the High Sierra in FULLTER does so within a more
substantive orographic cloud and precipitation region compared to NOSIERRA (cf. Figs. 4a,b
and 8a,b). This orographic cloud and precipitation region developed over the southern High
Sierra after 0900 UTC and subsequently contributes to airmass transformation [i.e., the acrossbarrier change in water vapor concentration and potential temperature due to orographic
precipitation, net latent heating, and the scrambling and reordering of air parcels vertically as
they seek their equilibrium level downstream of a mountain barrier (e.g., Varney 1920; Giorgi
and Bates 1989; Sinclair 1994; Smith et al. 2003; Smith et al. 2005)] in FULLTER.
Also during this period, confluence and convergence begin to concentrate the baroclinity
over northwest Nevada. Kinematic frontogenesis over northwest Nevada is broken into three
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bands along the aforementioned troughs and airstream boundary in FULLTER, but found solely
along the single trough in NOSIERRA (cf. Figs 4c,d). The baroclinity is also less organized and
concentrated in FULLTER compared to NOSIERRA. Nevertheless, because of the orographic
warm anomaly (Fig. 5a), a larger total temperature contrast exists across northwestern Nevada in
FULLTER (288-298K) than NOSIERRA (288-292K). At this time, diabatic frontogenesis is
weak and disorganized over northwest Nevada in FULLTER and weak but generally positive
near the trough in NOSIERRA (Figs. 4e,f).
b. Frontal development
By 1800 UTC the airstream boundary and two troughs in FULLTER merge into a single
intensifying trough and cold front, whereas in NOSIERRA the nascent cold front simply forms
along the solitary, pre-existing trough (Figs. 9a,b). In both simulations, a coherent kinematic
frontogenesis maximum lies along the frontal zone (cf. Figs. 9c,d) and the new cold front forms
well in advance of the remnants of the former occluded front. Thus, discrete propagation occurs
with or without the upper portion (>1500 m) of the Sierra Nevada. A larger cross-front potential
temperature contrast exists, however, in FULLTER because of the orographic warm anomaly,
which has strengthened and spread across central Nevada (Fig. 5b). In contrast, there is no major
difference between the simulations behind the frontal zone since the post-frontal airmass
traverses the relatively low northern Sierra and southern Cascades north of Lake Tahoe where
the crest-height difference between FULLTER and NOSIERRA is < 500 m (cf. Figs. 3a,b).
Although both simulations produce post-frontal precipitation, the direct effect of diabatic
frontogenesis remains relatively weak within FULLTER frontal zone and is not coherently
organized in NOSIERRA (Figs. 9e,f).
11
Cross sections from FULLTER and NOSIERRA further illustrate the influence of the
Sierra Nevada on the structure of the orographic warm pool at this time (Fig. 10). In FULLTER,
a substantive orographic cloud and an associated condensational heating maximum persist over
the windward slope of the Sierra Nevada, with downslope flow and a narrower region of
evaporative cooling to the immediate lee (Fig. 10a). The downstream Intermountain West is
largely cloud free with a well-mixed, ~100 mb deep convective boundary layer. In contrast,
NOSIERRA produces widespread cloud cover, localized areas of near surface evaporative
cooling (due to precipitation), and a shallower convective boundary layer (Fig. 10b). Surface
sensible heat are 50–150 W m-2 greater in FULLTER than NOSIERRA across much of the
region downstream of the Sierra Nevada (Fig. 11a). These contrasts in diabatic heating and
cooling, combined with advection in the southwesterly large scale flow, result in a lower
tropospheric warm anomaly that extends into northeast Nevada and through much of the cross
section (Figs. 5b, 10c).
A strong cold front extends across Nevada to the Sierra Nevada in both simulations by
2100 UTC (Figs. 11a,b). A coherent band of strong kinematic frontogenesis lies along the front
in both simulations, with only a weak break in northern Nevada (Figs 11c,d). The stronger
cross-front temperature contrast persists in FULLTER due to the orographic warm anomaly,
which has continued to spread and strengthen (Figs. 5c, 12c). At this time, post-frontal
precipitation lags the surface cold front and the diabatic frontogenesis is generally weakly
negative (i.e., frontolytical) within the frontal zone in both simulations (Figs. 11e,f).
