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Quarterly Journal of the Royal Meteorological Society
Q. J. R. Meteorol. Soc. 139: 1615–1631, July 2013 B
Dynamics of the Foehn Mechanism in the McMurdo Dry Valleys
of Antarctica from Polar WRF†
Daniel F. Steinhoff,a,b * David H. Bromwicha,b and Andrew Monaghanc
a
b
Polar Meteorology Group, Byrd Polar Research Center, The Ohio State University, Columbus, OH, USA
Atmospheric Sciences Program, Department of Geography, The Ohio State University, Columbus, OH, USA
c
Research Applications Laboratory, National Center for Atmospheric Research, Boulder, CO, USA
*Correspondence to: D. F. Steinhoff, Research Applications Laboratory, National Center for Atmospheric Research,
Boulder, CO, USA. E-mail: [email protected]
† Contribution 1423 of the Byrd Polar Research Center.
Foehn events over the McMurdo Dry Valleys (MDVs), the largest ice-free region
of Antarctica, promote glacial melt that supports biological activity in the lakes,
streams, rocks and soils. Although MDVs foehn events are known to depend upon
the synoptic-scale circulation, the physical processes responsible for foehn events
are unknown. A polar-optimized version of the Weather Research and Forecasting
model (Polar WRF) is used for a case study of a representative summer foehn
event from 29 December 2006 to 1 January 2007 in order to identify and explain
the MDVs foehn mechanism. Pressure differences across an elevated mountain gap
upstream of the MDVs provide forcing for southerly flow into the western, upvalley
entrance of the MDVs. Complex terrain over the elevated gap and the MDVs
leads to mountain wave effects such as leeside acceleration, hydraulic jumps, wave
breaking and critical layers. These mountain wave effects depend on the ambient
(geostrophic) wind direction. Pressure-driven channelling then brings the warm,
dry foehn air downvalley to eastern MDV sites. Brief easterly intrusions of maritime
air into the eastern MDVs during foehn events previously have been attributed
to either a sea-breeze effect in summer or local cold-pooling effects in winter. In
this particular case, the easterly intrusions result from blocking effects of nearby
Ross Island and the adjacent Antarctic coast. Temperature variability during the
summer foehn event, which is important for meltwater production and biological
activity when it exceeds 0◦ C, primarily depends on the source airmass rather than
differences in foehn dynamics.
Key Words:
mountain waves; gap flow; channelling
Received 27 February 2012; Revised 7 June 2012; 31 July 2012; Accepted 14 August 2012; Published online in
Wiley Online Library 12 November 2012
Citation: Steinhoff DF, Bromwich DH, Monaghan A. 2013. Dynamics of the Foehn Mechanism in
the McMurdo Dry Valleys of Antarctica from Polar WRF. Q. J. R. Meteorol. Soc. 139: 1615–1631.
DOI:10.1002/qj.2038
1. Introduction
The McMurdo Dry Valleys (MDVs) are the largest ice-free
region of Antarctica, characterized by extensive biological
activity in the soils, rocks, ephemeral steams and glacial melt
lakes (e.g. Fountain et al., 1999). As such, the MDVs have
been a National Science Foundation Long-Term Ecological
c 2012 Royal Meteorological Society
Research (NSF-LTER) site since 1993. The MDVs are also
of interest for soil properties (e.g. Campbell et al., 1998;
Ikard et al., 2009) and aeolian processes (e.g. Speirs et al.,
2008). Although glaciological processes are often cited as
being responsible for the ‘dry’ valleys (McKelvey and Webb,
1962; Chinn, 1990), the MDVs’ microclimate certainly
aids in perpetuating the arid conditions. The MDVs are
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D. F. Steinhoff et al.
(b)
Ross
Sea
(a)
Ross
Island
RIS MBL
M
TA
EAIS
McMurdo
Sound
Windless
Bight
McMurdo
Ross
Ice
Shelf
(c)
10 km
LBr
LF
ills
LH
Ku
kr
iH
LVa
As
ga
rd
ey
Vall
oria nge
Vict
a
R
us
mp
ey
Oly
Vall
t
h
Wrig
R
an
ge
LVi
LBo
r Valley
Taylo
Royal
Society
Range
Figure 1. Overview maps of (a) Antarctica (EAIS, is East Antarctic Ice Sheet; MBL, Marie Byrd Land; TAM, Transantarctic Mountains; RIS, Ross Ice
Shelf) and (b) MDVs and surrounding region, with terrain height (shading, m) from RADARSAT-1 Antarctic Mapping Project data (see text). (c) Inset
of (b), Landsat ETM+ image of the MDVs from 21 November 2001. Red dots indicate LTER AWS locations referenced in this study, with abbreviations
provided in Table 1.
a polar desert (e.g. Doran et al., 2002a), averaging between
3 mm and 50 mm (water-equivalent) precipitation annually.
Because the snow quickly sublimates (Chinn, 1981), MDVs’
hydrology is primarily driven by glacial melt and winddriven transport of snow from valley sidewalls to valley
floors (Fountain et al., 1998, 2010). The physical setting of
the MDVs is illustrated in Figure 1. The MDVs are located
between the relatively warm, maritime McMurdo Sound and
the cold, dry East Antarctic Ice Sheet (Figure 1(a) and(b)).
Besides the ice-free valleys (Figure 1(c)), the region also
features complex terrain, evidenced by mountain ranges
exceeding 2000 m elevation between valleys, and the Royal
Society Range (peaks upwards of 4000 m) just to the south.
The Royal Society Range imposes a precipitation shadow
effect on the MDVs (Monaghan et al., 2005).
The controlling factor on MDVs’ climate is the wind.
During quiescent conditions in the summer, a diurnal
cycle exists with easterly winds during daytime (Bull, 1966;
c 2012 Royal Meteorological Society
McKendry and Lewthwaite, 1992; Nylen et al., 2004). The
easterlies are assumed to be thermodynamically forced sea
breezes, resulting from differential heating between the bare
ground over the MDVs and the water (or sea ice) over
McMurdo Sound. The sea breeze circulation weakens at
night when the sun angle is lower, and light westerlies form.
During winter, the sea-breeze circulation is nonexistent,
and winds are generally light, with a localized katabatic
component down valley slopes (Riordan, 1975; Nylen et al.,
2004).
At any time of the year, instances of strong, gusty
southwesterly winds can occur on time-scales of a few
hours to a week. The strong wind events are characterized
by warming (especially during winter, when valley cold pools
are disrupted and temperatures can rise by up to 50◦ C) and
drying (relative humidity less than 10%). Although such
events are stronger and more common in winter, they are
especially important in summer, when temperatures can
Q. J. R. Meteorol. Soc. 139: 1615–1631 (2013)
Foehn Mechanism in the McMurdo Dry Valleys
rise above 0◦ C and significant melt can occur, enhancing
biological activity (Priscu et al., 1998). The forcing for the
strong wind events has been debated in previous literature.
Some researchers (Bull, 1966; Clow et al., 1988; Doran
et al., 2002a; Nylen et al., 2004) have referred to the winds
as ‘katabatic’, associated with katabatic surges from the
polar plateau (e.g. Parish and Bromwich, 1998). Although
a katabatic wind can be defined as any downslope wind,
in Antarctica it refers to negatively buoyant flow (Parish
and Cassano, 2003; van den Broeke and van Lipzig, 2003).
Others attribute a foehn effect to the events. The WMO
(1992) defines foehn as a ‘wind warmed and dried by
descent, in general on the lee side of a mountain.’ Thompson
et al. (1971a), Riordan (1975) and Bromley (1985) note the
deflection of ambient southerly to southwesterly flow aloft
into the MDVs. McKendry and Lewthwaite (1990, 1992),
based on the only observational study of MDVs atmospheric
vertical structure to date, support the foehn explanation,
as the strong winds are confined to the valleys, whereas
katabatic surges would be expected to be more widespread.
