The impact of soil moisture modifications on CBL characteristics in West Africa: A case-study from...

QUARTERLY JOURNAL OF THE ROYAL METEOROLOGICAL SOCIETYQ. J. R. Meteorol. Soc. 136(s1): 442–455 (2010)Published online 2 July 2009 in Wiley InterScience(www.interscience.wiley.com) DOI: 10.1002/qj.430

The impact of soil moisture modifications on CBLcharacteristics in West Africa: A case-study from the

AMMA campaign

M. Kohler,* N. Kalthoff and C. KottmeierInstitut fur Meteorologie und Klimaforschung, Forschungszentrum Karlsruhe, Germany

ABSTRACT: Within the framework of the AMMA campaign in 2006, the response of surface properties to precipitationand their effect on the state of the convective boundary layer (CBL) and on convective instability were analysed. Theobservation periods covered the pre-monsoon onset (SOP 1) and the mature monsoon phase (SOP 2) and were performedin southwest Burkina Faso. Precipitation caused a distinctive increase in the volumetric soil moisture content in the upper20 cm of the soil. Coupled with the increase in soil moisture, a continuous decrease of surface and soil temperature withtime was observed. Changes in surface temperature, albedo, and a higher availability of water affected the partitioning ofthe energy balance. Highest values of the Bowen ratio were found during SOP 1 when the surface was dry and vegetationsparse. In SOP 2, a higher vegetation cover made the albedo and Bowen ratio less sensitive to changes in soil moisture.Modifications of surface fluxes influenced the CBL conditions. The CBL height in SOP 1 was 1658 m and in SOP 2 877 m.The heat budget of the CBL was dominated by sensible heat flux convergence, whereas the moisture budget was controlledby both advection and latent heat flux convergence. It was confirmed by the measurements that the diurnal developmentof convective instability was dominated by the CBL evolution and controlled by changes in the mid- or upper troposphereto a minor degree only. Linear correlations were found between the near-surface equivalent potential temperature and bothconvective available potential energy and convection index. Copyright c© 2009 Royal Meteorological Society

KEY WORDS surface properties; energy balance; heat and moisture budget; convective instability

Received 5 December 2008; Revised 13 March 2009; Accepted 2 April 2009

1. Introduction

In the Sudanian (9–11◦N) and Sahelian (12–18◦N)climate zones south of the Sahara desert in West Africa,about 50% and 90%, respectively, of the precipitationis connected to convection and to travelling distur-bances such as mesoscale convective systems (MCSs;Hastenrath, 1994; Mathon et al., 2002). Due to its con-vective nature, the distribution of precipitation is inhomo-geneous and has a high spatial and temporal variability(Taylor et al., 2007). The objective of understanding themechanisms which modify or trigger convection and,hence, precipitation has been the focus of research fordecades. Apart from large-scale synoptic conditions, sur-face–atmosphere feedbacks that are not represented wellby large-scale models (Koster et al., 2004) play a keyrole in the evolution of convection.

To identify and understand the role of surface and soilparameters in West Africa, it is necessary to monitor theresponse of vegetation and soil properties to precipita-tion in different stages of the monsoon. Wet soils orvegetated areas influence, for example, the albedo andevapotranspiration, and hence the energy balance on theEarth’s surface. Different partitioning of the available

∗Correspondence to: M. Kohler, Institut fur Meteorologie undKlimaforschung, Forschungszentrum Karlsruhe, Postfach 3640, 76021Karlsruhe, Germany. E-mail: [email protected]

energy (net radiation minus soil heat flux) to the tur-bulent surface fluxes leads to a modification of the stateof the convective boundary layer (CBL). Inhomogeneousdistributions of soil moisture and vegetation cause spa-tial heterogeneities in CBL evolution, which favour thedevelopment of thermally induced circulations, so-callednon-classical mesoscale circulations (Pielke and Segal,1986; Segal and Arritt, 1992). These circulation sys-tems can provide regions of convergence, often triggeringdeep convection (Pielke, 2001). Modified CBL conditionsinfluence the location of convection (Garrett, 1982) andmay result in either a positive or negative feedback onconvection. On the one hand, convective initiation maybe suppressed when wet soils cause a shallower, moister,and cooler boundary layer (Taylor and Ellis, 2006). Onthe other hand, wet soil may have a positive feedbackon rainfall from MCSs, because wet soils can lead to ahigher convective available potential energy (CAPE) thandry soils (Taylor and Lebel, 1998; Taylor and Ellis, 2006).

Although numerous modelling studies are availableregarding the impact of surface properties, particularlysoil moisture and vegetation, on convection and precipi-tation in semi-arid regions (Chen and Avissar, 1994; Clarket al., 2004; Cheng and Cotton, 2004; Alonge et al.,2007), there still is a lack of observations. However,observations are crucial to improving the understandingof these feedback mechanisms (Taylor et al., 2007) and

Copyright c© 2009 Royal Meteorological Society

SOIL MOISTURE AND CBL CHARACTERISTICS IN WEST AFRICA 443

to help to confirm model results. Sophisticated fieldcampaigns like the Hydrological and Atmospheric PilotExperiment (HAPEX; Goutorbe et al., 1997), JET2000(Thorncroft et al., 2003), and others provided moreinsight into the role of surface–atmosphere feedbacksin West Africa. Based on this knowledge, the AfricanMonsoon Multidisciplinary Analysis (AMMA) campaignwas designed to improve the understanding of the WestAfrican monsoon system and its socio-economic impact(Redelsperger et al., 2006). A comprehensive overview ofthe AMMA observational program and scientific strategyis provided by Lebel et al. (2009).

