Local and non-local controls on a persistent cold-air pool in the
Arve River Valley
G. Arduinia,b,∗, C. Chemela,c and C. Staquetb
aCentre for Atmospheric & Instrumentation Research, University of Hertfordshire, College Lane, Hatfield, AL10 9AB, UK
bUniversite´ Grenoble Alpes, CNRS, LEGI, F-38000, Grenoble, France
cNational Centre for Atmospheric Science, Centre for Atmospheric & Instrumentation Research, University of Hertfordshire, College
Lane, Hatfield, AL10 9AB, UK
∗Correspondence to: G. Arduini, European Centre for Medium-Range Weather Forecasts, Reading, United Kingdom E-mail:
gabriele.arduini@ecmwf.int
A numerical model is used to simulate a persistent cold-air pool (PCAP) event that
occurred in the section of the Arve River Valley around Passy in the French Alps.
During this period, an upper-level ridge from the Atlantic moved over Europe, allowing
a PCAP to form and persist over time. The impact of the upper-level ridge on the PCAP
and on the dynamics within the valley section is quantified, by examining the mass and
heat budgets of the valley atmosphere. During the persistent stage, the magnitude of
the flow through the tributary valleys is enhanced by the large-scale flow. Also, the
direction of the flow through one of the tributaries is found to be determined by the
height of the PCAP with respect to that of the tributary above the valley floor. The
tributary flows, together with subsiding motions at the valley top, control by and large
the nighttime valley-scale circulation and the thermal structure of the upper part of the
PCAP, whereas thermally-driven valley flows control its lower part. When the upper-
level ridge passes over the Arve River Valley, warm air advection through the tributaries
continuously erodes the upper part of the PCAP during nighttime, thereby reducing
its depth, while down-valley flows export the air mass out of the valley. As the ridge
moves away from the valley, the near-surface air is found to be trapped within the
valley. This trapping results from the advection of warm air in the upper part of the
PCAP by the large-scale flow channelled through one of the tributary. This reduces the
thermally-induced pressure difference in the down-valley direction, thereby suppressing
the near-surface down-valley flow. The study therefore highlights the interplay between
the large-scale flow, the tributary flows and the thermal structure of the PCAP.
Key Words: PASSY-2015; cold-air pool; complex terrain; stable boundary layer
1. Introduction
A cold-air pool (CAP) is a stably stratified layer of air confined at
the bottom of a valley or basin. Diurnal CAPs form frequently
during nocturnal hours under clear-sky and weak synoptic
wind conditions and are destroyed following sunrise mainly by
convection resulting from surface heating. When relatively warm
air associated with anticyclonic conditions is transported right
above deep terrain over multiple days (Reeves and Stensrud
2009), convection is often insufficient to destroy the stable layer
within the valley, particularly during wintertime (see for instance
Vrhovec and Hrabar 1996; Lu and Zhong 2014), allowing a
persistent cold-air pool (PCAP) to develop. Since a PCAP is
characterised by a long-lasting (multi-day) valley inversion and
weak winds, air pollutants and moisture emitted within the PCAP
are trapped and accumulate from day to day, thereby affecting
air quality (Silcox et al. 2012; Chemel et al. 2016; Largeron
and Staquet 2016a) and visibility when fog forms. The removal
of the PCAP may occur because of the passage of a synoptic
disturbance over the valley (Zhong et al. 2001), turbulent erosion
of the PCAP from the top (Lareau and Horel 2015a) or the
displacement of the cold-air layer out of the valley (Flamant et al.
2006; Lareau and Horel 2015b). While the formation, persistence
and destruction stages are common to PCAP in many different
regions and conditions (Zhong et al. 2001; Za¨ngl 2005a; Lareau
et al. 2013; Lu and Zhong 2014; Largeron and Staquet 2016b),
the characteristics of the PCAP (depth and strength) during each
stage depend on the characteristics of the surrounding terrain and
variability of the synoptic conditions. Thereafter a distinction is
made between the processes that control the characteristics of
the PCAP and are governed by those of the terrain at the valley
or basin scale, which are referred to as local controls (such as
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This article has been accepted for publication and undergone full peer review but has not been through the copyediting, typesetting, pagination and proofreading process which may lead to differences between this version and the Version of Record. Please cite this article as doi: 10.1002/qj.3776 
 
local thermally-driven flows), and processes affecting the PCAP
characteristics but occurring at scales larger than that of the valley
(such as the synoptic wind above the valley boundary layer and
flows from surrounding tributary valleys modified by the large-
scale flow), which are referred to as non-local controls.
Previous work has considered local controls on CAPs. Vrhovec
(1991) showed using results from numerical simulations of a
diurnal CAP in the Slovenj Gradec basin that a dynamical
decoupling can take place between the flow within the valley
boundary layer and the synoptic flow aloft. As a result, at the
top of a CAP an abrupt change in wind speed and direction was
observed. In such decoupled condition thermally-driven along-
slope and along-valley flows are the key dynamical features of
the valley atmosphere (Largeron and Staquet 2016a). However,
much remains to be understood on their role in the formation
and evolution of CAPs (Bodine et al. 2009). In a numerical
modelling case study of the formation of a CAP in a ∼ 100-m
deep narrow valley, Vosper et al. (2014) concluded that the
principal mechanism responsible for the formation of the CAP
was the sheltering effect of the valley. The sheltering was shown
to increase the turbulent heat flux divergence close to the ground
compared to a less sheltered region (i.e. hilltop sites), thereby
increasing near-surface cooling and promoting the formation of
a ground-based inversion. However, in such a shallow valley,
thermally-driven flows are rapidly suppressed after the evening
transition (Clements et al. 2003; Vosper et al. 2014). In a deep
and narrow valley dynamically decoupled from the atmosphere
above, downslope flows can persist for a longer time, reaching
a quasi-steady state if a down-valley flow can develop during
the night (Burns and Chemel 2015; Arduini et al. 2016). Under
these conditions, downslope flows are the main driver of the
sensible heat flux divergence, which enhances the cooling of the
atmosphere in the valley with respect to a flat region nearby.
Downslope flows are also important in the evolution of the thermal
structure of the upper part of a CAP, through the vertical transport
induced by the air flowing down the slopes (Whiteman 1982;
Kiefer and Zhong 2011; Burns and Chemel 2015; Arduini et al.
2016).
In addition to thermally-driven flows, the variations of the
orography and surface properties affect the evolution of a CAP.
Because of the reduction in the volume of atmosphere in a valley
or basin when compared to that over a flat terrain, a valley or
basin experiences a cooling that is stronger than if it were a flat
terrain for the same amount of heat loss (Whiteman 1990). A
differential cooling can also take place between different sections
of the same valley because of along-valley geometrical variations
(McKee and O’Neal 1989). This differential cooling can lead
to along-valley flows that are up-valley instead of down-valley
during the night (Arduini et al. 2017), affecting the mass and
heat budgets of the valley atmosphere. If the ground is covered by
snow (as is common in winter in alpine regions), surface albedo is
increased and the heat conductivity to the ground is significantly
decreased, leading to stronger ground-based inversions (Za¨ngl
2005b; Billings et al. 2006; Neemann et al. 2015).
Non-local controls on PCAPs have been examined for a number
of case studies. Zhong et al. (2001) showed that the formation
and maintenance of a PCAP event in the Columbia basin in the
Pacific Northwest region of North America was primarily the
result of the continuous advection of warm air immediately above
the basin by strong westerly winds descending the lee slopes of
the surrounding Cascade mountains. Similar results as regards
the importance of large-scale advection of warm air above the
terrain in the formation and maintenance of a PCAP in the Salt
Lake Valley, Utah, USA, were reported by Wei et al. (2013). Lu
and Zhong (2014) quantified the respective contributions of the
different terms determining the heat budget of the atmosphere
during this PCAP event. Large-scale advection was found key to
the formation, maintenance and destruction of the PCAP.
While these studies have demonstrated the essential role of
warm air advection in the formation and maintenance of a PCAP,
the dynamical controls of large-scale flows on the temporal
evolution of CAPs have received less attention. In a numerical
modelling case study Za¨ngl (2005a) showed the importance of
the direction of the synoptic pressure gradient and of the along-
valley terrain geometry, with respect to the Alpine foreland,
in the maintenance of CAPs in a deep Alpine valley system.
Model results indicated that the direction of the synoptic pressure
gradient controlled the direction of the along-valley drainage
flows. The author concluded that, for that particular case study, the
ideal conditions for the removal of a cold-air pool are met when
the synoptic pressure gradient is orientated down a constriction
in the along-valley direction. In a following work based on semi-
idealised numerical simulations and observations, Za¨ngl (2005c)
extended this analysis to the shallower Bavarian Danube basin.
Model results showed that advection resulting from the synoptic
pressure difference (pressure-driven channelling) and advection
down the direction of the synoptic wind (forced-channelling)
can both act to remove the PCAP, depending on the direction
of the synoptic wind with respect to the orientation of the
highest mountain range. Another dynamical mechanism that may
affect the temporal evolution of PCAPs is shear-induced turbulent
mixing at the top of the PCAP due to strong winds aloft. Some
studies questioned the effectiveness of such turbulent erosion for
the destruction of CAPs over time-scales shorter than one day (see
for instance Zhong et al. 2001, 2003). However, semi-idealised
numerical simulations analysed by Za¨ngl (2005c) suggested that
this mechanism can contribute significantly to the destruction of
a CAP in the relatively shallow Danube basin. In deep valleys,
idealised simulations by Lareau and Horel (2015a) indicated that
CAPs can be removed by turbulent erosion on a time-scale of less
than a day, provided that the speed of the flow aloft increases
continuously with time such that the Froude number exceeds a
threshold value of about 2. A more systematic study on the link
between the size of the valley and the synoptic conditions have
been performed by Sheridan (2019) by means of idealised 2D
simulations. The author showed that for deep ( 1000 m-deep)
valleys embedded within a plateau, for increasing background
stability, the CAP becomes increasingly more sheltered from
the external flow and so more prone to being persistent. This
behaviour was found to be well described by a non-dimensional
valley depth parameter (Vosper and Brown 2008). However, in
the case of a valley embedded between two hills, a more common
situation in the Alpine environment, the relationship between the
non-dimensional valley depth and CAP persistency was found
more complex, and a very strong background stability (in a real
case associated for instance with synoptic warm air advection)
was required to maintain the CAP persistent in time.