During this period, the orographic warm anomaly continues to deepen and spread in the
southwesterly prefrontal flow (Fig. 13c). Cloud cover and precipitation remain limited over the
Intermountain West in FULLTER, but widespread in NOSIERRA (cf., Figs. 13a,b). Stronger
12
sensible heating over southern Nevada (Fig. 11b) continues to contribute to a deeper convective
boundary layer in FULLTER that is advected northweastward into northeast Nevada and
northern Utah. Comparison of soundings taken near the center of the orographic warm anomaly
(point A, Figs. 5c and 11c,d) confirms the warmer conditions in FULLTER below 600 hPa, and
deeper, drier surface-based convective boundary layer (Fig. 13).
In summary, several interrelated processes contribute to the development and
intensification of the orographic warm anomaly, including:

Blocking of the penetration of Pacific air into the Intermountain West by the High
Sierra;

Subsidence in the lee of the High Sierra;

Airmass transformation;

Increased sensible heating and decreased near-surface diabatic cooling within the
downstream cloud and precipitation shadow.
We now use trajectory analysis to further quantify these effects and the development of the
orogographic warm anomaly and cold front.
c. Trajectory analysis of the mature orographic warm anomaly and cold-front development
Nine-hour three-dimensional backwards trajectories that begin at 1200 UTC and end at
2100 UTC provide a Lagrangian perspective on the development of the orographic warm
anomaly and Intermountain cold front (Fig. 10). A dense grid of trajectories encompassing the
FULLTER and NOSIERRA cold fronts was examined, but for clarity we present a grid that is
one fourth as dense, with trajectories grouped based ending location and trajectory history.
FULLTER and NOSIERRA trajectories with corresponding numbers have the same Earth-
13
relative termination point.
Group A trajectories in both simulations (dark blue, Fig. 10) originate over north-central
California, cross the northern Sierra Nevada, and terminate in the postfrontal airmass. Compared
with NOSIERRA, the FULLTER trajectories are influenced by a stronger frontal trough,
experience greater deflection, and approach the front more directly (cf. Figs. 10a,b). Most
FULLTER and NOSIERRA trajectories that terminate at a common location have similar
beginning and ending potential temperatures (i.e., within 2.5 K) (Table 1), indicating little
difference in postfrontal airmass modification between the two simulations. Exceptions include
trajectories 9 and 10, which terminate in the vicinity of the post-frontal precipitation band (cf.
Figs. 7a, 7b, 10) and experience different diabatic heating and cooling histories due to contrasts
in band position, and intensity between the two simulations.
Group B trajectories in both simulations (light blue, Fig. 10) originate near the southern
periphery of the Sierra Nevada and terminate in the western portion of the prefrontal airmass.
The FULLTER trajectories are more strongly influenced by the High Sierra and experience
greater deflection and are more elongated due to stronger flow on the southern periphery of the
range (cf. Figs. 10a,b). Downstream of the High Sierra, the FULLTER trajectories also approach
the front more directly, consistent with the stronger lee and frontal troughs. The increase in
potential temperature along several of these trajectories (19, 23, 24, 27, 30) is more than 3K
greater in FULLTER than NOSIERRA, which reflects decreased cloud cover and increased solar
heating. The only trajectory with a smaller increase in potential temperature in FULLTER than
NOSIERRA is trajectory 15, which terminates very near the frontal surface.
Group C trajectories in FULLTER (magenta, Fig. 10a) originate near the Pacific coast
and are forced up the windward slopes of the High Sierra, undergoing diabatic heating within the
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orographic cloud.
They then descend rapidly into the lee of the High Sierra where they
experience boundary layer heating and terminate in the prefrontal airmass. As a result, the
potential temperature along the two Group C trajectories (18 and 26) increases 12.7 and 10.3 K,
respectively (Table 1). In NOSIERRA, these orographic effects are greatly reduced by the
elimination of the High Sierra (Fig. 10b). Instead, the NOSIERRA trajectories travel beneath the
orographic cloud over the plateau (rather than within it), experience reduced downstream
boundary layer heating due to more extensive cloud cover (e.g., cf. Figs. 8a,b), and thus observe
smaller increases in potential temperature (2.1K and 4.9K). These differences contribute to the
large magnitude of the orographic warm anomaly in the direct lee of the High Sierra.
