High-resolution simulations of convective cold pools

over the northwestern Sahara

P. Knippertz,1 J. Trentmann,1,2 and A. Seifert2

Received 9 October 2008; revised 22 January 2009; accepted 11 February 2009; published 24 April 2009.

[1] Cooling by evaporation of convective precipitation in the deep and dry subcloud layerover desert regions can generate intense downdrafts and long-lived and extensiveatmospheric density currents. The strong gusts at their leading edges can cause so-calledhaboob dust storms. Despite their importance for the dust cycle, the ability of state-of-the-art numerical weather prediction models to realistically simulate the associated convectivecold pools has been investigated very little to date. During the first field campaign ofthe Saharan Mineral Dust Experiment in southern Morocco in May/June 2006, severaldensity currents were observed. They were triggered by deep moist convection over theAtlas Mountains during the afternoon and propagated into the foothills in the course of theevening. Here we present numerical simulations of three of these density currentsusing the nonhydrostatic Consortium for Small-Scale Modelling model with 2.8-kmhorizontal grid spacing, which allows an explicit treatment of deep convection. The modelis capable of simulating the timely initiation of convective cells over the Atlas Mountainsand the subsequent formation of long-lived, extensive cold pools with a realistic three-dimensional structure. Deviations from available surface and satellite observations areclosely related to model deficiencies in simulating precipitating convection over theAlgerian Sahara. Sensitivity studies with modified microphysics reveal a large influence ofraindrop size distributions on evaporation and surface rainfall but a rather moderateinfluence on the cold pool evolution. Decreasing the length scale for turbulent verticalmixing in the boundary layer leads to more widespread but weaker precipitation, moreevaporation, and a faster and more extended cold pool.

Citation: Knippertz, P., J. Trentmann, and A. Seifert (2009), High-resolution simulations of convective cold pools over the

northwestern Sahara, J. Geophys. Res., 114, D08110, doi:10.1029/2008JD011271.

1. Introduction

[2] The evaporation and melting of convectively gener-ated hydrometeors sedimenting into the subsaturated sub-cloud layer can lead to deep, intense downdrafts and theformation of large cold pools at the surface that spreadlaterally as density currents [Simpson, 1997]. This processcan be enhanced through the evaporation of cloud water dueto lateral entrainment of subsaturated ambient air [Knuppand Cotton, 1985]. Conditions favorable to produce strongdowndrafts penetrating to the surface include a very deep,dry-adiabatic mixed layer, high rain water mixing ratios atcloud base, and small raindrop sizes [Kamburova andLudlam, 1966; Srivastava, 1985; Proctor, 1988, 1989;Takemi and Satomura, 2000]. The former condition is ofmajor importance, since more stable environmental lapserates allow descending parcels to attain neutral buoyancythrough compressional heating. Downdraft intensity can beenhanced through precipitation load, in particular in wet

microbursts often related to short-lived, small-scale, heavilyprecipitating thunderstorms in weakly sheared environments[Fujita, 1985; Goodman et al., 1988], and then dependsupon the width, type, and duration of precipitation [Proctor,1989]. Typical vertical velocities in convective downdraftsare 5–10 m s�1, but up to 20 m s�1 have been observed[Knupp and Cotton, 1985].[3] The high momentum and density of the cold-pool air

causes a strong, divergent wind field pattern at the ground,usually with a marked convergent gust front. In the case ofdry-adiabatic deep subcloud layers even shallow high-basedclouds producing small quantities of precipitation cangenerate strong downdrafts and surface gusts [Knupp andCotton, 1985]. The evolution and depth of the cold pool issensitive to the ambient vertical wind shear at lower levels[Weisman and Rotunno, 2004] and other factors such asconvective momentum transports. For a given densitydifference, a balance between the cold pool and the shearcirculations results in deep lifting at the leading edge of thedensity current [Rotunno et al., 1988; Xu et al., 1996; Xue,2000], which can trigger the formation of shallow arcclouds, or, if the atmosphere is conditionally unstable, ofnew convective cells [Knupp and Cotton, 1985]. This waycold pools can become a major ingredient in the organiza-

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1Institute for Atmospheric Physics, Johannes Gutenberg UniversityMainz, Mainz, Germany.

2German Weather Service, Offenbach, Germany.

Copyright 2009 by the American Geophysical Union.0148-0227/09/2008JD011271$09.00

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tion of moist convection from the meso-g to the meso-bscale and even meso-a scale [Takemi, 1999; Weisman andRotunno, 2004].[4] Over deserts, the deep, well-mixed, hot, and dry

boundary layer allows for a substantial amount of latentcooling and the formation of extraordinarily long-lived andextensive cold pools [Takemi, 1999; Knippertz et al., 2007;Miller et al., 2008]. The associated density currents prop-agate at speeds on the order of 10 m s�1 over hundreds ofkilometers, sometimes far away from the parent convectiveclouds [Freeman, 1952; Miller et al., 2008]. The passage ofthe leading edge is usually associated with abrupt changesin temperature, dew point, pressure, visibility, and withstrong, gusty winds and high turbulence. The latter can liftfine material from the ground leading to fast-moving,dramatic, billowing walls of dust that for this reason attractmore attention than similar systems that move over vege-tated surfaces. These so-called ‘‘haboobs’’ have been docu-mented for the Sahel [Sutton, 1925; Farquharson, 1937;Freeman, 1952; Lawson, 1971; Williams et al., 2008], thenorthwestern Sahara [Knippertz et al., 2007], semiarid partsof the USA [Idso et al., 1972; Chen and Fryrear, 2002],northwest China [Takemi, 1999], and the Arabian Peninsula[Membery, 1985; Miller et al., 2008]. Typical temperaturedrops are on the order of 7 K, while relative humiditytypically increases by 15% [Idso et al., 1972; Miller et al.,2008]. A typical depth of the dust layer found in theliterature is 1–2 km [Sutton, 1925; Freeman, 1952], butvertically pointing Doppler radar measurements by Williamset al. [2008] indicate substantial dust loadings well above2 km also. Most of the suspended dust is contained withinthe cold pool [Lawson, 1971] where near-surface wind gustsof 15 m s�1 and even 25 m s�1 are commonly observed[Sutton, 1925; Freeman, 1952; Miller et al., 2008]. Ideal-ized modeling work by Takemi [2005] has shown thatturbulent mixing can entrain dust into the system-relativefront-to-rear flow at midlevels and even into the cloudupdraft [Takemi, 2005, Figure 7]. How much dust remainsin the atmosphere after the decay of a convective cold poolis still a matter of debate [Williams, 2008]. Naturally, coldpools that quickly separate from the parent convection canmore efficiently mobilize dust as they avoid scavenging byprecipitation. In this sense the optimal condition for thegeneration of a haboob is enough moisture to produce rain,but not quite enough to allow it to reach the ground before itis evaporated. This leads to a distinct annual cycle ofhaboob occurrence that is closely, but not exactly, tied tothe climatology of convective activity [Farquharson, 1937;Freeman, 1952; Miller et al., 2008]. The amounts of dustbeing mobilized by a system with a given intensity alsodepend on annual changes in vegetation and soil moisture.[5] Despite their importance for the dust cycle and their

potential threat to aviation safety, most research on haboobsto date has been restricted to observational analyses ofsingle cases or multiyear statistics at single observationsites, while numerical modeling studies are rare. Miller etal. [2008] developed a simplified haboob deflation modelfor dust budget analysis, while Takemi [2005] used a three-dimensional cloud resolving model with a dust scheme in anidealized setup. Reinfried et al. [2009] employed theCOSMO (Consortium for Small-Scale Modelling) model,the operational nonhydrostatic, limited area model of the