They note that a mid-tropospheric inversion above the
ridge satisfies an important precondition of downslope
windstorms (e.g. Durran, 1990), and find evidence of
a particular synoptic-scale sequence enabling downward
deflection of flow into the western MDVs. Speirs et al. (2010)
confirm a foehn effect using surface observations and output
from the Antarctic Mesoscale Prediction System (AMPS;
Powers et al., 2012), while also finding that foehn events
correspond with strong pressure/height gradients over the
Ross Ice Shelf, associated with synoptic-scale cyclones off the
coast of Marie Byrd Land. Furthermore, significant katabatic
wind activity is not possible during summer due to weaker
surface temperature inversions (Parish and Cassano, 2003).
Warming and drying associated with foehn winds are
typically attributed to saturated adiabatic ascent and dry
adiabatic descent over a mountain barrier (resulting in
net warming and drying). A broader definition, applied
by Seibert (1990), Zängl (2003), and Speirs et al. (2010),
includes forced descent as a warming mechanism. Forced
descent can be associated with mountain waves (e.g. Klemp
and Lilly, 1975; Jiang et al., 2005) or by blocking of lowlevel winds windward of mountain barriers (e.g. Parish,
1983; Gohm et al., 2004). Foehn and the related bora (cold
downslope wind, e.g. Smith, 1987) occur throughout the
world with different underlying physical processes. Chinook
winds along the front range of the Rocky Mountains (e.g.
Lilly and Zipser, 1972; Doyle et al., 2000) are attributed to
stationary large-amplitude mountain waves and the effects
of wave breaking aloft. Over parts of the European Alps,
and with the Taku wind near Juneau Alaska, foehn forcing
is best described as a gap flow, associated with an alongflow pressure gradient through a channel or mountain pass,
primarily forced by upstream blocking and subsidence in
the lee (e.g. Overland and Walter, 1981; Zängl, 2002a; Bond
et al., 2006; Mayr et al., 2007). Mountain wave effects are
still important with gap flow, especially over elevated gaps
(e.g. Pan and Smith, 1999; Gohm et al., 2004; Gaberšek
and Durran, 2004). Other complicating factors include
atmospheric moisture (Durran and Klemp, 1982; Doyle and
Shapiro, 2000), either mean-state or wave-induced critical
levels (e.g. Clark and Peltier, 1984; Smith, 1985; Durran
and Klemp, 1987; Colle and Mass, 1998a; Bond et al.,
2006; Fudeyasu et al., 2008), and boundary-layer effects
(e.g. Ólafsson and Bougeault, 1997; Smith et al., 2002, 2006;
c 2012 Royal Meteorological Society
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Vosper, 2004; Jiang et al., 2006, 2008; Jiang and Doyle, 2008;
Smith and Skyllingstad, 2009).
Although a foehn effect is responsible for strong wind
events over the MDVs, and certain synoptic-scale patterns
have been attributed to foehn conditions, the physical
mechanisms responsible for foehn in the MDVs have not
been identified. In the past, lack of detailed meteorological
observations prevented such a study from being undertaken.
Here we use a mesoscale model optimized for use in polar
environments to study foehn conditions in the MDVs during
a case study from 29 December 2006 to 01 January 2007.
A sequence of mesoscale features, including gap flow,
mountain waves and pressure-driven channelling, are found
to connect the synoptic-scale circulation to foehn in the
MDVs. In the next section, the mesoscale model and
surface observations are described. After a case overview,
subsections then follow to explain each mesoscale effect,
including a new explanation of easterly intrusions during
foehn events. Temperature variability during the event is
then discussed, followed by conclusions.
2. Polar Weather Research and Forecasting model and
observations
2.1. Polar Weather Research and Forecasting model
The Advanced Research Weather Research And Forecasting
model (WRF-ARW; Skamarock et al., 2008) solves for the
non-hydrostatic Euler equations on a spatially discretized
grid in the horizontal and a terrain-following hydrostatic
pressure vertical coordinate. The WRF is used for
atmospheric science applications on a variety of spatial
scales, with version 3.2.1 used in this study. Several
physical parametrization schemes are available in WRF,
and those chosen for this study are described below.
The WRF single-moment 6-class microphysics scheme
(WSM6; Hong and Lim, 2006) is a bulk scheme, which
includes a predictive equation for graupel. This presumably
improves prediction at higher spatial resolution compared
with schemes with less prognostic variables. For longwave radiation, the rapid radiative transfer model (RRTM;
Mlawer et al., 1997) uses the correlated-k method that
compares favourably with a full line-by-line radiative
transfer model. The empirically based Dudhia short-wave
radiation scheme (Dudhia, 1989) uses an inexpensive binary
representation of cloud cover. To estimate subgrid-scale
turbulent fluxes, the Mellor–Yamada–Nakanishi–Niino
(MYNN; Nakanishi, 2001; Nakanishi and Niino, 2004,
2006) level-2.5 scheme is used. Changes to the master
length scale from the Mellor–Yamada–Janjić (MYJ; Janjić,
1994) scheme should alleviate underestimation of mixedlayer depth and turbulent kinetic energy (TKE) magnitude.
The corresponding MYNN surface layer scheme is also used.
The Noah land surface model (LSM) (Chen and Dudhia,
2001) has four soil layers and is able to handle sea-ice and
polar conditions through modifications described below.
The Grell–Devenyi cumulus scheme (Grell and Devenyi,
2002) is used on coarser domains.
Modifications of the WRF for polar environments, termed
Polar WRF, were done by the Polar Meteorology Group at
the Byrd Polar Research Center, The Ohio State University
(Hines and Bromwich, 2008; Bromwich et al., 2009; Hines
et al., 2011; http://polarmet.osu.edu/PolarMet/pwrf.html).
These modifications primarily encompass the land surface
Q. J. R. Meteorol. Soc. 139: 1615–1631 (2013)
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D. F. Steinhoff et al.
(a)
(b)
B’
A
A’
Box1
B
Box2
Figure 2. (a) The 32 km Polar WRF domain and terrain height (m, shaded). Solid-outlined boxes show 8 km, 2 km and 0.5 km nested domains.
Dashed-outlined box shows area used in Figure 3. (b) The 0.5 km Polar WRF domain and terrain height (m, shaded). Dashed black box shows area used
in Figure 7; red dots show LTER AWS locations used in this study (Figure 1; Table 1); Box1 is used in Figure 6(b); Box2 in section 4.2; transect A–A’ in
Figure 5; and transect B–B’ in Figure 6.
model, including optimal values of snow thermal properties and improved heat flux calculations. Time-variable
fractional sea ice is represented by separate calls to the surface layer scheme for ice and open water, and the surface
heat fluxes are aerially averaged to obtain the final values
for the fractional sea-ice grid box. Sea-ice thickness and
snow depth on sea ice are also allowed to be specified on a
grid-point basis. Further modifications by the authors were
implemented to use fractional sea ice with the MYNN PBL
scheme.
A 32–8–2–0.5 km nesting structure is used for the
spatial domains (Figure 2(a)), covering all of Antarctica
and surrounding ocean in the 32 km domain down to the
MDVs region only for the 0.5 km domain. For the 32, 8
and 2 km domains, 40 vertical levels are used, with the
lowest level approximately 13.4 m above the surface. For
c 2012 Royal Meteorological Society
the 0.5 km domain (Figure 2(b)), which is used primarily
in this study because terrain features in the MDVs are in
the order of a few kilometres (Figure 1(c)), 55 vertical
levels are used and the lowest model level is about 26.5 m
above the surface. A higher lowest model level is needed
for the 0.5 km domain for computational stability. Under
conditions of forced mechanical turbulence with stable
stratification there can be large differences in boundary
layer wind speed, temperature and flow separation from the
surface, depending on the height of the lowest model level
(Zängl et al., 2008). A sensitivity simulation with the lowest
model level at 13.4 m (performed on a smaller domain
to promote computational stability) yielded similar results
for wind speeds and mountain wave structure over the
MDVs compared with having the lowest level at 26.5 m (not
shown). This indicates that our boundary layer is somewhat
Q. J. R. Meteorol. Soc. 139: 1615–1631 (2013)
Foehn Mechanism in the McMurdo Dry Valleys
insensitive to the chosen height of the lowest model level. In
all simulations, a model top of 50 hPa is used, and all nested
domains are one-way coupled.