The contribution of tributary flows in the mass budget of a
valley was studied extensively as part of the Advanced Studies
in Complex Terrain (ASCOT) programme (Clements et al. 1989).
However, the computation of mass fluxes through tributaries relied
solely on limited observations, and so several assumptions and
approximations had to be made to estimate their contributions
to the mass budget. Porch et al. (1989) calculated that the mass
flux through a single tributary may account for 5 to 15% of the
drainage out of Brush Creek Valley, Colorado, USA. Assuming a
similar contribution from all the tributaries of Brush Creek Valley,
the authors estimated that tributary flows control by and large the
mass budget of the valley, and so that the contribution of slope
flows and subsiding motions at the valley top are only minor.
Numerical model simulations reported by O’Steen (2000) for an
idealised valley-tributary system under dynamically decoupledThis article is protected by copyright. All rights reserved.
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Figure 1. (a) Topographic map of domain d02 with the positions of the nested domains d03, d04 and d05, represented by continuous-line boxes. The dashed-line box
represents the region covered by the semi-idealised simulation, WRF-I. (b) Topographic map of the innermost domain d05; the dashed black polygon indicates the
horizontal extent of the PASSY control volume, the white line indicates the path of cross-sections used in the paper; mass and heat fluxes are calculated across ‘gates’
represented by the blue lines; the position of selected sites are labelled in bold: site radio soundings (RS), downstream part of the valley (DS), Mege`ve (MGV), Saint-
Gervais-les-Bains (STGV), Chamonix (CHMNX) and Marnaz (MRNZ). (c) Terrain height along the polygon defining the PASSY control volume; the dotted line indicates
the mean height of the terrain along the polygon; the dashed-dotted line is the height of the RS site; see text for details. (d) Snow height in the innermost domain d05 on
8 February 2015 at 1200 UTC; the position of selected sites are labelled in bold as in (b).
conditions confirmed that the flow through the tributary can
increase the along-valley mass flux by 5 to 10% compared to an
idealised valley without a tributary. However, some of the complex
interactions between the tributary flow and the down-valley flow
were not captured with this idealised configuration. Coulter
et al. (1989) showed that the mass flux through the tributary is
extremely sensitive to the external flow. Under particular synoptic
conditions the drainage from the tributary was found 20% larger
than that expected if the mass flux was assumed proportional
to the drainage area. Coulter et al. (1991) examined the mass
fluxes through three tributaries of Kimball Creek, Colorado,
and showed that the mass flux is largest for the tributary most
closely aligned with the valley axis. Despite the efforts led in
the ASCOT programme, large uncertainties remain about the
effects of tributary flows on the heat and mass budget of a valley
atmosphere.
In this paper, we investigate the role of local and non-local
controls, the latter including the synoptic wind above the valley
boundary layer and flows from the surrounding tributary valleys
affected by the large-scale flow, in the evolution of a PCAP using
numerical model simulations. The PCAP event occurred in the
section of the Arve River Valley around Passy in the French Alps,
in February 2015, during the first intensive observation period
(IOP1) of the PASSY-2015 field campaign (Staquet et al. 2015;
Paci et al. 2016). To reach this aim, the mass and heat budgets
in the Passy valley are analysed and compared to the ones for
a semi-idealised simulation in which the synoptic flow in the
initial condition is at rest. In Sect. 2, we present the design of
the numerical simulations. Section 3 provides an overview of the
synoptic conditions during IOP1. A detailed description of the
episode and the evaluation of model data with field observations
are reported in Sect. 4. Model results are analysed in Sect. 5 and 6.
A summary and conclusions are presented in Sect. 7.This article is protected by copyright. All rights reserved.
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2. Methodology
2.1. The Passy Valley
The Arve River Valley is located in the French Alps to the south
of the Swiss Plateau, a relatively flat area located between the
Jura mountain range and the Swiss Alps (see Fig. 1a). Hereafter
the Passy Valley refers to the section of the Arve River Valley
enclosed by the dashed black polygon in Fig. 1b. The mean height
of the terrain along the polygon defining the PASSY control
volume is zm = 1560 m a.s.l. (see Fig. 1c), and the highest massif
surrounding the Arve River Valley is Mont-Blanc, culminating
at 4808 m a.s.l. The valley-floor elevation decreases gradually
down-valley from 588 m a.s.l. at Passy [upstream of the radio
soundings site (RS)] to 488 m a.s.l. at Marnaz (MRNZ) over a
distance of about 23 km. The width of the valley floor along this
section of the valley increases gradually up-valley from just about
200 m upstream of MNRZ, where the valley presents a major
constriction, to approximately 2000 m near the location of the
RS site, and hardly decreases from there towards the escarpment
leading to Chamonix (CHMNX). In the context of this work, the
section of the Arve River Valley leading to CHMNX is considered
a tributary of the Passy Valley. Therefore, the Passy valley presents
three major tributaries leading to Mege`ve (MGV), Saint-Gervais-
les-Bains (STGV) and CHMNX (see Fig. 1c).
2.2. Terrain representation and snow initialisation
An accurate representation of the terrain is essential to numerical
modelling in complex terrain such as alpine valley systems.
Most alpine valleys are characterised by steep slopes that are
a challenge for numerical models formulated using terrain-
following coordinates, as is the case for the Weather Research
and Forecasting (WRF) model used here. Even using a horizontal
mesh size of the order of 3 km, slope angles exceed 25◦across
the Alpine chain of western Europe. With a horizontal mesh size
of 100 m, the maximum slope angle in the Arve River Valley is
about 75◦. In the present work, 90-m terrain elevation data from
the Shuttle Radar Topography Mission (SRTM, Farr et al. 2007)
were interpolated at the horizontal mesh size of the nested model
domains (see Fig. 1a). A smoothing filter was used to reduce
the largest slope angles of the terrain to a maximum slope angle
αmax = 45◦ while minimising any changes in the overall shape and
small-scale features of the valley geometry.
Several studies pointed out that the presence of snow
on the ground surface alters significantly surface–atmosphere
interactions (Za¨ngl 2005b; Tomasi et al. 2017; Neemann et al.
2015). The initialisation and modelling of the snow pack are
difficult tasks. Numerical model analysis products are currently
available at too coarse of a resolution (∼ 15 km) to be used
to provide an appropriate representation of the snow cover
at sub-kilometre scales. Furthermore, the initialisation of the
snow density performed by the WRF Preprocessing System
(WPS) assumes climatological values that are not representative
for a specific case study. In the present work, MODIS/Terra
(MOD10 L2) satellite products at a spatial resolution of 500 m
(Hall et al. 2006) were used to initialise snow cover and albedo.
The snow mask was directly interpolated at the horizontal grid
points of the nested model domains. To initialise snow albedo,
we followed the methodology outlined by Tomasi et al. (2017).
Snow albedo is a function of the age of the snow (see Livneh
et al. 2010), which is initialized as fresh snow (i.e. with the largest
value for snow albedo, based on an analysis of climatological
data, Robinson and Kukla 1985). A snow age corresponding
to the satellite-retrieved albedo was calculated by inverting the
functional relationship proposed by Livneh et al. (2010), which
is implemented in the current version of the Noah land-surface
Table 1. Spatial and temporal resolutions for the domains used for the real-
case numerical simulations: number of points nx, ny and nz in the east, north
and vertical directions, respectively, horizontal mesh size ∆x = ∆y, vertical
grid spacing near the ground surface ∆zmin, time step ∆t and frequency fbu at
which the lateral boundaries of the nested grids were updated.
Domain nx, ny, nz ∆x = ∆y ∆zmin ∆t fbu
d01 202, 202, 46 15 km 42 m 30 s 6 h
d02 246, 246, 46 3 km 42 m 6 s 30 s
d03 340, 340, 46 1 km 42 m 2 s 6 s
d04 406, 406, 92 333 m 21 m 0.6 s 10 min
d05 382, 382, 92 111 m 21 m 0.1 s 5 min
model. The calculated snow age was then used as initial condition.
The snow water equivalent was diagnosed by assuming that
the density of the snow is, as a zero-order approximation, only
function of the age of the snow (Meløysund et al. 2007). Hence,
given the snow depth, the snow water equivalent can be specified
at the initial time. Snow depth was assumed to depend only on
the elevation of the underlying terrain, as in previous studies
(e.g. Schmidli et al. 2009). Snow depth was initialised to increase
linearly with height from a value of 0.05 m at sea level to a value
of 0.15 m at 500 m above sea level (a.s.l.), and, from 500 m a.s.l.,
to a value of 1.5 m at 4500 m a.s.l. (see Fig. 1d).
2.3. Model setup
The numerical simulations were performed using the WRF model
with the advanced research core (ARW, Skamarock et al. 2008),
version 3.5.1. Five nested domains were used (see Table 1 for the
spatial and temporal resolutions that were used for the domains).
The simulations were performed in three steps.
Firstly, a simulation for the three outermost domains (d01,
d02 and d03) was performed using online one-way nesting and
a grid ratio of 5 and 3 for domains d02 and d03, respectively.