German Weather Service (Deutscher Wetterdienst, DWD;see section 2.1), with an additional dust module to simulatea density current that occurred during the first field cam-paign of the Saharan Mineral Dust Experiment (SAMUM)(see Heintzenberg [2009] for an overview) in southernMorocco in May/June 2006. Such systems are usuallytriggered by deep moist convection over the High AtlasMountains during the afternoon and then spread into thesouthern foothills in the course of the evening and night[Knippertz et al., 2007]. During SAMUM their occurrencewas closely tied to the presence of upper tropospherictroughs, which help to destabilize the atmosphere and toadvect hydrometeors toward the down-shear Saharan side ofthe mountain range. The observed systems had lifetimes ofup to 10 h and leading edge extensions of several hundredkilometers as identified from infrared (IR) satellite imagery.Reinfried et al. [2009] were able to satisfactorily reproducethe most important characteristics of the cold pool whenusing a horizontal grid spacing of 2.8 km that allows anexplicit treatment of deep convection. The existence of arealistic cold pool in the simulations is closely linked to thequality of the precipitation forecast, which in turn is highlysensitive to the treatment of convection.[6] Here, the work of Reinfried et al. [2009] is extended

by conducting convection permitting COSMO simulations(without dust module) of three SAMUM density currents.The goal of this paper is (1) to further corroborate the abilityof the model to generate realistic cold pools with respect totheir temporal evolution and three-dimensional structure,and (2) to test the sensitivity of the cold pool evolution tochanges in the microphysics and turbulence schemes of themodel. Previous idealized studies suggest an influence ofraindrop size on downdraft intensity and low-level cooling[e.g., Srivastava, 1985], but to our knowledge this has neverbeen tested with a state-of-the-art numerical model in arealistic setup. Changes to the treatment of boundary layerturbulence can be expected to affect both the precipitationgeneration and the density current propagation, for example,through the modified entrainment of ambient air [Proctor,1988]. In our view, convective cold pools in the northwest-ern Sahara are an ideal test bed for such a study as thecausative precipitation is strongly related to forcings byupper level troughs and the Atlas Mountains, and is there-fore more realistically represented in numerical models thanconvection resulting from air mass instabilities alone. Inaddition the deep dry boundary layer in the Sahara favorsthe formation of large, long-lived cold pools, whose mac-roscopic characteristics are well determined. The remainderof the paper is structured as follows: In section 2 theCOSMO model and the observational data used for modelevaluation are briefly introduced. Section 3 discusses refer-ence simulations for three cases including a comparison toobservations. The results from four sensitivity experimentsare presented in section 4. Section 5 contains a summaryand conclusions.

2. Model and Data

2.1. COSMO Model

[7] All simulations for this study were conducted withthe COSMO model version 4.0 [Steppeler et al., 2003;Schattler et al., 2008], the nonhydrostatic limited-area

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weather prediction model used for operational weatherforecasts in Germany, Switzerland, Italy, Greece, Romania,Poland and Russia. The results shown in this paper are fromsimulations with 2.8-km horizontal grid spacing nested intoCOSMO runs with 7-km horizontal grid spacing, which arein turn nested into operational analyses of the global modelGME (this model has been named GME as it replaced theoperational global model (GM) and the regional model(EM) for central Europe in 1999) of the DWD [Majewskiet al., 2002]. The configuration of the COSMO model usedhere applies an efficient split-explicit Runge-Kutta solver[Wicker and Skamarock, 2002], a Lin-type one-momentcloud microphysics scheme that predicts cloud water, rainwater, cloud ice, snow, and graupel [Lin et al., 1983;Reinhardt and Seifert, 2006], and a boundary layer schemeusing a prognostic turbulent kinetic energy (TKE) equationbased on a Mellor-Yamada level 2.5 turbulence closure[Mellor and Yamada, 1974; Raschendorfer, 2001]. In thehigh-resolution simulation with 2.8-km grid spacing deepmoist convection is explicitly treated, but shallow convec-tion with cloud depth below 300 hPa is parameterized usinga simple mass-flux formulation based on the Tiedtke [1983]scheme [Doms and Forstner, 2004]. The simulations areinitialized at 0000 UTC on 3 days during the SAMUM fieldcampaign, on which density currents were observed (i.e., 07and 31 May, and 03 June 2006), and then run for 30 h. If notnoted otherwise, the employed time step was 30 s. Themodel has 50 unevenly spaced vertical levels up to a heightof 22 km and the horizontal domain comprises 471 � 401grid points spanning the area from the Moroccan Atlanticcoast to the Algerian Sahara Atlas in the west–east andfrom the southern tip of Spain to northern Mauritania in thenorth–south direction (Figure 1). Close to the center of thedomain are the High Atlas Mountains with a maximumelevation of 3509 m in the model (4167 m in reality). Toevaluate the COSMO-simulated clouds with satellite obser-

vations, model-derived pseudo satellite images are used[Keil et al., 2006].

2.2. Observational Data

[8] For the evaluation of model-simulated near-surfacetemperature, dew point, pressure, wind, and precipitation,observations from automatic weather stations (CampbellScientific Inc., Logan, Utah) operated by the Germanresearch initiative IMPETUS (An Integrated Approach tothe Efficient Management of Scarce Water Resources inWest Africa) [Speth and Diekkruger, 2006] are used. Wewill concentrate on the station El Miyit (EMY hereafter)that is located in the Draa Valley in southern Morocco(30.36�N, 5.63�W, 792 m; see Figure 1). More details onstation locations and instrumentation are given by Knippertzet al. [2007]. In addition, standard 3-h SYNOP reportsfrom the Moroccan stations Ouarzazate (WMO 60265,30.93�N, 6.90�W, 1140 m) and Errachidia (60210,31.93�N, 4.40�W, 1042 m), and from the Algerian stationsBechar (60571, 31.62�N, 2.23�W, 816 m), Beni Abbes(60602, 30.13�N, 2.17�W, 505 m), Timimoun (60607,29.25�N, 0.28�E, 317 m), and Adrar (60620, 27.88�N,0.28�E, 283 m), also indicated in Figure 1, are used.Model-simulated brightness temperatures in the IR channel(10.8 mm) are compared to corresponding values fromMeteosat Second Generation. For the evaluation of model-simulated precipitation, the ‘‘Tropical Rainfall MeasuringMission (TRMM) and Other Rainfall Estimate’’ (3B42 V6)in 0.25� � 0.25� horizontal resolution as described byHuffman et al. [2007] are used. These data are 3-h accumu-lated, combined TRMM precipitation radar-microwave-IRestimates (with gauge adjustment) and were downloadedfrom http://disc2.nascom.nasa.gov/Giovanni/tovas/ operatedby the National Aeronautics and Space Administration(NASA).