Several input datasets are used in order to optimize the
simulations. Input conditions for the atmosphere, some
surface fields and lateral boundary conditions for the 32 km
domain are supplied by ERA-Interim, European Centre for
Medium Wave Weather Forecasts’ (ECMWF) most recent
global reanalysis product (Dee et al., 2011). ERA-Interim
features 4D-Var data assimilation at 12 h intervals at a
native spatial resolution of T255 (∼0.70◦ ). Important data
sources for high-latitude regions ingested into ERA-Interim
from 2001 onwards include Moderate-Resolution Imaging
Spectroradiometer (MODIS) winds, Atmospheric Infrared
Sounder (AIRS) radiances and Constellation Observing
System for Meteorology, Ionosphere, and Climate (COSMIC) GPS radio occultation soundings (Andersson, 2007).
ERA-Interim has the best representation of precipitation
(and therefore likely atmospheric circulation) among global
reanalyses for Antarctica (Bromwich et al., 2011). Here
we use ERA-Interim reanalyses on a regular 512 × 256
N128 Gaussian grid provided by the National Center for
Atmospheric Research (NCAR) Data Support Section.
Additionally, four-dimensional data assimilation (FDDA;
e.g. Stauffer and Seaman, 1994) is used to ‘nudge’ the
interior grid points of the outermost 32 km domain to
six-hourly ERA-Interim output throughout the simulation.
Grid-point FDDA includes an additional time-tendency
term in the solution for a given variable that nudges the
solution to a gridded analysis point-by-point, and allows
for simulations to be run for more than a few days without
significant model drift and costly restarts. In particular,
model drift is a problem during periods of cyclonic forcing
over the Ross Sea (Nigro et al., 2011). The top 27 vertical
levels are nudged with a 1 h time-scale parameter.
Sea-surface temperature (SST) data are supplied from
daily optimally interpolated 0.25◦ advanced very high
resolution radiometer (AVHRR) and advanced microwave
scanning radiometer for Earth Observing System (AMSR-E)
output from the National Climatic Data Center (NCDC) and
National Environmental Satellite, Data, and Information
Service (NESDIS) of National Oceanic and Atmospheric
Administration (NOAA) (Reynolds et al., 2007). For seaice concentration, we use AMSR-E output processed using
the ARTIST sea-ice (ASI) algorithm (Spreen et al., 2008).
This product uses 89 GHz channels, providing up to four
times greater horizontal resolution than other algorithms.
Data are obtained from Universität Bremen on a polar
stereographic grid at 6.25 km grid spacing. For sea-ice
thickness and snow thickness on top of sea ice, we use
climatological values from the Antarctic Sea-ice Processes
and Climate program (ASPeCt; Worby et al., 2008). The
ASPeCt program is comprised of over 23 000 ship and
aircraft sea-ice observations from 81 voyages from 1981 to
2005. An annual climatology was produced at a 2.5◦ latitude
by 5.0◦ longitude grid spacing. Terrain height is interpolated
from the 200 m RADARSAT-1 Antarctic Mapping Project
(RAMP) digital elevation model (Liu et al., 2001). This field
is smoothed twice for the 32, 8 and 2 km domains, and
three times for the 0.5 km domain, to avoid computational
instability with steeply sloping terrain. A special bare-ground
land use dataset for the MDVs is used, interpolated from
9’’ US Geological Survey GeoTiff files by Kevin Manning of
NCAR.
c 2012 Royal Meteorological Society
1619
Three WRF name-list options were changed to deal
with the complex terrain. Rather than being calculated
on model vertical levels, horizontal diffusion is calculated
in physical space over the 0.5 km domain. This is done to
prevent temperature gradients along valley slopes from being
erroneously diffused (Zängl, 2002b). Zängl et al. (2004a,b)
performed sensitivity experiments with horizontal diffusion
in the fifth-generation Pennsylvania University-National
Center for Atmospheric Research Mesoscale Model (MM5)
and found that calculating horizontal diffusion on model
surfaces led to incorrect timing of foehn initiation in
the Wipp and Rhine Valleys in the Alps. In order to
use diffusion in physical space, the 0.5 km domain had
to be run separately from the other nests, using WRF’s
‘ndown’ program, and input from the 2 km domain at
hourly intervals. Additionally, slope and shadowing effects in
regions of complex terrain influence incident solar radiation
in the MDVs (Dana et al., 1998). These effects are accounted
for in the short-wave parameterization in WRF (Garnier
and Ohmura, 1968). Finally, an implicit Rayleigh damping
of vertical velocity is used in the upper levels of the model,
to prevent unphysical downward reflection of gravity waves
from the model top (Klemp et al., 2008).
Further modifications are made for MDVs’ environment,
separate from the WRF name-list options. Any snow input
to the model from ERA-Interim over bare ground points
is removed before model initialization because the MDVs
are known to be snow-free nearly perennially (Fountain
et al., 2010). Furthermore, we restrict WRF-simulated waterequivalent snow depth over bare ground to 0.2 cm, as
WRF cannot simulate the removal of snow by strong
foehn winds. Since there is no pre-set soil categorization
for Antarctica in the WRF model, we specify bare ground as
‘loamy sand’, consisting of 82% sand, 12% silt and 6% clay.
This soil distribution compares favourably with available
observations (Campbell et al., 1998; Northcott et al., 2009).
Similar initialization issues exist for the roughness length and
albedo. Lancaster (2004) measured roughness length at 17
sites in the MDVs and found values ranging from 0.0005 to
0.036 m. As we cannot confidently specify roughness length
on a grid-point basis, we set the roughness length to 0.005 m
for all bare-ground points. Bare-ground albedo values in
the MDVs range from 0.06 to 0.26 (Thompson et al., 1971a;
Campbell et al., 1998; Hunt et al., 2010). The base (no snow)
albedo for the barren-ground category is set to 0.18. Realistic,
scale-consistent profiles of soil temperature and moisture
are necessary for ground heat-flux calculations but are not
provided by the ERA-Interim input conditions. Therefore,
the 32–8–2 km Polar WRF simulations were run for 1 yr
(1 November 2005 to 31 October 2006), and the resulting
‘spun up’ soil fields at the end of this simulation were used
as input for simulations beginning 1 November of any year.
The 32–8–2 km grids were run for the entire summer season
(November 2006 through February 2007), providing ample
spin-up for the 0.5 km case study simulation. To initialize the
soil spin-up simulations, soil temperature observations of
Thompson et al. (1971b) are extrapolated both horizontally
(using a function based on distance from the coast, Doran
et al., 2002a) and vertically (to the Noah LSM levels). Soil
moisture is similarly initialized based on observations from
Campbell et al. (1998), based on soil porosity of 0.41 (McKay
et al., 1998). To account for enhanced soil moisture values
near the shores of melt lakes (Ikard et al., 2009), volumetric
soil moisture is set to 0.15 for these grid points.
Q. J. R. Meteorol. Soc. 139: 1615–1631 (2013)
1620
D. F. Steinhoff et al.
Table 1. MDV LTER AWS locations.
Station
Code
Valley
Latitude
(◦ N)
Longitude
(◦ E)
Elevation
(m)
Lake Fryxell
Lake Hoare
Lake Bonney
Lake Brownworth
Lake Vanda
Lake Vida
LF
LH
LBo
LBr
LVa
LVi
Taylor
Taylor
Taylor
Wright
Wright
Victoria
−77.6109
−77.6254
−77.7144
−77.4335
−77.5168
−77.3778
163.1696
162.9004
162.4641
162.7036
161.6678
161.8006
19
78
64
279
296
351
2.2.
