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Keywords:

  • density current;
  • Mesoscale Alpine Programme;
  • south foehn

Abstract

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Description and preparation of dataset
  5. 3. Results
  6. 4. Discussion
  7. 5. Conclusions
  8. Acknowledgements
  9. References

The propagation of a cold front and its interaction with foehn winds is investigated in an Alpine valley, based on observations collected during the field campaign of the Mesoscale Alpine Programme (MAP) on 6 November 1999. The key instrument of this study is a Doppler lidar that had been operated in the Wipp Valley (Austria). The cold front approached the European Alps from the northwest, became distorted at the mountain barrier and entered the east– west aligned Inn Valley near the town of Innsbruck primarily via two passes. It continued to propagate towards Innsbruck from both valley directions as two separate fronts that eventually collided east of Innsbruck after part of the cold air had entered the adjacent north– south aligned Wipp Valley.

A synthesis of Doppler lidar measurements with conventional meteorological data, including automatic weather stations and radiosondes, leads to the conclusion that the cold front in the Wipp Valley was an atmospheric density current characterized by an elevated head, a front-relative feeder flow and a typical propagation speed of 7 m s−1. The foehn flow on top of the density current caused strong wind shear and triggered shear-flow instability that led to the formation of a turbulent wake behind the head. As the density current propagated towards the Brenner Pass, it slowed down. The shape of the frontal surface varied in time. Its inclination of about 10°– 20° is steeper than previously reported for the Inn Valley but is consistent with other observations of atmospheric density currents. Copyright © 2010 Royal Meteorological Society


1. Introduction

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Description and preparation of dataset
  5. 3. Results
  6. 4. Discussion
  7. 5. Conclusions
  8. Acknowledgements
  9. References

On 6 November 1999, at the end of the Mesoscale Alpine Programme (MAP) Special Observing Period (SOP) (Bougeault, et al.2001) and after a fruitful season of gap-flow measurements in the Wipp Valley (Austria) (Mayr, et al.2004; Mayr, et al.2007), scientists witnessed a spectacular event: the interruption of foehn winds by a cold front passage. Although such an event is nothing special per se and occurs frequently in the Alpine valleys, the dataset collected during this happening was unique because of the availability of a Doppler lidar in the Wipp Valley. Most of the previous MAP studies have focused on the dynamics of foehn during the mature stage of its life rather than on its decay. However, this frontal event has stimulated new questions: how do the foehn flow and cold front interact, how is the cold front modified by terrain, which pathways does it choose to enter Alpine valleys, what is the vertical structure of the cold front and what are the characteristics of propagation? The goal of this article is to address these questions, focusing on the Inn Valley and Wipp Valley (Figure 1).

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Figure 1. Topography of the target area: (a) the central part of the European Alps and (b) close-up of the Wipp Valley and Inn Valley (Austria). Elevation contours (km ASL) are shown as grey shadings. Markers with three-letter labels indicate the location of weather stations (see also Table I). A box in (a) illustrates the subdomain shown in (b). A thick line in (b) shows the location of the slope profile and two thin lines the orientation of two lidar scans.

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The influence of the European Alps on cold fronts has been investigated in the framework of two major field campaigns, the Alpine Experiment (ALPEX) in 1982 (Kuettner 1982a) and the German Front Experiment (GFE) in 1987 (Hoinka and Volkert 1987). The focal point of these campaigns was mainly on the orographic modification of fronts over the Alpine foreland and foothills and over relatively smooth (idealized) barriers. Some of the results of ALPEX and GFE– including earlier work, theoretical concepts and modelling studies – are summarized in Kuettner (1982b), Hoinka and Volkert (1992), Egger and Hoinka (1992) and Blumen (1992).

The deformation of cold fronts at the Alpine barrier has been documented in several observational studies (Bergeron, 1928; Steinacker, 1981; Kurz, 1990; Volkert, et al., 1991). Frontal analyses of Steinacker (1982, 1987) show that, as a result of low-level blocking at the Alps, cold fronts may advance faster at upper levels than near the surface. Davies (1984) uses a single-layer shallow-water model to reveal analytically that the strength of orographic retardation of a cold front depends on the non-dimensional depth of the cold-air layer (front– mountain aspect ratio) as well as on a rotational Froude number. Schumann (1987) extends this study with a numerical model to three dimensions and concludes that the retardation is strong for shallow fronts, high and steep mountains, strong temperature contrasts between the pre- and post-frontal air mass, small velocities normal to the front and very stable stratification. Further, he notices that in three dimensions stable stratification enhances cold front deformation due to flow splitting. Some modelling studies have focused on the interaction of foehn winds with cold fronts over the northern Alpine foreland (Hoinka 1987; Egger 1989; Heimann 1990; Heimann 1992). They indicate that strong pre-frontal foehn is able to increase the temperature contrast between the pre- and post-frontal air mass. The pre-frontal warming induces a mesoscale pressure depression ahead of the cold front that accelerates the front. However, if the foehn is weak and not able to form a mesoscale low ahead of the front or if the large-scale west– east pressure gradient is strong, the cold front slows down as it has to move towards higher pressure.

Only few observations exist that document the behaviour of cold fronts in Alpine valleys. Krenn (1949) notices that post-frontal cold air does not pass the Alps in the same direction as the front impinges on the barrier, but follows natural paths of the terrain to reach the inner Alpine region. Steinacker (1987) notes that cold fronts approaching the Alps from the west and northwest may enter valleys that are aligned in a west– east direction parallel to the Alps (such as the Inn Valley) through their eastern main entrance, similarly to a backdoor cold front. The cold wind following a backdoor cold front– a term most commonly used in North America– blows from the northeast instead of the typical northwest quadrant (Glickman 2000). Kaufmann (1989) analyzes ALPEX data and concludes that the actual path chosen by the cold front to reach the Inn Valley depends on the depth and stability of the cold air mass. One of the following three cases may occur. (1) If the post-frontal cold air is deep and well-mixed, it spills over the mountain range and causes northerly foehn in the Inn Valley. (2) If the cold air is less deep, it mainly flows over the mountain passes north of the valley. (3) In the case of a very shallow cold front, the cold air enters the Inn Valley via its mouth. The importance of these mountain passes for the penetration of north foehn into the Inn Valley has been shown by Zängl (2006) based on numerical simulations. This model study also reveals that north foehn in the Inn Valley can only develop without significant precipitation.

Based on the GFE dataset, Freytag (1990) classifies the observed cold fronts into two main groups. The first group is characterized by pre-frontal south foehn and deformation of the front at the Alps. The post-frontal cold air enters the lower Inn Valley via its eastern mouth. In the second group, pre-frontal foehn does not occur and the cold front passes the Alps without appreciable deformation. Hence, the cold air enters the Inn Valley from the west and northwest by spilling over the mountain ridges. Freytag (1990) notices that the front propagates more slowly in the valley than over the foreland and that the steepness of the front increases. From these observational surveys it has become obvious that several potential pathways exist that allow the cold air to enter the Inn Valley. However, their actual role appears to be case sensitive and has been inferred in the past based on a rather coarse resolution dataset.