In view of the geographical size of West Africa, itis very difficult to cover the whole region by observa-tions. Spatial distribution of surface properties is avail-able from satellites, but remote-sensing data often suf-fer from deficits in temporal and spatial resolution. TheAMSR-E (Advanced Microwave Scanning Radiometerfor EOS) instrument, for instance, covers the upper fewcentimetres of the soil only, and the soil moisture signalis restricted by clouds or vegetation cover. The couplingof the CBL state with the heat and moisture fluxes onthe surface are not represented well by satellite informa-tion. There still is a need for in situ measurements. Fullyequipped sites can cover all parameters from soil mois-ture, even in deeper layers of the soil, to surface fluxesand boundary-layer conditions, to atmospheric conditionsof the free troposphere. Within this context, the scope ofthis study is to monitor the soil conditions and the asso-ciated energy balance components on the Earth’s surfaceduring the different phases of monsoon, from pre-onsetto mature monsoon (Sultan and Janicot, 2003), to identifyrelations between the soil and surface properties, and toinvestigate the contribution of energy transformation ofthe underlying surface to the state of the boundary layerand atmospheric stratification that is essential for convec-tion. The paper is structured as follows. Measurements onwhich the study is based are described in section 2. Theinfluence of insolation and rainfall on soil moisture andsurface temperature is presented in sections 3.1 and 3.2,coupling of soil moisture, vegetation, and surface tem-perature with albedo and energy balance in section 3.3,and the influence of the energy balance on CBL con-ditions in section 3.4. Finally, the impact of the CBLconditions on convective instability is analysed (section3.5). The results will be summarised and discussed insection 4.

2. Site and instrumentation

2.1. Location

Within the framework of the AMMA field campaignin summer 2006, energy balance and soil moisturemeasurements were performed near Bontioli,Burkina Faso (10◦53′02.5′′N, 03◦04′04.0′′W, 280 mabove sea level, ASL). Radiosondes were launched atDano (11◦09′45.4′′N, 03◦04′34.2′′W, 350 m ASL) whichis about 40 km north of Bontioli (Figures 1(a) and (b)).

On a temporal scale, the measurements covered thepre-onset of monsoon from 1 to 15 June 2006 (SpecialObserving Period, SOP 1) (Parker et al., 2007) as wellas the mature monsoon from 25 July to 20 August 2006(SOP 2). The different phases of the monsoon were clas-sified according to the study of Janicot and Sultan (2007).

The terrain on the energy balance site was flat withouthills or hollows within a radius of 1 km. The area was ofa dry savannah type with sparse trees and scattered fieldsof millet, maize and cotton. Land cover between thetrees changed from bare soil during SOP 1 to vegetationheights from 30 cm to 100 cm depending on the typeof vegetation and growth at the end of SOP 2. Dueto the spatial homogenous distribution of the leaf areaindex (LAI), around Bontioli (Figure 1(b)) it can beassumed that the measurements are representative of theinvestigation area.

2.2. Energy balance instrumentation and radiosoundings

The energy balance site was equipped with a CM 14 byKipp and Zonen to monitor solar and reflected irradianceand a Schulze net radiometer. Three heat flux platesmade by Middleton served to measure soil heat flux ata depth of 0.02 m. Temperature and relative humiditywere measured by two Vaisala HMP35A instrumentsat heights of 1.7 m and 3.7 m. Surface temperaturewas determined using the infrared thermometer, modelKT 15 by Heimann. Precipitation was measured with arain gauge, type Ombrometer HP by Thiess (1 m). Thepressure was monitored with a barometer, model 270 bySetra, at a height of 0.5 m. Temporal resolution of thesedata was 10 min. If not specified explicitly, measuringheight was 4 m.

Sensible and latent fluxes were observed with anultrasonic anemometer/thermometer (Solent R1012, GillInstruments) and a fast infrared hygrometer (LI 7500open-path, LI-COR), both at a height of 4 m and witha sampling rate of 20 Hz. To calculate turbulent fluxes as30-minute means, the software package TK2 describedby Mauder and Foken (2004) was used.

Additionally, the site was equipped with soil mois-ture sensors. Two different types of sensor were used,eight SISOMOP sensors (Schlaeger et al., 2005), and twotime-domain reflectometry (TDR) probes, each of 1 m inlength. The measurement principle of SISOMOP sensorsis based on frequency-domain reflectometry. The sensorconsists of a transmission line in the feedback loop of aring oscillator. The frequency of the oscillator is influ-enced by the water content of the surrounding material.With an individual calibration function derived from testswith soil probes in the laboratory, the volumetric watercontent of the soil, �vol, was calculated. The sensor hada length of 15 cm. Hence, the water content measurewas an integral along this distance. Soil temperature wasmeasured with an integrated temperature sensor.