The LTER project observations
The MDVs LTER project provides automatic weather
station (AWS) data at sites across the MDVs (Doran et al.,
1995). Observation sites used in this study are shown in
Figure 1(c) and Table 1. Observations are at 3 m height, and
relative humidity is calculated with respect to ice when air
temperatures are below freezing. To facilitate comparison
with AWS observations, selected variables are output at
3 m height in Polar WRF by adjusting the height at
which exchange coefficients are calculated in the surface
layer scheme. Wind is sampled at 4 s intervals (except at
Lake Bonney AWS, which samples at 1 s), and all other
variables are sampled at 30 s intervals. All variables are then
averaged at 15 min intervals and quality controlled by LTER
investigators. Taylor Glacier observations are not available
for the case study period, and relative humidity data are not
available at Lake Bonney.
Care is taken to choose proper Polar WRF grid points
for representing LTER AWS locations (to avoid so-called
‘representativeness errors’, Jiménez and Dudhia, 2012). Due
to the complex terrain and land use, no spatial averaging of
grid points is used, nor is a strict determination of nearest
grid point used. Out of the four grid points surrounding the
true LTER AWS location, the one chosen was required to
have the correct land use and be properly located within the
valley geometry. All Polar WRF station points are located
within 500 m distance and 80 m elevation of the actual
station.
3.
Sea by 0000 UTC 1 January (Figure 3(d)). Through this time
period, the region illustrated in Figure 3 features a variety
of flow orientations (both wind speed and wind direction)
at the surface and aloft, which will be shown to largely
determine meteorological conditions in the MDVs.
To show conditions in the MDVs during this event, as
well as validate the model simulations, time-series plots of
LTER AWS and Polar WRF output at several sites are shown
in Figure 4. The foehn event begins around 1200 UTC
29 December, with increasing wind speeds and wind
directions shifting to the southwest. Temperatures surpass
0◦ C during daytime hours on 30 December, 31 December
and 1 January, with a muted diurnal temperature cycle.
Note that the MDVs’ ‘local time’ is UTC+13. Relative
humidity values are low, generally under 30% during foehn
conditions. Interruptions in foehn conditions at eastern
MDV sites (e.g. Lake Fryxell and Lake Brownworth) lead to
lighter easterly winds, decreased temperature and increased
relative humidity.
Polar WRF output interpolated to observation locations
is also shown in Figure 4. Temperature errors are generally
small across all sites, but there is a clear negative bias in
relative humidity. This negative bias is masked by negative
temperature biases at some sites, and may be related to
a negative moisture bias over Antarctica in ERA-Interim
(Nicolas and Bromwich, 2011). The model well-represents
onset of foehn conditions, but undersimulates easterly
intrusions at Lake Fryxell and Lake Brownworth early
30 December. The WRF wind speed is higher than observed
during foehn conditions, similar to other WRF studies over
valleys (e.g. Jiménez et al., 2010). The positive wind speed
bias (discussed more in section 4.2) may result from the
smoothed topography and the drag generated by unresolved
topographic features (e.g. Beljaars et al., 2004; Jiménez and
Dudhia, 2012), which can then result in misplaced hydraulic
jumps that cause large wind speed biases (e.g. Steinhoff
et al., 2008). Short sensitivity tests with different boundary
layer and surface parametrization schemes and less terrain
smoothing did not show substantially improved results.
However, the model does capture much of the variability in
wind speed and direction during the event, and therefore we
find it suitable for the following analysis.
Case overview
4.
The MDVs foehn event from 29 December 2006 to
1 January 2007 is chosen for study because it occurs
at all LTER AWS sites (using the foehn identification
scheme of Speirs et al., 2010) and it features varying flow
orientations that allow for a more complete explanation
of foehn structure. A synoptic overview, consisting of sealevel pressure, 3 m temperature, 3 m wind vectors, 500 hPa
geopotential height, and 500 hPa relative vorticity, from
1800 UTC 29 December 2006 to 0000 UTC 1 January 2007
every 18 h is shown in Figure 3. At 1800 UTC 29 December,
an upper-level low is located over the Ross Sea, with a weak
low-level cyclone over the Ross Ice Shelf and a stronger
cyclone off the Marie Byrd Land coast (Figure 3(a)). By
1200 UTC 30 December, the low-level cyclone over the Ross
Ice Shelf has dissipated, as the system over the southeastern
Ross Sea intensifies over a region of 500 hPa cyclonic
vorticity advection (Figure 3(b)). This system exhibits an
equivalent barotropic structure as it weakens and moves
westward to the centre of the Ross Ice Shelf by 0600 UTC
31 December (Figure 3(c)), then equatorward into the Ross
c 2012 Royal Meteorological Society
4.1.
Mesoscale features
Gap flow
As noted in the introduction, gap flow is forced by the
pressure difference along (through) the gap. This alonggap pressure gradient is often generated by flow blocking
upstream. The western Ross Ice Shelf is well-known for
barrier winds, associated with low-level cyclonic flow being
blocked against the Transantarctic Mountains and directed
northward by an evolving mesoscale pressure gradient
(O’Connor et al., 1994; Steinhoff et al., 2008). This is clearly
the case in Figure 3, where easterly flow on the poleward
sector of the cyclones over the Ross Ice Shelf and near
Marie Byrd Land impinges on the Transantarctic Mountains.
Figure 5 shows a profile of the terrain upstream of the MDVs
along transect A–A’ in Figure 2(b). There is a gap about
30 km wide, located west of the Royal Society Range and
south of western Taylor Valley. The terrain within the gap
is complex and there is no straight pass of lowest terrain
height through the gap.
Q. J. R. Meteorol. Soc. 139: 1615–1631 (2013)
Foehn Mechanism in the McMurdo Dry Valleys
1621
18 UTC 29 DEC
(a)
12 UTC 30 DEC
(b)
06 UTC 31 DEC
(c)
00 UTC 1 JAN
(d)
Figure 3. Synoptic overview plots of (left) sea-level pressure (hPa, contours, interval 2 hPa), 3 m temperature (◦ C, shading) and 3 m wind vectors
(arrows), and (right) 500 hPa geopotential height (m, contours, interval 60 m), relative vorticity (10−5 s−1 , shading), and wind vectors (arrows) for
(a) 1800 UTC 29 December, (b) 1200 UTC 30 December, (c) 0600 UTC 31 December and (d) 0000 UTC 1 January, from 32 km Polar WRF domain.
Sea-level pressure not plotted over areas exceeding 500 m elevation. Star symbol denotes approximate location of MDVs.
To estimate the pressure difference across the gap, neither
surface pressure nor sea-level pressure can be used, due to the
complex terrain and the poor assumption of a constant lapse
rate for the Antarctic boundary layer. Instead, the reduced
pressure along a transect through the gap is calculated.
Pressure is reduced to a constant height, defined here to be
the highest point along the transect. A reduction method
similar to that of Mayr et al. (2002) is used, but because
we have vertical profiles of virtual temperature from model
c 2012 Royal Meteorological Society
output, pressure is simply adjusted from the surface up to a
constant height level. Reduced pressure differences between
the starting and ending points of the transect are computed
to give the along-gap pressure difference.
Figure 6(a) shows the reduced pressure difference across
the gap (transect B–B’ in Figure 2(b)). The pressure
difference increases during 29 December to about 2.8 hPa
by early 30 December, before dropping under 2 hPa later in
the day. Then pressure differences increase substantially to
Q. J. R. Meteorol. Soc. 139: 1615–1631 (2013)
1622
D. F. Steinhoff et al.
Figure 4. Time series of 3 m temperature (◦ C, purple), relative humidity (%, red), wind direction (degrees, orange), and wind speed (m s−1 , green)
from LTER observations (dark colours) and Polar WRF (light colours).
3.5–4.5 hPa after 1200 UTC 31 December. To see if there
is a connection between this pressure difference and the gap
wind speed, the area-averaged near-surface total wind speed
through the gap (using Box1 in Figure 2(b) is shown in
Figure 6(b)). Increases in both pressure difference and wind
speed during 29 December and at the peak of the foehn
event around 0000 UTC 1 January are consistent. However,
strong gap winds (around 20 m s−1 in the model) exist
late 30 December when the pressure difference is 2.8 hPa.