The computations were made on 46 vertical layers. The grid was
stretched along the vertical axis, with a vertical spacing of about
40 m near the ground surface (so that the first mass point above
the surface is approximately at 20 m). The use of a relatively
low vertical resolution in the outermost domains was required
to keep both the model numerically stable and the simulated
terrain height as close as possible to the real orography. Lateral
boundary conditions for d02 and d03 were updated every parent
domain time step. Initial conditions other than those for the snow
fields and lateral boundary conditions for the outermost domain
d01 were derived from the European Centre for Medium-range
Weather Forecasts (ECMWF) gridded analyses available every
6 h with a horizontal mesh size of 0.25◦. The simulation was
run continuously for 7 days from 7 to 14 February 2015. To
minimise the deviation of the model results from the ECMWF
analyses, spectral nudging was used for d01 and d02. Only the
upper part of domain d03 (above 6000 m a.s.l.) was nudged
towards the analyses. Snow cover and albedo were initialised
from MODIS/Terra (MOD10 L2) satellite products (see Sect 2.2)
averaged in time between 5 and 10 February 2015.
Secondly, a simulation for domain d04 was performed using
initial and lateral boundary conditions extracted from the results
for d03, using offline nesting. The extent of d04 was chosen to
cover the main regional-scale orographic feature surrounding the
Passy Valley (see Fig. 1a). The horizontal mesh size was set to
333 m and the number of vertical levels increased to 92, compared
to the outermost domains, with a vertical mesh size of about 20
m near the ground surface (so that the first mass point above the
surface is approximately at 10 m) and 29 levels in the lowestThis article is protected by copyright. All rights reserved.
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1000 m above the surface. The lateral boundary conditions were
updated every 10 min. This simulation was run continuously from
8 to 14 February 2015.
Finally, a simulation for domain d05 was performed using a
methodology similar to that used for domain d04. Differences
are that the lateral boundary conditions were updated every
5 min in order to minimise spurious effects at these boundaries.
Also, this simulation was reinitialised every day at 1200 UTC
to facilitate the comparison with semi-idealised simulations
(discussed below). The coverage of d05 is shown in Fig. 1b. The
number of vertical levels was the same as that of d04 but the
horizontal grid mesh size used was refined to 111 m. Results from
additional sensitivity simulations revealed negligible differences
when a domain size twice as large as that of d05 is used. Results
from this sensitivity study also showed that a horizontal mesh size
as fine as 111 m was required to resolve the constriction in the
downstream part of the Passy valley. The constriction is closed for
a coarser resolution which was found to significantly modify the
dynamics of the near-surface valley atmosphere (not shown).
The simulation for d05 is the reference simulation and is
referred to as WRF-R hereafter.
All nested domains were centred at the airport in Sallanches
(45.935◦N, 6.636◦E) and shared the same ‘dynamics’ and
‘physics’ options, apart from that for the planetary boundary-
layer (PBL) and subgrid-scale turbulence. A third-order Runge-
Kutta scheme was used to integrate the model equations and
a time-splitting technique was used to integrate the acoustic
modes. The advection terms were discretised using a fifth-order
Weighted Essentially Non-Oscillatory (WENO) scheme with
positive definite filter. The PBL was parameterised for d01, d02
and d03 using the Yonsei University (YSU) parameterisation
scheme (Hong 2010). In the higher-resolution simulations for d04
and d05 a 3D turbulent kinetic energy (TKE) diffusion scheme is
used (and no PBL scheme), modified to enable the gravitational
settling of fog, following the formulation proposed by Nakanishi
(2000). Land surface processes were modelled using the Noah
land surface model (Chen and Dudhia 2001) using four soil layers.
A few modifications of the land surface model from the WRF
model version 3.7 were implemented in the model used in the
present work. These are the computation of the ground thermal
conductivity, to prevent discontinous behavior during the melting
of frozen water in the soil; and the calculation of the latent heat
flux, to take into account dew deposition processes over snow
covered surfaces. The Monin-Obukhov similarity theory (Jime´nez
et al. 2012) is used to couple the land-surface to the atmosphere.
Shortwave and longwave radiation were parameterised using
the Rapid Radiative Transfer Model (Mlawer et al. 1997).
Shadowing effects were included to take into account the effect
of the mountains on the radiative fluxes. Microphysics was
parameterised using the scheme developed by Morrison et al.
(2005), modified for the treatment of ice fog following Neemann
et al. (2015). The number of cloud condensation nuclei was
increased to 500× 106, a value more representative of continental
environments.
A set of semi-idealised numerical simulations was performed
to extract the effects of the large-scale flow (simulation WRF-I).
WRF-I differs from WRF-R in that the domain is twice as large
(see Fig. 1a), and that at the initial time the temperature and
humidity fields are horizontally homogeneous and the wind speed
is set to zero. The use of a larger domain in WRF-I compared to
WRF-R was necessary to minimise spurious effects at the lateral
boundaries, at which the normal velocity is set to zero. Hence, the
main difference between WRF-I and WRF-R lies in the absence of
large-scale flow for WRF-I. The latter simulation was run for the
period between 10 February 2015 at 1200 UTC and 12 February
2015 at 1200 UTC and re-initialised on 11 February 2015 at 1200
UTC. The initial temperature and humidity profiles for WRF-I
were extracted from the real-case simulation for d04 at the centre
of the domain at the time at which WRF-I was initialised, namely
10 and 11 February 2015 at 1200 UTC. The skin temperature was
initialised by extrapolating the temperature of the first three air
layers above the ground surface.
3. Overview of the large-scale circulation during IOP1
The PCAP that formed in the Passy Valley during IOP1 was the
result of the passage of an upper level ridge over the region.
The geopotential height at 500 hPa showed a dipole pattern over
Europe between 2 and 8 February 2015, characterized by an upper
level ridge adjacent to a trough. The upper level ridge of the dipole
pattern was located immediately to the west of the British Isles on
9 February 2015 at 0000 UTC (see Fig. 2a). The PCAP formed
in the evening on 9 February 2015, when the upper level ridge
moved over Northwest Europe (see Fig. 2b). This synoptic flow
pattern led to the advection of warmer air above the western Alps,
setting the conditions for the PCAP to develop in the Passy Valley.
From 10 to 12 February 2015 the upper level ridge continued
to move across Europe and was centred over Central Europe on
12 February 2015. The displacement of the ridge resulted in a
weakening of the wind speed and a veering of the wind with time
over the western Alps from north-easterly on 9 February 2015 to
south-easterly on 12 February 2015 (see Fig. 2b to d), allowing
the CAP to persist over time. On 13 February 2015 the trailing
upper level trough moved across Northwest Europe, leading to
a southerly flow and an increase in wind speed over the western
Alps (see Fig. 2e). On 14 February 2015 a large-scale perturbation
reached the western Alps (see Fig. 2f), thereby removing the
PCAP and putting an end to IOP1.
4. Life-cycle of the persistent cold-air pool during IOP1
4.1. Vertical structure of the cold-air pool
Figure 3 shows the temporal evolution of the vertical structure of
the PCAP obtained by compiling potential temperature profiles
from the radiosonde ascents and wind profiles retrieved from
LiDAR and wind profiler data at the RS site during IOP1, and
extracted from the WRF-R simulation. The temporal evolution
is overall well captured in WRF-R. However discrepancies can
be noted, in particular before the PCAP event during the night
from 8 to 9 February 2015 and at the end of the episode on 14
February 2015 when the simulated temperature and wind speed
are overestimated above zm. Vertical profiles of the root-mean-
squared-error (RMSE) and bias of potential temperature and wind
speed between 9 February 2015 at 1200UTC to 14 February
2015 at 1200UTC are also displayed to quantitatively evaluate the
numerical results of the simulations. Results before 9 February at
1200UTC were not included in the computation of the statistics,
because the simulation in the outer domain d04 was not fully spun
up yet for these times. Statistics for the wind direction are not
displayed as at low wind speeds, the fluctuations of this quantity
are large, leading to large RMSE values (Chow et al. 2006). The
RMSE of the potential temperature averaged between the valley
floor and zm is about 1.6 K, decreasing above zm to about 1 K. The
value of the bias is−0.32 K averaged between the valley floor and
zm, indicating a slight systematic underestimation of the potential
temperature in the simulated PCAP above the valley floor. This
statistical analysis indicate that the model is able to represent
the temporal variability of the observed temperature. Regarding
the wind speed, the model shows a nearly zero bias compared to
observations between the valley floor and zm, whereas the RMSE
is in between 1 and 2 m s−1.This article is protected by copyright. All rights reserved.
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Figure 2. Synoptic weather conditions from 9 to 14 February 2015 at 0000 UTC (a) to (f) extracted from the European Centre for Medium-range Weather Forecasts
(ECMWF) gridded analyses with a horizontal grid point spacing of 0.25◦. The colour contours indicate geopotential height at 500 hPa and wind barbs indicate the wind at
700 hPa. The green dot in the centre of the weather charts indicates the position of the Passy Valley.
The PCAP formed in the evening on 9 February 2015, when
warm air was advected above the Passy Valley as the upper
level ridge approached the western Alps (see Sect. 3). Note the
weakening and veering of the flow associated with the passage
of the ridge (see Fig. 3c to f and Fig. 2). The formation of the
PCAP led to a rapid increase of the vertically-integrated valley
heat deficit (VHD, Whiteman et al. 1999) from 3 to 12 MJ m−2,
which is well reproduced in WRF-R (see Fig. 4a).
Between 9 February 2015 at 1200 UTC and 12 February 2015
at 1200 UTC, the structure of the PCAP in the Passy Valley
varied substantially, while the wind speed below 2500 m a.s.l.
was in the order of a few metres per second. During this period
the height of the PCAP decreased gradually with time, and so
the VHD. This decrease was most rapid during the night from
10 to 11 February 2015, by 4–5 MJ m−2. Despite the VHD
kept decreasing with time, the nighttime near-surface potential
temperature gradient computed over the lowest 100 m above the
valley floor, ∆θ/∆z100, was found to increase from day to day (see
Fig. 4b). This behaviour is well reproduced in WRF-R, but with
an underestimation of its amplitude. The different behaviour of
VHD and ∆θ/∆z100 suggests that a thermal decoupling between
the near-surface atmosphere and the atmosphere aloft occurs,
which will be analysed in details in the following Sections. The
maximum value of ∆θ/∆z100 was found on 12 February 2015 at
0000 UTC, as a result of the continuous advection of warm air
within the valley atmosphere below z = zm.