Cold fronts have often been described as atmospheric density (or gravity) currents (Carbone, 1982; Bond and Fleagle, 1985; Seitter and Muench, 1985; Shapiro, et al., 1985; Hakim, 1991; Darby, et al., 1999; Koch and Clark, 1999; Neiman, et al., 2001; Geerts, et al., 2006). A conceptual model of such a density current is shown in Figure 2. Basic properties are an elevated head, a turbulent wake and a cold-air feeder flow towards the nose that is faster than the propagation speed of the density current. Smith and Reeder (1988) provide a critical discussion on the applicability of the density current model to cold fronts with specific regard to the propagation speed. In the Alps, the density current model has been applied to cold fronts moving over the foreland (Hoinka, et al.1990), forming barrier jets along the foothills (Hoinka 1987; Heimann 1988) and entering an idealized valley (Egger 1992). The latter study was stimulated by observations collected during the GFE in the Loisach Valley (Germany) (Müller and Sladkovic 1990). We will demonstrate that in our case the cold front exhibits density-current characteristics too.

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Figure 2. Schematic representation of an atmospheric density or gravity current. Adapted from several sources including Charba (1974), Goff (1976), Koch (1984), Smith and Reeder (1988), Simpson (1997) and Geerts, et al. (2006). According to Simpson (1997), two types of instability may occur at the head, which lead to the formation of billows or to lobes and clefts, respectively. The wake is characterized by turbulent mixing. The feeder flow of cold air is also called rear-to-front flow, and the backflow is also called the undercurrent. An arc or rope cloud may form above the head. The depth of the body of the density current hb is about half the depth of the head. The profile of potential temperature θ shown is typical for the thermal structure behind the head.

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Recently, Gohm, et al. (2009) used data from an airborne aerosol lidar in order to describe the interaction between a density-current-like flow and downvalley winds. Ground-based remote sensing systems, such as Doppler lidars, have proved extremely useful for the description of the kinematic structure of cold fronts over complex terrain (Darby, et al., 1999; Neiman, et al., 2001; Darby and Poulos, 2006). During the MAP SOP such a Doppler lidar had been operated in the Wipp Valley to investigate south foehn (Flamant, et al.2002, Durran, et al.2003, Gohm and Mayr 2004, Gohm, et al.2004, Weissmann, et al.2004, Zängl and Gohm 2006, mari09Aag) as well as thermally driven flows (Rucker, et al.2008). In our cold front study we will use data from the same lidar.

Our dataset allows us to extend previous cold front investigations in the Inn Valley to a major tributary, the Wipp Valley, and to focus on a much higher temporal and spatial resolution which covers the micro-α and meso-γ scale (Orlanski 1975). First analyses of the 6 November 1999 MAP cold-front case have been performed by Darby, et al. (2000), Gohm, et al. (2000) and Vill (2002). We extend these preliminary studies by focusing on the Doppler lidar dataset. Observations from a dense network of weather stations also play an important role. Our goal is to draw a comprehensive picture of the temporal evolution and spatial structure of a cold front in an Alpine valley. Further, we want to illuminate the propagation characteristics of the cold front and the interaction with foehn winds. The article is organized as follows: section 2 describes the dataset, the results of the data analysis are presented in section 3, a discussion is provided in section 4 and conclusions are drawn in section 5.

2. Description and preparation of dataset

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Description and preparation of dataset
  5. 3. Results
  6. 4. Discussion
  7. 5. Conclusions
  8. Acknowledgements
  9. References

2.1. Automatic weather stations

A dense network of automatic weather stations (AWS) was operated in the Inn Valley and Wipp Valley during the MAP SOP (Mayr, et al.2004). About half of these were equipped with standard instruments including temperature, humidity, pressure and wind sensors. They were mainly installed along the valley floor with a horizontal spacing of about 3– 10 km (cf. Figure 1(a) and (b)). The other half used temperature sensors only– so-called HOBO data loggers (Whiteman, et al.2000). They were installed along several slope profiles. The profile along the slope of the mountain Patscherkofel (PAK) is used herein (cf. Figure 1(b)). In conjunction with a fully equipped AWS collocated at the slope profile, potential temperature was derived from HOBO data by applying the hydrostatic relationship. The vertical spacing of HOBO data loggers is about 100– 200 m. In our study we utilize a larger subset of the whole MAP– AWS network. In addition to the MAP stations we integrate in our analysis observations from permanently installed AWS of the Austrian Weather Service ZAMG. Our AWS dataset is based on 10 minute averages. Hence, 10 minutes is also the temporal resolution for estimating time lags between the frontal passage at different sites. Some SYNOP stations with an observational time interval of one hour are used for deducing the time of the cold front passage over the Alpine foreland. Table I provides a list of relevant stations.

Table I. Description of stations
AcronymNameElevation (m ASL)Type
  1. Station type: M for MAP-AWS, Z for ZAMG-AWS, S for SYNOP, R for radiosonde, L for lidar. See text for further explanation.

ACHAchenkirch912Z
BREBrenner1373M
CHIChieming553S
ELBEllboegen1080M
FELFeldkirch439Z
GARGarmisch-Partenkirchen720S
GEDGedeir1084M/R/L
GENGenauen865M
GRIGries1278M
HALHall565M
HOHHohenpeissenberg986S
IBKInnsbruck612M
JENJenbach539Z
KEMKempten705S
KUFKufstein493Z
MUNMunich489R
PAKPatscherkofel2252Z
PATPatsch913M
PONPontigl1235M
REUReutte852Z
SEFSeefeld1182Z
STESteinach1116M
STZSterzing944M
SWZSchwaz543M
TIETienzens1121M
VOLVolders554M
WARWarth1471Z
ZENZenzenhof715M
ZIRZirl600M

2.2. Doppler lidar

The NOAA/ESRL scanning Doppler lidar TEACO2 was located in the Wipp Valley at Gedeir (GED) between the Brenner Pass and Innsbruck (see Figure 1(b)). The system is described in Post and Cupp (1990). Durran, et al. (2003) provide a more MAP-related description including an estimation of the accuracy of the system. The lidar emits pulses of infrared light at 10.59 µm. The signal is backscattered from aerosols that move with the flow. The Doppler-shifted frequency of the backscattered signal provides the radial wind velocity component along the direction of the lidar beam. The signal of each beam is split into 300 m range gates. In our study we use scans performed on a vertical plane [range– height indicator (RHI)] at a constant azimuth angle upvalley (174°) and downvalley (319°) of the lidar site (see Figure 1(b)). The increment in the elevation angle from one beam to the next is 0.5°. Hence, the data spacing normal to the beam is about 10 m at a distance of 1 km from the lidar site and 100 m at a distance of 10 km, respectively.