Gap wind speeds then decrease early 31 December while
the cross-gap pressure difference is increasing. Although
the broad distribution of wind speed through the gap scales
with the along-gap pressure difference, there are fluctuations
c 2012 Royal Meteorological Society
in gap wind speed that are independent of the pressure
difference.
Figure 6(c) shows 600 hPa wind direction, averaged for
all points in the model domain south of Box1 in Figure 2(b).
This level represents ambient flow, above the boundary
layer across the domain. Late on December 30, when
there is a discrepancy between wind speed and pressure
difference over the gap, south-southwesterly winds occur
aloft, meaning that flow is off-continent, and a strong
pressure difference would not be expected. The upstream
flow originates at a higher level (due to higher-elevation
terrain to the west), and is less apt to be blocked. The dashed
line in Figure 6(c) approximates the gap axis (the wind
direction directly through the gap) at 203◦ , and Figure 6(d)
Q. J. R. Meteorol. Soc. 139: 1615–1631 (2013)
Foehn Mechanism in the McMurdo Dry Valleys
1623
(a)
RSR
MF
EAIS
(b)
Gap
(c)
A
A’
Figure 5. Profile of Polar WRF terrain height above sea level along transect
A–A’ in Figure 2(b): EAIS, the East Antarctic Ice Sheet; MF, Mount
Feather (2985 m); RSR, the Royal Society Range; and ‘Gap’ represents the
approximate width of the gap discussed in the text, of which the lowest
point and vertical profile varies with latitude.
shows the 600 hPa domain-averaged wind speed. Late on
30 December, ambient flow of about 16 m s−1 is nearly
parallel to the gap axis, with a strong flow through the gap of
almost 20 m s−1 . Around 0000 UTC 1 January, ambient flow
is south-southeasterly, with a stronger pressure difference,
but similar flow speed through the gap. The gap flow set-up,
and the importance of the along-gap pressure difference,
depends on the ambient wind direction. When ambient
flow has an easterly component, the ambient along-gap
wind component is smaller (flow is more oblique to the
gap axis), but a strong along-gap pressure difference due
to blocking forces the gap flow. When ambient flow has
a westerly component, the pressure difference is smaller,
implying that other effects are important for along-gap flow
acceleration.
4.2. Mountain waves
Mountain wave activity in both level and elevated gaps
significantly affects the resulting flow via momentum fluxes
and pressure drag (Gaberšek and Durran, 2004; Gohm and
Mayr, 2004; Gohm et al., 2008). The combined influence
of upstream blocking and pressure drag from mountain
waves is at least as important as the ambient synopticscale pressure gradient for gap flow (Gaberšek and Durran,
2006). Colle and Mass (1998b) show that along-gap pressure
gradients can be enhanced by 25–50% from leeside wave
amplification. Bromley (1985) and Fountain et al. (2010)
observe wind-blown snow into the MDVs from valley walls
to the south, supporting the concept of mountain waves
over the MDVs. Although mountain waves over the MDVs
were identified by Speirs et al. (2010), here we provide more
detail as to the structure and features.
At 1800 UTC 29 December, large-scale flow towards the
MDVs region is from the south, where it is blocked upstream
by the Transantarctic Mountains, so that the low-level flow
into the gap is weak (Figure 7(a)). This has ramifications
for the flow structure over the complex terrain of the gap.
Figure 7(b) shows a cross-section of potential temperature
c 2012 Royal Meteorological Society
(d)
Figure 6. Polar WRF values along transect B–B’ in Figure 2(b) of
(a) reduced pressure difference (hPa), (b) 3 m total wind speed averaged
over Box1 in Figure 2(b), (c) 600 hPa upstream wind direction, where
dashed line at 203◦ indicates gap axis, and (d) 600 hPa upstream wind
speed.
and wind speed through the gap along transect C–C’ at
1800 UTC 29 December. Interpretation of wave effects in
complex multiscale terrain is difficult, but some features
can still be identified and explained. The flow in Figure 7(b)
appears to be nonlinear, evidenced by large wave amplitudes,
strong leeside acceleration and wave breaking. To quantify
the nonlinearity, we calculate the non-dimensional mountain
height (also known as the inverse Froude number) for a
volume upstream of the gap:
M=
Nhm
U
(1)
where U is the mean-state layer-average wind speed, N is the
mean-state Brunt–Väisälä frequency and hm is the terrain
height. Here we calculate M over Box2 in Figure 2(b) up
to 2000 m above ground level(AGL). As recommended by
Reinecke and Durran (2008), the average N in this layer is
used to estimate the mean-state in Eq. (1), rather than the
bulk difference between the top and bottom of the layer.
Variable M is a measure of the nonlinearity of the flow, with
an established threshold of 1.1 for continuously stratified,
hydrostatic flow over a three-dimensional axisymmetic
Gaussian hill (Smith and Grønås, 1993). This measure is
valid only for the upstream, relatively undisturbed flow,
and not over the complex terrain of the MDVs. A value
of 1.62 is found at 1800 UTC 29 December, indicating
significant nonlinearity. In Figure 7(b), blocking is apparent
upstream of the first ridge (Gap), with transitional flow at
the crest and acceleration in the lowest 1 km in the lee. Wave
steepening occurs over the lee slope, associated with weak
flow at 4–5 km above seal level (ASL). The flow pattern
Q. J. R. Meteorol. Soc. 139: 1615–1631 (2013)
1624
D. F. Steinhoff et al.
(a)
8.0
(b)
300
D’
C’
290
Height (km)
6.0
4.0
280
2.0
C D
Gap
KH
AR
FG
0.0
20
40
OR
TV
VV
80 WV
60
100
Distance (km)
(c)
(d)
8.0
300
C’
Height (km)
6.0
290
4.0
280
2.0
C
Gap
KH
AR
FG
0.0
(e)
8.0
80 WV
60
40
20
OR
TV
VV
100
Distance (km)
(f)
310
C’
300
E’
6.0
E
Height (km)
290
4.0
280
2.0
C
Gap
KH
AR
FG
0.0
C
-6
80 WV
60
40
Distance (km)
20
0
6
12
OR
TV
VV
100
C’
18
24
30
−1
Figure 7. (a) Polar WRF 3-m wind speed (m s , shading), vectors and terrain height above sea level (m, contours, interval 300 m) at 1800 UTC
29 December 2006. Green dots represent LTER AWS locations. (b) Polar WRF vertical cross-section of potential temperature (K, contours, 1 K interval),
along-transect wind speed component (m s−1 , shaded), turbulent kinetic energy (white dashed lines, starting at 2.5 m2 s−2 , 5 m2 s−2 interval), and
circulation vectors along transect C–C’ in (a) at 1800 UTC 29 December 2006. ‘Gap’ is the elevated gap; KH, Knobhead; FG, Ferrar Glacier; TV,
Taylor Valley; AR, Asgard Range; WV, Wright Valley; OR, Olympus Range; VV, Victoria Valley. Reference vectors in lower left represent 27.5 m s−1
in the horizontal and 7.21 m s−1 in the vertical. ((c) and (d)) Same as (a) and (b) except at 1800 UTC 30 December, reference vectors in lower left of
(d) represent 28.1 m s−1 in the horizontal and 7.75 m s−1 in the vertical, and no turbulent kinetic energy plotted in (d). ((e) and (f)) Same as (a) and
(b) except at 0300 UTC 1 January 2007 and reference vectors in lower left represent 32.2 m s−1 in the horizontal and 10.4 m s−1 in the vertical. Vertical
cross-sections along transects D–D’ and E–E’ are presented in Figure 8.
resembles a hydrostatic vertically propagating wave regime,
due to the relatively long length scale of the gap, and single
wave pattern over the gap ridge.
Flow accelerates in the lee of the gap ridge in Figure 7(b),
underneath a region of steepened isentropes. The wave
amplification is associated with the aforementioned wave
steepening/breaking, and the strong lee winds are best
explained by hydraulic theory (e.g. Long, 1954; Durran,
1986, 1990). Even without a mean-state stratification
c 2012 Royal Meteorological Society
discontinuity or a mean-state critical level, wave breaking
can result in a transition to supercritical flow (Durran,
1986; Durran and Klemp, 1987). A wave breaking region
also forms over Knobhead (KH), but at about 3 km ASL.