To record a soil moisture profile, five SISOMOPs wereinstalled horizontally at depths of 5, 10, 20, 50, and100 cm. Another three probes were deployed verticallynear the surface and covered a vertical layer from

Copyright c© 2009 Royal Meteorological Society Q. J. R. Meteorol. Soc. 136(s1): 442–455 (2010)

444 M. KOHLER ET AL.

Figure 1. (a) Location of the measurement sites Dano and Bontioli, Burkina Faso (rectangle) during the AMMA campaign 2006. The figure isbased on data from the GLOBE Digital Elevation Model (Globe Task Team, 1999). (b) Leaf area index composed at 10 June 2006 from a 1 kmresolution in the vicinity of Dano and Bontioli. The data were obtained from the level-4 MODIS global Leaf Area Index (LAI) V005 product.

5 cm to a depth of 20 cm. The temporal resolutionof the SISOMOP measurements was 40 s. The resultswere aggregated to hourly means. The measurementswith TDR probes were used to check the SISOMOPmeasurements. The checks were performed daily.

At the Dano site, a Graw radiosonde system, DFC-90ground station, was deployed. During SOP 1 and SOP2, 33 and another 61 radiosondes of the type DFM-97,respectively, were launched, most of them due to theappearance of MCSs.

3. Results

3.1. Insolation and precipitation

Surface–atmosphere feedbacks depend on two types ofsoil properties, fixed surface parameters, such as ele-vation, slope, and soil type, and on variable parameterslike water table height, vegetation, and soil moisture(Warner, 2004). Controlled by insolation and rainfall,the variable parameters change on a temporal scale. Toevaluate the soil and surface properties observed, it is

first necessary to analyse the characteristics of insolationand precipitation during both observation periods.

In SOP 1 as well as in SOP 2, precipitation wasconnected with the passage of MCSs, with the exceptionof three rainfall events in SOP 2 that were due tonon-organised deep convection. During SOP 1 and SOP2, 64.2 mm and 145.8 mm of rain, respectively, wereobserved at Bontioli. This was equivalent to 4.28 mmd−1 in SOP 1 and 5.0 mm d−1 in SOP 2. The time spanbetween two precipitation events in both periods wasabout 2 days, with a tendency to more intense rainfallin SOP 2. The highest amount of rainfall was observedon 11 August with 45.8 mm d−1. Observations ofthe TRMM multi-satellite precipitation analysis, TMPA,(Huffman et al., 2007) for the same periods agreedvery well with the observations at Bontioli (Figure 2).TMPA data indicated an inhomogeneous distributionof precipitation in the area. In SOP 1, Bontioli waslocated at the northeastern edge of a wet pattern (9◦N,10◦W to 12◦N, 2◦W) and north of a second areawith high precipitation amounts. In SOP 2, precipitationextended to the north, but the precipitation pattern in thesouthwest still remained. Different precipitation amounts



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Figure 2. Accumulated precipitation (colour shading) from TRMM multi-satellite precipitation analysis (TMPA) (a) between 1 June and 15 June(SOP 1) and (b) between 25 July and 20 August (SOP 2) and precipitation measurements at Bontioli (red dot). Precipitation is normalised to a

15-day period.

in both periods were normalised to a 15-day period forcomparison.

The trend of solar irradiance, G, indicated a decreasewith time in both periods from mean daily values of264 W m−2 in the beginning of SOP 1 to 227 W m−2 atthe end of SOP 1 and from 231 W m−2 at the beginning ofSOP 2 to 186 W m−2 at the end of the observation period(Figure 3(a)). The reduction of insolation was strongestfrom late morning to early afternoon (Figure 3(b)), whichwas due to a higher number of cloudy days in SOP 2.

3.2. Soil moisture, soil temperature, and surfacetemperature

The effect of the onset of precipitation on the soilproperties is investigated next. The coupling of thevolumetric soil moisture content and soil temperaturewith precipitation is obvious from Figure 4. During thepassage of MCSs in SOP 1 and SOP 2, an increaseof soil moisture was observed in the upper 20 cm ofthe soil (Figure 4(a)). Soil moisture content decreasedafter precipitation (Figure 4(b)) due to evapotranspiration

and lateral and vertical drainage. As a result of thefrequent rainfall in both periods, however, soil moistureincreased with time. Lower layers were almost unaffectedby single precipitation events and reacted more slowly.Nevertheless, a continuous increase of soil moisture wasobserved at 1 m depth from the beginning of SOP 1 tothe end of SOP 2.

Analysis of soil samples taken from the energy balancesite confirmed that the soil on the site was dividedroughly into two different layers. A top layer, which hada thickness of about 25 cm, was classified as loamy sandwith a porosity of 35%. In deeper layers solid laterite, aproduct of chemical decomposition (Ultisol), with lowpermeability prevailed. For these layers, a porosity ofabout 20% to 25% was determined.

