a case of cyclogenesis

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Meteorol. Appl. 11, 97113 (2004) DOI:10.1017/S1350482704001161

A case of cyclogenesis over the westernMediterranean Sea with extraordinaryconvective activity

M. Kurz1 & A. Dalla Fontana21Wittelsbacherstr. 53a, D-67434 Neustadt/Weinstrae, GermanyEmail: [email protected] Meteorologico di Teolo, via Marconi 55, 35037 TEOLO (PD), ItalyEmail: [email protected]

An interesting case of cyclogenesis over the western Mediterranean Sea, accompanied by extraordinaryconvective activity, is described in this paper. The event took place between 18 and 21 July 2001. Itssynoptic features are examined on the basis of quasi-geostrophic theory using numerical analyses from

the global model of Deutscher Wetterdienst GME and from the ECMWF. Particularly remarkable is thelarge amount of latent heat released in the condensation process, which produces significantmodifications of the flow. These effects are evident both in the fields of relative vorticity as well as

potential vorticity in the upper levels. The infrared Meteosat satellite pictures are used to follow theevolution of the upper-level cloudiness associated with the process and its connection with the values ofvorticity and winds at 300 hPa. Radar pictures, reflectivity and radial velocity from Dopplersurveillance radar in the Veneto region allowed us to outline the structure of the mesoscale convectivesystem a squall line that developed during the night of 1920 July over Italy, in the Padana Valley.The features of this system were compared with a conceptual model proposed by Houze et al. (1989).The vertical wind shear, as inferred from the radial velocity pictures, was likely to play an importantrole in the development of the storm. Finally, an analysis of the vertical stability in the pre-storm

environment was performed using sounding data from Udine, 20 July 2001, 00 UTC. The analysisshows the importance in the storm development of the potential instability released by ascent during thetransit of the baroclinic system, a process that was investigated using the vertical profile of equivalent

potential temperature and an estimate of ascending motion given by the ECMWF model.

1. Introduction

Between 18 and 21 July 2001 a cyclogenesis eventtook place over the western Mediterranean Sea. Thecyclone moved quickly north-eastwards influencing

the northern parts of Italy. The strong ascendingmotions ahead of and above the cyclone led to theformation of a convective system during the night of1920 July that struck the Veneto region, in north-east Italy, bringing hail, stormy winds and heavyrainfall. This kind of phenomenon is not unusual inthe Padana Valley during the summer season when theatmosphere is almost always potentially unstable. Thus,when external forcing occurs, for example with a coldfront or a trough, it is quite likely that convectiveoverturning will follow. The case presented here is richin interesting features at all scales of observation. On asynoptic scale the cyclogenesis took place following a

typical type B development, as defined by Petterssen &Smebye (1971), when an upper level vorticity maximum

approached a quasi-stationary frontal zone in the lowertroposphere. On the mesoscale, as already pointed out,a squall line developed during the nocturnal hours,triggered by the passage of the cyclone over northeastItaly. This system presents some interesting features

in accord with a conceptual model by Houze et al.(1989) that have been highlighted through the analysisof radar pictures. In the description of this case weused numerical analyses from the global model ofDeutscher Wetterdienst GME and from ECMWF andsatellite images in the infrared channel from Meteosat.Furthermore, to describe the vertical stability of theatmosphere we used sounding data from Udine stationlocated in Friuli Venezia Giulia, a border regionnorth-east of Veneto. To outline the structure ofthe convective system we used reflectivity and radialvelocity data obtained from the surveillance C-bandDoppler radar owned by the Meteorological Office of

Teolo, located on a hilltop near Padua, in the Venetoregion.

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Figure 1. Surface analyses with isobars, fronts and synoptic observations from 18 July 2001, 12 UTC and 19 July 2001, 00 and12 UTC (from left to right).

2. Synoptic description and diagnosis

As outlined above, a cyclogenesis event over the westernMediterranean took place following the scheme ofa typical type B-development. The surface cyclonecrossed northern Italy and reached its strongest in-tensity there before moving further north-eastwardstowards the Balkans. The passage of the low wasconnected with extraordinarily heavy convective over-turning above northern Italy.