In relation to a continuously stratified hydraulic analogue,
the layer of ‘fluid’ (the layer of strong stability just above
the surface layer) thins and accelerates in the lee of KH,
corresponding to supercritical flow. This ‘shooting flow’
rapidly adjusts to subcritical conditions downstream in a
Q. J. R. Meteorol. Soc. 139: 1615–1631 (2013)
Foehn Mechanism in the McMurdo Dry Valleys
(a) 8.0
300
24
290
Height (km)
6.0
18
12
4.0
280
2.0
6
RSR
0
CR
AR
KU
FG
20
0.0
40
D
60
TV
80
100
D’
Distance (km)
(b) 6.0
36
290
30
24
Height (km)
4.0
18
12
280
2.0
6
0
-6
KU
0.0
20
E
40
FG
60
Distance (km)
TV
E’
Figure 8. (a) Polar WRF vertical cross-section of potential temperature (K,
contours, 1 K interval), wind speed (m s−1 , shaded) and circulation vectors
along transect D–D’ in Figure 7(a) at 1800 UTC 29 December 2006: RSR,
Royal Society Range; CR, Cathedral Rocks; FG, Ferrar Glacier; KU, Kukri
Hills; TV, Taylor Valley; AR, Asgard Range. Reference vectors in lower
left represent 24.5 m s−1 in the horizontal and 8.65 m s−1 in the vertical.
(b) Same as (a) except along transect E–E’ in Figure 7(e) at 0300 UTC
1 January 2007, where FG is Ferrar Glacier, KU is Kukri Hills and TV is
Taylor Valley, and reference vectors in lower left represent 31.0 m s−1 in
the horizontal and 12.3 m s−1 in the vertical.
turbulent hydraulic jump just upstream of the Asgard Range.
This feature was interpreted incorrectly as a katabatic jump
by Doran et al. (2002a). Farther downstream, flow separates
from the surface, and low-level wave breaking occurs over lee
slopes of the Asgard Range (Wright Valley) and the Olympus
Range (Victoria Valley). Wave breaking is enhanced over
such steep lee slopes (Miller and Durran, 1991), and lowlevel wave breaking occurs with slow moving flow near
the ground (Jiang and Smith, 2003), as occurs over the
windward slopes of the Asgard and Olympus Ranges. Wave
breaking over the MDVs is supported by McKendry and
Lewthwaite (1990), who observe stagnant flow at ridge
level above Wright Valley during foehn conditions. Wave
amplitudes are dampened above low-level wave breaking
regions.
Figure 7(a) shows that flow is weak in the lee (north)
of the Royal Society Range, creating what appears to be a
c 2012 Royal Meteorological Society
1625
wake region that extends north into eastern Taylor Valley.
Figure 8(a) shows cross-section D–D’ east of the gap at
1800 UTC 29 December. Upstream flow is clearly blocked
when it encounters the ∼ 2.5 km high Royal Society Range, as
evidenced by abruptly lower downstream wind speeds above
the ∼ 1 km thick surface layer. Blocked flow results in a
lower effective mountain height, decreasing wave amplitude
(Jiang et al., 2005). Low-level wave breaking occurs in the
lee, and the layer of strong flow thins to below 2 km AGL. A
hydraulic jump occurs in the lee of Cathedral Rocks (CR),
and only weak wave effects exist downstream. Thus, the
eastern MDVs do not directly experience strong downslope
winds, except in rare instances when the ambient flow turns
more easterly (which will be discussed later in this section).
In contrast to the case for flow over a high ridge shown in
Figure 8(a), the gap flow that reaches western Taylor Valley
(Figure 7b) avoids Bernoulli loss associated with hydraulic
jumps and low-level wave breaking over lee slopes (Pan and
Smith, 1999; Gaberšek and Durran, 2004).
The flow set-up changes just 24 h later, as shown in
Figure 7(c). A strong cyclone offshore of Marie Byrd Land
sets up southwesterly flow over the Ross Ice Shelf. With a
large-scale westerly wind component, the nondimensional
mountain height at 1800 UTC 30 December is 0.78. This
means that nonlinear effects such as flow stagnation and
wave breaking will not be as significant as 24 h earlier.
Figure 7(d) shows a cross-section of potential temperature
and wind speed along transect C–C’ at 1800 UTC
30 December. The near-surface flow at the upstream end
of the cross-section is faster than 24 h earlier. Over Ferrar
Glacier (FG) and Taylor Valley (TV), trapped waves extend
up to about 7 km ASL, underneath a layer of strong wind
speed. Near-surface flow stagnates just upstream of the
Asgard Range, resulting in increased nonlinearity, and strong
leeside flow north of the Asgard Range occurs beneath a wave
breaking region above 4 km ASL. Low-level wave breaking
is once again present over Wright and Victoria valleys.
Note that flow over western Wright and Victoria Valleys
originates west of the primary gap, but mountain wave
dynamics are essentially the same. Due to the westerly largescale wind component, key differences from 24 h earlier are
stronger ambient flow, less upstream blocking, less upstream
low-level wave breaking, and amplified wave activity aloft.
A hydraulic jump does not form upstream of the Asgard
Range, and with a more westerly wind component, strong
winds reach farther down valley. Observed wind speeds
at this time are the strongest of the event at most sites
(Figure 4).
Through 31 December, the ambient wind direction turns
southerly and then southeasterly, as the cyclone that was
located near Marie Byrd Land moves westward and then
north into the Ross Sea (Figure 3). By 0300 UTC 1 January,
strong wind speeds exist in the gap and in the lee of
several terrain features, with large-scale southeasterly winds
and southerly flow through the gap (Figure 7(e)). Due to
the strong winds, the upstream non-dimensional mountain
height is only 0.46, indicating minimal nonlinear effects.
Figure 7(f) shows a cross-section at 0300 UTC 1 January. A
deep (almost 3 km) layer of strong winds (>25 m s−1 )
extends downstream from the gap. The strongest total
wind speeds through the gap occur when the ambient
(geostrophic) flow is directed oblique to the gap axis,
pointing to the importance of the along-gap pressure
difference, which at this time is enhanced by blocked flow
Q. J. R. Meteorol. Soc. 139: 1615–1631 (2013)
1626
D. F. Steinhoff et al.
upstream of the MDVs and an ambient pressure gradient
directed across the MDVs (Figures 6(a) and 7(e)).
The considerable vertical wind shear exhibited in
Figure 7(f), from terrain-induced southerly flow near the
surface to ambient easterly flow aloft, results in a critical
layer, where the along-gap wind speed goes to zero. Negative
vertical shear promotes wave breaking, and the resulting
critical layer effectively traps wave energy below. The
critical layer acts as a dividing streamline between strong
winds below and stagnant air in the wave-breaking region
(Smith, 1985), in a manner analogous to a free surface in
shallow-water hydraulic theory (Durran and Klemp, 1987;
Bacmeister and Pierrehumbert, 1988). In Figure 7(f), the
critical layer occurs above 7 km ASL, and forces wave
breaking and a stagnant region that descends to about 5 km
ASL over the gap. The layer of strong near-surface winds
remains coherent (no hydraulic jumps) across the transect,
and a layer of weak stability between 2 and 4 km ASL traps
wave energy near the surface over the MDVs, with only
weakly propagating waves aloft.
