The influence of mesoscale circulation systems on triggering convective cells over complex terrain

Atmospheric Research 81 (2006) 150 – 175 www.elsevier.com/locate/atmos

The influence of mesoscale circulation systems on triggering convective cells over complex terrain Christian Barthlott *, Ulrich Corsmeier, Cathe´rine Meißner, Frank Braun, Christoph Kottmeier Institut fu¨r Meteorologie und Klimaforschung, Forschungszentrum Karlsruhe/Universita¨t Karlsruhe, POB 3640, 76021 Karlsruhe, Germany Received 12 June 2005; accepted 19 November 2005

Abstract The project VERTIKATOR (Vertical Exchange and Orography) focused on the initiation and transport processes as well as the model representation of convective systems above low mountain ranges in central Europe. Secondary circulation systems developing during daytime seem to be responsible for triggering convection and subsequent precipitation. The 7 km version of the Lokal-Modell (LM) using the Tiedtke scheme for convection parameterization does not correctly resolve convection initiation forced by slope and valley winds as well as the time, amount, and location of precipitation. All results indicate that the operational version of the LM produces precipitation averaged over large areas without any internal structure over longer time periods than in reality. For simulating single cells with more realistic precipitation concerning time and amount of rainfall, a higher horizontal and vertical resolution is necessary. This is because of better resolved orography and convection initiation by boundary layer circulations. Furthermore, inhomogeneities of the temperature and humidity fields, which are essential for the generation of convection and subsequent rainfall events, are expected to be much too smoothed when using the 7 km grid. The operationally used closure for convection parameterization depends on the large-scale humidity convergence and on turbulent fluxes of latent heat at the surface. Both processes are decisive for the onset and intensity of convection. It has been shown that humidity convergence over the mountain ridges is described by the LM with a 2.8 km grid resolution in a much more realistic way. The better representation of the flow patterns in mountainous areas improves forecasting in these regions. D 2005 Elsevier B.V. All rights reserved. Keywords: Deep convection; Orography; Slope and valley wind regimes; Mass budget; Lokal-Modell

* Corresponding author. Tel.: +49 7247 82 2833; fax: +49 7247 82 4377. E-mail address: [email protected] (C. Barthlott). 0169-8095/$ - see front matter D 2005 Elsevier B.V. All rights reserved. doi:10.1016/j.atmosres.2005.11.010

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1. Introduction Convective processes significantly determine the momentum, heat and water vapor fluxes between the ground and the upper troposphere. In particular during cases with deep convection, polluted air masses are dispelled from the atmospheric boundary layer. Convection therefore is one major factor controlling the air quality. Due to multiple risks from hail, lightning, strong winds, and heavy precipitation, convective storms also have strong economic impacts. With the current horizontal resolution of numerical weather prediction models (about 10 km), convective processes cannot be described explicitly (Weisman et al., 1997). Consequently, parameterizations for its effects are needed (Arakawa and Schubert, 1974; Tiedtke, 1989; Donner et al., 2001). For the evolution of convection, three conditions have to be fulfilled: (1) a sufficient amount of humidity in the lower troposphere, so that the convective condensation level (CCL) can be reached and clouds can develop; (2) an appropriate density stratification of the troposphere; and (3) a small-scale triggering process at the ground by heating of the land surface and/or largerscale lifting of the air. The supply with humidity and lifting of air masses result from synoptic systems (1000 km) on a larger scale, from local systems (50 km), and turbulent transfer processes on the scale of a few kilometers. A number of observational and numerical studies have noted that mountain areas are frequently characterized by enhanced convection (Orville, 1965; Toth and Johnson, 1985; Tripoli and Cotton, 1989; Tian and Parker, 2002; Stein, 2004). The turbulent and convective processes over complex terrain differ significantly from those found over flat terrain (Banta, 1984). According to Founda et al. (1997), the variances of the vertical wind component were larger on the summit of a hill in complex terrain than over flat terrain. Particularly, thermally driven local wind systems seem to play an important role in the initiation of convection over mountainous terrain (Raymond and Wilkening, 1980). Pielke and Segal (1986) pointed out that the upward velocities associated with upslope flows provide a trigger mechanism for cloud formation and afternoon showers in many mountainous areas. Furthermore, the improvement of local weather forecasting is linked with an improved understanding of mesoscale circulations through both observations and models. Mesoscale circulations due to sensible heat flux discontinuities may produce a trigger mechanism of vertical motion for deep convection by lifting boundary-layer air to their level of free convection (Pielke, 2001). Hanesiak et al. (2004) stated that mesoscale circulations play a key role in the development of deep convection over highland areas under weak synoptic flows. As a consequence, a high variability of the formation of convection can be observed above complex terrain, even if the conditions on a larger scale are almost the same. In particular, boundary-layer wind convergence associated with the local wind circulations is considered to be a main factor controlling the initiation of deep cumulus convection (Pielke et al., 1991; Pielke, 2001). Homar et al. (2003) pointed out that mountain breezes could favor the mesocyclogenesis by convergence under certain mountain ridge shapes. To improve the predictability of strong precipitation and thunderstorms, a better understanding of the triggering effects for convection caused by orographic influences is required. In order to detect the triggering effects of complex terrain for the initiation of convection as completely as possible, extensive measurements were performed in May and June 2002 in the northern Black Forest in southwest Germany under the AFO2000 project VERTIKATOR (Vertical Exchange and Orography). In addition, numerical simulations with the Lokal-Modell (LM) of the Deutscher Wetterdienst (DWD) were run using different model settings. During the intensive