3. Reference Simulations

3.1. Cold Pool on 03–04 June 2006

[9] The first case to be studied is the rather extendeddensity current that formed on 03 June 2006 [Knippertz etal., 2007, section 5.2], which was also investigated byReinfried et al. [2009].3.1.1. Temporal Evolution[10] Figure 2 shows the evolution of the cold pool in the

course of the afternoon and evening of 03 June as simulatedby the COSMO model with 2.8 km horizontal grid spacing.At 1600 UTC the model simulates four precipitating con-vective cells slightly to the southeast (i.e., downwind withrespect to the westerly to northwesterly flow at upper levels)of the highest peaks of the Atlas Mountains (Figure 2a).Surface precipitation accumulated over the previous hourclearly exceeds 5 mm at some grid points. The 10-m windvectors in Figure 2a indicate a convergence between themainly southwesterly flow from the Sahara and the north-westerlies from the Atlantic Ocean over the mountain crestthat has helped to trigger the convection. Each of the fourstorms is associated with wind speeds exceeding 10 m s�1,usually at the Saharan, i.e., the down-shear side of thestrongest precipitation (red lines in Figure 2a). This value isabove the mobilization threshold for typical dust sources inthe region [e.g., Chomette et al., 1999]. The grid point

Figure 1. Model domain with orography. The 1000-misohypse is highlighted. Meteorological stations andgeographical terms used in the text are indicated.

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maximum at this time is 16.5 m s�1. These strong winds areassociated with convective downdrafts of up to 4.9 m s�1 inagreement with typical values in the literature [Knupp andCotton, 1985]. The black dots in Figure 2a mark grid pointswith ascent of more than 3 m s�1 below 700 hPa. Over thecloudless Sahara, where daytime heating is very intense,they indicate the turbulent upward motions related to deepdry boundary layer convection. At the southern and easternedges of the convective cells over the Atlas they indicatethat the cold pools begin to spread laterally and lift ambientair. This criterion will be used in the following to identifythe dynamically active part of the leading edges of densitycurrents (see discussion at the end of this subsection). At1600 UTC the leading edges are mostly oriented perpen-dicular to the low-level flow as has been observed for many

other cases [e.g., Weisman and Rotunno, 2004]. Maximumuplift is 10.2 m s�1 at this time, which is in the upper part ofthe range of vertical motions observed by Williams et al.[2008] for the Sahel.[11] Three hours later, at 1900 UTC, the precipitation

zone has shifted further downwind and also down-slope tothe Moroccan-Algerian border (Figure 2b). The decreasingrainfall amounts at the ground might be an indication ofincreasing evaporation. The precipitation regions and coldpools of the three western cells have merged, while thelarger and more intense cell to the northeast is still some-what isolated. The strong low-level ascent at the leadingedges to the south, which still reaches 7.6 m s�1 at this time,has supported the regeneration of convection. Winds behindthe leading edges largely exceed 10 m s�1 with a maximum

Figure 2. Temporal evolution of the cold pool on 03–04 June 2006 as simulated by the COSMOmodel.Shown is the precipitation accumulated over the previous hour in millimeters (shading according to scale), 10-mwind vectors together with 10 m s�1 isotachs (red lines) at (a) 1600 UTC 03 June, (b) 1900 UTC 03 June,(c) 2200 UTC 03 June, and (d) 0000 UTC 04 June. The black dots mark grid points where the verticalvelocity exceeds 3 m s�1 below 700 hPa. The thin black lines are political borders, and the thickblack lines mark the 1000-m isohypse of the COSMO model orography. The green line with the twocrosses in Figure 2c indicates the location of the cross section and the vertical profiles, respectively,shown in Figure 3.

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of 17.5 m s�1 to the south of the eastern cell. Owing to thedecreased insolation shortly before sunset the boundarylayer convection over the Sahara has ceased as indicatedby the absence of black dots in Figure 2b. Another 3 h later,by 2200 UTC, the precipitation zone has continued shiftingsoutheastward into Algeria (Figure 2c). The initial cellshave merged, but there are still indications of convectiveregeneration farther to the west with uplift reaching maxi-mum values of 6.6 m s�1 at this time. Precipitation intensityhas generally decreased. The leading edge of the cold poolis now clearly separated from the parent precipitation andforms an almost straight line that stretches some 500 kmfrom 29�N, 6�Wto 30�N, 1�W.Right behind the leading edgeis an extended area with 10-m winds exceeding 10 m s�1,indicating potential for substantial dust mobilization. Amaximum of 18.2 m s�1 is reached at this time. The windvectors indicate a continuation of the cold-pool spreadingto the northeast and west of the leading edge identifiedwith the vertical motion criterion. Finally at 0000 UTC 04June the precipitation has almost ceased and the cold poolbegins to weaken (Figure 2d). While the active leadingedge with strong horizontal and vertical velocities stretchesfor about 600 km across the Sahara, the area withindications of a density current in the wind field reachesmuch farther west illustrating the impressive distances acold pool can spread in this desert environment. Maximum10-m wind speeds have decreased to 15.6 m s�1, themaximum ascent at the leading edge is now 5.4 m s�1.[12] Between 1900 UTC 03 June and 0000 UTC 04 June

(Figures 2b and 2d) the leading edge has traveled between220 km in its western part and 270 km farther to the east,where the source of evaporative cooling is. This correspondsto propagation velocities between 12 and 15 m s�1. The factthat the maximum low-level winds behind the leading edgeare considerable faster than the leading edge movementitself is consistent with observations of haboobs in otherplaces [Lawson, 1971] and more theoretical arguments ondensity current behavior (see discussion by Smith and Reeder[1988]). It is remarkable that the maximum horizontal windspeed behind the leading edge, its propagation velocity, andthe maximum uplift above the leading edge vary rather littleover this period of 5 h indicating a stable dynamical evolu-tion. After midnight, however, the leading edge becomesmore and more diffuse, and is not detectable at 0300 UTCanymore, giving a total lifetime of the cold pool of 10 h.Potential reasons for the decay are the following: Radiativecooling of the surrounding desert surface at night stabilizesthe atmosphere (see Figure 3b) and impedes the lifting ofenvironmental air above the approaching density current andtherefore its propagation [Proctor, 1989]. It also reduces thedensity difference between the cold pool and the ambient air.The reduced lifting and the penetration of the density currentinto a very dry environment suppress the regeneration ofmoist convection at the leading edge, which cuts off thesource of evaporative cooling as indicated by the decayingprecipitation. Without this source, vertical mixing quicklyreduces the density contrast between the cold pool and theless dense environmental air.3.1.2. Structure[13] Figure 3 shows the three-dimensional structure of the

density current at 2200 UTC 03 June 2006, when the systemis fully developed (see Figure 2c). The cold pool can be