From Figure 7(e), with an easterly component to the
large-scale forcing, strong winds directly reach the eastern
MDVs. To examine the vertical structure of the flow
into the eastern MDVs, Figure 8(b) shows a transect of
potential temperature and wind speed along transect E–E’ at
0300 UTC 1 January. Flow accelerates through this gap into
Ferrar Glacier (FG), where a wave-steepening region forms
above 3 km ASL. The flow decelerates into Taylor Valley
(TV), with trapped waves above the narrow ridges between
valleys. Wind speed variations in eastern Taylor Valley result
from wave characteristics over the Kukri Hills, particularly
the placement and strength of hydraulic jumps. Polar WRF
captures the shift to southerly winds at Lake Fryxell, and
also at Lake Vanda and Lake Vida, but wind speeds are too
strong (Figure 4). The large positive wind speed bias may
result from wave dynamics over the smoothed terrain. To
the west, over higher portions of the Kukri Hills, the flow
separates from the surface over the lee slope. As a result, Lake
Hoare and Lake Bonney experience variable wind directions
with low wind speeds that Polar WRF correctly simulates
(Figure 4).
To summarize, the mountain wave patterns across the
gap and into the MDVs vary with upstream ambient wind
direction (similar to Zängl (2003) for Wipp Valley in
the Alps). When the ambient wind direction is southerly
(Figure 7(b)), blocking along the Transantarctic Mountains
upstream of the MDVs leads to lower near-surface wind
speeds, greater nonlinearity and low-level wave breaking.
Hydraulic jumps occur upstream of Taylor Valley, leading
to lower observed wind speeds at most MDVs sites. When
the ambient wind direction is more westerly (Figure 7(d)),
upstream blocking and nonlinearity decrease, meaning less
low-level wave breaking and greater vertical propagation of
wave activity. Stronger winds result at most MDVs sites. For
an easterly component to the ambient flow (Figure 7(f)),
gap flow forcing is strong, a critical layer forms aloft (which
is primarily responsible for wave amplification) and foehn
directly reaches the eastern MDVs.
4.3.
Pressure-driven channelling and easterly intrusions
Over eastern Taylor Valley, flow is dictated by the pressure
gradient force, described as ‘pressure-driven channelling’
(e.g. Whiteman and Doran, 1993; McGowan et al., 2002).
c 2012 Royal Meteorological Society
(a)
*
*
(b)
*
*
Figure 9. (a) Polar WRF sea-level pressure (hPa, shaded), terrain height
above sea level (m, contours, interval 150 m) and 3 m wind vectors over Ross
Island area at 1300 UTC 30 December 2006. Green dots represent LTER
AWS sites, black asterisks represent downvalley pressure gradient calculated
in Figure 10. (b) Same as (a) except at 1800 UTC 30 December 2006.
Winds are forced by the component of the pressure gradient
along the valley axis, with the force balance best described
as antitriptic. Because of the wake zone formed by the
Royal Society Range, downward momentum transport from
mountain waves would not be of substantial influence in
Taylor Valley. Figure 9(a) and (b) shows localized regions
of high pressure at the western end of Taylor and Wright
Valleys, associated with mountain wave pressure drag across
the adjacent slopes, which force the downvalley flow.
Foehn events are often interrupted for brief periods,
especially at eastern MDVs sites, by easterly intrusions.
These intrusions lead to a decrease in temperature, increase
in relative humidity and usually a decrease in wind speed.
Easterly intrusions have previously been attributed to either
a sea-breeze circulation (summer) or localized inversion
effects (winter) by Nylen et al. (2004). Here we present
evidence of easterly intrusions formed by blocking effects of
Ross Island and the Antarctic coast.
During periods of strong southerly flow along the western
Ross Ice Shelf – generally associated with passing synoptic
systems such as that depicted here (Figure 3) – Ross Island
exerts considerable control on the low-level flow in the
surrounding region (Slotten and Stearns, 1987; O’Connor
and Bromwich, 1988; Seefeldt et al., 2003; Steinhoff et al.,
2008). The LTER AWS observations at Lake Fryxell and Lake
Brownworth show easterly intrusions between 0900 UTC
and 1300 UTC 30 December. Polar WRF also has a
representation of easterly intrusions at this time. Figure 9(a)
shows easterly flow exceeding 8 m s−1 over McMurdo Sound
Q. J. R. Meteorol. Soc. 139: 1615–1631 (2013)
Foehn Mechanism in the McMurdo Dry Valleys
1627
by the Wilson Piedmont Glacier to the east and mountain
wave activity over the Asgard Range to the south. Doran
et al. (2002a) note events where strong winds occur at
Lake Brownworth with calm conditions at Lake Vanda.
They attribute this behaviour to foehn ‘riding over’ Lake
Vanda, possibly associated with cold-air pooling. Such an
argument is not valid during summer, and instead weak
flow at Lake Vanda results from an upstream hydraulic
jump, with pressure-driven channelling bringing strong
winds downvalley to Lake Brownworth.
5. Temperature variability
Figure 10. Polar WRF reduced pressure difference (hPa) between the two
asterisks in Figure 9(a) and (b), hourly from 0600 UTC to 1800 UTC
30 December 2006. Positive values represent pressure higher offshore than
in Taylor Valley.
at 1300 UTC, resulting from a mesoscale high-pressure
region over Windless Bight, as the southerly flow is forced
around Ross Island. As this easterly flow approaches the
Antarctic coast, it is also blocked, resulting in increased
surface pressure along the shoreline (Figure 9(a)). Flow in
eastern Taylor Valley is light and variable (Lake Fryxell in
Figures 4 and 9(a)).
Figure 10 shows the downvalley pressure gradient between
the two points indicated in Figure 9(a). This pressure
gradient increases from slightly negative at 0600 UTC to
3.4 hPa by 0900 UTC and 2.5 hPa at 1300 UTC. The
increase of this downvalley pressure gradient coincides with
the observed easterly intrusion, indicating that the increased
sea-level pressure offshore is responsible. One interesting
aspect of this case is that the same synoptic system that
forces the strong southwesterly downslope winds in Taylor
Valley also forces the easterly upslope intrusion.
Later in the day on 30 December, ambient flow shifts
westerly, and foehn is again recorded at all LTER AWS
stations. Flow around Ross Island is directed primarily to
the east, with a negligible high-pressure region over Windless
Bight, and weak flow over McMurdo Sound (Figure 9(b)).
The high-pressure region along the Antarctic coast is also
weaker. The downvalley pressure gradient is small after
1400 UTC (Figure 10). With no significant positive pressure
perturbations to the east, pressure-driven channelling results
in southwesterly flow down the entire Taylor Valley and
offshore, weakening over a convergence zone over McMurdo
Sound (Figure 9(b)).
The preceding analysis shows the influence that terraininduced pressure perturbations have on the MDVs. The
easterly intrusion illustrated here occurs around 2200 hours
local time and extends well into the night, so that even
just after the summer solstice, sea-breeze effects would be
expected to be weak. During periods of strong dynamic
forcing and foehn conditions, there is no distinct diurnal
wind cycle (McKendry and Lewthwaite, 1992). While the
example here focuses on Taylor Valley, similar arguments
can be made for Wright Valley, where flow in the western
sector of the valley is blocked by the Olympus Range.
However, the pressure-driven channelling is complicated
c 2012 Royal Meteorological Society
From Figure 4, the temperature at all LTER AWS sites
increases during the case study period, independent of the
diurnal cycle. Maximum observed temperatures occur early
2 January (UTC time), which rules out a stronger foehn
effect being responsible for warming, because the foehn
event was weaker at that time than a day earlier. Instead,
warmer conditions result from the advection of warmer
air into the region. To show this, we calculate air parcel
trajectories backwards from a point near the MDVs at 3 km,
4 km and 5 km ASL for 84 h using the 32 km domain. Foehn
descent and the vertical coherency of the flow between the
surface and mid-troposphere (Figure 7) motivate our choice
of trajectory heights.