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operation period (IOP) on June 19, a single thunderstorm was observed over the Murg valley in the northern Black Forest. This IOP is also subject of the articles by Bertram et al. (2004) and Meißner et al. (submitted for publication). The latter deals with adjusted humidity profiles which led to a better reproduction of convective initiation in the LM, whereas Bertram et al. used a 3D cloudresolving model in order to investigate the microphysical development of the storm observed. The purpose of this paper is to determine the influence of secondary flow systems on the initiation of deep convection above complex orography by comparing LM simulations with observations from VERTIKATOR. For this reason, the LM was used in its operational version (horizontal resolution 7 km, with convection parameterization) and with a higher grid resolution of 2.8 km without convection parameterization. These shall hereinafter be referred to as LM7 and LM2.8. The focus lies on thermally induced wind systems and their influence on the mass budget of convective systems represented by selected control volumes above the investigation area. 2. Measurements and numerical simulations 2.1. The VERTIKATOR project The VERTIKATOR project focused on orographic influences on convection and its model representation. Within VERTIKATOR, the complex structure of the atmosphere prior to and during the convective processes was recorded with a number of different measurement systems. The measurement area chosen was the northern Black Forest located in southwestern Germany. With heights up to 1500 m above sea level, the Black Forest can be considered a typical lowmountain range. The field phase was conducted in May and June 2002, including a total of 7 intensive operation periods comprising cases with shallow (May 31, June 1) and deep convection (June 19). The measurements included the operation of five research aircraft between 100 m agl and 6000 m asl that flew on individually designed flight patterns and several groundbased stations (Fig. 1). The ground stations were concentrated mainly along a cross section perpendicular to the main mountain ridge from the Rhine valley to the Neckar valley. Besides a network of automatic weather stations, two energy balance stations, four turbulence-resolving stations, two sodars, a C-band Doppler radar, an aerosol lidar, a disdrometer, a tethered balloon, and a wind-temperature profiler were used. In addition, vertical profiles of wind, temperature, and humidity were measured at two radiosonde stations from 04:00 to 20:00 UTC (every 2 h) and with dropsondes (Kottmeier et al., 2001). A more detailed description of the ground-based stations employed is given by Kalthoff et al. (submitted for publication). 2.2. Synoptic controls on June 19, 2002 As indicated by the satellite images in Fig. 2, the IOP of June 19 was a case of deep convection in the area of investigation. Large parts of south-western Germany were covered by the cirrus shields of thunderstorms in the late afternoon. On this day a trough over the eastern Atlantic led to a south-westerly upper air flow, whereas winds from north to northeast prevailed in the near-surface area (Fig. 3). Weak surface pressure gradients allowed for the formation of thermally induced wind systems in the area of interest. A single thunderstorm was observed over the Murg valley in the northern Black Forest between 12:55 and 14:30 UTC. By means of a Cband Doppler radar located at the Forschungszentrum Karlsruhe (Fig. 1, station 12), the development of this isolated cell could be documented (Fig. 4). For a detailed description of the most important stages of the storm’s development, see Bertram et al. (2004).

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Fig. 1. Location of ground-based stations during VERTIKATOR.

In order to indicate the potential of severe weather on this day, the convective available potential energy (CAPE) and the convective inhibition (CIN) were calculated on the basis of the data from the two radiosonde stations Freistett (Rhine valley, Fig. 1, station 2) and Horb (eastern part of northern Black Forest, Fig. 1, station 10). CAPE is a measure of the conditional stability of the troposphere to a finite vertical displacement, as it occurs during thunderstorms (Emanuel, 1994). It is defined as the vertical integral of a lifted air parcel’s buoyancy between its level of free convection (LFC) and the equilibrium level (EL) Z EL Tparcel  Tenv CAPE ¼ R dp ð1Þ p LFC where R is the universal gas constant, T parcel is the air parcel temperature, Tenv is the sounding environmental temperature, and p is pressure. CAPE values below 1000 J kg 1 indicate weak instability, values from 1000 to 2500 J kg 1 moderate instability, and values

Fig. 2. NOAA satellite images of June 19, 2002 (from left to right: 06:00, 12:00, 18:00 UTC). The black rectangle marks the research area.

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Fig. 3. Surface pressure (hPa, white isobars) and 500 hPa geopotential (gpdm, colour code) on June 19, 2002 (00:00 UTC). Data: NCEP reanalysis; Source: www.wetterzentrale.de.

higher than 2500 J kg 1 indicate strong instability. CIN represents the amount of negative buoyant energy available to inhibit upward vertical acceleration and is calculated as the integral of the negative portion below the LFC, where the parcel temperature is cooler than

Fig. 4. Radar reflectivity measured by the Doppler radar at the Forschungszentrum Karlsruhe on June 19, 2002. Maximum reflectivity of the storm over the Murg valley was 60 dBZ. The cloud top was 12 km and precipitation measured by the rain gauge was 30 mm. The duration of the storm was less than 90 min.

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that of the environment. The larger the CIN is, the stronger must be the amount of forced lift to bring the parcel to its LFC. Radiosonde observations at Freistett and Horb reveal CAPE values between 1200 and 2000 J kg 1 and CIN values between 51 and 142 J kg 1 from 12:00 to 18:00 UTC (Fig. 5). This indicates that with appropriate forcing the available potential energy can be released. Due to the simultaneous decline of CIN and rising of CAPE from 08:00 to 12:00 UTC, the initiation of convection is most likely after 12:00 UTC. By the release of dropsondes from the research aircraft DO-128, 14 additional soundings are available in the vicinity of the storm from 13:39 to 14:03 UTC. Concerning temperature and wind profiles, the measurements with the radiosonde and the dropsondes are almost identical. However, dew point data differ considerably in the region of interest. Following an approach by Bertram et al. (2004), the lowermost 1500 m agl dew point data of the Freistett sounding were replaced by the nearest dropsonde profile available in order to have reliable dew point values for the storm’s environment. In the vicinity of the storm, humidity was distinctly higher than in the Rhine valley, leading to a lower lifting condensation level (LCL) and LFC (Fig. 6). As a consequence, the CAPE value increased to almost 2500 J kg 1. The height of the mixed layer was limited by a weak inversion at 1000 m agl. The wind profile displayed weak near surface winds from north to northeast, turning with height to southwest and indicating a moderate shear flow with warm air advection in the lowest 3 km agl. Other thermodynamic variables are given in Table 1. 2.3. The Lokal-Modell of the Deutscher Wetterdienst Since 1999, the LM has been part of the operational weather forecast system of the DWD. It is a non-hydrostatic, limited-area atmospheric prediction model for applications on the meso-g and meso-h scales (Doms and Scha¨ttler, 1999). The calculations are based on the unfiltered Euler equations, with the three-dimensional wind vector, temperature, perturbation pressure, specific humidity, and cloud water content as prognostic variables. Optionally, cloud ice content as well as turbulent kinetic energy can be computed. The Arakawa C-grid is used for horizontal differencing on a rotated latitude/longitude grid with a horizontal resolution of 7 km. The number of vertical layers in a generalized terrain-following coordinate is 35. The subgrid-scale turbulence is parameterized using a diagnostic second-order closure of hierarchy level 2 for vertical turbulent fluxes according to Mellor and Yamada (1974). A scheme according to Louis (1979) is used to

CAPE Freistett CAPE Horb CIN Freistett CIN Horb

2000

J kg -1

1600 1200 800 400 0 3

4

5

6

7

8

9

10 11 12 13 14 15 16 17 18 19 20 21

UTC Fig. 5. Diurnal cycle of CAPE and CIN at the radiosonde stations of Freistett and Horb.