On 13 February 2015 the wind speed in the valley atmosphere
increased above 5 m s−1 as a large-scale perturbation approached
the western Alps (see Sect. 3). On 14 February 2015 the wind
speed increased also at the valley bottom, leading to mixing and
the full destruction of the PCAP.
4.2. Regional-scale circulation
A more detailed account of the interplay between the local (valley-
scale) flow and the regional-scale flow at z = zm, simulated
in domain d04, is given in Fig. 5. Timeseries of wind and
temperature at selected sites in the Passy Valley, RS and DS, and in
the main tributary valleys, MGV and CHMNX (see Fig. 1b for the
location of the sites), are displayed in Fig. 6 to show the response
of these fields (in WRF-R) to the large-scale flow.
On 10 February 2015 at 0000 UTC, the north-easterly large-
scale flow was channelled through the downstream part of theThis article is protected by copyright. All rights reserved.
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Figure 3. Temporal evolution of the vertical structure of the cold-air pool as obtained from the WRF-R simulation (left column) and from observational data (right column).
(a) and (b): potential temperature; root-mean-squared-error and bias of the potential temperature (c) and (d): wind speed (ws), (e) and (f): wind direction. The potential
temperature data are obtained from the radiosonde (RS) ascents, the wind speed and wind direction are retrieved from the LiDAR (below 1050 m a.s.l.) and the wind
profiler (above 1050 m a.s.l.) at the RS site during IOP1. The dashed lines indicates zm = 1560 m a.s.l. (see Fig. 1c). The RS site is shown in Fig. 1b. The middle panels
display the root-mean-squared-errors (continuous line) and bias (dashed line) of potential temperature (top row) and wind speed (middle row), with respect to observations.
Figure 4. Timeseries of the valley heat deficit, vertically integrated from the ground surface to the mean height of the terrain zm = 1560 m a.s.l. (a), and of the near-surface
potential temperature gradient over the lowest 100 m above the surface (b), at the RS site (see Fig. 1b), derived from data from the radiosonde ascents and computed from
the WRF-R and WRF-I outputs.
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Fi
gu
re
5.
H
or
iz
on
ta
lm
ap
of
th
e
w
in
d
(v
ec
to
rs
),
po
te
nt
ia
lt
em
pe
ra
tu
re
(c
ol
ou
re
d
co
nt
ou
rs
)a
nd
te
rr
ai
n
he
ig
ht
(b
la
ck
lin
e
co
nt
ou
rs
)fi
el
ds
in
do
m
ai
n
d0
4
at
z
=
z m
=
15
60
m
a.
s.
l.
on
10
Fe
br
ua
ry
20
15
at
00
00
U
T
C
an
d
12
00
U
T
C
(a
)a
nd
(b
),
re
sp
ec
tiv
el
y,
11
Fe
br
ua
ry
20
15
at
00
00
U
T
C
an
d
15
00
U
T
C
(c
)a
nd
(d
),
re
sp
ec
tiv
el
y,
12
Fe
br
ua
ry
20
15
at
12
00
U
T
C
(e
)a
nd
14
Fe
br
ua
ry
20
15
at
00
00
U
T
C
(f
).
A
re
as
fo
rw
hi
ch
th
e
te
rr
ai
n
he
ig
ht
is
ab
ov
e
th
e
di
sp
la
ye
d
he
ig
ht
le
ve
lz
=
z m
ar
e
m
as
ke
d
ou
tt
o
be
tte
rv
is
ua
lis
e
th
e
flo
w
ch
an
ne
lin
g.
T
he
sp
ee
d
an
d
di
re
ct
io
n
of
th
e
la
rg
e-
sc
al
e
flo
w
at
z
=
60
00
m
a.
s.
l.
av
er
ag
ed
ov
er
th
e
lo
ca
tio
ns
of
th
e
R
S,
D
S,
M
G
V
an
d
C
H
M
N
X
si
te
s
(s
ee
Fi
g.
1b
fo
r
th
e
lo
ca
tio
n
of
th
e
si
te
s)
ar
e
in
di
ca
te
d
on
th
e
to
p
ri
gh
tc
or
ne
r
of
th
e
pl
ot
s.
T
he
gr
ee
n
do
ts
in
th
e
ce
nt
re
of
ea
ch
pa
ne
li
nd
ic
at
e
th
e
po
si
tio
n
of
th
e
Pa
ss
y
V
al
le
y.
A
zo
om
on
do
m
ai
n
d0
5
is
pr
ov
id
ed
in
th
e
bo
tto
m
ri
gh
tc
or
ne
ro
ft
he
pl
ot
.
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Passy Valley, and split between CHMNX and MGV at z= zm (see
Fig. 5a and 6e). Warm air advection occurred over night across the
entire region, where potential temperature θ at z = zm increased
by 7 K in 12 hours (see Fig. 6f), while it decreased by 2–3 K at
z= 1000 m a.s.l. (see Fig. 6i). This process, that is warming at mid
levels and cooling at lower levels, led to the initial strengthening
of the PCAP within the Passy Valley, and it is well represented
in the numerical simulation. At 1200 UTC, the large-scale flow
rotated to a south-westerly direction, reaching a minimum wind
speed of less than 1 m s−1 at the different sites (see Fig. 6a and
6b). A strong and deep down-valley (north-easterly) wind formed
in the Rhone Valley, located north-east of the Chamonix Valley
near Sion, which was channelled through the gap that connects the
two valleys, strengthening the down-valley (north-easterly) flow
in CHMNX (see Fig. 5b).
Potential temperature θ at z = zm continuously increased on
10 February 2015, reaching a maximum value at 2100 UTC (see
Fig. 6f), that is when the upper level ridge just passed over the
Passy Valley. This can be inferred by noticing the rotation and
weakening of the wind aloft (at z = 6000 m a.s.l., see Fig. 6b and
Fig. 6a, respectively). In the early night between 10 to 11 February
2015 the near-surface flow from MGV reversed to down-valley
(south-westerly) as the upper level ridge moved away from the
western Alps, and remained down-valley (south-westerly) until
the end of the simulated time period (see Fig. 6k).
During the night between 10 to 11 February 2015 θ at z =
1000 m a.s.l. increased, while it stayed approximately constant
at z = zm (see Fig. 6f and i). The large-scale flow at this time
was northeast-easterly, i.e. aligned with the Rhone and Chamonix
valleys, thereby favouring the channelling of the flow down these
two valleys (see Fig. 5c). Throughout the night the large-scale flow
slowly rotated clockwise and strengthened as the upper level ridge
moved towards the north-east (see Fig. 6a and 6b).
Later on, on 11 February 2015 at 1200 UTC, the large-
scale flow veered to south-easterly, and the resulting channelling
through the Swiss Plateau gave rise to a regional south-westerly
flow at mid-level (see Fig, 5d), leading to a channelling of the
flow through the downstream part of the Passy Valley. During
the night between 11 and 12 February 2015 θ at z = 1000 m
a.s.l. and z = zm hardly varied (Fig. 6f and i), while the near-
surface temperature dropped throughout the night, as during the
previous night (Fig. 6l). During this night the amplitude of the
diurnal range of θ at z = zm is underestimated in WRF-R, but the
overall temporal evolution is well captured.
From 10 to 12 February 2015 the near-surface temperature at
night was non-homogeneous across the valley floor, as indicated
by differences of 4–5 K between the DS and RS sites that
are approximately at the same height (see Fig. 6l), while the
wind speed there was very weak (less than 1 m s−1 on average
during this time period). It is worth noting that simulated near-
surface temperatures are slightly underestimated at night, though
they fall within the observed range when spatially averaged
within a radius of a few hundred metres around the sites (not
shown). A cold bias of about 3 K is also present in the
afternoon. The errors in the representation of the amplitude of the
diurnal cycle of the near-surface temperature over snow-covered
surfaces in anticyclonic conditions can be related to issues in
the representation of turbulent surface fluxes and conductive heat
flux into the snowpack. Between 10 February at 1200UTC and
12 February at 1200UTC, the RMSE of near-surface potential
temperature, wind speed and direction are 2.9 K, 0.3 m s−1
and 65◦, respectively. The values for the potential temperature
are comparable or smaller than the ones reported in previous
modeling studies of CAPs and PCAPs (see for instance Page`s
et al. 2017). It should be noted that the RMSE of the wind
speed and direction are computed on a smaller number of data.
As before, the relatively large value of the RMSE of the wind
direction can be related to the large fluctuations of this quantity
occurring at low wind speeds.
During 13 and 14 February 2015, advection of cold air by
the large-scale perturbation led to the destruction of the PCAP,
strengthening of the near-surface flow with prevailing southerly
winds, and a homogenisation of the temperature across the valley
floor (see Fig. 5g, 5h and 6l).
The overview of the life-cycle of the PCAP presented above
leads to the identification of three main time periods, which
correspond to the three stages of a PCAP:
P1 The formation stage – from 8 to 10 February 2015, when
the temperature inversion formed in the Passy Valley and a
change in wind direction was observed aloft;
P2 The persistent stage – from 10 to 13 February 2015, when
the temperature inversion persisted in the Passy Valley;
P3 The destruction stage – on 14 February 2015.
Since the large-scale flow shows substantial variations with
time in wind speed and direction throughout the anticyclonic
period, the persistent stage is divided in two sub-periods:
P2a From 10 February 2015 at 1200 UTC to 11 February 2015
at 1200 UTC, when the upper level ridge moved over the
Passy Valley, and the wind speed at z = 6000 m a.s.l.
increased over night from a minimum value;
P2b From 11 February 2015 at 1200 UTC to 12 February
2015 at 1200 UTC, when the large-scale flow was veering
from south-easterly to southerly, and the wind speed at
z = 6000 m a.s.l. hardly varies with time.