The cyclogenesis occurred south of a large surfaceand collocated upper air low over western Europecontaining deep cold air. The frontal zone in thelower troposphere ran initially from southern Spainto the Alps, i.e. from southwest to northeast. It wasembedded in a broad trough in the surface pressure field(Figure 1) and connected with vorticity maximum at

850 hPa above southern Spain and the Gulf of Genoa.It was a shallow feature reaching vertically upwardsonly to about 700 hPa. The upper vorticity maximum(300 hPa) was at first positioned above western Biscay,i.e. more than 1300 km from the lower frontal zone(Figures 2 and 3). It was embedded in a small troughat the cyclonic flank of a jet streak running northwestto southeast, i.e. perpendicular to the lower front.The jet streak was connected with an upper frontalzone (500300 hPa) lying ahead of an upper ridge andrunning in the same direction.

At the beginning there was no defined connection

between the lower and the upper frontal zones. Withisotherms running from west to east in the mid-troposphere, a clear component of the thermal windwas towards the lower front. That was accordinglyvalid also for the upper current so that a quick relativemovement of the upper vorticity maximum towards thelower frontal zone was possible.

The approach of the upper trough with its vorticitymaximum took place during the 24 hours between18 July, 12 UTC, and 19 July, 12 UTC. Owing tothe vertically increasing positive vorticity advection(PVA) ahead of the maximum, an ascending motion

was released connected with divergence aloft andconvergence in the middle and lower levels leading to

the production there of cyclonic vorticity. As shownby the vertical cross section in Figure 4, a sloped zonedeveloped in which the vorticity steadily increased andextended downwards to engage finally with the surfacelevel. This process was connected with pressure fall at

all levels. At the surface, a closed low developed with acentral pressure below 1005 hPa east of the Balearics on19 July, 12 UTC.

Together with the upper trough, thecoldair in the upperand middle levels moved forward in the direction of thefrontal zone in the lower troposphereprimarilydue tothe cold air advection (CA) working in nearly all levelsat first, but surely also due to the effect of the ascendingmotion being effective ahead of the trough. As a result,the baroclinicity increased strongly, especially in themiddle levels, and a continuous frontal zone developedaheadofthetroughwithaslopeof1:150atthebeginning

and 1:80 by 19 July, 12 UTC (see Figure 4). Connectedwith this change, the upper jet streak moved forwardrelative to the trough, finally reaching maximum speedahead of it.

The vertical coupling between the lower frontal zoneand the upper vorticity maximum was completed on19 July, 12 UTC. The analyses of the events on 19 July inFigures 13 show the typical features of a baroclinicallyunstable wave, namely the surface low ahead of theupper trough in the region with strong upper PVA anddivergence, ascending motion in the middle levels and

convergence in the lower troposphere, and the uppertrough, on the other hand, above the region with strongCA west of the surface low giving rise to a forcing ofdescent with lower divergence and upper convergence.

In such a system a further intensification has tobe expected in all levels. Indeed, the surface lowdeepened to a central pressure below 1000 hPa duringthe following 12 hours (Figure 5), as it moved tonorthern Italy following the lower warm air advection(WA). Also, the vorticity at 850 and 700 hPa increasedfrom 10 to 16105 s1. At 500 hPa, however, thevorticity remained the same (21105 s1), whereas it

significantly decreased at 300 hPa (from 30 down to24 105 s1; compare Figures 3 and 7).

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Cyclogenesis over the Mediterranean Sea

Figure 2. GME analyses of geopotential (solid lines, in gpdam) and temperature (dashed lines, in C) for 850, 700, 500 and300 hPa (from bottom to top) from 18 July 2001, 12 UTC and 19 July 2001, 00 and 12 UTC (from left to right).

The missing vorticity increase in the mid-troposphereand the decrease in the upper levels can be tracedback to the effect of the release of large amountsof latent heat in the area together with ascendingmotion, cloud formation and precipitation. The coolingdue to the ascent might thus be strongly reducedor even fully compensated. On the other hand, theforcing for the ascent is enforced giving rise to strongerconvergence below and stronger divergence above thelevel with the maximum vertical velocity. Consideringthe temperature at the different levels in the region ofthe moving upper vorticity maximum, it has to be statedthat the temperature at 300 hPa itself remained nearly

constant (around 40 C) apart from the first time,whereas a warming by 6 K took place at 500 and 700 hPa

in the period between 19 July, 12 UTC, and 20 July,12 UTC (Figures 2 and 6). The temperature at 500 hPaincreased from 21 to 15 C, and at 700 hPa from3 to 3 C. Since there was ascent ahead of the vorticitymaximum during the whole period, the warming can beexplained only by the effects of either WA relatively tothe moving maximum or release of latent heat, or byboth. Here we can assume that the latter effect providedthe biggest contribution to the temperature increase inthe mid-troposphere.