Group D trajectories in FULLTER (yellow, Fig. 10a) originate over the lower Colorado
River Basin and terminate in the pre-frontal airmass over eastern Nevada. Surface sensible
heating is limited only by some high clouds, and the potential temperature along these
trajectories increases 2.4–7.6K (Table 1). The southernmost FULLTER trajectories (21, 25, 29,
32) are similar to those in NOSIERRA (cf. Figs. 10a,b), with comparable tracks and beginning
and ending potential temperatures. The northernmost FULLTER trajectories (6, 11, 16, 20),
however, originate over southern Nevada, whereas the comparable NOSIERRA trajectories
originate over central and southwestern Nevada (cf. Figs. 10a,b) This trajectory difference
reflects the stronger lee trough in FULLTER, which leads to a stronger pressure gradient that
rotates the flow counter clockwise. As a result, the FULLTER trajectories originate further to
the south over the lower Colorado River Basin (cf. Figs. 3a,b and 6a,b). The potential
temperature increase along these trajectories is also greater than in NOSIERRA (Table 1).
Group E trajectories in FULLTER (purple, Fig. 10a) originate near the southern
periphery of the Sierra Nevada. In contrast, the NOSIERRA trajectories originate in cooler air to
15
the north (Fig. 10b, Table 1). This largely explains the somewhat higher potential temperatures
found along the Utah-Nevada border where the Group E trajectories terminate.
In summary, these trajectories further illustrate the role of the Sierra Nevada in the
development of the orographic warm anomaly and Intermountain cold front. Specifically, the
High Sierra block the eastward penetration of low-level Pacific air, dry and warm parcels that
cross the barrier by generating orographic precipitation, and lead to enhanced surface sensible
heating and reduced evaporative cooling in the downstream cloud and precipitation shadow over
the Intermountain West. In contrast, orographic modification of the post-frontal airmass is more
limited since it traverses the relatively low northern Sierra Nevada. This leads to a larger crossfront temperature contrast than would exist if crest of the Sierra Nevada was uniform and/or the
High Sierra were not present. The trajectories also illustrate that front-relative flow in FULLTER
is stronger in both the post- and pre-frontal environments, which is consistent with a more
intense lee trough, frontal trough, and collapse, the latter of which is discussed below.
d. Frontal collapse
In both simulations the cold front reaches maximum intensity at 0000 UTC 26 March
(Figs. 11c,d) as the post-frontal precipitation band phases with the surface front (Figs. 11a,b) and
the pre-frontal convective boundary layer becomes fully developed (not explicitly shown, see
Steenburgh et al. 2009). During this period, kinematic (Figs. 11c,d) and diabatic frontogenesis
(Figs. 11e,f) become stronger and better organized within the FULLTER frontal zone than
NOSIERRA, leading to a stronger frontal collapse. In particular, the magnitude of the surface
potential temperature gradient (i.e., ||) is ~25% larger in FULLTER than NOSIERRA (cf.
Figs. 12a,b). Finally, although the orographic warm anomaly is strongest directly downstream of
16
the High Sierra where cold air has encircled the southern end of the barrier forming a seclusion
of warm air in FULLTER, a neck of the orographic warm anomaly persists just ahead of the
front (Fig. 4d). Farther downstream of the Sierra Nevada, the cold front moves into the plane of
cross-section XY in both simulations (cf. Figs. 11c,d and 13a,b). As this occurs, the orographic
warm anomaly ascends into the middle and upper troposphere and is removed from the surface
(Fig. 13c).
The stronger frontal collapse in FULLTER compared to NOSIERRA (as inferred from
the magnitude of the lowest-half-η level potential temperature gradient, (cf. Figs. 12a,b) also
occurs in FKDRY simulations that do not include the diabatic effects of cloud and precipitation
processes (cf., Figs. 8c,d). Thus, the more intense frontal collapse in FULLTER does not simply
reflect more intense sub-cloud diabatic cooling in the post-frontal environment.