Soil temperature and surface temperature decreasedwith time due to two different effects (Figure 4(c)).Firstly, increasing soil moisture led to increased thermalcapacity and conductivity of the soil. Analyses of soilsamples showed that the conductivity changed from0.62 W m−1 K−1 under dry conditions to 3.7 W m−1 K−1

in saturated soil. Secondly, after the rainfall there was a



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Figure 3. (a) Mean daily solar irradiance, G, at Bontioli during SOP 1 and SOP 2 and (b) mean diurnal cycle of G during SOP 1 compared toSOP 2.

decrease in radiation due to cloudiness and an increasein evapotranspiration. In SOP 1, when the vegetationwas sparse, these effects worked directly on the soilsurface. In SOP 2 when vegetation was growing, soiltemperature decreased further, because in addition tothe impact of soil moisture on temperature, the soilwas shaded by the foliage and, additionally, much ofthe absorbed radiation is used to evaporate water. Theimpact of the vegetation cover on the soil temperaturewas clearly obvious in a ‘dry period’ (1 to 10 August) inSOP 2. During this period �vol decreased continuouslyin the upper soil layers and reached the lowest value ofboth SOPs (Figure 4(a)). Therefore, without vegetationcover, a similar soil temperature as at the beginning ofSOP 1 could be expected. However, the mean daily soiltemperature at 5 cm depth increased by about 3◦C from29◦C to 32◦C only and was considerably lower than in thebeginning of SOP 1 when values of 35◦C were common.

3.3. Albedo and energy balance

The increase of soil moisture and the decrease of sur-face temperature as well as the reduced solar irradianceaffected net radiation and the partitioning of the availableenergy into turbulent fluxes. The wet surface after precip-itation caused a reduction of the albedo (Figure 5(a)). InSOP 1, a decrease of the mean albedo from 0.218 in thebeginning of the observation period to 0.18 at the end ofthe period was observed, whereas in SOP 2 no trend inthe albedo was found. In both periods a strong correlation

existed between the soil moisture content at a depth of5 cm and albedo (Figure 5(c)). In SOP 2, however, thesensitivity of albedo to soil moisture modifications wasless pronounced than in SOP 1. This was explained bythe increasing influence of vegetation cover on the albedo.Vegetation has often a lower albedo than bare soil and, inthe case of drying, the albedo of the vegetation preventsan increase of the albedo to the high values observed inSOP 1.

Compared to SOP 1, net radiation in SOP 2 wasreduced by 13.5 W m−2 on average (averages were cal-culated from daily values between 0900 and 1700 UTC)and by 54.8 W m−2 at peak values (Table I). The lossof net radiation was caused by a loss of solar irradiance,which was only partly compensated by a lower outgoinglong-wave radiation and a lower reflected irradiance. Theformer was related to lower surface temperatures, thelatter mainly to a reduction of the albedo by the foliage.

The mean soil heat flux was reduced from SOP 1 toSOP 2 by 54.6 W m−2 and by 76.0 W m−2 at noon onaverage. Mean sensible heat flux decreased by 13.8 Wm−2 and mean latent heat flux increased by 23.4 W m−2.Evapotranspiration in SOP 1 was mainly governed bythe water content of the first few centimetres of the soilat least. In contrast to this, evapotranspiration in SOP 2mainly depended on vegetation, because a bigger reser-voir of water was made available for evapotranspirationas a result of the depths reached by the plant roots. Addi-tionally, plants effectively transpire until they reach theirwilting point. Consequently, the day-to-day variability of



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Figure 4. (a) Daily means of volumetric water content, �vol, at different depths (5, 10, 20, 50 and 100 cm) at Bontioli during SOP 1, pre-onsetof monsoon (left) and SOP 2, mature monsoon (right). (b) Accumulated daily precipitation, at Bontioli during SOP 1 (left) and SOP 2 (right).(c) Daily means of surface temperature, TSUR, and soil temperatures at the same five depths at Bontioli during SOP 1 (left) and SOP 2 (right).

evapotranspiration (Figure 5(a)) in SOP 2 was reducedcompared to SOP 1.

Evapotranspiration increased strongly after precipita-tion in combination with the passages of MCSs (Fig-ure 5(a)). Maximum daily sums of evapotranspiration of4.8 mm d−1 (SOP 1) and of 4.0 mm d−1 (SOP 2) wereobserved. The total amount of evapotranspiration in SOP1 was found to be 38 mm which was equivalent to 59%of the precipitation observed. For SOP 2, a similar ratioof 53% (77.3 mm) was determined. Hence, more than50% of the water input was fed back into the atmospherein both periods, while the remainder caused the increaseof soil moisture at lower levels from SOP 1 to SOP 2(Figure 4(a)). A specific feature was observed during the‘dry phase’ (1 to 10 August) in SOP 2. At the beginningof this period, evapotranspiration decreased and reachedsimilar values as in the beginning of SOP 1 (1.4 mm d−1).When Bontioli was hit by rainfall from non-organisedconvection, evapotranspiration increased up to 3 mm d−1,which is in the same range as the monitored precipitation(3.1 mm d−1, Figure 4(b)). Consequently, no precipitationsignal was observed in the soil moisture content (Fig-ure 4(a)). Eventually, the volumetric soil moisture content

observed during the ‘dry phase’ was the lowest of bothSOPs.