The enhanced divergence above the level of themaximum ascent (up to 5105 s1) provides a logical

explanation for the vorticity decrease at 300 hPa(Figure 7). This is confirmed when considering analyses

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Figure 3. GMEanalyses of a) winds and relativevorticity (in units of 3 105 s1 , negative valuesdashed) for 300 hPa (top) and

850 hPa (bottom) and b) divergence (in units of 5

106

s1

, negative values dashed) for 300 hPa (top) and 950 hPa (bottom)from 18 July 2001, 12 UTC and 19 July 2001, 00 and 12 UTC (from left to right).

of the isentropic potential vorticity (IPV) which isdefined by

IPV= g( f+ )

p

where g is gravity acceleration, f is the Coriolisparameter, is the potential temperature and is relative vorticity at surfaces of constant potential

temperature. This property remains constant forparticles under adiabatic conditions. Here, analyses of

the IPV at the 320 K surface of constant potentialtemperature (as shown in Figure 8) are considered. Thisisentropic surface had a pressure height of between 300and 500 hPa in the area of interest. Corresponding to the300 hPa vorticity maximum there was an IPV maximumnorthwest of theBay of Biscay at the beginning, movingsouth-eastwards and approaching the lower frontalzone and the developing surface low until 19 July,

12 UTC. During this time the amount of the IPVremained nearly constant at 1113 units, indicating the

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Figure 4. Vertical cross sections of: (a) potential temperature (in C) and (b) relative vorticity (in 105 s1) from 18 July 2001,12 UTC and 19 July 2001, 00 and 12 UTC (from left to right). (c) Geographical positions of the cross-sections at different times.

nearly adiabatic movement of the air particles formingthis maximum, which had a position between 300 and350 hPa. Afterwards, however, a significant decreaseof the IPV is reflected in the analyses: down to only5 or even 4 units in the region where the 300 hPa

vorticity maximum was situated. As result, the originalIPV maximum disappeared fully for a while (20 July,

00 UTC), only becoming visible again on 20 July,12 UTC, as a very weak feature compared with the nextextreme following upstream.

This decrease in IPV has to be traced back to non-

adiabatic processes, i.e. the release of latent heat leadingto a stabilisation below and a destabilisation above

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Figure 5. Surface analyses with isobars, fronts and synoptic observations from 20 July 2001, 00 and 12 UTC (from left to right).

the level where its contribution to the temperatureincrease is strongest. According to the definition of IPV,a diabatically forced stabilisation causes an increase inthis otherwise conservative property of the air, whereasa diabatically forced destabilisation causes a decrease.The latter statement can be applied to the significantIPV decrease reflected by the IPV analyses on the 320 Ksurface in our case.

The release of latent heat and the associated anti-cyclogenetic effect in the upper levels surely also contri-

buted to the changes that affected the vorticity fielddownstream of the upper trough. With the formationof a ridge in the isohypses ahead of the trough, a re-gion with increasing negative vorticity developed northof the trough between the main vorticity maxi-mum and another one downstream of it which wasquickly advected north-eastwards. The negative relativevorticity in this area reached close to 10105 s1

indicating conditions with very small absolute vort-icity and therefore small dynamic stability or even indif-ference. It was this area that saw the main cloud forma-tions and the strongest convective activity of the system.