identified by the elevated 306-K isentropic surface to thesoutheast of the Atlas, which is almost vertical at thesouthern and eastern flanks reflecting the very strongbaroclinicity along the leading edge (Figure 3a). Fartherto the northeast (i.e., to the right-hand side of the plot) the306-K isentrope gently slopes upward marking a weakbaroclinic zone. The isentropic surfaces are color-codedwith specific humidity showing that particularly the westernpart of the density current is substantially moister than thecool air to the northeast. The isolated nature of the cold pooland its high moisture content are clear signs of a localgeneration as opposed to an advection into this region.While the isentropic surface indicating the baroclinic zonein the northeast is relatively smooth, the cold pool has amore jagged structure, which points to effects of updraftsand downdrafts in the precipitation/evaporation zone, andhigh turbulence in the regions of strong horizontal winds(see also Figure 3b). Turbulent mixing in the wake of thehead of a density current is a typical element found inlaboratory and idealized numerical experiments [Simpson,1997]. Details of this process, however, are not wellresolved with 2.8-km horizontal grid spacing. The verticalcross section in Figure 3b shows that a vertical orientation isfound for all isentropes between 306 and 312 K, with thelatter spanning a vertical range from near the surface toabout 680 hPa at the leading edge. The height minimum ofthe isentropes some 15 km from the leading edge might bean indication of descent behind the head of the densitycurrent [cf. Williams et al., 2008, Figure 3]. Vertical profilesof temperature and dew point on either side of the leadingedge show that the regions with substantial cooling reachfrom the surface to about 800 hPa, while the moisteningaffects levels up to 750 hPa (Figure 3c). A height of 1.5 kmof the main bulk of the cold pool agrees well withobservations found in the literature [Sutton, 1925; Freeman,1952], while weaker effects of the density current can beseen as high as 4 km as also observed by Williams et al.[2008].[14] The air mass that the density current is intruding into

(left side of Figure 3b and black profile in Figure 3c) ischaracterized by a deep well-mixed layer up to 600 hPa withan almost constant potential temperature q of about 312–313 K, which is on the order of maximum daytime 2-mtemperatures in this region. Winds in this layer are fromeasterly directions and the water vapor mixing ratio isalmost constant. As the sun set almost 3 h before the timeshown in Figure 3, radiative cooling has created a shallowstable layer at the surface with potential temperature being2–3 K cooler than the mixed-layer air aloft, which is likelyto contribute to the decay of the cold pool in the course ofthe night (see discussion at the end of section 3.1.1). Thetop of the boundary layer near 600 hPa is marked by anincrease in q by 4 K within 50 hPa with no apparentdifferences between the region of the cold pool and theundisturbed surroundings (Figure 3b). The neutral stratifi-cation of the ambient air up to a level well above the heightof the cold pool is an important prerequisite for its longevity[Takemi and Satomura, 2000]. The free troposphere differsonly little between the regions ahead of and behind theleading edge with winds coming predominantly from west-erly directions (Figure 3c). The high relative humidity andalmost moist-adiabatic lapse rate above the 600-hPa level

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indicate the cloud layer. This profile clearly fulfills the ‘‘dry/wet’’ condition (i.e., dew point temperature depressiongreater (smaller) than 8�C at 700 hPa (500 hPa)) favorablefor supporting dry microbursts as defined by Miller et al.[2008]. The wind profile shows that the cold pool is mainlyspreading in the direction of the moderate westerly shearbetween the subcloud and the cloud layer. As discussedabove there is a distinct maximum of near-surface horizon-tal winds behind the leading edge (Figure 2). Figures 3b and3c show that the abrupt jump in wind speed and the change

in direction from easterly to northerly winds prevails up toabout 750 hPa with a maximum of more than 20 m s�1 ataround 900 hPa. This leads to massive convergence in thislayer and to strong ascent of more than 3 m s�1 along anextended stretch of the leading edge and throughout most ofthe mixed layer (Figures 3a and 3b). Maximum values ofvertical wind reach 7 m s�1 at around 740 hPa in closeagreement with modeling results by Takemi [2005] andobservations by Williams et al. [2008]. This marked low-level ascent region is so unique that the authors decided to

Figure 3. Structure of the cold pool at 2200 UTC 03 June 2006 as simulated by the COSMO model.(a) Three-dimensional view of the area 27�N–35�N, 7.5�W–2.5�E looking from SE toward NW.Displayed are the model orography, the 306-K isentropic surface colored with specific humidityaccording to the scale, and regions with vertical velocity exceeding 4 m s�1 are indicated by whiteisosurfaces. (b) Vertical cross section along the green line in Figure 2c showing potential temperature(shading according to scale), horizontal wind speed (black contours every 5 m s�1), and the 3 m s�1

vertical velocity isotach in white. (c) Vertical profiles of temperature, dew point, and wind depicted in theform of a skew-T log-p diagram. The black (red) lines and arrows show a location ahead of (behind) theleading edge as indicated by the crosses in Figure 2c.

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use it as a criterion to identify the active part of the leadingedge in all horizontal distributions shown in this paper (e.g.,Figure 2).[15] The vertical profiles shown in Figure 3 allow an

estimate of the density difference across the leading edge,which is closely related to the propagation velocity c (andalso varies across the cold pool). In a simple dry, frictionlesstwo-fluid system c is proportional to

ffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffigHDq=q0

p, where g

is acceleration of gravity, H is the height of the current, andq is potential temperature [e.g., Smith and Reeder, 1988]. Ina moist system with vertically varying density contrasts it ismore appropriate to use the integrated buoyancy to calculatec in the following form [Weisman and Rotunno, 2004]:

c2 ¼ 2

Z H

0

�gDqq0

þ 0:61 q� q0ð Þ� �

dz; ð1Þ

where q is the mixing ratio of water vapor. Subscript zerostands for undisturbed background values. The effects ofprecipitation load can be assumed to be small in this dryenvironment and has been neglected. If the two profiles inFigure 3c are used as representatives of the disturbed andundisturbed environments, a theoretical propagation speedof 23.4 m s�1 is obtained, which is largely dominated by thetemperature effect. This value is substantially larger than the15 m s�1 determined from the propagation of the leadingedge analyzed in Figure 2 (see section 3.1.1). Possiblereasons for this discrepancy are friction and three-dimen-sional effects, both neglected in equation (1).3.1.3. Comparison to Observations[16] The foregoing discussion reveals that the COSMO

model is able to simulate an extended, long-lived cool poolwith a realistic three-dimensional structure. In the followingthese results will be compared with available satellite andground-based observations. In IR satellite imagery, theleading edge is subjectively identified by tracing the bound-ary between the cold pool and the ambient air in cloud freeregions or, even better, by tracing arc clouds forming owingto ascent at the leading edge similar to the method used byKnippertz et al. [2007] (e.g., red line in Figure 4d).[17] During the initial stage at 1600 UTC 03 June 2006,