Because the vertical structure and bulk properties of the PCAP
show a different behaviour during the two sub-periods of the
persistent stage, the link with the variations of the large-scale flow
during these periods is analysed in the following sections.
4.3. Valley-scale circulation during the persistent stage
In this section the valley-scale circulation during the persistent
stage P2 is described by comparing results from WRF-R and
the semi-idealised simulation WRF-I. A mass and heat budget
analysis and a mechanistic description of the flow behaviour
presented in this section will be discussed in Sect. 5.
4.3.1. Sub-period P2a
Sub-period P2a (from 10 February 2015 at 1200 UTC to
11 February 2015 at 1200 UTC) is characterised by a maximum
VHD around 1800 UTC, followed by a gradual decrease of the
VHD during the course of the night. The wind speed near the
surface and at z = 1000 m a.s.l. was at most 1 m s−1 at the
different sites considered in Fig. 6, except MGV, where it was
maximum and in the order of 3–4 m s−1 near the surface during
P2a. The non-local control of the large-scale flow on the near-
surface circulation can be extracted by comparing results from
WRF-R and WRF-I (see Fig. 7). Between 1800 UTC and 0600
UTC on the following day, the circulation at the bottom of the
Passy Valley is qualitatively similar in WRF-R and WRF-I. There
is a clear down-valley flow in the downstream part of the valley
and down-valley flows from CHMNX and STGV mostly detach
above the valley floor. At the valley exit the along-valley flow
is found to accelerate because of the narrowing of the valley.
The along-valley flow in MGV is up-valley at 1800 UTC both
in WRF-I and WRF-R, and reverses to down-valley at the surface
during the night (see Fig. 7b and d). However this flow is weaker in
WRF-I than in WRF-R, with wind speeds of about 0.3–0.5 m s−1
and 3–4 m s−1, respectively.This article is protected by copyright. All rights reserved.
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P2
a
P2
b
P2
a
P2
b
P2
a
P2
b
Fi
gu
re
6.
Ti
m
es
er
ie
s
of
th
e
ho
ri
zo
nt
al
w
in
d
sp
ee
d
(l
ef
tc
ol
um
n)
an
d
di
re
ct
io
n
(m
id
dl
e
co
lu
m
n)
,a
nd
po
te
nt
ia
lt
em
pe
ra
tu
re
(r
ig
ht
co
lu
m
n)
at
z
=
60
00
m
a.
s.
l.,
z
=
z m
=
15
60
m
a.
s.
l.,
z
=
10
00
m
a.
s.
l.
an
d
z
=
10
m
a.
g.
l.
(t
op
to
bo
tto
m
ro
w
s,
re
sp
ec
tiv
el
y)
,
at
th
e
lo
ca
tio
ns
of
th
e
R
S,
D
S,
M
G
V
an
d
C
H
M
N
X
si
te
s
(s
ee
Fi
g.
1b
fo
r
th
e
lo
ca
tio
n
of
th
e
si
te
s)
,i
n
W
R
F-
R
du
ri
ng
th
e
si
m
ul
at
ed
tim
e
pe
ri
od
.O
bs
er
va
tio
na
ld
at
a
at
th
e
R
S
si
te
ar
e
su
pe
ri
m
po
se
d
(b
la
ck
do
ts
):
w
in
d
da
ta
de
riv
ed
fr
om
th
e
w
in
d
pr
ofi
le
r
m
ea
su
re
m
en
ts
at
z
=
z m
an
d
z
=
10
00
m
a.
s.
l.
an
d
fr
om
an
in
st
ru
m
en
te
d
m
et
eo
ro
lo
gi
ca
l
m
as
t
at
z
=
10
m
a.
g.
l.,
an
d
po
te
nt
ia
l
te
m
pe
ra
tu
re
da
ta
at
z
=
z m
an
d
z
=
10
00
m
a.
s.
l.
de
riv
ed
fr
om
th
e
ra
di
os
on
de
m
ea
su
re
m
en
ts
an
d
fr
om
an
in
st
ru
m
en
te
d
m
et
eo
ro
lo
gi
ca
lm
as
ta
tz
=
10
m
a.
g.
l..
Si
m
ul
at
ed
da
ta
at
th
e
R
S
si
te
ar
e
di
sp
la
ye
d
w
ith
a
th
ic
k
bl
ac
k
lin
e
to
fa
ci
lit
at
e
co
m
pa
ri
so
n
w
ith
ob
se
rv
at
io
ns
.
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Figure 7. Simulated wind and potential temperature fields at z= 10 m a.g.l. in WRF-R (top row) and WRF-I (bottom row) on 10 February 2015 at 1800 UTC (left column)
and 11 February 2015 at 0600 UTC (right column). The speed and direction of the large-scale flow at z = 6000 m a.s.l. are indicated on the top right corner of the plots for
WRF-R.
The discontinuity in the magnitude of the wind vectors in
Fig. 7b where the air flows from MGV to the Passy Valley
indicates that the flow detaches from the surface because of
the strong stratification of the PCAP. This detachment is clearly
visible in the composite vertical cross-sections that approximately
follows the flow (along the white line in Fig. 1b), shown in
Fig. 8. The capping inversion at the top of the PCAP lowers down
progressively and ‘sharpens’ (i.e. its depth decreases) as the flow
through MGV strengthens during the course of the night (see
Fig. 8a,c and e). A hydraulic jump forms as the air descends into
the Passy Valley (see Fig. 8e). Conversely, although the flow from
MGV in WRF-I reverses above the near-surface layer only around
0000 UTC, the upper part of the PCAP cools steadily throughout
the night (see Fig. 8b,d and f), resulting in a deeper PCAP by
the end of the night compared to WRF-R. The differences in the
flow from the tributaries between WRF-R and WRF-R explain
the different behaviour of the VHD in the two simulations during
the night of P2a (see Fig. 4a). In WRF-R, the lowering of the
PCAP height due to the flow through the tributaries leads to the
reduction of VHD throughout the night, whereas in WRF-I the
VHD remains about constant with time due to the steady cooling
of the atmospheric column.
The lower part of the PCAP displays a temporal evolution of
wind speed and thermal structure similar to that in WRF-R, as can
be inferred from the isentrope θ = 276 K in Fig. 8. This result
suggests that the flow from MGV mostly controlled the increase
in temperature around z = 1000 m a.s.l. between 0000 and 0600
UTC, and more generally the thermal structure of the upper part of
the PCAP, while the lower part of the PCAP was mostly controlled
by local processes. This will be quantified in Sect. 5, where the
heat and mass fluxes through the tributaries of the Passy Valley
atmosphere are analyzed.
At 0600 UTC, the near-surface flow in the part of the
Passy Valley upstream of RS is weak and almost stagnant
in both WRF-R and WRF-I (see Fig. 7b and 7d). However,
the near-surface temperature there is lower in WRF-I than in
WRF-R, which explains the larger near-surface vertical gradient
of potential temperature that develops in WRF-I when compared
to WRF-R during the second part of the night (see Fig. 4b).
4.3.2. Sub-period P2b
During the sub-period P2b (from 11 February 2015 at 1200 UTC
to 12 February 2015 at 1200 UTC) the evolution of the bulk
thermal properties of the Passy Valley atmosphere are more
similar in WRF-R and WRF-I than during P2a (see Fig. 4). At
1800 UTC, the near-surface flows through MGV, CHMNX and
STGV are down these tributaries and the near-surface flow in the
Passy Valley is down-valley (see Fig. 9a and 9c). Hence, the near-
surface flow and the flow in the upper part of the PCAP are in
opposite directions, i.e. they are dynamically decoupled (see Fig. 5
and Sect. 4.2).
The speed of the flow from MGV increases with time in
WRF-R, reaching a maximum value in the morning on 12
February 2015, while it remains constant in WRF-I (see Fig. 9b
and 9d). The large scale flow slightly rotates during the night,
being more aligned with the MGV valley axis in the morning,
suggesting that the increase of wind speed at MGV is an effect of
the channelling of this large scale flow. The flow from CHMNX
is weaker than that from MGV and comparable in WRF-R and
WRF-I. In contrast to P2a, the near-surface temperature in theThis article is protected by copyright. All rights reserved.
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Figure 8. Composite vertical cross-sections following the flow (along the white line in Fig. 1b) of horizontal wind speed (in the plane of the cross-section) and potential
temperature (K) in WRF-R (left column) and WRF-I (right column) on 10 February 2015 at 1800 UTC, and 11 February 2015 at 0000 and 0400 UTC (top to bottom rows,
respectively).
Passy Valley is similar in WRF-R and WRF-I during P2b (see
also Fig. 4b).
5. Mass and heat fluxes through the Passy Valley
Sect. 4.3 highlighted non-local controls of the large-scale flow
on the PCAP in the Passy Valley during the persistent stage
(P2), in particular in determining the respective contributions of
the tributary flows to the mass and heat budgets of the valley
atmosphere. In this section, these contributions are quantified.
5.1. Mass fluxes
The mass budget of the Passy Valley atmosphere is performed
over the PASSY control volume, defined in Sect. 2.1. Assuming
incompressibility, the mass budget reads
∑
i
∫
Ai
ρ vi ni dS︸ ︷︷ ︸
Mi
=
MDS +MMGV +MCHMNX +MSTGV +MTOP = 0,
(1)
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Figure 9. Simulated wind and potential temperature fields at z= 10 m a.g.l. in WRF-R (top row) and WRF-I (bottom row) on 11 February 2015 at 1800 UTC (left column)
and 12 February 2015 at 0600 UTC (right column). The speed and direction of the large-scale flow at z = 6000 m a.s.l. are indicated on the top right corner of the plots for
WRF-R.