Ageostrophic circulations are essential for frontal collapse, especially on such rapid time
scales (<12 h), and result from geostrophic adjustment as large-scale deformation concentrates a
horizontal temperature gradient (Hoskins and Bretherton 1972). In some cases, this frontal
collapse may be enhanced by terrain-induced flows (Volkert et al. 1991; Colle et al. 1999;
Steenburgh and Blazek 2001), differential sub-cloud evaporative and sublimational cooling
(Schultz and Trapp 2003), and differential surface heating (e.g., Koch et al. 1995, 1997; Gallus
and Segal 1999), and Segal et al. 2004).
For the 25 Mar 2006 cold front, terrain-induced flows appear to play the greatest role in
the immediate lee of the Sierra Nevada where the pressure trough in FULLTER is much stronger
than found in NOSIERRA (cf. Figs. 11a,b). This results in enhanced front-relative flow in the
pre-frontal and postfrontal environments (Fig. 4c), with the resulting deformation and
convergence producing a front that is much stronger in FULLTER than NOSIERRA (cf. 12a,b).
17
Sub-cloud evaporative and sublimational cooling contributes to frontal collapse in areas
more removed from the Sierra Nevada (i.e., over northern Nevada), where a pronounced postfrontal precipitation band develops in both simulations. In this region, simulations that do not
include the diabatic effects of cloud and precipitation processes (FULLTER-FKDRY and
NOSIERRA-FKDRY) produce weaker frontal temperature gradients than their full physics
counterparts (cf. Figs. 12a,c and Figs. 12b,d). Related cooling in the post-frontal environment is
larger in FULLTER than NOSIERRA (Figs 4c,d), which, combined with the reduced pre-frontal
cloud and precipitation, suggests that the indirect influence of the Sierra Nevada is to enhance
the differential diabatic cooling and its influence on frontal development in this case.
Finally, the frontal collapse is strongly influenced by sensible heating of the pre-frontal
environment, which reinforces the evaporative and sublimational cooling effects above and is
greater in FULLTER because of the cloud and precipitation shadow downstream of the High
Sierra. Although this ultimately leads to stronger diabatic frontogenesis (FD) in FULLTER
compared to NOSIERRA (cf. Figs. 7e,f), it likely also contributes to the stronger kinematic
frontogenesis (FW) since differential heating arising from inhomogeneous cloud cover is known
to generate a thermally forced circulation that enhances the cross-front ageostrophic circulation
(e.g., Koch et al. 1995, 1997, Gallus and Segal 1999, and Segal et al. 2004). Thus, the influence
of the Sierra Nevada on this event extends beyond their direct influence on wind and temperature
and includes their indirect downstream influences on clouds, precipitation, sensible heating, and
subsequent frontal collapse.
4. Conclusions
Recent studies show that the Intermountain West is a common breeding ground for
18
strong, rapidly developing cold fronts. Numerical simulations of the 25 Mar 2006 Intermountain
cold front event described here illustrate one way that flow interaction with the Sierra Nevada
can enhance cold-front development and potentially contribute to the high frequency of strong
cold-frontal passages over the Intermountain West. Specifically, a comparison of a full-terrain
simulation (FULLTER) with one in which the height of the Sierra Nevada and southern
Cascades of California is restricted to no more than 1500 m (NOSIERRA) shows that the
interaction of southwesterly flow with the formidable southern High Sierra creates a pre-frontal
orographic warm anomaly that enhances the cross-front temperature contrast. Several processes
generate this orographic warm anomaly including the blocking of Pacific air by the High Sierra,
airmass transformation (i.e., the across-barrier change in water vapor concentration and potential
temperature due to orographic precipitation, related diabatic processes, and airmass scrambling),
and subsequent surface sensible heating of the largely cloud-free downstream airmass. The
direct and indirect effects of the Sierra Nevada also contribute to a stronger frontal collapse,
although the role of pre-frontal sensible heating within the downstream cloud and precipitation
shadow is likely maximized in this event since the front is moving across the Intermountain West
during the day. This heating would not occur at night, when radiative cooling might prove
frontolytical.