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11 12 10 9 8 13 7

150 2

200

EL

1

4 3 5 0

6 13 km

SFC

10ms-1

250

20ms-1

12km 11km 10km

300

p in hPa

9 km 8 km

400 7 km 6 km

500

5 km

600 4 km

700

3 km

LFC

850

2 km

LCL

1 km

1000 -80

0 km

-60

-40

-20

0

20

40

T in °C Fig. 6. Sounding of June 19, 2002, 14:00 UTC at Freistett. Dew point data of the original sounding (dashed line) and after modification with dropsonde data (solid line). Hodograph points are plotted at 1 km intervals from the surface.

simulate the surface layer and a mass flux scheme for cumulus parameterization developed by Tiedtke (1989) is included. The impacts of short- and long-wave radiation are parameterized according to Ritter and Geleyn (1992). The connection between soil and atmosphere is established through the energy balance equation, whereas the required soil parameters (temperature and moisture) are calculated with the soil model Terra (Doms and Scha¨ttler, 1999). Apart from the simulations with the operational version of the LM, model runs with LM2.8 were performed to determine possible improvements of the representation of convection by the Table 1 Bulk thermodynamic variables for the modified 1400 UTC sounding at Freistett Thermodynamic parameter

Value

Lifting condensation level LCL (hPa) Level of free convection LFC (hPa) Equilibrium level EL (hPa) Convective available potential energy CAPE (J kg 1) Convective inhibition CIN (J kg 1) Lifted index (K) Convective temperature (8C) Precipitable water (mm)

872.8 771.7 184.8 2467.5 31.9 6.26 32.85 41.06

Air parcel indices are computed using surface values.

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Table 2 Model settings

Grid resolution Convective parameterization Time step Vertical layers

LM7

LM2.8

7 km yes 40 s 35

2.8 km no 8s 40

model. For models with resolutions of less than 4 km, the major fraction of deep convection with subsequent precipitation can be calculated explicitly (Weisman et al., 1997). The 2.8 km resolution is sufficient to give up a convection parameterization (Adrian, 2004; Walser et al., 2003). Models of this scale use one or more new prognostic equations for cloud water, cloud ice, rain water, snow, and other hydrometeor types. However, parameterizations of the cloud microphysics and of small-scale turbulence are still necessary. LM7 and LM2.8 model runs were initialized at 00:00 UTC with the analysis data of the Global-Modell (GME). See Table 2, for the main model settings. 3. Experimental findings and model results 3.1. Density stratification, convective initiation, precipitation and radar observations Atmospheric density stratification is one major factor controlling the formation and further evolution of convective systems. In order to compare the large-scale conditions in the area under investigation, the measured and simulated profiles of potential temperature h, equivalent potential temperature h e, and saturated equivalent potential temperature h es measured at Freistett and Horb are given in Fig. 7. As can be seen, the LM7 detects the stratification of the atmosphere, since the model simulates the profiles well, even if the fine structure cannot be 400

400

Freistett measured simulated

500

500

600

p in hPa

600

p in hPa

Horb measured simulated

700

700

800

800

900

900

θ

1000 300

320

340

K

θes

θe

θ

θes

θe

1000 360

380

300

320

340

360

380

K

Fig. 7. Measured and simulated profiles of potential temperature h, equivalent potential temperature h e, and equivalent potential temperature assuming saturation h es on June 19, 2002 (14:00 UTC).

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modeled. The results show that the air was conditionally unstable up to a level of about 550 hPa, and release of conditional instability may occur when rising parcel reaches its LFC. In order to investigate the initiation of convection in the LM, the convective temperature Tconv was determined for each grid point. On the basis of hourly LM results, the exceeding of the convective temperature can thus be investigated by comparison with the simulated temperatures at 2 m height. The results show that the initiation of convection occurs over a large part of the northern Black Forest, beginning at 11:00 UTC (not shown). On the other hand, the measurements show that the convective temperature is reached only at the highest mountain ridges because of high humidity and temperature values (Meißner et al., submitted for publication). Due to widespread initiation, the model simulates a wide-stretched cloud formation, leading to precipitation in the whole region (Fig. 8). However, measurements with the IMK precipitation radar located at the Forschungszentrum Karlsruhe reveal strong differences from the accumulated precipitation rates simulated with the LM7 model run: (i) the simulation, including the Tiedtke scheme for parameterization of convection, shows 13 mm rainfall only in the area of the convective cell over the Murg valley, east of Strasbourg; (ii) the simulated precipitation covers too large an area; and (iii) the rainfall measured by radar is up to 30 mm in the center of the cell with an area smaller than 5 km by 7 km in size (Fig. 9). The radar observations only reveal a single, isolated thunderstorm over the Murg valley with a lifetime of about 90 min. 3.2. Wind systems and model topography It is known from many studies that small-scale secondary flow systems exist over mountainous terrain, leading to strong inhomogeneities in the flow structure, e.g. thermally induced wind systems play an important role in the initiation of convection over complex terrain. 24 mm

49.4

22 mm 49.2

20 mm

Karlsruhe

18 mm

Enz

16 mm

Rhi ne

ig nz

7.6

7.8

8.0

8.2

8 mm

AB I 8.4

8.6

6 mm 4 mm 2 mm

SW

48.0

10 mm

AN

BLAC

48.2

ch

n Re

48.4

12 mm

JU RA

Strasbourg

14 mm

K FO RES

48.6

g

T

Mur

48.8

Ki

northern latitude in °

49.0

0 mm 8.8

9.0

eastern longitude in ° Fig. 8. Total precipitation on June 19, 2002 as simulated by LM7. Isolines represent the model topography.

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Fig. 9. Total precipitation on June 19, 2002 from 12:00 to 24:00 LT, as derived from radar observations. The white rectangle marks the discussed thunderstorm over the Murg valley.

Under weak lower tropospheric synoptic forcing, small scale upslope winds lead to convergence zones above the mountain ridges, which provide vertical motion due to the conservation of mass and subsequently initiate cumulus convection (Raymond and Wilkening, 1982; Pielke, 2001). In addition, valley winds caused by stronger heating over the mountains compared to the free atmosphere at the same elevation carry humid air towards the mountain ridges. Consequently, the initiation of such systems is a combination of hydrodynamic and thermodynamic processes, where boundary-layer processes play a key role: By differential heating of the earth’s surface, humidity and heat are transferred to upper layers, where the release of latent heat due to condensation and other microphysical cloud processes induces enhanced convection. Recent investigations of Kossmann and Fiedler (2000) and Kalthoff et al. (2000) document the presence of slope and valley winds in the area under investigation. The lower the mean wind speed, the greater is the impact of the interaction processes between soil surface and atmosphere on the near-surface air temperature. The temperature at a specific place is dependent on the soil-type, soil moisture, land use, and topographical features like steepness and exposure to the sun. The primary forcing term in the momentum equation of slope flows is the along-slope component of the buoyancy force, with an additional along-slope pressure gradient term. Valley winds are caused by a stronger heating of the air over the mountains compared to the free atmosphere at the same elevation. By means of a scanning infrared radiation thermometer, the soil temperature was measured on board of the research aircraft DO-128. The aperture of the instrument is 1.158, leading to a diameter on the ground of 10 m, if the flight height is 500 m above ground. Flights over the Murg valley (Fig. 1, between 7 and 9) in the afternoon show that the slope towards the west exhibits soil temperatures from 24 to 30 8C, which are considerably higher than those found on the slope facing the east of the Murg valley (18–23 8C). Under these circumstances, the evolution of the boundary layer and near-surface temperatures mainly depends on the partitioning of the available solar energy into the turbulent fluxes of the sensible heat H and the latent heat LE. This is expressed by means of