At the beginning of SOP 1 when vegetation was sparseand soil moisture was low, a high sensible heat flux, H0,and a small latent heat flux, E0, resulted in a maximumvalue of the Bowen ratio of 1.6 (Figure 5(b)). After thefirst precipitation events, the Bowen ratio was reduced to0.78 on dry days and 0.27 after rainfall. The evaporativefraction, EF, which is a indicator to quantify the changefrom dry (<0.5) to wet (>0.5) seasons, and is defined as

EF = E0

E0 + H0,

was about 0.39 at the beginning of SOP 1 and alreadyreached a value of about 0.78 at the end of SOP 1. InSOP 2 the Bowen ratio ranged from 0.28 to 0.89 and EFfrom 0.78 to 0.55. The maximum of the Bowen ratio andminimum of EF in SOP 2 was reached on 10 Augustat the end of the ‘dry’ period. Then, soil moisture waslowest, but transpiration from vegetation prevented highvalues of the Bowen ratio similar to those at the beginningof SOP 1 from being reached. The impact of vegetationon the Bowen ratio and on surface temperature can also be



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Figure 5. (a) Daily means (0900–1700 UTC) of albedo and daily sums of evapotranspiration, E0/L, at Bontioli during SOP 1, pre-onset ofmonsoon (left) and SOP 2, mature monsoon (right). (b) Daily means (0900–1700 UTC) of Bowen ratio, β, and daily accumulated precipitationat Bontioli during SOP 1 (left) and SOP 2 (right). (c) Albedo as a function of volumetric water content, �vol, (left), β as a function of �vol

(middle), and surface temperature, TSUR, as a function of �vol (right) at Bontioli during SOP 1 and SOP 2.

seen from Figure 5(c). Both quantities are less sensitiveto changes of �vol in SOP 2 than in SOP 1. Analysisof the dependence of the turbulent fluxes and surfacetemperature on the soil and surface properties revealedthat, in contrast to pre-monsoon onset, the vegetationcover had a compensating effect so that the day-to-dayvariability was significantly reduced.

3.4. Boundary-layer conditions

After the impact of soil properties and vegetation coveron the modification of surfaces fluxes had been identified,the next step was to study how variations of the energybalance influenced the boundary-layer conditions. Theevolution of the CBL during the different states ofthe West African monsoon is displayed in Figure 6,where mean morning profiles of potential temperature, θ ,specific humidity, q, wind speed, vh, and wind direction,wd, are compared to mean afternoon profiles. Meanprofiles included radiosoundings performed between 0600UTC and 0900 UTC and between 1200 UTC and 1700UTC, respectively. Because of different launching times,the mean morning sounding in SOP 1 was at 0700 UTC

and the afternoon sounding was at 1400 UTC. In SOP2 the correspondent profiles were at 0900 UTC andat 1300 UTC. In both periods, three different layerswere identified. The lowest layer was the CBL, wheretemperature exhibited a strong diurnal variation. Whenthe top of the CBL, zi, was defined by the height upto which a strong diurnal variation of θ was observed(Melgarejo and Deardorff, 1974), in SOP 1 a mean CBLheight of zi = 1658 m above ground level (AGL) wasdetermined (Figure 6(a)). In the morning a low-level jetwas established at typical heights of 600 m AGL. Thewind maximum at low levels vanished due to a strongervertical exchange of momentum during daytime. Abovethe CBL, an entrainment layer (1658–2600 m AGL)was found, in which diurnal variations of θ and q wereobserved. The top of the monsoon layer (ML) and theinterface to the Saharan air layer (SAL) are reflected bythe 313 K threshold of the virtual potential temperature(Parker et al., 2005). In SOP 1 it was located at 1870 mAGL. Hence, the entrainment layer included the upperpart of the monsoon layer (1658–1870 m AGL) as wellas lower parts of the SAL (1870–2600 m AGL). As isobvious from the specific humidity profile, dry air was



Table I. Comparison of mean radiation and energy balance terms (all W m−2) at Bontioli, Burkina Faso in SOP 1 and SOP 2.Mean and maximum values are calculated from daily periods between 0900 and 1700 UTC.

SOP 1 SOP 2 SOP 1 SOP 2Variable mean mean SOP 2 – SOP 1 max max SOP 2 – SOP1

Net radiation 424.0 410.5 −13.5 627.6 572.8 −54.8Solar irradiance 616.4 533.1 −83.3 879.2 728.2 −151.0Long-wave incoming radiation 471.8 479.4 7.7 492.4 495.7 3.47Long-wave outgoing radiation 542.8 505.2 −37.6 570.8 524.8 −46.0Reflected irradiance 121.4 96.9 −24.5 173.1 126.3 −46.8Soil heat flux 103.6 49.0 −54.6 158.1 82.1 −76.0Sensible heat flux 104.6 90.8 −13.8 160.8 128.5 −32.3Latent heat flux 184.6 208.0 23.4 230.8 267.3 36.5