This study adds to our growing understanding of the influence of the Sierra Nevada on
the Intermountain cyclone and front evolution. Although it is clear that the Sierra Nevada
enhance frontal development and collapse, the underlying mechanisms responsible for the
development of pressure troughing, confluence and convergence, and frontogenesis over the
Intermountain West in this and other cases still need to be fully elucidated. West and Steenburgh
(2010) show that a region of confluent deformation and convergence downstream of the High
19
Sierra, which they dubbed the Great Basin Confluence Zone (GBCZ), became a locus for
frontogenesis during the 15 Apr 2002 Tax Day Storm. While a coherent prefrontal GBCZ is not
evident in 25 Mar 2006 event examined here, confluent deformation does develop downstream
of the Sierra Nevada (e.g., Fig. 3c) and terrain-induced confluence and convergence contribute to
frontal development and collapse over the portion of the front in the direct lee of the Sierra
Nevada. Further, although Steenburgh et al. (2009) show that a moderately strong front can
develop in the absence of cloud diabatic processes in this case, removal of both cloud diabiatic
effects and the High Sierra produces a dramatically weaker front. This illustrates the need to
better understand the complex nonlinear interactions at play during Intermountain frontogenesis.
Acknowledgements. We thank Trevor Alcott, Lance Bosart, Larry Dunn, John Horel, Jan
Paegle, and Ed Zipser for their suggestions and comments, and gratefully acknowledge the
provision of datasets, software, or computer time and services by NCEP, NCAR, Unidata, the
Center for Ocean–Land–Atmosphere Studies, and the University of Utah Center for High
Performance Computing. This research was supported by National Science Foundation grant
ATM-0627937 and the National Weather Service C-STAR program. Any opinions, findings, and
conclusions or recommendations expressed in this material are those of the authors and do not
necessarily reflect the views of the National Science Foundation or the National Weather
Service.
20
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24
Tables
Table 1. Beginning, ending, and change in potential temperature (, K) for trajectories plotted in
Fig. 10.
Trajectory
Group / #
FULLTER
NOSIERRA
FULLTER
NOSIERRA
FULLTER
NOSIERRA
12 UTC 
12 UTC 
21 UTC 
21 UTC 


A/1
286.4
286.1
291.3
289.8
4.9
3.7
A/2
286.4
286.4
293.2
292.7
6.8
6.3
A/3
289.3
287.6
291.4
293.9
2.1
6.3
A/4
291.1
289.9
292.6
292.4
1.5
2.5
A/7
286
285.2
290.5
289.7
4.5
4.5
A/8
286.6
286.2
293.6
292.6
7
6.4
A/9
289.1
288.3
294.4
291.2
5.3
2.9
A / 10
296.3
289.4
298
295
1.7
5.6
A / 13
284.5
286.8
295
292.8
10.5
6
A / 14
286.9
287.9
294.7
292.9
7.8
5
B/5
299.6
297.8
302.9
300.3
3.3
2.5
B / 15
300
294.5
304.9
301.2
4.9
6.7
B / 19
298
297
305.5
300.2
7.5
3.2
B / 22
299
293.2
305.8
297.9
6.8
4.7
B / 23
294.9
299.4
304.7
298.8
9.8
-0.6
B / 24
299.3
299.4
304.7
298.8
4
0.3
B / 27
293.8
293.5
302.9
299.5
9.1
6
B / 28
298.9
297.8
302.7
300.8
3.8
3
B / 30
292.2
296.1
301.2
300.4
9
4.3
B / 31
295.6
296.5
302.2
301.7
6.6
5.2
C / 18
290.9
294.4
303.6
296.5
12.7
2.1
C / 26
294.7
293.2
305
298.1
10.3
4.9
D/6
300
298.5
302.8
298.9
2.8
0.4
25
D / 11
298.1
299.3
304
301
5.9
1.7
D / 16
296.6
297.1
303.6
298.4
7
1.3
D / 20
296
296.8
303.6
298.1
7.6
1.3
D / 21
298.3
297.4
300.7
299.6
2.4
2.2
D / 25
298.8
297.4
301.4
299.7
2.6
2.3
D / 29
297.6
297
300.5
300.7
2.9
3.7
D / 32
298.4
298
302.1
302.1
3.7
4.1
E / 12
302.1
298.9
303.9
301.8
1.8
2.9
E / 17
300.9
299.6
302.1
299.7
1.2
0.1
26
Figure Captions
Figure 1.