For the derivation of the ‘full’ wind field, radial velocities are interpolated on to a Cartesian grid with a horizontal and vertical mesh size of 100 m. We use a similar procedure to that applied in previous studies (Wakimoto, 1982; Intrieri, et al., 1990; Banta, et al., 1992) to derive the along-valley wind component u and the vertical wind component w from the radial wind velocity ur. We assume that the flow is along the valley in the direction of the lidar beam with a cross-valley wind component v = 0. This assumption is valid as long as we focus on the channelled flow below crest level, which is located at about 3 km above sea level (ASL), and neglect flow from and into tributaries. Hence, we can derive u from

  • equation image(1)

Here, ε is the elevation angle of the lidar beam and x the horizontal coordinate along the valley with x ≥ 0 in the northern semicircle. In our case, ur is positive when the flow is directed towards the lidar. Positive u means southerly flow. Strictly speaking, Eq. (1) is only a good approximation for relatively small ε, say below 30°, or for nearly horizontal flow. Then, assuming an incompressible fluid, we can derive w from the continuity equation:

  • equation image(2)

Here, we basically neglect vertical variations in the density. As shown by the scale analysis of Wakimoto (1982) this is a suitable approximation as long as the depth of the considered layer is not much more than 2 km. In order to solve Eq. (2) on a Cartesian grid iteratively from the bottom upward, we have to specify a lower boundary condition. For simplicity, we choose

  • equation image(3)

with h being the terrain height. This assumption is valid as long as the valley floor is nearly flat. We have tested other boundary conditions, e.g. terrain-following flow with w(z = h) = u(∂ h/∂x), however we did not observe significant differences in our results. Some sensitivity on the boundary condition is found at places where the lidar scan intersects steep terrain, e.g. near the Brenner Pass. We will not use our estimates of u and w for a quantitative analysis but for a qualitative discussion of the kinematic structure of the cold front and foehn in section 3.5.

2.3. Additional datasets

The vertical thermodynamic structure of the cold front is assessed inter alia from radiosonde soundings conducted at Gedeir in the Wipp Valley and at Munich over the northern Alpine foreland. The operational analyses of the European Centre for Medium-Range Weather Forecasts (ECMWF) and the Vienna Enhanced Resolution Analysis (VERA) (Steinacker, et al.2000) are used to compile a synoptic and mesoscale overview of the event. ‘Quicklooks’ of the operational forecast of the Canadian Mesoscale Compressible Community Model (MC2) (Benoit, et al.2002) available from the MAP Data Centre are used for the interpretation of some observations.

3. Results

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Description and preparation of dataset
  5. 3. Results
  6. 4. Discussion
  7. 5. Conclusions
  8. Acknowledgements
  9. References

3.1. Synoptic and mesoscale-scale overview

We start with a description of the synoptic-scale situation between 4 and 7 November 1999. This period was characterized by south foehn in the Alps and by the passage of a cold front that caused the decay of foehn.

On 4 November 1999 a weak low-pressure system over southern France was responsible for southerly flow over the Alps and the onset of south foehn in the Wipp Valley. At the AWS Ellboegen (ELB) in Figure 3 this onset occurred at 0950 UTC and was associated with a rapid warming and a shift to southerly (downvalley) winds with average wind speeds exceeding 10 m s−1. A shaded region in Figure 3 indicates the period of south foehn. On the morning of 5 November, the depression moved further southward towards the island of Sardinia. As a result, the flow at the Alpine crest level temporarily turned to a northwesterly direction and foehn winds in the Wipp Valley weakened (cf. Figure 3(b)). Foehn winds increased again until noon as the flow near crest level became more southerly again.

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Figure 3. Time series of data from the automatic weather station Ellboegen (ELB) from 0000 UTC on 4 November 1999 to 0000 UTC on 7 November 1999: (a) potential temperature (K) as a solid line and pressure (hPa) as a dashed line, (b) 10-minute average wind speed (m s −1) as a solid line and wind direction (deg) as dots. A grey shaded region in (a) and (b) indicates the period of south foehn.

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Figure 4 shows the ECMWF analysis for the 700 and 850 hPa levels. The situation at 0000 UTC on 6 November 1999 is depicted in Figure 4(a) and (d). The remnant of the forementioned depression is visible over the Mediterranean. A cyclogenesis east of the British Isles (Figure 4(a)) was associated with a cold front over France (Figure 4(d)). This cold front approached the Alps from the northwest and arrived at the mountain range at around noon (Figure 4(e)). The front caused the breakdown of south foehn in the Wipp Valley with a rapid cooling, pressure increase and wind shift to a northerly direction (Figure 3). An upper-level jet stream (not shown) initiated, at its left exit region, a lee cyclogenesis in the Gulf of Genoa (Figure 4(b) and (c)). This cyclone was associated with the outbreak of cold air over the Mediterranean with northwesterly Mistral winds south of France (Figure 4(f)). However, over the eastern Alps the depression inhibited the southward propagation of the cold front. One part of the front became quasi-stationary near the main Alpine crest. Another part of the cold air was deflected around the eastern edge of the Alps on 7 November, where it was responsible for northeasterly bora winds along the Adriatic coast. At that time and on the following two days north foehn occurred in some valleys south of the Alpine crest.

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Figure 4. ECMWF analysis at (a, d) 000 UTC and (b, e) 1200 UTC on 6 November 1999, and (c, f) 0000 UTC on 7 November 1999. Fields for (a)– (c): 700 hPa of geopotential height (m) as contour lines with increments of 20 m and wind as vectors. Fields for (d)– (f): 850 hPa of temperature (°C) as colour contours with increments of 2°C and horizontal wind as vectors. For a reference vector see the lower left corner. The 800 m elevation contour of the topography is shown by grey shading in (a)–(c) and as a thin line in (d)–(f). The coast is shown as a thick line.

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The mesoscale analysis in Figure 5 illustrates the situation at the surface at 1200 UTC on 6 November 1999. Shown are VERA fields of potential temperature and sea-level pressure as well as individual observations from routine weather stations. The cold front is located close to the main Alpine crest (see white line), with colder air on the northern side of the mountain range. In our target area, the Inn and Wipp Valley, the front is located between Innsbruck (IBK) and the Brenner Pass (BRE).

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Figure 5. Mesoscale surface analysis at 1200 UTC on 6 November 1999: gridded VERA data of potential temperature (K) as colour contours, sea-level pressure (hPa) as black contour lines and horizontal wind vectors (see reference vector in the lower left corner). Contour increments are 1 K and 1 hPa, respectively. The location of the cold front is indicated as a thick white line. Surface observations of routine weather stations are shown as circles (the colour indicates potential temperature) and wind barbs (half barb, full barb and triangle denote 2.5, 5 and 25 m s−1). The coast and the 1000 m elevation contour of the topography are shown as thick black lines.