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the Bowen ratio b = H/LE. In elevated terrain, sunlight strikes the slopes of the mountains more directly than the base, and increased heating drives an upslope flow and increases sensible and latent heat fluxes (Fig. 10). In the Rhine valley, low Bowen ratios of around 0.25 prevail, indicating that the available radiation energy is mainly transformed into latent heat and humidity is accumulated within the lower planetary boundary layer. This air can be included in convective processes over the mountain ridges by slope and valley wind regimes. At higher elevated terrain like Brandru¨ttel, Hornisgrinde or Leimiss, higher Bowen ratios of 0.5–1 during the day indicate that more energy is transformed into sensible heat than in the Rhine valley. Under such conditions, the evolution of a high and well-mixed boundary layer is enhanced. The interaction of low Bowen ratios in the Rhine valley with higher ones above the mountain region indicates that elevated terrain is a preferred location for the formation of convective clouds. Higher Bowen ratios in the Black Forest (between 1 and 2) than in the Rhine valley exist for most of the year (Wenzel et al., 1997). The results show that the sensible heat flux is dominating at higher elevated terrain and drives local circulation systems and wind convergence. On the other hand, latent heat flux is dominating in the Rhine valley, giving the humidity supply for the mountain region by upslope and valley winds. Freistett 6 2.5

Brandrüttel

8

10

8

10

Hornisgrinde

Leimiss

12

14

16

18

12

14

16

18

Bowen ratio

2.0 1.5 1.0

Sensible heat flux in Wm-2

Latent heat flux in Wm-2

0.5 0 400 350 300 250 200 150 100 50 0 250 200 150 100 50 0 -50 6

UTC Fig. 10. Bowen ratio, sensible and latent heat flux on June 19, 2002 at stations Freistett (Rhine valley, 2), Brandru¨ttel (slope facing west, 5), Hornisgrinde (mountain ridge 1163 m asl, 6), and Leimiss (slope facing east, 7). Locations of sites are given in Fig. 1.

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If real-world thermally induced wind systems shall be represented by the LM, the model topography has to exhibit the orographical structures of the real topography in a simplified form at least. In comparison with a high-resolution topographical data set (Fig. 1), the model topography of the operational LM version shows strong smoothing of the orographical complex terrain (Fig. 11a). The Murg and Rench valleys are even not included in this version. The highest elevation of the northern Black Forest (Hornisgrinde) with a height of 1163 m above msl is not located at the highest point of the model topography and the difference in height from the model is striking, reaching a value of 562 m. The model height of another instrumented site on the eastern slope at Leimiss (Fig. 1, site 7) is even higher than the model height of Hornisgrinde. A much more realistic description is given by the LM2.8 (Fig. 11b). The Hornisgrinde mountain is the highest model point and the valleys of Murg and Rench are resolved quite well. Similarly, the Kinzig valley is more distinct and the orographic structures in general are described in more detail. As a consequence of the strong smoothing of the orography in the 7 km version, it is questionable, whether the model can reproduce the local wind systems in the absence of a dominating mean flow. The diurnal cycles of the wind directions observed and simulated at four stations for June 1 is given in Fig. 12. On this day, weak synoptic pressure gradients, together with a cloud-less sky until the afternoon, allowed for the formation of slope and valley winds in a well-defined way. Shallow cumulus convection was observed above the mountain tops in the afternoon only (Kalthoff et al., submitted for publication). The model simulates the wind directions in the Rhine valley and on top of the Hornisgrinde quite well. Due to the channelling effect (Kalthoff and Vogel, 1992), the wind direction prevailing in the Rhine valley (Freistett) is northeast. On the eastern (Leimiss) and the western slopes of the northern Black Forest (Brandru¨ttel), upslope winds evolve during daytime. It is significant that LM7 does not resolve the upslope wind on the western slope of the mountain. The same applies to the valley wind in the Kinzig valley at Gengenbach, where the model again is not capable of reproducing the corresponding wind direction. Due to the north-easterly near-surface flow prevailing, the upslope wind on the eastern slope is simulated. However, the down-slope winds during the night are not resolved by LM7. These characteristics of the simulated and measured wind field are also valid for the IOP of June 19. However, due to the superposition of local wind systems with the intense convective flow patterns, these data become less distinct. Due to the poor reproduction of the local wind systems by the LM7, additional simulations were performed with a higher horizontal grid resolution of 2.8 km. The mass flux approach to the parameterization of convection (Tiedtke, 1989) used in the LM is conceptionally not designed for high-resolution modeling, where convective processes are either resolved explicitly or can occur because of the parameterization scheme. The comparison of the measured surface wind field with LM2.8 simulations is given in Fig. 13 for 11:00 UTC. The measurements reveal two main wind systems: Valley winds in the Kinzig valley and upslope winds on the eastern and western slopes of Hornisgrinde, leading to a convergence zone above the mountain ridge. Sodar measurements reveal the vertical extension of the valley winds of up to 400 m, while slope winds typically reach an extension of some tens of meters. As was pointed out by Meißner et al. (submitted for publication), the location of the convergence zones basically agrees with the location of cumulus clouds on the satellite images (Fig. 2) and, thus, indicates their influence on the initiation of convection. Both wind systems are not simulated by the LM7 (not shown). However, the wind field of the LM with a higher spatial resolution agrees much better with the observations. The valley winds in the Kinzig and Murg valleys are simulated well and the upslope winds from the east and west on the Hornisgrinde are also resolved, leading to a strong

162

altitude above sl

altitude above sl

49.1

900 m

800 m

Freistett

48.1 48.0

400 m

100 m 8.2

8.4

8.6

eastern longitude in °

8.8

48.3 48.2

48.0

47.8 8.0

48.4

300 m

200 m

7.8

48.5

48.1

47.9

7.6

northern latitude in °

600 m

500 m

JU RA

48.2

r Necka

SW AB IA N

Rhi

ne

48.3

Horb

Strasbourg

48.6

ig nz Ki

ig nz Ki

48.4

700 m

9.0

900 m

Freistett

h nc Re

h nc Re

48.5

Hornisgrinde

B LA C K F O R E S T

Strasbourg

48.6

rg

48.7

ne

48.7

1000 m

Mu

48.8

800 m

Hornisgrinde

Horb

r Necka

700 m 600 m

JU RA

rg

48.9

Rhi

48.8

1100 m

Enz

500 m 400 m

SW AB IA N

Mu

Karlsruhe

B L A C K F O R E S T

Enz

48.9

b)

49.0

300 m

47.9

200 m

47.8

100 m 7.6

7.8

8.0

8.2

8.4

8.6

eastern longitude in °

Fig. 11. Model topography of southwest Germany: (a) LM7; (b) LM2.8.