Thedescribed effects were even much more pronouncedin the relevant analyses of ECMWF. Comparing thevorticity analyses for 300 hPa from GME and ECMWFfor 18, 19 and 20 July, 12 UTC (Figure 9), goodagreement can be obtained for the first and the lastdays apart from the fact that the amounts ofrelative vorticity are generally somewhat higher in theECMWF analyses according to their higher resolution.By contrast, the analyses for 19 July, 12 UTC, aresignificantly different: instead of the strong maximum inthe trough and the rapid vorticity decrease downstreamin the GME analysis, the ECMWF analysis shows an

elliptic area with strong negative values of vorticityjust ahead of the trough connected with bands of high

positive vorticity southeast and northwest of it. This isdue to the strong anticyclonic curvature of the analysedwind field immediately ahead of the trough togetherwith a splitting of the jet stream around a distinct areaof lower velocity. This analysis is very similar to the12-hourly forecast and points to the fact that the ascentsimulated in the ECMWF model was so strong thatthe upper divergence to which the ascended air wassubjected,influencedthe wind and vorticitydistributionmuch more than in the other model. Owing to theinadequate resolution of the upper-air observations it isnot possible to determine which analysis was the morerealistic.

3. Satellite images description

This section describes how the cyclogenetic develop-ment over the Mediterranean was reflected by thedistribution and change over time of cloudiness in themiddle and upper levels as shown by the infrared imagesobtained from METEOSAT. In order to allow a directcomparison, the GME analyses of winds and relative

vorticity at 300 hPa were superimposed on the imagesin Figure 10 covering the interval between 18 July 2001,12 UTC and 20 July 2001, 00 UTC.

At noon on 18 July (Figure 10a) the cloud distributionwas characterised by two large cyclonically bent cloudbands winding around the centre of the surface andupper low above the English Channel. The easternband was situated downstream of an upper vorticitymaximum (i.e. in the area of PVA), whereas the westernband lay not ahead but on the cyclonic flank of thevorticity maximum above the western Bay of Biscaymoving south-eastwards. There was nearly no cloud

connected with the quasi-stationary frontal zone in thelower troposphere above the western Mediterranean.

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Figure 6. GME analyses of geopotential (solid lines, in gpdam) and temperature (dashed lines, in C) for 850, 700, 500 and300 hPa (from bottom to top) from 20 July 2001, 00 and 12 UTC (from left to right).

Some 12 hours later (Figure 10b), the image showsmedium and upper cloud both ahead of and behindthe trough and the vorticity maximum, which bynow had reached northern Spain. This means that thecloud distribution was not only determined by thePVA ahead of the trough. Above the lower front to

the southeast of Spain, there was still only very littlecloud.

By 19 July, 12 UTC (Figure 10c), however, the cloudat the rear of the eastward-swinging trough hadprogressively disappeared, whereas an impressive cloudshield had developed in the area of strong PVA aheadof the trough. The nearly circular shape of this shieldemphasises the very strong divergence of the flow in the

upper levels. This is in good agreement with the modelanalysis of the divergence at 300 hPa in Figure 3b.

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Figure 7.GME analyses of a) winds and relative vorticity (in units of 3

10

5

s

1

, negative values dashed) for 300 hPa (top)and 850 hPa (bottom) and b) divergence (in units of 5 106 s1) for 300 hPa (top) and 950 hPa (bottom) from 20 July 2001,00 and 12 UTC (from left to right).

According to surface observations, this cloud shieldformed the upper part of a deep cloud mass that wasproducing rain and also thunderstorms north of thesurface low that had formed to the east of the Balearesand its warm front. The rapid development of thiscloud mass indicated that the cyclogenetic couplingbetween the upper trough and the lower frontalzone had now been completed, and the whole tropo-

spheric air was being subjected to a strong ascendingmotion.

As already mentioned, the ECMWF analysis for300 hPa on 19 July was quite different in showingan elliptic area with negative vorticity just ahead ofthe trough, and two bands of high positive vorti-city southeast and northwest of it. As demonstrated byFigure 11, the area with negative vorticity coincides toa great extent with the upper cloud shield.

During the following 12 hours, the cloud shieldexpanded a little bit and extended band-like in a

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Figure 8. GME analyses of a) isentropic potential vorticity (IPV, in 5107 K m2kg1 s1 ) at 320 K surface of potentialtemperature from 18 July 2001, 12 UTC, 19 July 2001, 12 UTC, 20 July 2001, 00 and 12 UTC (from top left to bottom right),thick arrows indicate the maximum value associated with the upper level trough, and b) pressure of this surface (in units of25 hPa) from 19 July 2001, 00 and 20 July 2001, 00 UTC (left to right).