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Radiosonde profiles from Munich (MUN) in Figure 6 demonstrate the depth of the low-level cold air associated with the cold front over the northern Alpine foreland on 6 November 1999. The location of MUN is indicated in Figure 5. The pre-frontal sounding at 0600 UTC is characterized by relatively warm, southwesterly (cross-mountain) winds above the low-level inversion and below ≈ 2 km ASL. These winds are most likely related to large-scale foehn subsidence north of the Alps. With the cold front passage, low-level winds shift to a westerly (mountain-parallel) direction with speeds exceeding 15 m s−1 (see profile at 1200 UTC), which is an indication for the formation of a barrier jet (Neiman, et al., 2009). The top height of the cold-air layer increases from about 2.8 km ASL at 1200 UTC to about 3.6 km ASL at 2100 UTC.

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Figure 6. Radiosonde profiles from Munich (MUN) valid for 0600 UTC (solid line), 1200 UTC (dotted line) and 2100 UTC (dashed line) on 6 November 1999: (a) potential temperature (K), (b) wind speed (m s −1) and (c) wind direction (degrees).

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Figure 7 illustrates the situation along the northern rim of the Alps close to the foreland. Shown are time series of potential temperature and relative humidity for six AWS. The stations are aligned along a east– west cross-section that covers a distance of about 200 km (see Figure 1(a)). The east– west spacing between two sites is about 40 km. The stations are either close to the mouth of a valley or at a major pass. The cold front arrives first at the most western station Feldkirch (FEL) in the Rhine Valley (0640 UTC). It passes Warth (WAR) near the Arlberg Pass 1.5 hours later and Reutte (REU) at the entrance of the Lech Valley 2 hours later. Finally, at about 1000 UTC, the front passes the saddle of Seefeld (SEF), the pass at Achenkirch (ACH) and the mouth of the Inn Valley at Kufstein (KUF). Hence, flow deflection at the Alps has caused considerable deformation and eastward acceleration of the front along the foothills (cf. Figure 4(d) and (e)) so that the cold air enters the Inn Valley via SEF, ACH and KUF at nearly the same time. The southeastward propagation of the front is also documented by hourly SYNOP observations (not shown) from stations located along the northern Alpine foothills: the front passes Kempten (KEM) and Hohenpeissenberg (HOH) at about 0800 UTC and Garmisch-Partenkirchen (GAR) and Chieming (CHI) at about 1000 UTC (see Figure 1(a) for location of stations).

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Figure 7. Time series of potential temperature in K (solid line; left axis) and relative humidity in per cent (dashed line; right axis) from 0500– 1500 UTC on 6 November 1999. Data are taken from six weather stations located along the northern rim of the Alps from east to west (top to bottom panel): Kufstein (KUF), Achenkirch (ACH), Seefeld (SEF), Reutte (REU), Warth (WAR) and Feldkirch (FEL). See Figure 1(a) for the location of stations. A vertical line in each panel indicates the time of frontal passage.

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3.2. Temporal evolution in the Inn Valley

The temporal evolution of the thermal structure and wind field in the Inn Valley near the surface is illustrated in Figure 8(a). Shown is a Hovmöller-type diagram of near-surface potential temperature (θ) and horizontal wind with time on the abscissa. The ordinate represents a west– east cross-section from Zirl (ZIR) to Jenbach (JEN) over a distance of about 45 km based on five AWS (see Figure 1).

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Figure 8. Temporal evolution of potential temperature (K) and wind field at the surface on 6 November 1999: as a function of horizontal distance (km) (a) along the Inn Valley and (b) along the Wipp Valley, and (c) as a function of altitude (km ASL) along the slope of the mountain Patscherkofel (PAK). The contour increment is 1 K. Vectors represent horizontal wind speed and direction (see the reference vector in the lower right corner, which indicates flow from the west). Locations of automatic weather stations are indicated on the right ordinate with labels (cf. Figure 1). The distance in (a) is relative to Innsbruck (IBK) with positive values eastward and that in (b) is relative to the lidar site (GED) with positive values northward. Solid and dashed lines indicate the location and displacement of the cold front assuming a constant speed of the front in (a) of uf = 6.5 m s−1 (solid) and uf = 5.5 m s−1 (dashed), in (b) of uf = 7.5 m s−1 (solid) and uf = 4.5 m s−1 (dashed) and in (c) of wf = 1 m s−1.

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In the morning of 6 November 1999, the temperature distribution along the valley is heterogeneous. Two cold pools can be identified in the transect (equation image K, bluish colours), one west of IBK and the other one east of Hall (HAL). They have formed by radiative cooling during the previous night. However, between IBK and HAL, i.e. near the exit of the Wipp Valley, the air is significantly warmer by several degrees as a result of warmer foehn air that is advected and mixed from the Wipp Valley into the Inn Valley. There, a maximum of equation image K occurs between about 0800 and 0930 UTC. However, the foehn period at IBK and HAL is rather transient and characterized by several shorter periods of consecutive warming and cooling. This pattern essentially represents a ‘battle’, i.e. strong mixing, between two air masses: the southerly foehn flow from the Wipp Valley and the two cold air pools in the Inn Valley west and east of IBK. The wind field is rather complicated. Winds are weak and unsteady, blowing either up- or downvalley. However, during the period when θ values are highest and when the air mass is essentially the same as in the Wipp Valley (compare Figure 8(a) and (b)), the flow is from the southwest at IBK and from the northwest at HAL, i.e. winds have a cross-valley component there. This flow pattern is consistent with numerical simulations (Gohm, et al., 2004; Gohm and Mayr, 2004), which shows that part of the foehn jet from the Wipp Valley that passes IBK as a southwesterly flow becomes deflected at the mountain range north of IBK and returns as a northwesterly flow near HAL. Cooling with a shift to easterly (upvalley) winds at about 0930 UTC at HAL and one hour later at IBK indicates the end of the foehn period in the Inn Valley but not yet the arrival of the cold front. Instead this cooling illustrates the back-flow of air from the easterly cold pool.

The actual passage of the cold front in the Inn Valley is indicated in Figure 8(a) as two lines (solid and dashed). The front arrives first at the eastern and western boundaries of the shown cross-section, i.e. at JEN (1050 UTC) and ZIR (1120 UTC), and later in between, e.g. at 1150 UTC at IBK. This pattern suggests that the cold air enters the Inn Valley both east and west of IBK and subsequently flows towards IBK from both valley directions as two separate fronts. This pattern is consistent with case studies of Kaufmann (1989) and Freytag (1990) (see also section 8.2 in Egger and Hoinka, 1992). The propagation speed of these fronts is approximately uf = 6.5 m s−1 for the eastern front and uf = 5.5 m s−1 for the western front, respectively. For one of the GFE87 cases, Freytag (1990) estimates a typical propagation speed in the Inn Valley of 5 m s−1, which is therefore close to our estimate. In our case, the two cold fronts collide at a location between IBK and HAL (see meeting point of solid and dashed line in Figure 8(a)) at about 1200 UTC. Already half an hour earlier, the cold air has entered the Wipp Valley (see Figure 8(b)). The existence of a convergence zone east of IBK as a result of two colliding air masses has also been documented for other cold-front cases (Kaufmann 1989).