8.8

9.0

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Karlsruhe

a)

49.0

northern latitude in °

1200 m

49.1 1000 m

measurements 360

Leimiss

360

330

300

300

270

270

270

210 180 150 120

wind direction in °

330

300

wind direction in °

330

240

240 210 180 150 120

210 180 150 120

90

90

60

60

60

30

30

30

0

0 3

6

9

12

15

18

21

24

0 0

3

6

UTC

Freistett

48.65

nc

Re

Brandruettel Hornisgrinde Leimiss

h

48.60 48.55 48.50

ig

nz Ki

48.45 48.40

Gengenbach

12

15

18

21

24

7.9

8.0

8.1

8.2

8.3

eastern longitude in °

8.4

8.5

3

6

330

300

300

270

270

240 210 180 150 120

9

12

15

18

21

24

18

21

24

UTC Brandruettel

360

330

240 210 180 150 120

90

90

60

60

30

30

0

48.35

0

wind direction in °

ine Rh

wind direction in °

48.70

9

UTC Freistett

360

Murg

48.75

northern latitude in °

240

90

0

Gengenbach

360

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wind direction in °

model simulations (LM-7km)

Hornisgrinde

0 0

3

6

9

12

UTC

15

18

21

24

0

3

6

9

12

15

UTC 163

Fig. 12. Comparison of wind directions measured and simulated on June 1, 2002 at 3 mountain sites (Brandru¨ttel 5, Hornisgrinde 6, Leimiss 7), 1 valley site (Gengenbach 11), and 1 Rhine valley site (Freistett 2).

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49.1 49.0

3ms-1

a) En

48.9

z

3ms-1

ch en

Mur g

48.8 48.7 48.6 48.5 48.4

g

48.1 7.6

48.2

7.8

8.0

8.2

8.4

eastern longitude in °

8.6

8.8

48.1 7.6

Rh

Rh

ine

48.3

ine

g

48.2

2ms-1

ch en

R i nz Ki

48.3

z

R

northern latitude in °

48.6

48.4

En

i nz Ki

northern latitude in °

48.7

48.5

49.0

2ms-1

b)

48.9

Mur g

48.8

49.1

7.8

8.0

8.2

8.4

8.6

8.8

eastern longitude in °

Fig. 13. Surface wind field on June 19, 2002 (11:00 UTC): (a) LM2.8 (every second vector is plotted), (b) observations.

convergence zone in the southern part of the northern Black Forest, which is some kilometers more to the southeast than in reality. 3.3. Convection in the model It was already pointed out that the area of precipitation simulated by the operational LM covers the whole Black Forest and therefore does not agree with the radar observations (Figs. 8 and 9). The widespread exceeding of the convective temperature leads to the activation of the convection parameterization, but the simulated convective precipitation is too weak. Meißner et al. (submitted for publication) compared the temporal evolution and spatial distribution of the observed precipitation with the precipitation simulated by the LM7. Regarding the isolated thunderstorm in the Murg valley, it was found that precipitation was modeled 4 h earlier than observed. The LM7 simulated precipitation for about 9 h in several parts of the model domain starting from 09:00 UTC, while observations only reveal about 4 h of precipitation starting from 12:30 UTC. In order to reveal the occurrence of convection of the LM2.8 simulations, Fig. 14 gives the vertical cross sections of liquid water content and wind vector a little south of the Hornisgrinde for the time of the thunderstorm observed over the Murg valley. In contrast to the simulations with the 7 km version (not shown), the LM2.8 reflects the circulation systems observed above the much more realistic orography quite well: A single convective cell develops at 12:30 UTC, which rapidly grows to a height of 7 km above ground until 13:00 UTC. The largest vertical wind speeds in the core of the cell of about 8–10 m s 1 are reached at 13:30 and 14:00 UTC, when the convective system covers the entire troposphere, which is in good agreement with the radar observations. At much lower vertical velocities, the LM7 does not exhibit any features of deep convection. Due to the intense vertical exchange, including cloud formation, precipitation is simulated by the LM2.8 (Fig. 15). Comparing the temporal and spatial distributions of the observed precipitation, it is found that the simulated amount of rainfall (30–35 mm) is close to the measurements of 30 mm by the Karlsruhe C-band radar (Fig. 16) and that the center of the cell is shifted to the south by a few kilometers only. The duration of precipitation of about 60–90 min

height in m

800

800

600

600

400

400 model orography

200

10 3 10

8

8

6

6

4

4

2

2

1230 UTC

3

1300 UTC

wmax = 1.24 m s-1

wmax = 5.96 m s-1

0

ms -1 10

0 7.9

8.0

8.1

8.2

8.3

8.4

8.5

8.6

8.7

8.8

10

8.0

8.2

8.4

8.6

8.8

10 3 10

8

8

6

6

4

4

2

2

1330 UTC

1400 UTC

wmax = 9.73 m s-1

wmax = 8.18 m s-1

0

0 7.9

8.0

8.1

8.2

8.3

8.4

8.5

eastern longitude in °

8.6

8.7

8.8

7.9

8.0

8.1

8.2

8.3

8.4

8.5

eastern longitude in °

8.6

8.7

1.4 g kg-1 1.3 g kg-1 1.2 g kg-1 1.1 g kg-1 1.0 g kg-1 0.9 g kg-1 0.8 g kg-1 0.7 g kg-1 0.6 g kg-1 0.5 g kg-1 0.4 g kg-1 0.3 g kg-1 0.2 g kg-1 0.1 g kg-1 0 g kg-1

C. Barthlott et al. / Atmospheric Research 81 (2006) 150–175

height above ground in km

10

height above ground in km

model orography

200

8.8

Fig. 14. Vertical cross section of the wind vector in the x–z plane and liquid water content south of Hornisgrinde on June 19, 2002, as simulated by LM2.8. Only every second wind vector is plotted. 165