Figure 9. ECMWF analyses of winds and relative vorticity (in units of 1105 s1) for 300 hPa from 18 July 2001, 12 UTC,19 July 2001, 12 UTC and 20 July 2001, 12 UTC.

north-easterly direction. By 20 July, 00 UTC (Fig-

ure 10d), it covered the whole of northern Italy, theAlps, Austria and Czech Republic and as far as southern

Poland. Its southern part still lay in the area of strong

PVA and related divergence between the upper troughand the ridge which began to form north of it. This was

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Figure 10. METEOSAT infrared images superimposed to GME analyses of relative vorticity (in units of 3105 s1) and windsat 300 hPa. (a) 18 July 2001, 12 UTC; (b) 19 July 2001, 00 UTC; (c) 19 July 2001, 12 UTC; (d) 20 July 2001, 00 UTC.

the phase of the most intense convective activity. The

radar pictures show that an organised convective systemdeveloped inside the circulation of the cyclone, cross-ing the Venetoregion from southwest to northeast. Butthis was not discernible from the satellite imageryalone.

4. Description and diagnosis of the radarpictures

During the afternoon of 19 July, near the quasi-stationary surface front, many thunderstorms struckthe Veneto region, particularly the flat plain. In the

night, as the deepening cyclone and the upper vorticitymaximum approached, an organised convective system,

linear in shape, developed to the southwest and crossed

the region north-eastwards, bringing hail, stormy windsand heavy rainfall. The C-band Doppler surveillanceradar owned by the Meteorological Centre of Teolo,located on the top of a hill near Padua, followedthis event with volumetric scans every 5 minutes. Thepictures of reflectivity and radial velocity exhibit thestructure of the convective system and its evolution.It was an example of a squall line, with features inagreement with a conceptual model of a squall linein its mature stage proposed by Houze et al. (1989).Each reflectivity picture presented here is composedof a horizontal picture (CAPPI) at a height of 1.5 km,together with a vertical section in the direction of storm

motion, projected along the two Cartesian axes. Therange explored by the radar beam is 120 km with a

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Figure 11. METEOSAT infrared image at 19 July 2001,12 UTC superimposed to ECMWF analysis of relative vorticity(in units of 1 105 s1) and winds at 300 hPa.

spatial resolution of 1 km. No clutter filter was appliedin these pictures, but looking at the vertical profile itis possible to differentiate the echoes of precipitation.Furthermore, you will note two large shadow coneslocated to the west and to the south which are due tothe blocking effect of higher hills surrounding the radarsite.

At 2130 UTCthe squall line crossed the border betweenVeneto and Emilia-Romagna, moving north-eastwards(Figure 12a). Stronger reflectivity values at this stagewere generally between 50 and 55 dbz. The verticalsection shows the maximum reflectivity in the medium

layers at a height of between 3 and 7 km, as is typical forthe initial stages of development. Another convectivecell has just crossed the radar site. Fifteen minutes later(Figure 12b) reflectivity also increased in the lowerlevels, precipitation at the ground was intensifying and,most likely, so was the descending motions inside thestorm. From the vertical section the cell appears to besloping downstream; this could be due to the presenceof a vertical wind shear but it could also simply be aresult of the movement of the storm during the scaninterval. In fact, the antenna speed was somewhat low,two rounds per minutes, and the time interval between

the lowest (0.5

) and the highest (15

) elevation of theradar scan was about 5 minutes. A rough calculation canbe made to see how this affects the radar observationof the storm. Since the storm motion based on theobserved displacement between subsequent pictureswas estimated at about 40 km/h, the calculated shift ofthe top of the convective cell during the scan interval(5 minutes) should be about 3 km. This is not enoughto explain the observed shift which was about 6 km. Soa real tilting downstream of the cell has to be assumed;the presence of a sheared environment will be clearerfrom the analysis of the radial velocity pictures below.