(a) (b)
Figure 10. Timeseries of the mass fluxes through DS (MDS), MGV (MMGV), CHMNX (MCHMNX), STGV (MSTGV), vertically integrated from the ground surface to
z = zm, and through the top surface at z = zm (MTOP) for the period from 10 February 2015 at 1200 UTC to 12 February 2015 at 1200 UTC in WRF-R (a) and WRF-I
(b); see Fig. 1b and 1c for the location of the gates. The nighttime periods (from sunset to sunrise) are displayed as grey shading. The bottom panel shows the mass fluxes
normalised by the sum of all mass fluxes out of the Passy Valley (i.e. a positive quantity). Positive and negative values of the mass flux correspond to a flux out of and into
the Passy Valley, respectively. Note the different scale used between panel (a) and (b).
where Ai = ADS, AMGV, ACHMNX, ASTGV, are the lateral surfaces of
the control volume (see the blue surfaces in Fig. 1c), Ai = ATOP
is the top surface at z = zm, ρ is the air density, vi is the wind
component normal to the surface Ai orientated by the unit vector
ni (defined positive outwards). Mi = MDS, MMGV, MCHMNX, MSTGV
are the mass fluxes across the surface Ai. To compute the massThis article is protected by copyright. All rights reserved.
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fluxes, the model fields have to be interpolated on a cartesian
grid. For the vertical mass flux, this can lead to a misrepresented
contribution from the slope flows from the terrain surrounding the
PASSY control volume. Hence, to take into account these effects,
in the following the vertical mass flux is computed as the residual
of the mass budget.
5.1.1. Sub-period P2a
As one may expect, DS is the main exit for the air mass
within the valley (see Fig. 10a). During P2a, both CHMNX and
STGV contribute an inflow into the valley, while the mass flux
through MGV can be in or out of the valley. When averaged
during nighttime (from sunset to sunrise), MCHMNX and MSTGV
account for 35% and 18% of the total flow of mass out of the
valley, respectively. When combined together, the flow from these
tributaries provide about half of the mass flowing out of the
valley at night during P2a. The other half is coming from MMGV
and/or MTOP. MMGV > 0 (outflow) when the upper level ridge
passed over the valley (i.e. when the wind speed aloft decreased
to its minimum value, see Fig. 6a), and subsiding motions at
the valley top account for about 40% of the inflow of mass into
the valley. As the upper level ridge moves away from the region
(and the wind speed increases at upper level, see Fig. 6a), most
of the air mass out of the valley flows through DS, and MTOP
decreases continuously while the flow through MGV reverses
(MMGV changes from positive to negative at about 2200 UTC). In
the late night from 0300 to 0600 UTC, the inflow of mass from
MGV exceeds that from CHMNX and STGV and the total mass
originating from the tributaries exceeds that flowing through DS,
resulting in an export of mass across the top of the valley volume.
The mass fluxes in WRF-I are one order of magnitude smaller
than those in WRF-R (see Fig. 10b). The nighttime-averaged
contribution to the total flow of mass out of the valley from
CHMNX is smaller in WRF-I than in WRF-R (25% compared
to 35%), implying that the down-valley flow at CHMNX in WRF-
R is affected by the channelling of the flow from the Sion valley.
Overall, the mass flux analysis demonstrates that the large-scale
flow controls the magnitude of the tributary flows and allows to
quantify the relative contribution of the latter flows to the total
outflow from the Valley during P2a.
5.1.2. Sub-period P2b
During P2b, between 1300 and 2000 UTC the large-scale flow
was channelled over the Swiss Plateau (see Fig. 6b and Fig. 5d),
resulting in a strong inflow into the downstream part of the Passy
Valley (see Fig. 10a). This inflow through DS is mostly balanced
by the outflow through MGV and the export of mass across the
top of the valley volume. MDS reverses from negative to positive at
2000 UTC when the large-scale flow rotates to south-southeasterly
and so was in the down-valley direction. At this time, MMGV > 0
and MTOP, MCHMNX and MSTGV contribute equally to the total
drainage out of the valley through DS and MGV. As the large-
scale flow rotated to southerly, the flow at MGV reverses from up-
to down-valley and strengthens, leading to a continuous increase
of MMGV after 0000 UTC. MCHMNX decreases during the course
of the night, contributing on average 15% of the total drainage
out of the valley, compared to 35% during P2a. Conversely, the
nighttime-averaged contribution to the total flow of mass out of
the valley from STGV is 16% and comparable to that during
P2a. This suggests that the flow through STGV is less sensitive
to variations in the large-scale flow than those in CHMNX and
MGV.
In WRF-I, each tributary contributes about 30% of the total
drainage out of the Passy Valley through DS during P2b.
Interestingly MMGV < 0 throughout P2b, as opposed to P2a,
implying that the vertical structure of the PCAP influences the
direction of the flow through MGV. This can be explained as
follows: since the PCAP at 1200 UTC is much shallower than
during P2a, and its top is below the height of MGV, there is no
horizontal temperature contrast between RS and MGV and hence
no thermally-driven up-valley flow from the Passy Valley to MGV.
5.1.3. Vertical and horizontal structure of mass fluxes
The analysis presented so far is further developed by examining
the vertical structure of the mass fluxes through DS. A significant
fraction of the total drainage out of the Passy Valley occurs in
the upper part of the PCAP as a result of the detachment of the
flow from the tributaries (see Fig. 11a). The mass flux through
DS in the lower part of the PCAP is more variable and has
a similar magnitude in WRF-R and WRF-I. This suggests that
despite the flow from the tributaries is a major control on the
total outflow from the valley, it mainly affects the upper part of
the PCAP with only an indirect effect on the lower part. Indeed,
the mass flux in WRF-R is characterized by temporal oscillations
between outflow (positive) and inflow (negative) in WRF-R. Such
oscillations are also visible in WRF-I but only during P2b (see
Fig. 11b). These oscillations can be triggered either by those
of thermally-driven flows or by internal gravity waves (IGWs)
created by the detachment of the flow through the tributaries at
the top of the CAP via a hydraulic jump (see for instance Largeron
et al. 2013). A closer inspection of Fig. 11a and 11b reveals that
the slope of the contours associated with the oscillations with
respect to the vertical is non zero, indicating that there is a phase
propagation in the vertical. As a consequence, the hypothesis
that oscillations could be triggered by those of thermally-driven
flows can be rejected, as those are pure oscillations in time with
no phase propagation (Largeron et al. 2013; Quimbayo-Duarte
et al. 2019). The horizontal cross-section of the vertical mass
flux at the height of STGV during P2a at 2200UTC confirms the
IGW mechanism behind the mass flux oscillations in WRF-R (see
Fig. 12a). In WRF-R the flow through STGV detaches at about z=
1000 m a.s.l., leading to the formation of internal gravity waves
that propagate within the Passy Valley atmosphere. In WRF-I,
internal gravity waves are excited by the detachment of the flow
from MGV and propagate downstream (the link with the waves
observed at DS needs to be clarified) (see Fig. 12b). A different
time period is selected for WRF-I compared to WRF-R because
the tributary flows are characterised by a different magnitude and
time of occurence in the two simulations.
5.2. Heat fluxes and impact on the CAP height
The previous subsection revealed the key role of the
inflows/outflows from the tributary valleys in the mass budget of
the Passy Valley atmosphere during the persistent stage of the
CAP and how those flows are affected by the large-scale flow.
This section examines their role on the heat budget of the Passy
Valley atmosphere.
The heat fluxes are computed from the non-linear advective
term of the volume-averaged density-weighted heat budget
equation over the PASSY control volume, that is, assuming
incompressibility,
ADV =
1
[Mass]P
∫
V
ρ v j
∂θ
∂x j
dV =∑
i
1
[Mass]P
∫
Ai
ρ
(
θ ′ v j
)
n j dS =
HDS +HMGV +HCHMNX +HSTGV +HTOP,
(2)
where [Mass]P is the mass of the PASSY control volume, θ ′ =
θ − θ and θ is the volume-averaged potential temperature inThis article is protected by copyright. All rights reserved.
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Figure 11. Time-height plot of the mass flux through the downstream (DS) gate for the period from 10 February 2015 at 1200 UTC to 12 February 2015 at 1200 UTC in
WRF-R (a) and WRF-I (b); see Fig. 1b and 1c for the location of the gate. The nighttime periods (from sunset to sunrise) are displayed as dashed vertical lines. Note the
different scale used between panel (a) and (b).
Figure 12. Horizontal cross-sections (across the polygon defining the PASSY volume, see Fig. 1b) of the vertical mass flux at the height of STGV on 10 February 2015 at
2200 UTC (during P2a) in WRF-R (a) and at the height of MGV on 11 February 2015 at 1900 UTC (during P2b) in WRF-I (b). Positive and negative values of the mass
flux correspond to a flux out of and into the Passy Valley, respectively. Note the different scale used between panel (a) and (b).
the PASSY volume (see for instance Lee et al. 2004), and Ai
the lateral and top surfaces of the volume, defined in Sect. 5.1.
Hi = HDS, HMGV, HCHMNX, HSTGV are the heat fluxes across the
surface Ai, divided by [Mass]P. The heat flux through the top
surface is computed as a residual of Eq. 2, as it is done for
the vertical mass flux (see Sect. 5.1). In order to quantify the
contribution of advective processes to the heat budget of the valley
atmosphere, the ADV term above is compared with the volume-
averaged density-weighted total tendency, that is
TEND =
1
[Mass]P
∫
V
ρ
∂θ
∂ t
dV . (3)
TEND is the variation of the heat content of the PASSY volume
due to the advection of heat through the lateral and top surfaces
of the volume, as expressed in Eq. 2, and the vertical divergence
of radiative and turbulent heat fluxes. As clear-sky conditions
persisted through P2 in WRF-R (expect for some sporadic fog
that formed near the surface during the night of P2a, not shown),
latent heat release/absorption associated to phase changes can be
neglected in the heat budget at the valley scale. The effect of
the heat fluxes through the tributaries on the thermal structure of
the PCAP is evaluated in WRF-R by investigating the temporal
evolution of the CAP height, CAPh, and the bottom height of the
capping inversion, CIh. CAPh is defined as the maximum height
where the vertical gradient of absolute temperature is positive and
CIh is defined as the height where the curvature of θ (z) presents
a positive maximum. During P2, when a capping inversion layer
exists within the valley atmosphere, CAPh coincides with the top
height of this layer, so that the depth of the capping inversion d
can be defined as d = CAPh−CIh.