166

1300 UTC g

25 mm h-1 20 mm h-1 15 mm h-1

48.6

10 mm h-1

Mur

15 mm h-1 10 mm h-1

48.5 Ki

0 mm h-1

ig

8.8

7.8

8.0

eastern longitude in ° 30 mm h-1

48.7

g

25 mm h-1

ine Rh

20 mm h-1 15 mm h-1

48.6

10 mm h-1

Mur

48.7

15 mm h-1

48.6

10 mm h-1

48.5

0 mm h-1

5 mm h-1

ig

8.8

20 mm h-1

nz

eastern longitude in °

8.6

25 mm h-1

ine Rh

Ki

8.4

g

ch

ch

ig

8.2

30 mm h-1

48.8

48.4

8.0

0 mm h-1

35 mm h-1

z En

1430 UTC

5 mm h-1

nz

7.8

8.8

en

en

Ki

48.4

8.6

R

48.5

48.9

northern latitude in °

1400 UTC Mur

8.4

eastern longitude in ° z En

48.8

8.2

35 mm h-1

R

northern latitude in °

48.9

5 mm h-1

nz

ig

8.6

ch

ch

nz

8.4

20 mm h-1

48.6

5 mm h-1 8.2

25 mm h-1

ine Rh

48.4

8.0

g

en

en

Ki

7.8

30 mm h-1

R

48.5

1330 UTC

48.8 48.7

35 mm h-1

z En

7.8

8.0

8.2

8.4

eastern longitude in °

Fig. 15. Hourly precipitation rates on June 19, 2002, as simulated by LM2.8.

8.6

8.8

0 mm h-1

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ine Rh

northern latitude in °

Mur

48.4

48.9

30 mm h-1

48.8 48.7

35 mm h-1

z En

R

northern latitude in °

48.9

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167

Fig. 16. Precipitation rates as calculated from radar observations on June 19, 2002.

fits well with the measured values. The observations show that the thunderstorm base is almost stationary, whereas the top is shifted with the mean wind into northeasterly direction. The simulations, however, reveal a different behavior: The cell is moving towards the southeast. But with respect to the initiation of convection, the vertical extension, and the temporal evolution of the cell, the simulations are close to reality. 4. Mass budget calculations 4.1. Method In order to assess the importance of thermally induced wind systems and the accompanying mass convergence to the initiation of convection above orographically structured terrain, mass budget calculations were performed for a number of control volumes in the area of investigation. The selection of reasonable control volumes for a low mountain region (Fig. 17) allows for a detailed analysis of transport and its contribution to the horizontal mass convergence which finally is considered to be the decisive mechanism for the initiation of lifting processes. An explicit comparison is only possible between the two model versions, since the estimation of the mass budget based on measurement data only is too uncertain. Although large measuring efforts have been conducted, too few data exist in a reasonable time window of about 30 min. As was documented e.g. by Whiteman et al. (1996), many assumptions are required to conduct mass budget calculations from measurements. On the other hand, numerical experiments provide information about the mass budgets, because they provide data on all grid points and represent dynamically balanced realizations of the flow. The further time series refer to control volumes ranging from the earth’s surface to a height of 3 km above ground. The integral form of the mass continuity equation is (Whiteman et al., 1996): Z Z Bq dV þ jd ðqvY ÞdV ¼ 0; ð2Þ V Bt V

168

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Fig. 17. Sketch of the control sectors for the mass budget calculations.

where q is the air density, t is time and vY is the velocity vector. The volume integral of the divergence jd ðqvY Þ is related to the surface integral of qvY by Gauss’ divergence theorem: Z Z jd ðqvY ÞdV ¼ ðqvY ÞnY d dA; ð3Þ V

BV

Y

where n is the unit vector normal to an infinitesimal element of surface area dA. In the absence of the creation or destruction of matter, the density within a volume can change only by flow into or out of the volume through its boundary. The surface integral may then be split into the mantle M (vertical boundary of the control volume) and top cap T (lid) (Kalthoff et al., 2002): Z Z Z ðqvY ÞnY d dA ¼ ðqv8 Þd dM þ ðqwÞd dT ð4Þ BV

BV

Bv

where only the perpendicular components of the wind vector v ?and w are relevant. Since the vertical velocity at the surface is 0, the bottom level can be neglected. In a final step, Eq. (3) can be written as: Z jd ðqvY ÞdV ¼ F M d Rðqv8 Þ þ F T dRðqwÞ; ð5Þ V

where F M and F T are the respective surface areas of the control volume. The perpendicular components of the wind vector are model output variables, whereas the air density can be calculated from other model output data. With the given temperature T and relative humidity, the water vapor pressure e can be calculated. The total pressure p consists of the pressure caused by dry air p d and the water vapor pressure. The ideal gas law yields: p ¼ pd þ e ¼ md Rd T =V þ mv Rv T =V ;

ð6Þ

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169

with the gas constants for dry air R d and for water vapor R v. The required variables therefore are given by: pd e and qv ¼ : ð7Þ qd ¼ Rv T Rd T Eq. (5) can then be calculated by replacing q by q d and q v. The wind vector is positive into the control volume and negative out of the control volume. Mass convergence will therefore be represented by a positive mass budget in our calculations. In order to simplify the calculations, horizontally homogeneous conditions are assumed. Under this assumption, horizontal turbulent fluxes may be neglected, because they become divergence-free. Since the upper boundary of the control volumes is higher than the height of the boundary layer, no vertical turbulent fluxes have to be considered. It must be stated that uncertainties exist particularly for the smaller control volumes of the lower resolved model because of the presence of model data at discrete points only. Due to the large surface area of the boundary, the order of magnitude of the mass budgets is mainly determined by the surface area of the boundary itself. The dimensions of all control volumes with a height of 3 km are given in Table 3. Due to the different horizontal resolutions, the positions of the corners differ between the two model versions. The control volume Hornisgrinde e.g. has a mantle area of 236.4 km2 and a top cap area of 384.4 km2 in the LM7 version, leading to a 7% smaller total area compared with the higher resolved model. This must be considered when discussing the results. 4.2. Time series of mass budgets The diurnal cycle of the mass budgets for dry air for both model versions is given in Fig. 18. It can be seen that both model versions qualitatively show the same characteristics for most of the control volumes. Hornisgrinde, northern Black Forest, and the Murg valley are characterized by a net outflow during the night, which turns into a mass convergence starting from 04:00 to 06:00 UTC. After the onset of larger-scale convection at 12 UTC, the curves of the higher resolved model show a high variability due to the large transport processes occurring. These variations do not appear in the LM7, since the convective processes are too weak and they occur over too widespread an area compared to the observations. Another difference between the model versions concerns the amplitude of the mass budgets: Hornisgrinde with LM2.8 shows an approximately 4 times higher mass convergence from 08:00 to 12:00 UTC than LM7. This different behavior is important to the initiation of convection in the respective model. Considering the Rhine valley, the LM7 does not exhibit any convergence or divergence, whereas the LM2.8 reveals a net outflow of air until the convective outflow regions of the mountains reach the valley in the afternoon.