At 2200 UTC (Figure 12c) a chimney with reflectivityvalues over 50 dbz, and with a maximum value near

the ground of over 60 dbz, developed from the groundto the highest levels. In this phase hail and stormywinds were detected at ground level. After 30 minutes(Figure 12d) the squall line wasalmost crossing the radarsite. Reflectivity values were on average over 55 dbzand reached values over 60 dbz. The strong attenuationof the radar beam behind the line of convectiondue to intense precipitation is very noticeable; here

the radar echo is very low or absent, and the nextpicture will show that precipitation is indeed presentand widespread. Notice also that the convective cellnorth of the radar site has evolved into another linear-shaped convective system. At 2315 UTC (Figure 13a)the storm had crossed the radar site. In this phasethe system was already in the dissipating stage, as isevident from the lower values of reflectivity and thelarge area of precipitation that is almost stratiformbehind the convective region. This area was previouslyhidden by the blocking effect of the convective front. Acomparison can here be made with a conceptual model

of mature stage squall lines proposed by Houze et al.(1989) (Figure 13b). In this model the convective lineformed by new developing cells is followed by the olddissipating cells and by the trailing area of precipitation,which is almoststratiform with lower reflectivity values.In the region of new cell formation an area with lowerreflectivity values is present in the medium-upper levels,spreading out with height. These features are easilyrecognised in the Figure 13a.

A further correspondence could be found by observingradial velocity pictures. The first picture considered is avelocity PPI at 4.5 of elevation (Figure 14b) with the

corresponding reflectivity (Figure 14a). Comparing thisimage with the conceptual model in Figure 13b we cannote the rear inflow descending towards the convectivefront, the inflow ahead of the convective front in thearea of new cell formation, and the outflow in the upperlevels ahead of the storm. An expected front-to-rearupper-level flow is not observed; however, the flowsdepicted in the conceptual model are storm relative,whereas we are observing the velocities with respectto the radar. If we subtract the storm velocity fromthe radar velocities we can find the right values. Notethat the inflow band ahead of the convective region is

broadening and the wind is intensifying southwards,favouring the propagation towards the southeast and ageneral rotation clockwiseof the storm (see thesequencein Figures 12 and 13a). Looking at the orientation ofzero isodop we see that in the pre-storm region thewind rotates clockwise with height, starting from east inthe lower levels up to slightly south-west in the upperlevels. This wind shear is likely to play a key role inthe development of this storm because of the low valueof convective available potential energy (see section 5below). On the other side, behind the convectivefront, the orientation of the zero isodop suggests acounterclockwise rotation of the wind with height, from

west in the lower levels up to just about south-eastin the upper levels. This leads to the convergence of

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Figure 12. (a), (b) and (c) CAPPI reflectivity pictures (dBz) with vertical sections parallel to the storm motion from 19 July2001, 2130 UTC, 2145 UTC and 2200 UTC respectively; (d) CAPPI reflectivity picture from 19 July 2001, 2230 UTC. Colourfigures are available on the internet at http://digilander.iol.it/atrok/metapps/index.html

the flow along the convective region and divergence

in the upper levels, in agreement with the conceptualmodel. The second picture considered is a velocity PPIat 1.3 of elevation (Figure 14d) with its correspondingreflectivity (Figure 14c). The region just ahead of theleading front behind the 30 km circle to the north-east and behind the 60 km circle to the south-east is noticeable. In this region the wind is oriented awayfrom the radar, probably owing to the outflow fromthe low-level divergence in the area with precipitation.This region of outflow corresponds to the gust frontin the conceptual model and plays an important role inoriginating new convective cells ahead of the storm.

During the following night, the storm tended to spreadand decay further, with other scattered thunderstorms

still affecting the region of Veneto until the morning of

20 July, though with decreased intensity.

5. Consideration of vertical stability and itschange over time

During the summer season thunderstorm activity is notunusual in the Padana Valley where the atmosphereis almost always, even if only to a small degree,potentially unstable, including during the hours ofdarkness. Statistics for the occurrence of thunderstormsin the summer season do indeed show a secondarymaximum during the night, when diabatic forcing from

the ground is absent. Thus, when external forcing isin operation, for example in a cold front or a trough

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Figure 13. (a) CAPPI reflectivity picture (dBz) and vertical section parallel to storm motion from 19 July 2001, 2315 UTC;(b) conceptual model of a squall line viewed in a vertical section oriented parallel to the storm motion (Houze et al. 1989). Colourfigures are available on the internet at http://digilander.iol.it/atrok/metapps/index.html

with ascending motion, it is quite likely that convectiveoverturning willoccur(Cacciamani et al. 1995). Lookingat the soundings, some qualitative observations could bemade about the genesis and intensity of the convectivephenomena in this case. The skew-T plot sounding fromUdine on 20 July 2001, 00 UTC (Figure 15a) is the