deep convection in the model is restricted to the region closeto the High Atlas and there are only a few cloud streaks overthe adjacent Sahara (Figure 4a). The corresponding Meteo-sat image equally shows deep convective cells over themountains, but also over the Algerian Sahara (Figure 4b),which is consistent with ground observations. Surfacestations Bechar and Errachidia (for station locations, seeorange-filled circles in Figure 4) report thunderstorms and,respectively, 2 and 3 mm of rain between 1000 and1800 UTC, and even the arid station Beni Abbes reportstraces of precipitation before 1800 UTC. This difference inprecipitation and therefore in the production of evapora-tively cooled air has a pronounced impact on the densitycurrent extent, orientation, and propagation as we will seebelow. The reason why the model struggles to produce theconvective cells over Algeria is unclear. Among the possiblecauses are too dry initial conditions in this extremelyobservation-sparse region, but midday dew point observa-tions from synoptic stations and from the model bothindicate a lifting condensation level around 600 hPa. It

appears possible that the model struggles to realisticallysimulate the evolution of the highly complex and very deepsummertime desert boundary layer with its high aerosolcontent. It is also conceivable that ascent in the weakbaroclinic zone described above (see Figure 3a) is under-represented in the simulation with the COSMO model.[18] In the mature stage of the density current at 2200

UTC clouds and precipitation have shifted southeastward(Figures 4c and 4d). The COSMO model simulates anextensive region of mostly convectively generated cirrusclouds stretching in a west–east direction across the Alger-ian Sahara. Single convective cells are hard to make out inthis pseudo satellite image, but precipitation rates at theground of more than 1 mm h�1 (blue lines in Figure 4c)indicate active convection, predominantly in the central partof this cloud band. The leading edge of the cold pool islocated below the southern and western flanks of the clouds.In contrast, the corresponding Meteosat image shows a muchsmaller cloud band with several single cold-cloud-top cells,notably a large system around 30.5�N, 1�W (Figure 4d).The tendency of the COSMO model to produce too wide-spread cirrus in this situation has already been noted byReinfried et al. [2009]. The observed leading edge is moreextended than in the model and spans from the Moroccan-Algerian border to near the Saharan station Adrar. The coldpool is clearly separated from the cloud band and ispropagating into southerly directions in contrast to thesoutheastward propagation in the model. This analysis issupported by station observations. Timimoun is reached bythe cold pool between 1800 and 2100 UTC indicated by adew-point increase of 8 K, a temperature drop of 7 K, apressure increase of 3.1 hPa, and a drop in visibility from10 km to 1.5 km, accompanied by observations of dustmobilization near the station. At 2100 UTC sustained 10-mwinds reach 15 m s�1 and light rain is observed. Adrarreports a similar evolution but without precipitation between2100 UTC 03 June and 0000 UTC 04 June 2006.[19] The differences between the model simulation and

observations become very apparent in 24-h accumulatedprecipitation amounts for 03 June. Close to the Atlas, modeland TRMM precipitation estimates agree well and there issome confidence for the latter through observations fromErrachidia, Bechar, Beni Abbes, and Timimoun (Figure 5).Precipitation records at the five high-mountain stations ofthe IMPETUS climate network (located near 31.5�N, 6.4�Wand between 1900 and 3850 m [see Knippertz et al., 2007])of 2.1–6.2 mm during the period from 1200–2300 UTC arealso in good agreement with the satellite data. Over theeastern part of the domain, however, discrepancies becomequite large. TRMM precipitation estimates exceed 10 mm inthis region, but unfortunately there is no ground truth toconfirm such amounts (Figure 5b). This rather intenseprecipitation over the desert consists an additional sourceof evaporationally cooled air that is not present in the modeland explains the larger, southward shifted, southward prop-agating cold pool in the observations (compare black dots inFigure 5a with black line in Figure 5b for 0000 UTC 04June). The observed propagation speed, subjectively esti-mated from IR imagery between 1900 UTC 03 June and0000 UTC 04 June, is about 11 m s�1 and thus considerablyslower than in the model (see section 3.1.1). This points to asmaller density difference between cold pools and environ-

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ment, and thus overall weaker evaporative cooling. Thelarger extent is therefore mainly the consequence of a largercold-air source. This comparison demonstrates that a correctinitiation of convection is an important prerequisite for thecorrect simulation of the associated cold pools.[20] Unfortunately available observations are not suffi-

cient to document the three-dimensional structure of thedensity current. However, observations from the surfacestation EMY in southern Morocco with 15-min temporalresolution can be used to compare observed and modeledcharacteristics of the leading edge. This station was passedby the cold pool in the late afternoon of 03 June. Withrespect to 2-m temperature both data sets show an abruptdrop on the order of 4–6 K in half an hour with the model

leading by about 45 min (Figure 6a). The model satisfac-torily reproduces the temperature evolution before this timewith a slight overestimation of the diurnal cycle. The dewpoint time series also agree well showing a jump of about6 K concomitant with the temperature drop (Figure 6b). Themodel output indicates a first (unverified) moisture jumparound 0200 UTC, leading to a moist bias of about 2 K overmost of the day. A jump and associated bias are also foundfor wind speed (Figure 6c). Closer inspection of horizontalwinds, clouds, and precipitation suggest that some verysmall convective cells are triggered during the model spin-up that cause these spurious signals (not shown). In themodel, the wind signal associated with the passage of theafternoon cold pool consists of a relatively short high-wind

Figure 4. Comparison of brightness temperatures in K (left) simulated by the COSMO model with(right) corresponding Meteosat data for (a and b) 1600 UTC and (c and d) 2200 UTC on 03 June 2006.Red lines in Figures 4a and 4c mark the 10 m s�1 isotach of the 10-m wind as in Figure 2, and blue linesindicate instantaneous precipitation rates of 1 mm h�1. The red line in Figure 4d is the subjectivelyanalyzed leading edge of the cold pool as in work by Knippertz et al. [2007]. Political borders, the 1000-mmodel orography isohypses, and synoptic stations (orange solid circles) are also indicated.

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phase with a peak velocity of 15 m s�1, while observationsshow more sustained winds of more than 10 m s�1 behindthe leading edge. This difference can be expected to reducedust emissions in the model owing to the relatively shorttime, during which the mobilization threshold is exceeded.Both data sets show an abrupt change from southwesterly toeasterly winds and a pressure increase by almost 2 hPa atthe leading edge (Figures 6c and 6d). The more abruptpressure rise in the model is consistent with the differencesin wind speed evolution. The strong decrease in pressureduring the afternoon is most likely caused to a substantialdegree by a combination of atmospheric tides and the strongdaytime heating of the boundary layer over the desert.[21] The previous discussion reveals that the COSMO

model is able to simulate realistically most of the generalcharacteristics of the density current such as size, propaga-tion speed, initiation time, and air mass characteristics.Problems with the triggering of convection over the desert,however, lead to differences in the precipitation distributionthat affect the location and propagation direction of thesystem.