Table 3 Dimensions of control volumes for the LM with a resolution of 2.8 km Control volume

Base line (km)

Area of mantle M (km2)

Area of top cap T (km2)

Hornisgrinde Northern Black Forest Rhine valley Kinzig valley Murg valley

83.6 222.3 137.0 81.8 100.8

250.8 666.9 411.0 245.4 302.4

419.2 2708.5 660.6 578.3 389.3

C. Barthlott et al. / Atmospheric Research 81 (2006) 150–175

40 20 0 -20 -40 -60 -80 -100

Murg valley

LM 2.8 km LM 7 km

0

2

4

6

8

10 12 14 16 18 20 22 24

UTC 300

Rhine valley

200 100 0 -100 0

2

4

6

8

10 12 14 16 18 20 22 24

mass budget in mtons s-1

mass budget in mtons s-1 mass budget in mtons s-1

170

80 60 Kinzig valley 40 20 0 -20 -40 -60 0 2 4 6 8

40

Hornisgrinde

20 0

400 Northern Black Forest 200 0

-200

-20 -40

10 12 14 16 18 20 22 24

UTC mass budget in mtons s-1

mass budget in mtons s-1

UTC

0

2

4

6

8

10 12 14 16 18 20 22 24

UTC

-400

0

2

4

6

8

10 12 14 16 18 20 22 24

UTC

Fig. 18. Time series of the mass budgets for dry air on June 19, 2002 for both model simulations (height of the box: 3000 m).

4.3. Vertical profiles of dry air and of humidity convergence In order to study the influence of thermally induced wind systems on mass convergence, vertical profiles of dry air and humidity convergence were determined from the model data of both LM versions. In addition to these profiles, Fig. 19 displays profiles of the mean vertical wind and its standard deviation calculated from all data points of each model level of LM2.8 within the control volume of Hornisgrinde. In order to manifest the initiation of convection, the profiles from 12:00 to 13:30 UTC only are displayed, since the description of the most important stages of the thunderstorm shows the climax of convective activity at 13:30 UTC (Fig. 14). Due to the fact that only the sum of qv ? is used to derive the individual values on each height level, the resulting profiles are not influenced by the order of magnitude of the area of the control volume itself. It can be seen in LM2.8 that prior to the initiation of the thunderstorm, the control volume Hornisgrinde reveals a near-surface convergence of dry air and humidity reaching up to a height of approximately 1000 m (Fig. 19b). As the storm evolution proceeds, this region does not exhibit any convergence at 13:00 UTC and changes into a region with divergence up to a height of 1500 m afterwards. The most important fact is the behavior of the near-surface region with time: Convergence prior to the thunderstorm development, no convergence during and directly after the initiation, and divergence afterwards. Another finding of Fig. 19 is that strong convergence at greater heights is present during intense convective activity: The profile from 13:00 UTC reveals a maximum at 1800 m above the ground. At the next time step, this

1200 UTC

1230 UTC

1300 UTC

LM 2.8 km

1330 UTC

standard deviation of w

mean value of w

humidity convergence

dry air convergence 6000 5500

C. Barthlott et al. / Atmospheric Research 81 (2006) 150–175

height above ground in m

5000 4500 4000 3500 3000 2500 2000 1500 1000 500

-20

0

20

kg m-2 s -1

40 - 400

- 200

0

g m-2 s -1

200

400

0

0.2

0.4

m s -1

0.6

-0.5

0

0.5

1.0

1.5

2.0

2.5

m s -1

Fig. 19. Vertical profiles of dry air and humidity convergence and profiles of mean vertical wind and its standard deviation for the control volume of Hornisgrinde (LM2.8).

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maximum broadens between 2 and 4 km height with a slightly reduced intensity, indicating strong entrainment processes during the lifetime of the thunderstorm. Concerning the characteristic shape, the profiles of humidity convergence show almost the same features as the profiles of dry air convergence. Only at greater heights, do the profiles approach zero due to the decreasing water content of the air. The mean vertical velocity curves exhibit relatively small mean values: No value higher than 0.2 m s 1 can be seen until 13:00 UTC. At the time of the strongest convective activity at 13:30 UTC only, a broad maximum of mean vertical wind is calculated between 2000 m and 6000 m with almost 0.7 m s 1. At this height level, the profile of standard deviation of vertical wind reaches also its maximum, showing greater variations this time. This indicates that the high upward motion in the center of the convective cell is concentrated on a rather small area (Fig. 14). During the course of convective activity, the peak value of standard deviation shifts towards greater heights with a simultaneous increase of the peak value itself. In particular, the shifting from 13:00 to 13:30 UTC is remarkable. Even with the large vertical wind speeds simulated by LM2.8 (9.73 m s 1 at 13:30 UTC, 8.18 m s 1 at 14:00 UTC, Fig. 14), the resulting mean vertical wind speed at each height level of the control volume remains small. This can be explained by two factors: (i) Subsidence in the vicinity of the thunderstorm and (ii) the small horizontal extent of the thunderstorm, which leads to a moderate mean vertical wind speed, even though the maximum values are quite high. The appropriate curves for the other control volumes reveal near-surface convergence for the northern Black Forest during the whole period (not shown). Besides the profiles of vertical wind speed averaged over the area of the control volume on each model level, time series of mean vertical wind speed of one height level provide insight into the temporal evolution of convective activity (Fig. 20). The height level chosen was 4300 m 0.8 Hornisgrinde Northern Black Forest Rhine valley Murg valley

0.6

__

w in ms-1

0.4

LM 7 km

0.2 0 -0.2 -0.4 2

4

6

8

10

12

14

16

18

20

22

0.8 LM 2.8 km

__

w in ms-1

0.6 0.4 0.2 0 -0.2 -0.4 2

4

6

8

10

12

14

16

18

20

22

UTC Fig. 20. Diurnal cycle of mean lifting for a selected level of both model versions inside the control volumes.