nearest available sounding being some 150 km north-east of Teolo in Friuli Venezia Giulia region. At thistime the storm was already influencing the Veneto flatplains but not yet Friuli, so this sounding could berepresentative of the environment before the arrival ofthe squall line. The vertical profile looks very stable:a strong thermal inversion is present at ground level,and another inversion layer is present between 800 hPaand 850 hPa. Referring to the moist adiabat from thelifted condensation level, the buoyancy is negative upto 600 hPa; above this there is a layer with smallpositive buoyancy until 400 hPa. This means that takinga particle from the lowest level and forcing its ascent,

its motion is not accelerated before 600 hPa; also, abovethis level, a small amount of energy is available. These

features strongly inhibit any convective vertical motionso in these conditions a trigger of convection, withoutany external forcing, is very unlikely. Some stabilityindices could be derived from the sounding data:

LIFTED INDEX: LIT500

obs

T500

lifted from surface

= 0.96

SHOWALTER INDEX: SI= T500obs T500

lifted from 850

= 2.12

CAPE: = 168.19 J/kg

KO Index: KOI=1

2

500eobs +

700eobs

850eobs

1000eobs

= 1

LI shows a weak latent instability and CAPE a

small convective energy available for developing severethunderstorms. In fact, values of LI less than 5 and of

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Figure 14. (a) PPI reflectivity (dBz) at 4.5 of elevation from 19 July 2001, 23 UTC; (b) PPI wind (m/s) at 4.5 at the samehour, also the winds are sketched accordingly to the orientation of the zero isodop (dashed line); (c) PPI reflectivity (dBz) at 1.3

of elevation from 19 July 2001, 23 UTC; (d) PPI wind (m/s) at 1.3 of elevation at the same hour, also the regions of outflow areshown. Colour figures are available on the internet at http://digilander.iol.it/atrok/metapps/index.html

CAPE over 1500 J/Kg are usually required to assess asignificant probability of severe thunderstorms. The SIvalue indicates a higher probability of thunderstorms,but values less than3 areusually associated with severeconvection (Showalter 1953, Galway 1956).

The vertical profiles of potential temperature and equi-valent potential temperature are shown in Figure 15b.An examination of the vertical profile of equivalentpotential temperature allows us to explore the potentialinstability that is associated with layers in whichequivalent potential temperature decreases with height(Kurz 1998). The atmosphere is seen to be potentiallyunstable between 750 and 650 hPa, with potentiallyunstable layers between 850 and 800 hPa and between500 and 450 hPa. The potential instability between500 and 1000 hPa was calculated in terms of the KOIndex; its value of1 indicates a significant potentialinstability for the whole layer. In fact, a KOI value of

less than 2 is usually required to indicate an appreciabledegree of potential instability in the medium-low

troposphere. Thus, where a suitable external upwardlifting is superimposed, a convective overturning has tobe expected.

To see how this potential instability is released due tothe passage of the baroclinic wave, it is useful to see

how the vertical profiles of temperature and dew-pointtemperature are modified after an upward displacement,for example by 100 hPa (Figure 16). As a rough estimateof this ascending motion we could take the valueestimated by the ECMWF model from the run of19 July 2001, 12 UTC, for 20 July 2001, 00 UTC(+12 h forecast). The model forecasts an area ofstrong ascending motion extending vertically from850 hPa to 500 hPa just over the flat plains of Venetoregion, reaching a maximum value of135 hPa/h at850 hPa.

The modified sounding (Figure 16) was created

graphically, applying the same vertical displacementof 100 hPa (i.e. the same vertical motion) to all levels,

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Figure 15. (a) Skew-T plot from sounding of Udine, 20 July 2001, 00 UTC; (b) vertical profile of potential temperature ( ande) and equivalent potential temperature (e) derived from sounding data.