5.2.1. Sub-period P2a
In WRF-R the temporal evolution of TEND during P2 follows
that of the advection contribution ADV (see Fig. 13a). TEND is
controlled by ADV during P2a while it is close to zero during
P2b (implying a balance between ADV and the contributions
from turbulent heat and radiative flux divergences). The large
variations in ADV during P2a are associated with the variability
of the advection contributions from the tributaries and top surface
of the valley volume. Conversely, in WRF-I the total tendency
and advection term have opposite signs and a steady temporal
evolution (see Fig. 13c). Advection contributes to a warming ofThis article is protected by copyright. All rights reserved.
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(c) (d)
(a) (b)
Figure 13. (a) Timeseries of the volume-averaged density-weighted total potential temperature tendency (TEND), advection contribution (ADV) to the total tendency and
that restricted to the tributaries (Trib, horizontal advection), averaged over the PASSY volume (see Fig. 1b), for the period from 10 February 2015 at 1200 UTC to 12
February 2015 at 1200 UTC in WRF-R; (b) Timeseries of the horizontal advection contribution to the total tendency from DS, MGV, CHMNX and STGV in WRF-R (see
Fig. 1b). (c) and (d): same as (a) and (b), but for WRF-I. Note the different scale used for the y-axis between panels (a), (b) and panels (c), (d). The nighttime periods (from
sunset to sunrise) are displayed as grey shading.
(b)(a)
Figure 14. (a) Timeseries of the heights of the cold-air pool (CAPh) and of the capping inversion (CIh), and potential temperature at the RS and MGV sites (see
Fig. 1b), vertically-averaged in the height range 1150 < z < zm = 1560 m a.s.l (θRS and θMGV, respectively), for the period from 10 February 2015 at 1200 UTC to
12 February 2015 at 1200 UTC in WRF-R. The nighttime periods (from sunset to sunrise) are displayed as grey shading. Note that CIh subsides down to the valley floor
after 11 February 2015 at 1500UTC; see the text for more details. (b) Vertical profiles of potential temperature at the RS site for selected times during P2a and P2b in
WRF-R: 10 February 2015 at 1230 UTC (continuous red line), 10 February 2015 at 1930 UTC (dashed red line), 11 February 2015 at 0430 UTC (dotted red line) and
11 February 2015 at 0900 UTC (dashed-dotted red line). In both panels, dashed lines indicate the height of the valley floor at MGV (1100 m a.g.l.) and zm = 1560 m a.s.l.
(see Fig. 1c).
the valley atmosphere during most of the analysed time period,
a behaviour that is typical in synoptically unperturbed conditions
and reported in previous studies of idealised CAPs (Schmidli and
Rotunno 2010; Arduini et al. 2017), hence demonstrating the
impact of the large-scale flow on the valley heat budget in WRF-R.
A more detailed analysis of the advection contribution in
WRF-R reveals the large contribution of tributary flows. In
general, the sum of the heat fluxes through MGV, STGV and
CHMNX is mostly balanced by an opposite heat flux through
DS (see Fig. 13b). However, the (un)balance between these fluxes
varies with time and responds to variations in the large-scale flow,
resulting in a net heat input (warming) or export (cooling) from
the tributaries, with a large impact on the evolution of the thermal
structure of the PCAP. In WRF-I, there is a close balance between
the horizontal advection through DS and the vertical advection,
with only a minor contribution from the other tributary valleys
(see Fig. 13b).
During the early night until 2200 UTC, the flow through MGV
is out of the valley volume (see Fig. 10a) and PCAPh decreases
from z = zm to z = 1300 m a.s.l (see Fig. 14a). This is mainlyThis article is protected by copyright. All rights reserved.
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attributed to the warming of the upper part of the PCAP due to
subsiding motions across the top surface of the valley volume
(70%) and ADV through the other two tributaries, CHMNX and
STGV (30%); see Fig. 13b. This result explains why later on the
direction of the flow in MGV reverses from up- to down-valley.
When it is up-valley, i.e. flowing from RS towards MGV, CIh
is about the height of the valley floor at MGV and d ' 500 m
(see Fig. 14b). Hence, the upper part of the PCAP above the
height of the valley floor at MGV is colder at RS than at MGV
(see Fig. 14a), leading to a pressure force between the two sites
directed out of the valley. As warm air is advected in the upper
part of the PCAP, the capping inversion sharpens (i.e. d decreases)
and the temperature in the upper part of the valley atmosphere
becomes higher at RS than at MGV (see Fig. 14a and Fig. 14b).
This chain of processes leads to the reversal of the pressure force,
which is then directed into the valley.
As the flow through the tributaries strengthens during the
course of the night of P2a, the heat flux through the tributaries
increases. From 0200 to 0600 UTC when the mass flux through
the tributaries is maximum (see Fig. 10a) and CAPh and CIh
decrease continuously with time, ADV is mostly controlled by
the heat flux through the tributaries (see Fig. 13a). The analysis
of the heat fluxes together with that of Fig. 8, suggest that the
upper part of the PCAP is eroded by mixing with the warm
air advected through the tributary valleys (mainly MGV after
0200 UTC). Another mechanism that could explain the increase
of the potential temperature in the upper part of the PCAP is
the downstream displacement of the cold air out of the valley
(Lareau and Horel 2015b). However, the comparison of potential
temperature profiles at DS and Marnaz does not show evidences of
a downstream displacement (not shown). The advection of warm
air leads to a continuous decrease of CAPh and CIh during the
course of the night, (see Fig. 14b) the latter decreasing to 150
m a.g.l. by 0900 UTC. In WRF-I, the CAP height hardly varies
with time (not shown) because of the negligible contribution of the
advection from the tributary flow to the heat budget at the valley
scale (see Fig. 13d).
5.2.2. Sub-period P2b
During P2b the heat fluxes through the tributary valleys are
generally weaker than during P2a (see Fig. 13b). The capping
inversion subsides down to the surface at 1500UTC, hence leading
to a change in the vertical structure of the PCAP. From 1200 to
1500 UTC the channelling of the large-scale flow into the valley
through DS led to a substantial inflow into the PCAP, leading
to a heat export (cooling) through the top surface of the valley
volume. This heat export results from upward vertical motions
induced by the convergence of the tributary flows in the upper
part of the PCAP and of the channelled flow through DS. Later
on, between 1500 UTC to 2000 UTC, turbulent and radiative heat
fluxes divergence explain most of the cooling within the PCAP,
as the net contribution from advection is close to zero. At 2000
UTC, HDS and HTOP change sign as was the case for the mass
fluxes (see Fig. 10a), becoming negative (cooling) and positive
(warming), respectively. From this time on and until 1200 UTC,
ADV is controlled by the interplay between subsiding motions at
the valley top and heat fluxes through MGV and DS, which are
modulated by the IGW oscillations of the subsiding motions.
6. Factors controlling the near-surface inversion layer
During a PCAP episode, thermally-driven valley flows are one of
the main mechanisms favouring the recirculation of air in the near-
surface layer (Quimbayo-Duarte et al. 2019). The basin-shape of
the Passy Valley constrains horizontal motions in this layer, for
which DS is the only gate through which the air mass can flow in
and out of the valley. Hence, to investigate the near-surface valley
flow, a metric is introduced based on the mass flux through the
gate DS (see for instance Whiteman et al. 1996):
τ−1 =
M∆z100DS
[Mass]∆z100
, (4)
where [Mass]∆z100 is the mass of the PASSY control volume
over the lowest 100 m above the valley floor and M∆z100DS is the
mass flux through DS, vertically-integrated over the lowest 100 m
above the valley floor. This height was chosen as it approximately
corresponds to the height of the nighttime ground-based inversion
layer during the persistent stage (see Fig. 14b). The dimension of
τ is time, however this parameter is a signed quantity, which sign
depends on the direction of the flow through DS. If τ > 0, the flow
is out of the valley and τ measures the time required to renew the
air mass within the PASSY volume below z = z100 through DS.
Conversely, if τ < 0 the flow through DS is into the valley and the
near-surface air is nearly stagnant within the valley, assuming that
in very stable conditions, vertical motions at z=100 m a.g.l. can be
neglected.
The parameter τ−1 is displayed in Fig. 15a but, in the following,
the behavior of τ is discussed to facilitate its interpretation. τ
displays temporal oscillations with a period of approximately 3–
4 hours (see Fig. 15a), which are the signature of the internal
gravity waves excited by the flow through the tributary valleys.
As pointed out in Sect. 5.1, such oscillations are not present during
P2a in WRF-I. The amplitude of the oscillations of τ in WRF-R
is fairly large compared to its mean positive value during both
P2a and P2b and makes τ to vary between positive (outflow)
and negative (inflow) values. During P2a, the average value of
τ over a period of oscillation increases with time during the
course of the night and the nighttime average is about 7 hours
in both WRF-R and WRF-I. The similar trend of τ in WRF-R and
WRF-I indicates that the near-surface drainage out of the valley
is controlled by local thermally-driven flows but is modulated by
the non-local control of the large-scale flow on the flows through
the tributaries. The temporal evolution of τ during P2b differs
in WRF-R and WRF-I. In WRF-R the average value of τ over
a period of oscillation is about 4.5 hours before 2200 UTC and
increases to more than 24 hours later on. By contrast, the nighttime
average of τ in WRF-I in about 10 hours. This suggests that the
non-local control of the large-scale flow is playing a role in the
near-surface drainage during P2b in WRF-R.