3.2. Other Simulated Cases

[22] In order to test the representativeness of these results,reference simulations for two other cases documented byKnippertz et al. [2007] were conducted. These are the

horizontally rather extended density current on 07 May2006 and the more localized event on 31 May 2006. For07 May the model quite realistically reproduces severalconvective cells that form over the High Atlas in the earlyafternoon and have propagated into the Algerian Sahara by2100 UTC, similar to the case discussed in section 3.1. Atthis time the vertical-velocity identification criterion (blacklines in Figure 7a) shows separate leading edges of fourcells (Figure 7a), while the satellite image indicates onlytwo already merged leading edges marked by arc clouds(Figure 7c). As in section 3.1, the COSMO model simulatesa too extended cirrus shield that covers the leading edge atthis time (Figure 7b). The propagation velocity of theleading edge between 1800 UTC 07 May and 0000 UTC08 May is on the order of 14 m s�1, again with a slowerpropagation toward the west. Maximum grid point windsassociated with the density current reach somewhat highervalues of 16 m s�1 at the time shown in Figure 7a. Bymidnight the cold pool reaches an west–east extension ofmore than 500 km (not shown). The southward displace-ment of the leading edge in the satellite data by approxi-mately 50 km is mainly due to an earlier initiation ofconvection by about an hour than by a faster propagationof the cold pool. Overall the size, location, and propagationof the observed system are quite realistically reproduced. Inagreement with this, the model-simulated precipitation

Figure 5. Comparison of precipitation accumulated over the 24-h period 0000 UTC 03 June to 0000UTC 04 June 2006 (a) simulated by the COSMO model with (b) TRMM estimates and selected stationobservations.The cross marks the station Adrar, where no precipitation was observed. The black and redlines in Figures 5a mark regions of, respectively, strong vertical and horizontal winds at 0000 UTC 04June 2006 as in Figure 2. The black line in Figure 5b is the corresponding, subjectively identified leadingedge as in Figure 4d.

Figure 6. Time series between 0000 and 2400 UTC 03 June 2006 of (a) temperature, (b) dew point, (c) wind speed (lines)and direction (diamonds), and (d) pressure observed at the IMPETUS station EMY (red lines and circles), and simulated bythe COSMO model and interpolated to the station location (blue lines and circles). The slightly higher model orography atthis point leads to the systematic differences in pressure in Figure 6d, which is measured with a rather coarse resolution of1 hPa at EMY. Temporal resolution is 15 min for both time series. The time of passage of the leading edge between 1700and 1900 UTC is indicated by orange shading.

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Figure 6

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Figure 7. As in Figure 2 but for the cold pools (a) at 2100 UTC 07 May 2006 and (d) at 2000 UTC 31May 2006. The corresponding brightness temperature images (b and e) simulated by the COSMO modeland (c and f) from Meteosat data are shown as in Figure 4.

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matches the TRMM estimates well over the Sahara (notshown). Maximum values are reached over Algeria close tothe southeastern corner of Morocco with storm totalsexceeding 20 mm in the model, TRMM estimates, andsurface observations.[23] Despite its smaller horizontal extent, the density

current on 31 May is also reasonably well reproduced,although with some time delay. In satellite imagery deepconvection over the Atlas is visible as early as 1200 UTC,while the model initiates the first precipitating cells about2 h later (not shown). By 2000 UTC the leading edge haspassed EMY [see Knippertz et al., 2007, Figure 7] andspreads into Algeria as indicated by brightness temperaturesbelow (above) 295 K within the cold pool (in the surround-ings) (Figure 7f). Owing to the delayed triggering ofconvection, the associated cirrus shields are less extendedin the model than in the observations (Figure 7e) and theleading edge lags behind by about 100 km (Figure 7d). Thepropagation direction to the south and a propagation speedon the order of 10 m s�1, however, are in good agreement,and the accumulated model-generated precipitation satisfac-torily matches TRMM estimates (not shown). In contrast tothe other two cases, there are hardly any grid points withlow-level ascent above 3 m s�1 along the leading edge (notethe absence of black dots in Figure 7d). The authors suspectthat this is due to the orientation of the leading edge almostparallel to the fairly strong westerly shear on this day, whichimpedes an optimal organization of the system [Weismanand Rotunno, 2004]. Weaker ascent is consistent withprecipitation being more confined to the mountains and aslower propagation than in the other two cases. Maximum

grid point winds are also weaker, reaching 15.4 m s�1 at thetime shown in Figure 7d.[24] This analysis corroborates the general ability of the

COSMO model with 2.8-km horizontal grid spacing togenerate cold pools with a realistic temporal evolution,structure, air mass characteristics, and propagation speed.The cold pool evolution reacts sensitively to temporal andspatial details of the precipitation distribution in the modelas will be shown in section 4. The authors believe that oneof the reasons for this overall satisfactory model perfor-mance is the robust orographic trigger provided by the AtlasMountains when compared to the sometimes very problem-atic initiation of convection over flatter terrain [Montmerleet al., 2006].

4. Sensitivity Studies

[25] In this section the sensitivity of the cold poolevolution over the northwestern Sahara to the treatment ofmicrophysical and turbulent processes in the COSMOmodel will be tested, again using the case of 03 Junedescribed in section 3.1. The following subsection containsinformation on the physical meaning of the parameters to bevaried in the sensitivity experiments, while the secondsubsection discusses the results.

4.1. Rationale

[26] Several studies have shown that convective down-drafts are fostered by the existence of relatively smallraindrops that can be more effectively evaporated in thesubcloud layer [Kamburova and Ludlam, 1966; Srivastava,1985; Proctor, 1989]. In the bulk microphysics scheme usedin the COSMO model, evaporation of rain is highly param-eterized and drop size distributions are represented by agamma distribution of the form

n Dð Þ ¼ N0Dm exp �lDð Þ; ð2Þ

where n(D) is the number density, D is the drop diameter,N0 the intercept parameter, l the slope, and m the shapeparameter. In one-moment microphysical schemes, N0 and mare set constant, and l is a unique function of the rain watercontent (RWC). With some additional assumptions, theevaporation rate can then be calculated from the RWC.[27] In the following simulations the impact of the choice

of m on the evolution of the density current is evaluated.In the standard version of the COSMO model m = 0.5 isused, which is to some extent motivated by the study ofSchlesinger et al. [1988], who found a value of about 0.4 tobe optimal on the basis of radar observations of evaporatingrain. Here, results from additional model simulations usingvalues of zero and unity are presented. For these runs, theintercept parameter, N0, is determined as a function of musing equation (27) of Ulbrich [1983],

N0 ¼ 6� 104 exp 3:2mð Þ: ð3Þ

[28] The corresponding raindrop size distributions areshown in Figure 8. The standard value of 0.5 used for thereference simulations gives a maximum near 0.4 mm. Thedistribution with m = 0 describes an exponential decreaseof drop numbers with diameter and corresponds to the

Figure 8. Raindrop size distributions for different valuesof the dimensionless shape parameter m. D is drop diameterin millimeters, and n(D) is the corresponding numberdensity in m�3 mm�1.