30-sec mean topography (m, shaded following scale at upper left) of the
Intermountain West and adjoining regions, with the height of Sierra Nevada crest
between points N and S at lower left (position of Lake Tahoe indicated by vertical line).
Abbreviations used for Lake Tahoe (LT) and the lower Colorado River Basin (LCRB).
Figure 2.
WRF topography [m, shaded following scale in (a)] for a subset of the (a)
FULLTER and (b) NOSIERRA 12-km nested domain.
Figure 3.
WRF-model analyses for 1500 UTC 25 Mar 2006. (a) FULLTER radar
reflectivity [dBZ, color shaded according to scale in (a)], cloud-top temperature (°C, grey
shaded according to scale in (b)], and 850-hPa geopotential height (solid contours every
10 m). (b) Same as (a) but for NOSIERRA. (c) FULLTER lowest half-η level potential
temperature (every 2 K), wind (vector scale at left), and kinematic frontogenesis [K (100
km h)−1, shaded following scale at left]. (d) Same as (c), but for NOSIERRA. (e) Same as
(c), but with diabatic frontogenesis. (f) As in (e) but for NOSIERRA.
Figure 4.
FULLTER lowest half-η level potential temperature (every 2 K), FULLTER-
NOSIERRA lowest half-η level potential temperature difference [K, shaded according to
scale in (a)], and FULLTER-NOSIERRA lowest half-η level vector wind difference
[scale in (a)] at (a) 1500 UTC 25 Mar, (b) 1800 UTC 25 Mar, (c) 2100 UTC 25 Mar, and
(d) 0000 UTC 26 Mar 2006.
Figure 5.
Cross sections along line XY of Figs. 3, 4, 6, and 7 at 1500 UTC 25 Mar 2006.
(a) FULLTER potential temperature (contours every 2K), total cloud water and ice
mixing ratio (g kg-1, shaded following inset scale), and vectors of along-section wind and
pressure vertical velocity (following inset scale). (b) Same as (a) except for NOSIERRA.
27
(c) FULLTER potential temperature [as in (a)] and FULLTER-NOSIERRA potential
temperature difference (K, shaded following inset scale). Shading over Sierra Nevada
indicates region where differences in the height of the half-η surfaces contribute to the
anomaly and introduce a false positive bias.
Figure 6.
Same as Fig. 3 except for 1800 UTC 25 Mar 2006.
Figure 7.
Same as Fig. 3 except for 2100 UTC 25 Mar 2006.
Figure 8.
Same as Fig. 5 except for 2100 UTC 25 Mar 2006.
Figure 9.
FULLTER (red) and NOSIERRA (blue) skew T–log p diagram (temperature and
dewpoint) for point A (see Figs. 4c and 7c,d for location) at 2100 UTC 25 Mar 2006.
Figure 10.
Three-dimensional nine-hour trajectories ending at 2100 UTC. (a) FULLTER and
(b) NOSIERRA.
Trajectory number indicated to right of trajectory ending arrow.
FULLTER-NOSIERRA lowest half-η level potential temperature difference (color filled
following inset scale) and terrain difference shaded (0–1000m gray, >1000m black).
Figure 11.
Same as Fig. 3 except for 0000 UTC 26 Mar 2006.
Figure 12.
Lowest half-η level potential temperature (thin contours every 2 K), potential
temperature gradient magnitude [thick contours every 5 K (100 km)-1 beginning with 10
K (100 km)-1], wind [vector scale at upper left of (a)], and kinematic frontogenesis [K
(100 km h)−1, shaded following scale in Fig. 4c] from (a) FULLTER, (b) NOSIERRA, (c)
FULLTER-FKDRY, and (d) NOSIERRA-FKDRY at 0000 UTC 26 Mar 2006.
Figure 13.
Same as Fig. 5 except for 0000 UTC 26 Mar 2006.
28