entrained from the SAL deep down (to about 500 mAGL) into the mixed layer. The entrainment layer wasassociated with a wind shear from southerlies in themonsoon layer to easterlies in the SAL. In the upperparts of the SAL (>2600 m AGL), no essential diurnalchange of temperature and specific humidity occurred. InSOP 2 a mean boundary-layer height of zi = 877 m AGLwas observed (Figure 6(b)). In SOP 2 much lower CBLheights than in SOP 1 were also observed close to Niameyby Said et al.(2009) and Kalapureddy et al.(2009). Themonsoon layer reached 2370 m AGL and the entrainmentlayer was restricted to 1400 m AGL. In contrast to SOP1, where the CBL nearly reached the top of the ML andthe entrainment layer exceeded the ML and interferedwith the SAL, the CBL in SOP 2 was significantly lowerthan the ML top and the entrainment layer was embeddedcompletely in the ML. Consequently, no exchange ofheat and moisture with the SAL occurred. In both SOPs,specific humidity increased mainly at the top of the CBLand above. In the lower part of the CBL (of depth about50 hPa), specific humidity did not change significantlywithin a day. This means that diurnal changes in CAPEand in the energetic imbalance between the CBL and themid-troposphere were mainly related to a change of θ inthe CBL only. Intrusions of dry air from the overlyingSAL down to at least 0.6 to 0.4zi were also observedin aircraft measurements performed during AMMA inthe vicinity of Niamey/Niger (Canut et al., 2009). Theseentrainment processes caused large moisture fluxes in theupper part of the CBL. Canut et al.(2009) also foundstronger entrainment during SOP 1 when the CBL topusually reached the SAL.

To investigate the evolution of the CBL heights inSOP 1 and SOP 2, radiosoundings from Dano between1200 UTC and 1700 UTC were analysed with respectto maximum daily CBL heights and were correlatedwith mean surface sensible heat fluxes (Figure 7). Meansensible heat fluxes were calculated from sunrise to thelaunching time of the particular radiosonde. In SOP 1 alinear correlation between mean daily sensible heat fluxand CBL height was found, while in SOP 2 the CBLheight was nearly independent of the mean daily sensibleheat flux. Averaged over SOP 1, a mean sensible heat fluxof 104.8 W m−2 was linked with a mean CBL height of

1658 m AGL, whereas averaged over SOP 2 the reducedsensible heat flux of 83.7 W m−2 resulted in a meanCBL height of 877 m AGL. Thus, during the maturemonsoon phase, the boundary-layer height is lower due tothe consumption of available energy in evapotranspirationinstead of CBL heating.

The impact of sensible and latent heat fluxes on theevolution of the CBL can be estimated by calculating theheat and moisture budgets. The budgets can be written as(e.g. Stull, 1988)

∂θ

∂t= − 1

cp ρa

∂H

∂z− v · ∇θ

Lcθ CoH Advθ (1)

and

∂q

∂t= − 1

L ρa

∂E

∂z− v · ∇q.

Lcq CoE Advq (2)

where L denotes the latent heat of evaporation, ρa theair density, v the wind vector, and cp the heat capacityat constant pressure. In Equation (1), Lcθ denotes thenet local change of temperature, CoH is the tempera-ture change caused by sensible heat flux convergenceand Advθ includes the contributions from advection.Diabatic processes from radiation and phase changesare neglected. Equation (2) includes the correspondingcontributions of the moisture budget. Integration ofEquation (1) from the surface to zi and using the meanvalue theorem of calculus yields

∂θ

∂t= − 1

zi cp ρa

[H0 − H(zi)] − v · ∇θ, (3)

where H(zi) denotes the sensible heat flux at the top ofthe CBL. The fluxes at CBL top were calculated, takingentrainment processes into account, according to

H(zi) = γH0 (4)

and

E(zi) = γEE0. (5)



Θm

Θd

qm

qd

vhm

vhd

wdm

wdd

0

500

1000

1500

2000

2500

3000

3500

4000

4500

5000

altit

ude

in m

AG

L

0

500

1000

1500

2000

2500

3000

3500

4000

4500

5000

altit

ude

in m

AG

Lal

titud

e in

m A

GL

0

500

1000

1500

2000

2500

3000

3500

4000

4500

5000

0

500

1000

1500

2000

2500

3000

3500

4000

4500

5000

altit

ude

in m

AG

L

0 5 10 15

q in g kg−1

0 5 10 15

q in g kg−1

Θm

Θd

qm

qd

300 310 320 330

Θ (K)

Θ (K)

0 5 10 15 20

vh in m s−1

300 310 320 330 0 5 10 15 20

vh in m s−1

0 90 180 270 360

wind direction in degree

0 90 180 270 360

wind direction in degree

vhm

vhd

wdm

wdd

(a)

(b)

Figure 6. Comparison of the mean evolution of CBL and lower troposphere at Dano, Burkina Faso, during (a) SOP 1 and (b) SOP 2. The stateof the atmosphere is represented by mean profiles of potential temperature, θ , and specific humidity, q (left), wind speed, vh, and wind direction,wd, (right). Subscripts m and d denote morning and daytime profiles, respectively. The bold solid horizontal line represents the mean CBL top,the thin solid line marks the top of the entrainment layer and the dashed line indicates the top of the monsoon layer and bottom of the Saharan

air layer.