In order to investigate the impact of non-local processes on the
thermal structure of the near-surface atmosphere, the tendency
(TEND) and advection (ADV) terms of the heat budget (see
Sect. 5.2), averaged over the lowest 100 m over the ground
surface, are shown for WRF-R and WRF-I in Fig. 15b. The
temporal evolution and magnitude of both TEND and ADV during
the morning and evening transitions (0700 to 0900 UTC and 1300
to 1500 UTC, respectively) is similar in WRF-R and WRF-I,
indicating that local processes control the heat budget of the
near-surface atmosphere during these transition periods. During
P2a TEND and ADV display a similar trend in WRF-R and
WRF-I during nighttime. During P2b, the trends are different:
in WRF-R TEND and ADV decrease with time during the night
while in WRF-I they increase. Note that as the consequence of
the stagnation of the air mass in WRF-R in the late night during
P2b, ADV fluctuates about zero. However, the total tendency
integrated throughout the night in P2b is very similar between
WRF-R and WRF-I (−5.4 K and −5.8 K from sunset to sunrise
for WRF-R and WRF-I, respectively), implying that even though
the advection contribution varies between the two simulations, the
net cooling is hardly affected by the change in the down-valley
flow.This article is protected by copyright. All rights reserved.
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Figure 15. (a) Timeseries of the inverse of the near-surface (z < 100 m a.g.l.) mass turnover time τ associated with the mass flux through the downstream (DS) gate (see
Fig. 1b) for the period from 10 February 2015 at 1200 UTC to 12 February 2015 at 1200 UTC in WRF-R and WRF-I. (b) Timeseries of the density-weighted total potential
temperature tendency (TEND) and advection contribution (ADV) to the total tendency, averaged over the PASSY volume over the lowest 100 m above the ground surface,
for the period from 10 February 2015 at 1200 UTC to 12 February 2015 at 1200 UTC in WRF-R and WRF-I. The nighttime periods (from sunset to sunrise) are displayed
as grey shading.
Figure 16. Timeseries of the pressure difference (black lines) between the RS and MRNZ sites (seeFig. 1b) at z= 600 m a.s.l. (∆P600) and at z= zm = 1560 m a.s.l. (∆P1560),
in (a) WRF-R and (b) WRF-I. Each panel also display the difference in potential temperature (blue line) between RS and MRNZ sites, vertically averaged in the height
range 600 < z < zm = 1560m a.s.l. (∆θRS-MRNZ). The nighttime periods (from sunset to sunrise) are displayed as grey shading.
This difference between WRF-R and WRF-I during P2b is a
result of the evolution of the large-scale flow. The strengthening
of the south-southeasterly large-scale flow during the course of
the night favours its channelling through MGV in the upper part
of the PCAP. The large PCAP stability prevents the channeled
flow to penetrate down at the valley bottom and so to enhance
the down-valley flow. The channelled flow transports warmer
air in the upper part of the PCAP, so reducing the horizontal
temperature difference between the valley centre and outside the
valley towards MRNZ (see Fig. 13a and Fig. 16a). This results
in a reduction in the near-surface pressure difference driving the
down-valley flow between RS and MRNZ between 2100 and
0000 UTC (see Fig. 16a), leading to the stagnation of the near-
surface air within the valley. This non-local effect is not present in
WRF-I, which accounts for the near-surface pressure difference
to continue to drive a down-valley flow throughout the night (see
Fig. 16b).
7. Summary and conclusions
The purpose of the present work is to examine local and non-
local controls on a persistent cold-air pool (PCAP) that formed
in the Passy Valley in February 2015 using numerical model
simulations. The effect of the non-local controls on the PCAP was
determined by comparing the real case simulation with a semi-
idealised simulation with initial quiescent synoptic conditions.
The Passy Valley is a region of Alpine terrain with steep
slopes and three major tributary valleys, Mege`ve, Saint-Gervais-
les-Bains and Chamonix (see Sect. 2.1). The numerical model
simulations were performed using the Weather Research and
Forecasting model at a horizontal mesh size of 111 m, carefully
initialised (see Sect. 2 for details). The simulation results are
found to realistically capture features of the persistent cold-air
pool by means of comparison with field observations. The model
realistically simulates the magnitude and temporal evolution of the
valley heat deficit, as well as the vertical structure of the potential
temperature, when compared to the available observations.
However, it overestimates the near-surface minimum temperature
at nighttime, which leads to an underestimation of the magnitude
of the temperature inversion over the lowest 100m.
The cold-air pool formed in the evening on 9 February 2015 as
a result of the radiative cooling at the valley floor and advection
of warm air above the Passy valley at the mean height of the
terrain surrounding the valley, which sets the height of the cold-
air pool during its formation stage. The advection of warm air
was associated with the passage of an upper level ridge over the
region. The effects of the large-scale flow on the cold-air pool
during its persistent stage, from 10 February 2015 at 1200 UTC
to 12 February 2015 at 1200 UTC, was investigated at the valley
scale and in the near-surface layer above the valley floor (lowest
100 m).
The analysis of the numerical simulations showed that the mass
fluxes at the valley scale through the tributaries are about one
order of magnitude larger than in the semi-idealised simulation,
which demonstrates the impact of the large scale flow on the
flow through the tributary valleys. The flow through Chamonix
and Saint-Gervais-les-Bains valleys can supply up to 50% of
the air mass that drained out of the valley during one of the
nights of the persistent stage of the CAP. As a consequence,
the heat budget at the valley scale differs between the real andThis article is protected by copyright. All rights reserved.
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idealised simulations: this budget is controlled primarily by the
flows through the tributary valleys and the vertical advection at
the valley top in the real case. This combination of horizontal
and vertical advection modified the height of the PCAP. This led
to a change in the direction of the flow in the Mege`ve tributary
valley, which substantially modified the dynamics and thermal
structure of the Passy Valley atmosphere. The results thus showed
that local thermally driven flows alone cannot explain the day-to-
day variability of the CAP height.
Conversely, the near-surface atmosphere was found to respond
differently depending on the position of the upper level ridge.
When the upper level ridge passed over the Passy Valley
(between 10 and 11 February 2015), the mass and heat budgets
computed for a near-surface atmospheric volume were found to
be comparable between the real and semi-idealised simulations.
This implies that the near-surface atmosphere above the valley
floor was nearly decoupled from the upper part of the cold-air
pool and controlled by and large by local processes, although
the flow was modulated by internal gravity waves generated
by the detachment of the flow from one of the tributary
valleys. As the upper level ridge moved away from the Passy
Valley (between 11 and 12 February 2015), the large-scale flow
strengthened, favouring channelling through the Mege`ve tributary
valley towards Passy. This flow transported warmer air (in terms
of potential temperature) in the upper part of the PCAP, reducing
the thermal gradient in the down-valley direction, while being
decoupled from the near-surface atmosphere because of the large
PCAP stability. This leads to the reduction of the surface pressure
difference driving the down-valley flow, eventually leading to
stagnant air over the valley floor. This mechanism is not present
in the semi-idealised simulation, in which a down-valley flow was
found to flow out of the Passy Valley throughout the night.
In summary, the main findings of the present work, which are of
relevance for other deep Alpine valleys under similar wintertime
anticyclonic conditions, are:
• The large-scale flow controls the magnitude of the mass
fluxes through the tributary valleys and their relative
contribution to the total flow out of the Passy Valley. The
flows from the tributaries can contribute more than 50% to
this outflow during the persistent stage of the episode.
• The direction of the flow through one of the tributaries is
found to be determined by the height of the PCAP with
respect to that of the tributary above the valley floor. This
result highlights the coupling between the large-scale flow,
the tributary flows and the thermal structure of the valley
atmosphere during a PCAP event.
• The tributary and large-scale flows modify the heat budget
of the valley atmosphere during the persistent stage of the
episode, when compared to the heat budget for quiescent
synoptic conditions. At the valley scale, the advection
contribution is found to be the main control on the total
potential temperature tendency.
• Hence, the effect of the non-local controls is key for
determining the change in the thermal structure of the CAP
from one day to the following and so the change in the heat
deficit. It follows that local thermally-driven flows alone are
not able to modify the thermal structure of the PCAP on the
time scale of this episode.
• In the present case, the near-surface atmosphere remains
almost decoupled from the flow aloft, and two mechanisms
are identified. When the upper level ridge passes over the
region, the near-surface atmosphere is subject to usual
down-valley flows during nighttime; when the ridge moves
away from the region, the near-surface atmosphere become
nearly stagnant. This is due to the warm air advection in the
upper part of the CAP by the large-scale flow.
The present work calls for a better characterisation of the
controls of tributary flows and of varying large-scale flows on
cold-air pooling. This would require an approach based on a
climatology of the area to identify the prevalence of weather
patterns. Future field campaigns should consider instrumenting
the main tributary valleys and collecting long-term datasets
to sample an extended range of synoptic forcings. Further
observations are also required in order to evaluate the role of
internal gravity waves in the modulation of the heat and mass
budgets of the valley atmosphere. Finally, future modelling studies
should also consider the role of fog, and more generally of
microphysics processes, in the evolution of a PCAP, as these were
not considered in the current work.
Acknowledgements
This work has been supported by a grant from LabEx
Osug@2020 (Investissements d’avenir – ANR10LABX56).
Numerical model simulations were performed using the UK
national supercomputing facilities (ARCHER), accessed through
the UK National Centre for Atmospheric Science (NCAS). The
Passy-2015 field experiment was supported by ADEME through
the French national programme LEFE/INSU and by METEO-
FRANCE. We thank the cities of Passy and Sallanches for their
kind support. The field experiment was led by CNRM while the
LEGI laboratory was the principal investigator of the LEFE/INSU
project. The rawinsonde data used in Fig. 3 are managed by
SEDOO at Observatoire Midi-Pyre´ne´es. We thank the people
involved in the launching of the rawinsonde and data acquisition
therefrom.
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