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Figure 16. Skew-T plot of the modified sounding. It is obtained from the sounding in Figure 15a lifting all the relevant pointsby 100 hPa. Profiles of dew-point temperature, temperature (solid lines) and moist adiabat from 900 hPa (dashed) are shown.

and so must be considered only an approximation ofthe real profile. The lifted values of temperature anddew point were found by following the dry adiabatand the line of constant mixing ratio, respectively,from the initial level until the level 100 hPa higher. Ifsaturation occurred, the changes of temperature anddewpoint above the saturation level were given by the

moist adiabat. As result of these changes the thermalinversion above 850 hPa has been destroyed and theenvironment is almost saturated from about 700 hPaup to the highest levels of the troposphere. Since thelowest level (900 hPa) is also saturated, a particle liftedfrom there will follow the moist adiabat depicted(dashed line). Above 700 hPa, a large area with positivebuoyancy is now present, and upward accelerationhas to be expected up to 250 hPa. Furthermore, dueto the saturation of the surrounding atmosphere,no breaking effects result due to entrainment ofsurrounding air, but only cooling due to the adiabaticexpansion and warming from delivery of latent heat.

As seen in the synoptic description, the result is a netwarming of air, at least in the middle levels. However,

below 700 hPa the environment is not completelysaturated and a layer with negative buoyancy is stillpresent, so more time is needed for a full release ofconvection. This delay could lead to a more intensedevelopment of convective motions because the upperlevels of the troposphere are cooling due to the as-cending motion working aloft. The vertical profile

becomes unstable and saturated from the lowest to thehighest levels of troposphere.

Hence, in this case, looking only at the thermodynamicinstability would have led to an underestimate ofthe thunderstorm potential and intensity. A moresatisfactoryresult is obtained by evaluating the potentialinstability, together with the ascending motionsestimated by the model, to see how the sounding couldbe timely modified. As an estimate of the verticalvelocity we can use the maximum value forecasted at850 hPa over Veneto by the ECMWF model, namely135 hPa/h. With this value, an ascent of 100 hPa needs

about 45 minutes, a time interval that is consistent withthe time scale of the phenomenon, in which the squall

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line crosses the region between 2130 UTC on 19 Julyand 00 UTC on 20 July 2001.

6. Conclusions

In this paper we have described an interesting case ofcyclogenesis over the western Mediterranean Sea thatled to the formation of intense convective overturningabove the flat plains of the Veneto region in northeastItaly, 1821 July 2001. The analysis has been conductedusing data from different sources, such as numericalmodels, vertical soundings, satellite and radar. Thisallowed us to observe the phenomena at the synopticscale as well as the mesoscale. On a synoptic scale thecyclogenesis took place following the classic Petterssenscheme: the interaction between the lower frontal zoneand the upper vorticity maximum led to the formationof a cyclone in which strong ascending motionswere accompanied by the release of great amounts oflatent heat. The cloudiness associated with this process

was clearly identified in the infrared METEOSATsatellite pictures. Extraordinary convective activity wastriggeredby thecyclone over theplains of Veneto, wherean organised convective system a squall line formedduring the evening and the night of 19 July. By analysingradar pictures, good agreements were obtained withthe conceptual model of mature-stage squall linesproposed by Houze et al. (1989). Furthermore, from theanalysis of the radial velocity pictures it was possibleto estimate the significant vertical wind shear in thepre-storm region that was likely to play an important

role in the development of the storm. In addition,the analysis of vertical stability executed on soundingdata from Udine, 20 July 2001, 00 UTC, shows theimportance of the potential instability released duringthe transit of the baroclinic wave. The ascendingmotions associated with the cyclone led to substantialmodification of the vertical profiles of temperature anddew-point temperature, initially very stable, and finally

to an environment conducive to severe convection withlarger convective available energy and high levels ofmoisture. The evaluation of potential instability is hencevery important in estimating thunderstorm potential,especially in the nocturnal hours when heating from theground is virtually zero.

References

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Galway, J. G. (1956) The lifted index as a predictor of latentinstability. Bull Am. Meteorol. Soc. 37: 528529.

Houze, R. A., Jr., Rutledge, S. A., Biggerstaff, M. I. & Smull,B. F. (1989) Interpretation of Doppler weather radardisplays of midlatitude mesoscale convective systems. BullAm. Meteorol. Soc. 70: 608619.

Kurz, M. (1998). Synoptic Meteorology. Deutscher Wet-terdienst, pp. 163170.

Petterssen, S. & Smebye, S. J. (1971) On the development ofextratropical cyclones. Q. J. R. Meteorol. Soc. 97: 457482.

Showalter, A. K. (1953) A stability index for thunderstormsforecasting. Bull. Am. Meteorol. Soc. 34: 250252.

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a case of cyclogenesis

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