γE = γβL

cp

�q

�θ

is the ratio of the vertical difference of specific humidityand temperature in the entrainment zone, γ is the entrain-ment factor for the sensible heat flux, and ß denotes theBowen ratio. For a typical γ value of −0.2 (Stull, 1988;Said et al., 2009) and with a mean ß = 0.74 for SOP 1 andß = 0.49 for SOP 2 an entrainment factor for latent heat,γE, of 0.5 and 0.2 was determined for SOP 1 and SOP 2,

respectively. A positive value of γE corresponds to a pos-itive latent heat flux at the boundary-layer top, which wasalso observed by Canut et al.(2009) during the AMMAcampaign further to the north (aircraft measurementsbetween 13 and 14◦N). In this dry region close to Niameyin SOP 1, the latent heat flux even increased with height.

With these entrainment factors, the heat and moistureflux convergences were calculated every hour, basedon the surface fluxes shown in Figure 8 and on theassumption that zi increases linearly with time between



0

500

1000

1500

2000

2500

3000

0 20 40 60 80 100 120 140 160

H0 in W m−2

CB

L h

eig

ht

in m

AG

L

Figure 7. Relation between mean daily (0600 UTC to time of radiosonde ascent) surface sensible heat flux, H0 and CBL height in SOP 1 (dotsand dashed line) and SOP 2 (crosses and solid line).

the morning and afternoon values. In the next step theheat and moisture flux convergences were averaged overthe period from morning to afternoon. The net localchanges of heat and moisture were calculated from themorning and afternoon profiles. The advection termsresulted as residues. The heat and moisture budgets werecalculated for the CBL and entrainment layers separately.The turbulent fluxes at the top of the entrainment zonewere set to zero. The results for both monsoon phasesare given in Table II. In SOP 1 the observed net localchange of θ was caused entirely by the convergenceof the sensible heat flux. The local change of moisture,however, was controlled in equal parts by latent heat fluxconvergence and by advection of moist air in the monsoonflow. The increase of moisture in the entrainment layerresulted mainly from turbulent mixing from below ratherthan by advection. In SOP 2 the warming of the CBLwas controlled again by the contribution of the sensibleheat flux, whereas cooling of the entrainment layer wasreduced by the advection of warm air. In contrast toSOP 1, the local increase of moisture in the CBL due tothe latent heat flux convergence in SOP 2 was reduced bythe advection of dry air. As the wind direction is from thesouth, dry air advection may have been associated withthe progress of precipitation further to the north and, atthe same time, with a drop of precipitation amounts inthe south (Figure 2).

3.5. Impact of the boundary-layer conditions onconditional and convective instability

Work was then dedicated to analysing the extent to whichthe boundary layer conditions and evolution influencedthe conditional and convective instability over the Sahel.This means that it was to be found out whether the diurnalevolution of convective indices depended on processes inthe CBL or on processes in the free troposphere. Twodifferent indices were used. CAPE describes the positivebuoyancy of an air parcel (conditional instability) and is

defined as

CAPE =zEL∫

zLFC

g{T vpar(z) − T venv(z)}

T venv(z)dz, (6)

where T vpar(z) is the virtual temperature of an air parcelraised dry-adiabatically from the surface to the levelof free convection, LFC, and moist-adiabatically furtherupwards. T venv(z) is the virtual temperature profile of theenvironment, g the acceleration due to gravity, zLFC theheight of the LFC, and zEL the height of the equilibriumlevel, EL.

The convection index, KO, quantifies the convectiveinstability. KO describes the energetic imbalance betweenthe boundary layer and mid-troposphere and is applied bythe German Weather Service

KO = θe,500 + θe,700

2− θe,850 + θe,1000

2, (7)

with the equivalent potential temperature, θe, at the 500,700, 850, and 1000 hPa pressure levels. Instead of the1000 hPa pressure level, a mean value of the lower 25 hPaof the boundary layer was used here. KO >6 K indicatea weak thunderstorm potential, values in the range of 2to 6 K a moderate potential, and thunderstorms are verylikely for KO <2 K.

From two consecutive radiosonde profiles at time t1 inthe morning and t2 in the afternoon, the evolutions ofCAPE and the KO index were determined. In order toanalyse the impact of the evolution of the CBL on theevolution of the convective indices, an expected CAPEvalue was calculated at t2, which was based on the heatand moisture change of the CBL between t1 and t2 only,CAPEmod. This means that the environmental profile att1 was kept constant, but the parcel was lifted with thesurface conditions at time t2. To calculate the conditionsof the lifted parcel on the surface, mean temperature andhumidity values of the lower CBL (50 hPa) were used.



−100

0

100

200

300

400

500

600

700

fluxe

s in

W m

−2

00:00 02:00 04:00 06:00 08:00 10:00 12:00 14:00 16:00 18:00 20:00 22:00

UTC

Q0,SOP1

B0,SOP1

H0,SOP1

E0,SOP1

Q0,SOP2

B0,SOP2

H0,SOP2

E0,SOP2

Figure 8. Mean diurnal cycle of the energy balance terms during SOP 1 and SOP 2. Q0 denotes the net radiation, B0 the soil heat flux, H0 thesensible heat flux and E0 the latent heat flux at Bontioli, Burkina Faso.