A shift to northwesterly winds at JEN and ZIR at about 1100 UTC indicates that the cold air has entered the Inn Valley via two shallow gaps in the northern mountain range: the pass at Achenkirch (ACH) and the saddle of Seefeld (SEF) (see Figure 1(a)). Northerly cold-air flow through the mountain gap at SEF is supported by NOAA P-3 aircraft in situ observations at a flight level of about 1.4 km ASL (see figure 3 in Darby, et al., 2000). As mentioned in section 3.1, part of the cold air has entered the Inn Valley via its mouth at Kufstein (KUF). It is also conceivable that part of the cold air has penetrated into the Inn Valley by spilling over the mountain range north of IBK, the so-called Karwendel. This range has typical peak heights of 2.3– 2.6 km ASL. Therefore, its crestline is lower by about 200– 500 m than the top height of the cold air mass estimated from the MUN sounding at noon in section 3.1. This height difference would allow the cold air to spill over the Karwendel. However, the approaching air mass is too warm near crest level (equation image K above 2.3 km ASL at 1200 UTC in Figure 6) to penetrate down to the very bottom of the valley and to replace the cooler pre-frontal air near IBK (equation image K at the surface at 1230 UTC in Figure 8(a)). Thus, the contribution of spillover to the initial low-level cooling is presumably small. This is supported by the MC2 forecast valid for 1200 UTC, which reveals that the top height of the cold air mass at Seefeld is about 2.4 km ASL and, hence, close to the crest height of the Karwendel (not shown).

3.3. Temporal evolution in the Wipp Valley

Figure 8(b) illustrates the evolution of the flow in the Wipp Valley near the surface on 6 November 1999, including the pre-frontal foehn period and the southward propagation of the cold front. Shown are near-surface potential temperature and horizontal wind with time on the abscissa and horizontal distance from the lidar site Gedeir (GED) on the ordinate. The whole south– north transect covers a distance of about 50 km. Observations are taken from a total of 12 AWS between Genauen (GEN) and Innsbruck (IBK) (see right ordinate and Figure 1). Note that GEN is located south of the Brenner Pass (BRE) and IBK north of BRE.

The foehn period in the morning is depicted by potentially warm air north of the Brenner Pass and southerly (downvalley) winds. Foehn air north of the pass is about 4– 5 K warmer than the nearly-blocked air mass in the basin of Sterzing (STZ). Foehn winds are strongest in the lower part of the Wipp Valley near Ellboegen (ELB). The south foehn period abruptly ends around noon with a rapid cooling and a shift to upvalley winds as the cold front arrives. The southward propagation of the cold front is illustrated by different times of frontal passage: e.g. 1150 UTC at IBK, 1200 UTC at ELB, 1230 UTC at Steinach (STE), 1310 UTC at BRE and 1350 UTC at STZ. Thus, the cold front takes approximately two hours to propagate from IBK to STZ over a horizontal distance of ≈ 43 km. However, the horizontal propagation speed of the front uf is not constant. It is higher in the northern part (equation image m s−1) and lower in the southern part (equation image m s−1) of the Wipp Valley. This behaviour is indicated by two straight lines with different slopes in Figure 8(b), which are each characterized by a constant but different speed. The solid (dashed) line is valid for the northern (southern) part of the valley. The slow-down of the cold front occurs near STE where these two lines intersect. In reality a gradual decrease in the propagation speed may have occurred south of STE instead of the abrupt decrease indicated by the kink between the solid and the dashed line. The latter is the result of a rather coarse station density.

3.4. Vertical structure in the Wipp Valley

The temporal evolution of the vertical thermodynamic structure of the air in the Wipp Valley on 6 November 1999 is depicted in Figure 8(c). Near-surface potential temperature is shown as a function of time and altitude based on observations from 13 AWS that are located along the slope of the mountain Patscherkofel (PAK) (see Figure 1). Wind data are only available and shown for two AWS sites (PAT and PAK). The slope profile covers the altitude from the valley floor at 708 m ASL (H01) to the peak of PAK at 2252 m ASL. The stations are aligned nearly perpendicular to the valley axis.

The foehn period is characterized by a nearly mixed valley atmosphere. The average vertical gradient of potential temperature between H01 and PAK is ∂θ/∂z ≈ 1 K km−1, corresponding to a buoyancy frequency of N = 0.006 s−1. After the frontal passage, the static stability increases to an average gradient of ∂θ /∂z≈ 4 K km−1 (N = 0.012 s−1). In fact, the valley atmosphere now has a two-layer structure: a less stable layer with ∂θ/∂z≈ 3 K km−1 below approximately 1.4 km ASL and a more stable layer with ∂θ/∂z≈ 5 K km−1 above this level. The frontal passage does not occur simultaneously at all levels. The arrival time of the front is 1140 UTC at H01 and 1210 UTC at PAK, for example. Thus, the time lag increases with altitude. This lag can be expressed as a vertical advance speed of the cold front of equation image m s−1. Its magnitude is approximately constant with height (see the straight line in Figure 8(c)). Here, a positive wf implies that the frontal surface has a backward inclination relative to the direction of propagation, which can be expected for the case of a cold front. The average slope of the frontal surface can be estimated with equation image; for uf we use the value estimated from Figure 8(b) (see section 3.3). This slope value corresponds to an inclination angle of approximately 8°. Freytag (1990) estimates a smaller value of Δzx = 1/20 for the inclination of the front in the Inn Valley based on his GFE87 observations, hence a less steep frontal surface than in our case. However, our estimated inclination of 1/7.5 is still smaller by a factor of about two, compared with the alternative and more realistic value derived from lidar data in section 3.5.

Figure 9 shows soundings from three radiosondes launched at the lidar site (GED) on 6 November 1999 three hours before, shortly after and one hour after the passage of the cold front, respectively. This sequence impressively illustrates the cooling (Figure 9(a)) and the wind shift (Figure 9(b) and (c)) associated with the transition from a downvalley foehn flow to an upvalley cold-air flow. The foehn flow at 0850 UTC is characterized by a southerly low-level jet with peak winds of about 20 m s−1 at 2 km ASL, a nearly mixed layer in the lower part of the jet and a stable layer in the upper part between 1.7 and 3.6 km ASL with ∂θ /∂z≈ 5 K km−1. Up to 1202 UTC, i.e. shortly after the frontal passage at GED, cooling has decreased potential temperature by about 2– 4 K between 2.5 and 4.5 km ASL compared with the earlier sounding. We believe that this elevated cooling is not caused by horizontal cold-air advection behind the cold front, since the flow is still southerly there. Rather it is the result of vertical lifting of the foehn air. Warming by subsidence is a specific attribute of foehn and becomes manifest in the Wipp Valley by descending isentropes (Gohm, et al., 2004). However, by 1202 UTC low-level foehn winds– and therefore also isentropes– have been lifted ‘back’ above the low-level cold air, which results in adiabatic cooling between 2.5 and 4.5 km ASL. This lifting results in a reduction of the gravity-wave activity over the Wipp Valley, with isentropes becoming nearly horizontal as illustrated in vertical transects of the MC2 forecast (not shown). Later, at 1308 UTC, horizontal air-mass advection behind the front has caused a distinctive cooling below 2.5 km ASL by up to 5 K. Between 1202 UTC and 1308 UTC the maximum speed of the cold-air flow has increased from about 10 to 13 m s−1 and its depth from about 600 m to 1300 m. Thus, the propagation speed of the front (cf. section 3.3) is smaller than the maximum wind speed behind the front, which is a typical feature of a density current (Simpson and Britter, 1980; Smith and Reeder, 1988). We will return to this feature in the next section.