Trajectories backwards from 1800 UTC 29 December
are shown in Figure 11(a). At all three levels, trajectories
originate over the central Ross Sea and traverse the Ross
Ice Shelf into the MDVs. Figure 11(b) shows time series
of several variables along the 5 km ASL trajectory. This
trajectory rises from about 3 km ASL at the beginning, and
potential temperature drops from 287 K to 284 K, probably
from mixing of colder air. Equivalent potential temperature
is slightly warmer at the beginning, at 290 K, associated
with the maritime environment of the Ross Sea. The
equivalent potential temperature approaches the potential
temperature over the western Ross Ice Shelf, as precipitable
water decreases. In contrast, Figure 11(c) and (d) shows
trajectories and time series of the 5 km trajectory ending at
0600 UTC 1 January. The 3 km trajectory originates in the
Amundsen Sea, but the 4 km and 5 km trajectories originate
north of the Ross Sea. All trajectories arrive over the MDVs
from due east. The 5 km trajectory rises from about 1.1 km
ASL, and is warm and moist, based on equivalent potential
temperature of about 301 K being almost 15 K warmer than
the potential temperature, and precipitable water values
of about 18 mm. Unlike the previous trajectory, potential
temperature increases markedly along the transect, when
the air parcel traverses Marie Byrd Land and rises moist
adiabatically, evidenced by the near-constant equivalent
potential temperature and decreasing precipitable water.
At the terminus, the potential temperature is 294 K, 10 K
warmer than the other trajectory.
Just upstream of the MDVs, temperature differences
between these two times are restricted to the 3–7 km
ASL layer, and air at the surface in Taylor Valley appears
to originate from the 2–2.5 km ASL layer upstream of the
MDVs (Figure 7(b) and (f)). However, large values of TKE
in the 3–5 km ASL layer above Knobhead, Ferrar Glacier,
and Taylor Valley on 1 January (Figure 7f) indicate that
vertical turbulent mixing aids in transporting air in the
elevated warm layer towards the surface (e.g. Zängl et al.,
2004a). In contrast, no elevated warm air layer and only
Q. J. R. Meteorol. Soc. 139: 1615–1631 (2013)
1628
D. F. Steinhoff et al.
(a)
(c)
(b)
(d)
Figure 11. (a) Polar WRF 32 km domain backward trajectories from 1800 UTC 29 December 2006 for 84 h at 3 km, 4 km and 5 km, with arrows plotted
every 24 h. (b) Potential temperature (K, red), equivalent potential temperature (K, blue), column-averaged precipitable water (mm, green) and height
(m, black) along the 5 km trajectory. ((c) and (d)) Same as (a) and (b) except at 0600 UTC 1 January 2007.
weak vertical turbulent mixing exist upstream of the MDVs
on 29 December (Figure 7(b)). The 3 K increase in potential
temperature on 1 January along the transect in Figure 7(f) to
Taylor Valley largely explains the 4–5◦ C increase in surface
temperature between 29 December and 1 January (Figure 4).
By 2 January, with only weak mountain wave effects, a deeper
layer of warm air results in surface temperatures exceeding
+7◦ C at several sites (not shown). The warmest foehn
conditions are associated with the advection of warm, moist
maritime air towards the MDVs, which gains enthalpy
as it ascends moist adiabatically and then descends dry
adiabatically as foehn into the MDVs. Turbulent vertical
mixing below wave breaking layers is effective in bringing
warm air aloft towards the surface.
6.
Conclusions
A case study from December 2006 to January 2007 is used to
illustrate the physical mechanisms associated with a MDVs
foehn event. Foehn events initiate over the western MDVs,
typically associated with southerly flow. This southerly flow
is a gap wind, forced by reduced pressure differences across
an elevated gap just south of the MDVs between the East
Antarctic ice sheet and the Royal Society Range. The alonggap pressure differences are set up by flow being blocked
by the Antarctic continent (for easterly flow) or the Royal
Society Range (for southerly and westerly flow). Mountain
c 2012 Royal Meteorological Society
waves are forced over the elevated gap and nearby terrain,
for which the nonlinearity and other characteristics of the
flow are determined by the incident ambient wind direction.
Southerly flow is blocked upstream of the gap, resulting in
higher nonlinearity and increased low-level wave breaking.
Flow with a westerly off-continent component features
vertically propagating mountain waves with stronger wind
speeds over the MDVs. Easterly flow features a strong alonggap pressure difference and a mean-state critical layer in the
upper troposphere that changes the wave dynamics. Except
for large-scale easterly flow, foehn does not have a direct
impact on the eastern MDVs, and instead pressure-driven
channelling from blocked flow to the west brings the warm,
dry air downvalley.
We believe that several generalizations about MDVs
foehn can be made based on this case. The MDVs foehn
can thus be explained through a sequence of mesoscale
processes – gap flow forces air into the MDVs region,
mountain waves accelerate the flow at low-levels and
pressure-driven channelling brings the warm, dry air down
the valleys. Easterly intrusions during foehn events are
independent of the diurnal cycle, and primarily result from
flow deflection around Ross Island, as localized pressure
gradients counter pressure-driven channelled westerly flow
over the MDVs. Just as the synoptic-scale circulation is
largely responsible for the flow characteristics over the
MDVs, temperature variability during foehn events depends
on large-scale advection and the source region of the flow.
Q. J. R. Meteorol. Soc. 139: 1615–1631 (2013)
Foehn Mechanism in the McMurdo Dry Valleys
Since the existence and the characteristics of MDVs foehn
depend on the synoptic-scale circulation, generalization of
the case study presented here to MDVs climate rests on the
incidence of foehn and cyclonic forcing over the Ross Sea
and Ross Ice Shelf. This association has already been made
by Speirs et al. (2010), who show that MDVs foehn occurs
concomitantly with synoptic-scale cyclones to the east, and
that MDVs foehn frequency increases with an increase in the
climatological pressure gradient over the Ross Ice Shelf. The
implication of an amplified synoptic-scale circulation for
anomalously warm conditions over the MDVs is not unique
to this region of Antarctica, as shown for a prominent warm
event in January 2002 by Turner et al. (2002) and Massom
et al. (2006). The importance of an amplified circulation for
anomalously warm foehn conditions is also shown for the
Alps by Zängl and Hornsteiner (2007).
The positive association between summer foehn and
meltwater generation in the MDVs is well-established
(Fountain et al., 1999; Doran et al., 2002b, 2008). However,
as MDVs melt depends primarily on incident solar radiation
and low wind speeds to reduce sensible heat flux loss
from the surface (Clow et al., 1988; Hoffman et al., 2008),
there is a paradox as to how melt increases during foehn
conditions, which are gusty, and associated with synopticscale cyclones that bring cloudy conditions. First, foehn
brings above-freezing temperatures, which typically do not
occur during cloud-free quiescent conditions due to the seabreeze circulation. Second, adiabatic warming associated
with foehn leads to drying that will clear some cloud cover
over the MDVs (and promotes sublimation too). Even with
cloud cover, there are instances when downward long-wave
heat fluxes can exceed what downward solar (short wave)
forcing would be without clouds (Hoffman et al., 2008),
presumably during twilight conditions. However, it is likely
that the greatest melt occurs during foehn events, but when
wind speeds are low, associated with a hydraulic jump
or along the boundary between channelled foehn and an
easterly intrusion. Another favourable time is when warm
air stagnates over the MDVs after a foehn event, in the order
of a few days (Nylen et al., 2004; Speirs et al., 2010).
Verification of the concepts put forth here by observations
is still necessary, as the limited surface observations and
the observational study of McKendry and Lewthwaite
(1990, 1992) are not sufficient. A temporally limited field
program with vertical wind profiling at a few locations
could aid in describing the gap flow and mountain wave
structure. Additional surface observations of atmospheric
pressure within the MDVs would verify the importance
of pressure-driven channelling to the downvalley extent
of warming. Improved numerical simulations, especially
higher horizontal and vertical resolution and more realistic
boundary layer processes, will also aid in the understanding
of MDVs’ meteorology.
Acknowledgements
This research is supported by the National Science Foundation via NSF Grant ANT-0636523. High-performance
computing support is provided by NCAR’s Computational
and Information Systems Laboratory, sponsored by the
National Science Foundation, and the Ohio Supercomputing Center. The McMurdo Dry Valleys LTER Program
provided automatic weather station data. Kevin Manning
provided the bare-ground land-use data for the MDVs.
c 2012 Royal Meteorological Society
1629
Comments by two anonymous reviewers improved the
manuscript. NCAR is sponsored by the National Science
Foundation and other agencies
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