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Marshall-Palmer distribution as used in most bulk micro-physical schemes [Lin et al., 1983]. Compared to m = 0.5 ithas a reduced (increased) number of large (small) drops. Forsuch a modification the evaporation in the subcloud layer isexpected to increase, which should decrease precipitation atthe ground and intensify the cold pool owing to theenhanced density differences to the environment. Theopposite can be expected for a change to m = 1, whichshifts the maximum of the distribution to about 0.8 mm(Figure 8). In the sensitivity experiments the changes to mwere only applied in the parameterization of evaporationand the sedimentation velocity of rain, but not in other partsof the microphysics package.[29] In addition, the more sophisticated two-moment

microphysics scheme of Seifert and Beheng [2006] (SBhereafter) is applied in one simulation. The SB two-momentscheme has recently been extended to include a separate hailcategory (U. Blahak, Towards a better representation ofhigh density ice particles in a state-of-the-art two-momentbulk microphysical scheme, paper presented at 15th Inter-national Conference on Clouds and Precipitation, Interna-tional Association of Meteorology and AtmosphericSciences, Cancun, Mexico, 2008), and a new parameteriza-tion of evaporation of raindrops with a diagnostic relationfor m [Seifert, 2008]. For the experiment with the SB two-moment scheme the time step had to be reduced from 30 to10 s. In order to rule out an influence of the change in timestep on the results, an additional simulation with thestandard configuration but a time step of 10 s was run,and only negligible differences were found (not shown). Itshould be stressed that the usage of the SB two-momentscheme affects all microphysical processes and not just theparameterization of evaporation as in the experiments withthe modified m parameter.[30] In another simulation the asymptotic turbulence

length scale l1 was modified. The TKE scheme of theCOSMO model uses the classic Blackadar-Deardorff for-mulation of the turbulent mixing length

lturb ¼kz l1

kzþ l1; ð4Þ

with height z and the von Karman constant k = 0.4. Thestandard value is l1 = 200 m. In one simulation a decreasedl1 of 60 m is used, which reduces vertical mixing. This canaffect the density current in two ways. First, reduced mixingof environmental air into the cold pool will slow its demise.Second, reduced mixing in the planetary boundary layerwill sustain larger vertical temperature gradients duringdaytime. The associated destabilization of the lower layerscan help to initiate convection and increase evaporation,which would also favor the formation of cold pools. Thereduction of l1 to 60 m was found to improve forecasts ofdeep convection over Germany (A. Seifert et al., Thechallenge of convective-scale quantitative precipitationforecasting, paper presented at 15th International Con-ference on Clouds and Precipitation, International Associa-tion of Meteorology and Atmospheric Sciences, Cancun,Mexico, Cancun, Mexico, 2008) and therefore thismodification has recently been made operational at theDWD.

4.2. Results

[31] Figure 9 shows accumulated precipitation, the10 m s�1 isotach and leading edge identification for the foursensitivity experiments described in section 4.1. The leadingedge from the corresponding figure for the reference sim-ulation, Figure 5a, is indicated in green. The drop sizedistribution used for the evaporation computation has asubstantial impact on the amount of precipitation reachingthe ground (Figures 9a and 9b). While the overall structureof the precipitation field remains almost unchanged,amounts over the Algerian desert are increased by severalmillimeters over a large area with m = 1, and reduced totraces with m = 0. The impact on the propagation of thedensity current, however, is moderate, resulting in differ-ences in leading edge position of up to 30 km as comparedto the reference simulation. As expected, the area of windsabove 10 m s�1 is extended (reduced) with maximum windsof 16 m s�1 (14.4 m s�1) for m = 0 (m = 1) demonstratingthe influence of the microphysics on the dynamics. This hassome implications for dust mobilization, which, being athreshold problem with an approximately cubic relationbetween wind and vertical dust flux, reacts rather sensitivelyto changes in peak winds [Cakmur et al., 2004].[32] The use of the SB two-moment scheme (Figure 9c)

produces a precipitation pattern similar to the referencesimulation (Figure 5a) with somewhat higher amounts overthe Sahara (up to 10 mm at some grid points). The leadingedge position agrees with the one-moment simulation with-in less than 10 km but extends more than 200 km farther tothe north according to the vertical motion criterion (comparegreen and black dots in Figure 9c). Surprisingly the near-surface winds behind the leading edge are weaker than in allother simulations with a maximum of only 12.4 m s�1 at0000 UTC 04 June (15.6 m s�1 in the reference simulation),pointing to a weaker density current. The explanation forthis unexpected result lies in a different temporal evolutionas compared to the reference simulation. The convectioninitiation is concomitant, but the simulation with the SBtwo-moment scheme subsequently generates more precipi-tation and a more intense density current during the after-noon and early evening. At 1900 UTC the run with the SBtwo-moment scheme generates maximum wind speeds of22.7 m s�1 as opposed to 17.5 m s�1 in the referencesimulation (not shown). At 2000 UTC the eastern part ofthe leading edge lies about 42 km ahead of the one in thereference simulation (not shown). According to equation (1),the faster propagation in the sensitivity experiment shouldbe related to a larger integrated density difference betweenthe cold pool and the surroundings and thus to a strongerevaporative cooling associated with a more localized (inspace and time) evolution of precipitation. Between thistime until midnight, precipitation rates in the simulationwith the two-moment scheme decrease significantly and theleading edge slows down. The described differences havemany potential reasons that are difficult to unravel from afully nonlinear model simulation. They include the usage ofhail in addition to low-density graupel, the different treat-ment of particle sedimentation (no gravitational sorting inthe one-moment scheme), and last but not least, the differentparameterizations of evaporation of raindrops.[33] One possible explanation of the comparably small

sensitivity of propagation velocity to changes in evaporation

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microphysics is related to the deep dry-adiabatic mixedlayer the density current is forming in (see Figure 3b).Idealized one-dimensional computations show that in thiscase downdrafts and surface cold pools form for a wholerange of drop sizes, while precipitation microphysicsbecome increasingly important for greater vertical stability[Kamburova and Ludlam, 1966; Srivastava, 1985]. A givenRWC distributed into smaller drops is generally more effi-cient, while bigger drops spread the cooling over a greaterdepths, which may become important for small RWCs at theedges of the simulated precipitation zones. Another moredynamical reason is the rather weak dependence of thetheoretical density current propagation velocity c on theinduced cooling (equation (1)). If we assume the height tobe mainly controlled by ambient factors such as shear, we areleft with a simple square-root dependence of c on Dq.[34] The reduction in l1 to 60 m has the by far largest

influence on precipitation and cold pool evolution (Figure 9d).