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Figure 9. Vertical profiles of (a) potential temperature (K), (b) wind speed (m s −1) and (c) wind direction (degrees). Lines show data from radiosondes launched at the lidar site Gedeir (GED) at 0850 UTC (solid), 1202 UTC (dotted) and 1308 UTC (dashed) on 6 November 1999. Markers in (a) illustrate observations taken along the slope of the mountain Patscherkofel between the stations H05 and PAK at 0900 UTC (squares) and 1300 UTC (circles) on 6 November 1999.

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Two selected slope profiles taken from Figure 8(c) between H05 and PAK are also shown in Figure 9(a) for comparison purposes. Especially after the frontal passage, the vertical structure of the slope profile agrees very well with the corresponding radiosonde profile, despite the fact that the horizontal location of the two profiles is different. During foehn, the lower part of the slope profile is about 2 K warmer than the radiosonde profile, which suggests that the temperature field in the Wipp Valley is substantially inhomogeneous. Slanted isentropes in along- and across-valley directions, which may explain this temperature difference, have been documented by Gohm, et al. (2004), for example.

3.5. Density current as observed by the lidar

Figure 10 depicts the wind field in the northern part of the Wipp Valley on a vertical plane along the valley axis based on a series of Doppler lidar RHI scans. This structure is very similar to that observed by Doppler lidar for other types of density currents, including thunderstorm gust fronts (Intrieri, et al., 1990; Darby, et al., 2002) and the sea breeze (Banta, et al., 2005). Shown is the horizontal along-valley wind component u derived from Eq. (1) together with wind vectors representing the two components u and w, the latter derived from Eq. (2). The azimuth angle of the scans is α = 319° (see Figure 1(b)) and the lidar site GED is on the abscissa at x = 0. The sequence of scans covers the fully evolved foehn stage on the morning of 6 November 1999 (Figure 10(a)) as well as the approach of the cold front towards the lidar site at around noon (Figure 10(b)– (h)).

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Figure 10. Wind field in the northern part of the Wipp Valley as revealed by a Doppler lidar at approximately (a) 0908 UTC, (b) 1130 UTC, (c) 1142 UTC, (d) 1145 UTC, (e) 1148 UTC, (f) 1152 UTC, (g) 1154 UTC and (h) 1158 UTC on 6 November 1999. The along-valley wind component is shown as colour shadings with increments of 1 m s−1 (positive values for southerly downvalley flow) and vectors representing winds along the cross-section (see the two reference vectors in the upper right corner). The lidar site GED is on the abscissa at x = 0 (with x increasing northward). The lidar scans downvalley at an azimuth angle of α = 319° (see also Figure 1(b)).

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The southerly downvalley foehn flow in Figure 10(a) at 0908 UTC is characterized by a jet-wind layer, the depth of which decreases and the speed of which increases along the valley north of the lidar site. Similar patterns of flow acceleration were also observed for other MAP foehn cases (Flamant, et al., 2002; Gohm, et al., 2004; Weissmann, et al., 2004; Marić and Durran, 2009). A maximum of about 18 m s−1 occurs at a distance between 5 and 6 km north of GED, which is consistent with the location of the speed maximum identified in Figure 8(b) based on AWS data. The speed decreases again as the foehn flow approaches the mouth of the Wipp Valley. By 1130 UTC foehn winds have weakened– the maximum has decreased to 14 m s−1 (Figure 10(b)). A first sign of a flow reversal becomes apparent at about 8 km north of the lidar site. This reversal illustrates the inflow of cold air at the mouth of the valley associated with the approaching cold front. The warmer foehn flow is lifted above this shallow cold-air layer. By 1142 UTC the cold front has advanced along the valley by about 2.5 km and the depth of the cold-air layer near the valley entrance has increased from about 300 to 700 m (Figure 10(c)). The top of the layer is slanted along the valley, yet without an elevated head at its leading edge. However, such a head starts to form a few minutes later (Figure 10(d)). From now on, shown by a sequence of scans with a time interval of about three minutes (Figure 10(d)– (g)), the cold front exhibits typical features of an atmospheric density current with a nose, an elevated head and a shallower trailing body (cf. Figure 2). Strong subsidence and turbulent mixing appears to occur on the rear side of the head. The spatial resolution of the lidar dataset is not sufficient to resolve turbulent eddies in a great detail. Hence, the flow field is much smoother compared with the higher-resolution Doppler radar observations of Geerts, et al. (2006). Our lidar observations indicate a relatively large eddy with a diameter of about 500– 1000 m at the head. The interface at the top of the trailing body is not flat but exhibits wave structures. Individual wave crests appear to move upvalley, i.e. in the same direction as the density current (compare Figure 10(g) and (h) at equation image km). The propagation of wave crests is illustrated better in a sequence of scans with a smaller interval of about one minute (not shown). Some of these waves appear to break, for example, at x = 3.5 km in Figure 10(h). Presumably, we see gravity waves that are affected by shear-flow instability.

From the change in the position of the nose between Figure 10(e) and Figure 10(g) we derive an average speed of the cold front of uf = 7.1 m s−1. This value is close to the estimate of 7.5 m s−1 in section 3.3 derived from AWS data for the northern part of the Wipp Valley. Based on Figure 10(g) we estimate equation image for the slope of the frontal surface, which corresponds to an inclination angle of 18°. The top heights of the head and body of the density current are about 2.3 km ASL and 1.7 km ASL, respectively. With an average elevation of the valley floor of about 1 km ASL, these two heights correspond to a layer depth of 1300 m for the head and 700 m for the body of the density current. Hence, the head is about twice as high as the body of the density current, which is in accordance with the conceptual model in Figure 2 and lies within the range of values determined from laboratory experiments (Britter and Simpson 1978; Simpson and Britter 1979). The maximum vertical wind shear at the interface between the density current and the foehn flow is about equation image s−1 at the head and about 0.04 s−1 at the trailing body, respectively. These values are based on a vertical mesh size of Δz = 100 m.