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(LM2.8) and 4500 m (LM7) above ground. In case of the control volume Hornisgrinde, weak lifting processes are modeled at 09:00 UTC already by the LM7, i.e. about 4 h earlier than observed. The start time of convective activity simulated by the LM2.8 is close to reality and the peak value is considerably higher than the one from the operational model. The idea of compensatory subsidence in the vicinity of regions with a strong lifting is reproduced, since the adjacent areas of Murg valley and Rhine valley show negative values of the mean vertical wind speed for the time from 12:30 to 14:00 UTC. 5. Summary and conclusions One main goal of VERTIKATOR was to determine the influence of orographically structured terrain on the triggering and development of convection. Therefore, an extensive data set for process studies was collected in the northern Black Forest and compared with results of the Lokal-Modell of the DWD. With respect to the terrain structure and height, the area under investigation can be considered a typical low-mountain range. Numerical simulations were carried out with the operational grid size of the model (7 km; with convection parameterization) and with a finer horizontal resolution (2.8 km; without convection parameterization). Simulations with LM7 using the Tiedtke parameterization for convection show a premature exceeding of convective temperature over a large area of the Black Forest. Summarizing the main deficiencies, it can be stated that (i) convective activity is too weak, starts too early, and is distributed over too large an area and (ii) as a consequence, precipitation with reduced intensity covers the whole Black Forest without any isolated convective cells. Due to the strong smoothing of the topography by the LM7, the model does not reproduce the near-surface wind field adequately. The observations show slope wind regimes on the east and west slopes of the Black Forest. These circulations are caused by locally enhanced surface temperatures that result from varying steepness and exposure of the slopes to the sun. The amount of total energy available (net radiative flux minus soil heat flux) for the turbulent fluxes therefore varies significantly depending on exposure, land use, and soil moisture. These thermally induced wind systems cause convergence lines over the main mountain ridge, which accounts for the initiation of orographically induced convective cells. In addition, valley wind regimes with a higher vertical extent exist in the region of interest. For the formation of deep convective systems, a high amount of humidity is required in the atmospheric boundary layer to foster the updrafts. Under the synoptic conditions prevailing, the air from the lower parts of the valley exhibits high humidity values of 18 g kg 1. By valley and slope wind regimes, this air is transported into the convergence area over the mountain ridges and can thus lead to a further destabilization by the release of latent heat through condensation. Another point is the reduction of the convective temperature due to a lower cumulus condensation level. Model calculations with the operational LM show that these mesoscale wind regimes are missing nearly completely. Thus, an essential process for the initiation of convection is missing. A much better result can be achieved by simulations with 2.8 km resolution, where convection is calculated explicitly and the convection parameterization can be switched off. Good agreement with measurements of the near-surface wind and the start time of convection can be found, although the location of the simulated cell is slightly south of the observed one. For the quantification of transport processes by deep convection and the investigation of the relation between thermally induced wind systems and the initiation of convection, mass budget calculations were performed for selected control volumes. It was shown that the presence of secondary circulation systems leads to horizontal mass convergence over the main mountain

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ridge prior to the formation of the thunderstorm. Vertical movements are triggered due to mass continuity, so that mesoscale circulation systems can be assumed as a factor controlling the initiation and development of convection above complex terrain under weak synoptic flows. Recent findings of Tian and Parker (2002) and Hanesiak et al. (2004) support these findings. This leads to the assumption that the real topography is of crucial importance to the triggering and modification of convection. The authors state that there is some evidence that these specific findings in the research area of the Black Forest are also valid for other low-mountain ranges with similar terrain structure (predominant valleys originating in narrow elevated terrain) and similar synoptic controls since the governing processes are expected to be the same or to run in a similar way. A higher horizontal resolution therefore is indispensable for a better representation of convective processes in the LM in order to obtain a more realistic description of the energetic exchange processes over mountainous areas. Acknowledgements This study was funded by the German Federal Ministry of Education and Research (BMBF) within the framework of the AFO2000 project VERTIKATOR under grant 07ATF45-TP3. The authors wish to thank the DWD, the LfU, and Meteocontrol for the supply of the meteorological data as well as K.-D. Beheng and J. Handwerker for providing of IMK precipitation radar data. We also are grateful to the Institut fu¨r Flugfu¨hrung of the Technische Universita¨t Braunschweig and its DO-128 crew with R. Hankers, T. Feuerle, H. Schulz, and G. Wende. References Adrian, G., 2004. Vertical mixing over mountains (VERTIKATOR). AFO2000—The German Atmospheric Research Programme—Result Abstracts, Final Symposium 22–24 March 2004, Bad To¨lz, Germany, p. 57. Arakawa, A., Schubert, M.H., 1974. Interaction of a cumulus cloud ensemble with the large-scale environment, Part I. J. Atmos. Sci. 31, 674 – 701. Banta, R.M., 1984. Daytime boundary-layer evolution over mountainous terrain: Part I. Observations of dry circulations. Mon. Weather Rev. 112, 340 – 356. Bertram, I., Seifert, A., Beheng, K.D., 2004. The evolution of liquid water/ice contents of a mid-latitude convective storm derived from radar data and results from a cloud-resolving model. Meteorol. Z. 13, 221 – 232. Doms, G., Scha¨ttler, U., 1999. The nonhydrostatic limited-area-model LM (Lokal-Modell) of DWD: Part I. Scientific documentation. Deutscher Wetterdienst, Gescha¨ftsbereich Forschung und Entwicklung. Donner, L.J., Seman, C.J., Hemler, R.S., 2001. A cumulus parameterization including mass fluxes, convective vertical velocities, and mesoscale effects: thermodynamic and hydrological aspects in a general circulation model. J. Clim. 14, 3444 – 3463. Emanuel, K., 1994. Atmospheric Convection. Oxford Univ. Press, New York. 580 pp. Founda, D., Tombrou, M., Palas, D.P., Asimakopoulos, D.N., 1997. Some measurements of turbulence characteristics over complex terrain. Boundary-Layer Meteorol. 83, 221 – 245. Hanesiak, J.M., Raddatz, R.L., Lobban, S., 2004. Local initiation of deep convection on the Canadian prairie provinces. Boundary-Layer Meteorol. 110, 455 – 470. Homar, V., Gaya, M., Romero, R., Ramis, C., Alonso, S., 2003. Tornadoes over complex terrain: an analysis of the 28th August 1999 tornadic event in eastern Spain. Atmos. Res. 67, 301 – 317. Kalthoff, N., Vogel, B., 1992. Counter-current and channeling effect under stable stratification in the area of Karlsruhe. Theor. Appl. Climatol. 45, 113 – 126. Kalthoff, N., Horlacher, V., Corsmeier, U., Volz-Thomas, A., Kolahgar, B., Geih, H., Mo¨llermann-Coers, M., Knaps, A., 2000. Infuence of valley winds on transport and dispersion of airborne pollutants in the Freiburg-Schauinsland area. J. Geophys. Res. 105, 1585 – 1597. Kalthoff, N., Corsmeier, U., Schmidt, K., Kottmeier, C., Fiedler, F., Habram, M., Slemr, F., 2002. Emissions of the city of Augsburg determined using the mass balance method. Atmos. Environ. 36 (Supplement No. 1), 19 – 31.

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