In contrast to the three other simulations the precipitationincreases again over the Algerian Sahara after an intermit-tent decrease close to the Moroccan-Algerian border (notshown). This way the total amount of precipitation isdistributed more evenly over a larger area, most likelyleading to more evaporation and a larger density contrast.Consequently, the cold pool propagates faster (13–18 m s�1)than in the reference simulation and generates the highestmaximum wind of 17.1 m s�1 at 0000 UTC 04 June 2006.According to the vertical motion criterion the cold poolextends some 70 km farther to the west than in the referencesimulation at this time (Figure 9d). This simulation agreesbest with the TRMM rainfall estimates and with the leadingedge identified from infrared imagery (Figure 5b), eventhough the initiation of precipitation over the AlgerianSahara during the afternoon is missed out as in all the othersimulations. As explained above, possible ways of howl1 affects the cold pool are (1) through changes in the

Figure 9. As in Figure 5a but for the sensitivity experiments with (a) m = 0, (b) m = 1, (c) the SB two-moment scheme and 10-s time step, and (d) l1 = 60 m. The position of the leading edge from thereference simulation with m = 0.5, l1 = 200 m, and 30-s time step (see Figure 5a) is marked in green.

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boundary layer structure that allow an easier initiation ofconvection and (2) through changing the turbulent mixingof environmental air into the density current.[35] These results show that the most important factor for

reproducing the position and propagation direction of thedensity current is a correct initiation of deep convectivecells. Changes to the microphysics and turbulence param-eterizations mainly modify the fraction of precipitationreaching the ground and thereby indirectly affect the tem-poral evolution and strength of the associated cold pool. Themoderate impact on near-surface winds has some relevancefor the modeling of dust emissions associated with suchsystems.

5. Summary and Conclusions

[36] High-resolution model simulations of three long-lived and extensive cold pools in the northwestern Saharaduring May/June 2006 have been conducted with the non-hydrostatic COSMO model with 2.8-km horizontal gridspacing, sufficient for an explicit treatment of deep convec-tion. These cold pools are initiated by evaporating rain fromdeep convection over the High Atlas in the deep, dry-adiabatic, well-mixed desert boundary layer. The associateddensity currents have lifetimes of 10 h and spread overseveral hundred kilometers far into the Algerian Sahara withfairly constant propagation speeds over several hours. At thesurface, the passage of the leading edge is accompanied byan abrupt drop in temperature and visibility, jumps in dewpoint, wind speed and pressure, and a change in winddirection [see also Knippertz et al., 2007]. Strong ascentat the leading edge of more than 7 m s�1 allows aregeneration of convective cells and thus an extension ofrainfalls into the desert in two cases, while one case withstrong westerly wind shear shows no such propagation. Themore sustained source of evaporatively cooled air in theformer two cases leads to a faster propagation and ulti-mately to a larger extent of the cold pool. Strong near-surface winds behind the leading edge of up to 20 m s�1

lead to dust emissions, as shown by the few availablesurface observations in this region [Knippertz et al., 2007]and by dust model studies [Reinfried et al., 2009]. These so-called haboob dust storms [e.g., Miller et al., 2008] are acommon feature of the warm season over the world’sdeserts, but comparably few attempts have been made tomodel these systems with a nonidealistic setup. The largedistance of the leading edge with the strongest near-surfacewinds from the active precipitation region during the latestages of the development in combination with the very lowprecipitation rates at the ground indicate that an immediatescavenging of the emitted dust as has been suggested in theliterature (see discussion by Williams [2008]) appears ratherunlikely.[37] Comparisons to available satellite and ground obser-

vations reveal that the COSMO model successfully repro-duces the formation and the propagation of all three densitycurrents with a realistic three-dimensional structure. Thetriggering of deep convection over the mountain range iswell represented, although at times with some delay com-pared to satellite observations, while convective initiationover the Algerian Sahara appears more problematic. Thelatter affects the orientation of the leading edge, the extent

of the cold pool, and its propagation direction and speed.This further corroborates and extends results by Reinfried etal. [2009] that a correct initiation of convection is the mostimportant factor for a satisfactory reproduction of the coldpools. The same authors have also demonstrated the need ofan explicit treatment of deep convection to achieve this.[38] Four sensitivity studies were conducted to test the

impact of changes to the model’s microphysics and turbu-lence schemes on the density currents. Changes in raindropsize distribution used for the evaporation parameterizationlargely affected the amount of precipitation reaching theground, but had a moderate effect on the cold pool propa-gation (leading edge position differences of less than 30 km).This is most likely related to the deep dry-adiabatic sub-cloud layer that is particularly favorable for convectivedowndraft generation and therefore less sensitive to micro-physics than more stable stratifications. Other potentialreasons are the effect of precipitation drag and the weakdependence of density current propagation velocity on theamount of cooling. The usage of the SB two-momentmicrophysical scheme changed the temporal evolution ofthe density current leading to more intense early stages withstrongest winds and an earlier decay. Presumably this is dueto changes in microphysical properties both inside andbelow the clouds. Finally a reduction in the turbulencelength used for vertical mixing in the boundary layerparameterization scheme resulted in more widespread andweaker precipitation, and a larger, faster propagating coldpool with higher winds. This change affects both theconvective initiation out of the boundary layer and thepropagation of the cold pool via changes in turbulentmixing.[39] In conclusion these results corroborate the conjecture

made by Freeman [1952] more than half a century ago thatforecasting the occurrence of haboob dust storms largelyreduces to forecasting moist convection. As shown byReinfried et al. [2009] and in this paper an explicit treatmentof deep convection appears to be an important prerequisiteto achieve this. Potential problems in this context are theavailability of reliable initial conditions in the commonlyextremely data-sparse desert regions and model deficienciesin the initiation of convection in areas without synoptic-scale or orographic forcing. Once the cold pool is initiated,the forecast of new cells is facilitated by the convectivetriggering through the ascent at the leading edge, which hassome implication for nowcasting. Though a secondaryeffect for the macroscopic behavior of the cold pool, thesensitivities to microphysics and turbulence treatment dem-onstrated here have some implications for dust emissionsimulations due to the threshold nature of the problem andthe cubic dependence of vertical dust flux on wind speed[Cakmur et al., 2004]. More cases studies and in particularbetter observations are urgently needed to improve thephysical parameterizations in the models for robust andrealistic representations of cold pools in desert regions.

[40] Acknowledgments. P.K. is currently funded through the DFGEmmy Noether program (grant KN 581/2-3). His participation in theSAMUM field campaign was supported through the Forschungsfond ofthe Johannes Gutenberg University Mainz. The data from the IMPETUSautomatic weather stations were provided by the climatology group(Winiger, Loeffler, Schulz) at the Department of Geography, Universityof Bonn, Germany. We acknowledge the DWD for supporting the use of the

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COSMO model and for providing the necessary external parameters. TheIR satellite data were provided by EUMETSAT and the TRMM data wereacquired using the GES-DISC Interactive Online Visualization ANd aNal-ysis Infrastructure (Giovanni) as part of the NASA’s Goddard EarthSciences (GES) Data and Information Services Center (DISC). We wouldlike to thank Florian Meier and Philipp Reutter for their help withprocessing these data and Earle Williams and Cristel Bouet for theirvaluable comments that helped to improve an earlier version of this paper.

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�����������������������P. Knippertz, Institute for Atmospheric Physics, Johannes Gutenberg

University Mainz, Becherweg 21, D-55099 Mainz, Germany. (knippertz@uni-mainz.de)A. Seifert and J. Trentmann, German Weather Service, Frankfurter

Strasse 135, D-63067 Offenbach, Germany.

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