Figure 11 shows the flow in the central part of the Wipp Valley after the cold front has passed the lidar site based on two consecutive upvalley scans at α = 174°. Note that the cross-section shown does not cover the whole southern part of the valley but only extends to about 2.5 km south of STE where it intersects terrain (see Figure 1(b)). At 1214 UTC (1221 UTC) the nose of the density current is located about 6.3 km (9.3 km) south of GED. From this change in the position of the nose we estimate an average speed of uf = 7.1 m s−1, which is the same value as derived from the downvalley scans. The two snapshots in Figure 11 show a well-developed elevated head. Figure 11(b) in particular indicates a turbulent wake with strong mixing behind the head (compare with Figure 2).

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Figure 11. As in Figure 10 but for the central part of the Wipp Valley, based on upvalley scans at an azimuth angle of α = 174° (see also Figure 1(b)) at approximately (a) 1214 UTC and (b) 1221 UTC on 6 November 1999.

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In contrast to the conceptual model in Figure 2, we do not see a backflow with reversed winds near the surface. Either the backflow does not exist in our case or we were not able to detect it due to two possible reasons. First of all, beam blocking by local terrain features at low elevation angles causes data gaps near the surface. Secondly, due to its location and scan strategy, the lidar is not able to capture the flow at the very bottom of the valley, i.e. in the narrow, winding gorge embedded in the broader valley. Similarly to the conceptual model and supported by visual observations, a rope cloud forms near the head of the density current. In some snapshots (Darby, et al., 2000) this cloud causes data gaps due to beam blocking. In the lower Wipp Valley (Figure 10(g)), the vertically averaged wind speed in the trailing body of the density current is about 7 m s−1. The maximum wind speed in the core of the body exceeds 10 m s−1. In the central part of the valley (Figure 11(a)), the vertically averaged wind speed in the cold air is about 9 m s−1 with a maximum exceeding 13 m s−1. Hence, winds behind the front are moving towards the front faster than the front itself. This ‘overtaking speed’ of the feeder flow is a typical feature of a density current and is the result of intense mixing between the warmer and the colder fluid at the front (Britter and Simpson, 1978; Simpson and Britter, 1980; Smith and Reeder, 1988).

3.6. Comparison with theory

Following Simpson and Britter (1980), the speed of the density current is given by

  • equation image(4)

with g′ = g(Δθ/θ) being the reduced acceleration due to gravity, hb the depth of the body, ua the speed of the ambient flow (in our case the foehn flow) and a and b positive empirical constants. Here, the sign of ua indicates whether ambient winds are blowing in the same direction (ua > 0) or the opposite direction (ua < 0) to the propagation direction of the density current. Thus, the effect of tail (head) winds is to increase (decrease) the density current speed. Based on laboratory experiments, Simpson and Britter (1980) suggest the values a = 0.91 and b = 0.62. With ht = 700 m, equation image K (Figure 9; 1308 UTC) and ua ≈ − 7 m s−1 (Figure 10(f)), we obtain ud = 7.2 m s−1, which is close to the observed speed (see section 3.5). However, as shown by Smith and Reeder (1988), there is a relatively large uncertainty in the estimation of the flow parameters in Eq. (4). This uncertainty results in a relatively large range of density current speeds, in our case approximately 2 ≤ ud≤ 12 m s−1. Further, slightly different values for a and b appear in the literature with a variation of about ± 10 to 20% (Thorpe, et al., 1980; Wakimoto, 1982; Smith and Reeder, 1988). Moreover, Eq. (4) assumes flat terrain, which is only approximately fulfilled in our case (see terrain profile in Figure 10), and laterally unconfined flow without additional friction due to valley side walls. Thus, the above-mentioned agreement with the observed speed is not a clear proof for the cold front being a density current. We believe that instead it is the qualitative anatomy of the cold front, illustrated in the previous section, that supplies this evidence. Furthermore, not only ud but also the speed of the front uf estimated from our observations in Figure 10 has a broad range of uncertainty. If we use several consecutive scans separated by about 3 min, the corresponding range of frontal speeds is about 3 ≤ uf ≤ 9 m s−1. This range of values may be either real as a result of unsteady propagation or caused by some uncertainty in determining the exact position of the front based on our lidar data (the range gate size is 300 m; with a speed of 7 m s−1 the front passes about four lidar range gates in 3 min).

Laboratory experiments have shown that mixing may occur through billows that develop at the front of the head, roll back above the head, and finally break down forming a turbulent wake (Britter and Simpson, 1978). Dynamic instability of a stratified shear flow occurs over the range equation image (Nappo, 2002), with equation image being the gradient Richardson number (Stull, 1988), defined as

  • equation image(5)

Here, θv is the virtual potential temperature. The condition for convective instability is equation image (Nappo, 2002). In order to estimate whether or not shear-flow instability occurs in our case, we calculate the Richardson number from the radiosonde sounding at Gedeir after the cold front has passed this site (cf. Figure 9). Following Geerts, et al. (2006), gradients in Eq. (5) are derived from data interpolated to equidistant levels with a vertical spacing of 100 m. The minimum Richardson number at the inversion layer separating the density current from the foehn flow is equation image at 1202 UTC at 1.8 km ASL and equation image at 1308 UTC at 2.4 km ASL. Hence, it is most likely that the turbulent structures near the head of the density current in Figures 10(g)– (h) and 11(b) are breaking shear-instability waves. It is also conceivable that some of the turbulence behind the head may be generated by convective instability as cold air is elevated by the convergence at the head over warm air behind.

4. Discussion

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Description and preparation of dataset
  5. 3. Results
  6. 4. Discussion
  7. 5. Conclusions
  8. Acknowledgements
  9. References

Data from the MAP gap-flow project (Mayr, et al.2004) have been used to study the passage of a cold front and the decay of pre-frontal south foehn in an Alpine valley. A dense network of automatic weather stations and in particular a Doppler lidar have proven to be valuable to describe the evolution and structure of a cold front over complex terrain in unprecedented detail.

The synoptic and mesoscale analysis in Figures 4 and 5 shows that the cold front approaches the Alps on 6 November 1999 from the northwest and becomes distorted at the orography, where it accelerates eastward. The acceleration may be enhanced by pre-frontal foehn (Hoinka, 1987; Heimann, 1990, 1992). Such frontal behaviour over the northern Alpine foreland is well known and has been documented in several earlier studies (Steinacker, 1982; Heimann, 1988; Kurz, 1990). However, less well known is the behaviour of the cold front in the inner Alpine region. For our specific case, Figure 12 summarizes the temporal evolution in the Inn and Wipp Valleys. The post-frontal cold air enters the Inn Valley shortly after 1000 UTC from the north via the saddle of Seefeld (SEF) and the Achenpass (ACH). Inflow occurs also via the mouth of the Inn Valley at Kufstein. A smaller part of the cold air spills over the mountain range but is still too warm to reach the valley floor. As a result, cold air flows towards Innsbruck (IBK) from both valley directions, which is consistent with earlier studies (Kaufmann, 1989; Freytag, 1990). The corresponding two cold fronts pass IBK and Hall (HAL) at 1150 UTC and collide in between these two sites at about 1200 UTC. Somewhat earlier, cold air has entered the Wipp Valley. The corresponding front propagates as an atmospheric density current southward and causes the abrupt change from warm southerly to cold northerly winds. With the frontal passage, the southerly foehn flow is lifted above the density current and is replaced at the surface. The front arrives at the Brenner Pass at 1310 UTC. It continues to move down into the basin of Sterzing (STZ) where it arrives at 1350 UTC.

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Figure 12. Schematic diagram of the cold front location in the Inn Valley and Wipp Valley at various times on 6 November 1999. Numbers represent the time in UTC. A dashed line shows the location of collision of two cold fronts. Markers and labels represent AWS sites. Orography is shown as grey-shaded elevation contours with an increment of 500 m starting at 1000 m ASL and as a white contour line at 2000 m ASL.

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Lidar observations in Figures 10 and 11 reveal that the cold front in the Wipp Valley exhibits clear features of a density current: an elevated head that is about twice as high as the trailing body, a flow in the body towards the front that is faster than the propagation speed of the front (i.e. a front-relative feeder flow) and a turbulent wake that has formed as a result of shear-flow instability. The latter is supported by a subcritical Richardson number (equation image) at the interface between the cold-air flow and the foehn. Such density current characteristics have been found for other cold fronts too (Garratt, 1988; Nielsen and Neilley, 1990; Koch and Clark, 1999; Geerts, et al., 2006). The inclination of the frontal surface is about Δzx = 1/3 as derived from lidar data and about 1/7.5 as estimated from a temperature slope profile. These values reveal a much steeper front than in an earlier study conducted in the Inn Valley (Freytag 1990) but are consistent with other observations of atmospheric density currents (Geerts, et al., 2006). However, lidar observations also show that the structure of the head of the density current and therefore also the steepness of the frontal surface is variable in time. It remains unclear whether the difference in the inclination of the frontal surface derived from two different data sets is a result of this temporal (and spatial) variability or the result of an insufficient temporal resolution of the slope-profile data.

Estimated from lidar data, the speed of the cold front is about 7 m s−1, which agrees with the speed predicted by theory for a density current. However, considerable uncertainty in the determination of parameters in the theoretical speed equation (Eq. (4)) raise doubts regarding the applicability of this formula for realistic fronts (Smith and Reeder, 1988; Hoinka, et al., 1990). AWS measurements reveal that the frontal speed is about 6 m s−1 in the Inn Valley (Figure 8(a)). Further, they show that the frontal speed is nearly constant in the northern and central part of the Wipp Valley but decreases from 7.5 m s−1 to 4.5 m s−1 in the southern part of the valley (Figure 8(b)). This decrease in speed may be explained by a decrease in the layer depth of the density current as the front approaches the Brenner Pass (note that the lidar does not capture this part of the valley). Such a reduction in layer depth is conceivable since the valley floor is inclined. The floor increases, especially in the southern part, from about 1000 m ASL at STE to about 1400 m ASL at BRE. According to Eq. (4), ud decreases with hb. Assuming that the top height of the density current stays nearly constant (also other parameters in Eq. (4) except for hb), the corresponding decrease in hb from 700 m to 300 m would reduce ud from 7.5 m s−1 to 3.4 m s−1, which is slightly more than observed. It is likely that the top height of the density current is slightly inclined in the same direction as the valley floor, which would result in a smaller reduction of hb and ud. Another argument for the decrease in propagation speed is the conversion of kinetic energy to potential energy as the density current moves up along the sloped terrain. These considerations are consistent with the numerical simulations of Bischoff-Gauss, et al. (1989), who showed that the effect of a hill on the advance of a density current is to decrease the propagation speed upstream and even more downstream of the obstacle.

5. Conclusions

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Description and preparation of dataset
  5. 3. Results
  6. 4. Discussion
  7. 5. Conclusions
  8. Acknowledgements
  9. References

The evolution and structure of a cold front and the decay of south foehn have been investigated in the Inn Valley and Wipp Valley (Austria) by analyzing in situ and remote sensing observations collected on 6 November 1999 during the MAP field campaign. The following conclusions are drawn from this case study.

  • The air mass behind the cold front is characterized by a relatively shallow layer of cold air that barely exceeds the height of the foothills. As a result the cold front becomes distorted as it impinges from the northwest on to the Alps.
  • The post-frontal cold air that arrives at Innsbruck enters the Inn Valley via two passes, about 45 km apart, in the mountain range north of Innsbruck at nearly the same time. Hence, cold-air advection occurs from both sides of the valley. The corresponding two fronts collide between Innsbruck and Hall.
  • The cold front in the Wipp Valley manifests itself as an atmospheric density current with an elevated head, a turbulent wake behind the head and a front-relative feeder flow in the body. Its propagation speed of about 7 m s−1 in the northern and central part of the Wipp Valley agrees with theory.
  • With the passage of the front, foehn winds decay at the surface but are lifted above the density current where they continue to blow for a while. The strong wind shear between these two opposing flows triggers shear-flow instability at their interface.
  • The frontal surface at the head of the density current has an inclination of about 10°– 20°. The decrease of the frontal speed towards the Brenner Pass by a factor of about two is presumably related to the increase in the height of the valley floor.

This study describes qualitatively the structure of a cold front but does not assess quantitatively the relative contributions of the inflow of cold air from various directions into the Inn Valley. Such an assessment, however, could be made with a mass-budget calculation in a predefined volume and could be tackled in a future study based on high-resolution numerical simulations. From a forecasting point of view, especially for air-traffic control issues at Innsbruck Airport, it would be desirable to derive prognostic rules that would help to assess the exact pathway(s) of a cold front in the Inn Valley and its time of arrival as a function of frontal properties over the foreland, e.g. cold-air layer depth, frontal propagation speed and direction of propagation. For such a study, the MAP dataset does not provide enough cases to obtain significant results in a statistical sense. However, with targeted observations based on standard instrumentation (e.g. AWS) operated over a longer period, such rules could be derived rather easily.

Acknowledgements

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Description and preparation of dataset
  5. 3. Results
  6. 4. Discussion
  7. 5. Conclusions
  8. Acknowledgements
  9. References

We thank the lidar crew from NOAA/ESRL, with Jim Howell, Mike Hardesty, Richard Marchbanks, Scott Sandberg, Ann Weickmann and Janet Intrieri, for operating the Doppler lidar at Gedeir during the MAP SOP. The installation of a dense network of AWS in the Inn Valley and Wipp Valley has been enabled especially by the efforts of Stephen Mobbs. Dave Whiteman is acknowledged for providing HOBO data from the slope profile and Peter Jackson for providing AWS data south of Sterzing. Reinhold Steinacker contributed with his VERA analyses. We are indebted to the Austrian Weather Service ZAMG for access to the database of routine observations. This work was supported by the Austrian Science Fund (FWF) under grants P13489 and P15077.

References

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Description and preparation of dataset
  5. 3. Results
  6. 4. Discussion
  7. 5. Conclusions
  8. Acknowledgements
  9. References
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