Table II. Heat and moisture budget of the CBL. Lcθ and Lcq denote the local change of heat and moisture in the CBL. CoHand CoE indicate the contribution of the sensible and latent heat flux convergence. Advθ and Advq are the residues of the heat

and moisture budget and z indicates the depths of the CBL and entrainment layer, EL.

Lcq CoE Advq Lcθ CoH Advθ

z (m) (g kg−1 h−1) (g kg−1 h−1) (g kg−1 h−1) (K h−1) (K h−1) (K h−1)

SOP1CBL 1658 0.20 0.10 0.10 0.41 0.41 0.00EL 942 0.23 0.20 0.03 −0.10 −0.13 0.03Total 2600 0.21 0.14 0.07 0.23 0.22 0.01

SOP2CBL 877 0.09 0.34 −0.25 0.55 0.63 −0.08EL 523 0.24 0.19 0.05 −0.11 −0.24 0.13Total 1400 0.14 0.30 −0.16 0.31 0.36 −0.06

0

500

1000

1500

2000

2500

3000

CA

PE

mod

in J

kg−1

0 50 1000 1500 2000 2500 3000

CAPE in J kg−1

−35

−30

−25

−20

−15

−10

−5

0

KO

mod

in K

−35 −30 −25 −20 −15 −10 −5 0

KO in K

(a) (b)

Figure 9. (a) Comparison between measured CAPE and CAPE values based on CBL modifications only, CAPEmod. (b) The same for KO values.Solid lines indicate the 1:1 ratio.



Compared to the CAPE values observed, there was astrong correlation between both values. This implied thatCAPE production during the day was primarily due to theboundary-layer evolution (Figure 9(a)). This estimationexplained 90% of the CAPE production in the major-ity of the cases studied. Only three cases were found,where CAPE production was connected to upper-levelmodifications of the troposphere. Similar results wereobtained for the energetic imbalance expressed by theKO index (Figure 9(b)). Based on the knowledge of thisclose link between CBL conditions and convective andconditional instability, it was argued that changes of con-vective indices like CAPE and KO were connected tochanges of θe in the lower CBL (25 hPa). An evaluationof all soundings available at Dano between 0900 UTCand 2000 UTC in both periods showed a linear correlationbetween CAPE and θe, (Figure 10) which confirms stud-ies by Williams and Renno (1993). In addition, KO andθe were found to be correlated linearly. The sensitivity ofCAPE to modifications of θe was found to be 159 J kg−1

K−1 which is in good agreement with values of 180 Jkg−1 K−1 reported by Parker (2002). The higher sensitiv-ity determined by Parker can be explained by the fact thatthe parcels used in his study had been released on the sur-face, while the parcels studied here had been released inthe lower 50 hPa of the CBL. A threshold value of 344 Kwas determined for zero-CAPE values, which is in goodagreement with thresholds between 347 and 350 K foundby Betts and Ridgeway (1989). The KO index thresh-old value was found at 339 K. Lower values of θe thenimplied positive values of KO, i.e. stable stratification.Low CAPE values and a positive KO index occurredafter the passage of MCSs. Based on this linear correla-tion and observations of temperature and moisture whichare straightforward to obtain near the surface, it would bepossible to predict the potential of a convective environ-ment favouring thunderstorms and, hence, precipitation.

4. Summary and discussion

Based on observations made within the framework of the2006 AMMA campaign at Dano and Bontioli, Burkina

Faso, the impacts of soil properties and vegetation coveron surfaces fluxes and, hence, boundary-layer condi-tions and convective instability were studied. The in situmeasurements covered the pre-onset phase of monsoon aswell as the mature monsoon phase. In both periods, pre-cipitation reached similar mean daily values and rainfallevents occurred with a mean frequency of 2 days. Inso-lation decreased with time due to an increasing numberof cloudy days while the monsoon progressed. Rainfallimmediately affected the soil moisture in the upper 20 cmof the soil. Lower layers (50–100 cm) reacted muchmore slowly. Increasing soil moisture and growing veg-etation reduced soil and surface temperature. Decreaseof albedo and surface temperature meant a decrease inupward radiative fluxes, which partly compensated forthe reduction of incoming solar radiation. Modificationof soil and surface properties affected the partitioning ofthe available energy. Bowen ratio was highest (1.6) atthe beginning of SOP 1 when soil moisture was low andvegetation still sparse. With the onset of precipitation andwith the increasing influence of vegetation, Bowen ratiodecreased to mean values of 0.49 and the evaporativefraction became greater than 0.5. In SOP 2 vegetationrather than soil moisture became the crucial factor forprocesses interacting with the boundary layer. This wasobvious during a more or less dry phase in SOP 2, when,despite some rainfall, soil moisture reached the lowestvalues observed at Bontioli during the AMMA campaign.The amount of rainfall in this period was less than theevapotranspiration rate. Hence, precipitation could notfeed the soil moisture. Even then, however, the Bowenratio did not reach high values similar to those at thebeginning of SOP 1 when vegetation was sparse.

In SOP 1, strong sensible heat fluxes and high Bowenratios resulted in a warmer and deeper CBL than in SOP 2under conditions with higher latent heat fluxes and lowerBowen ratio. The state of the CBL was not only affectedby turbulent flux convergence; entrainment processes andadvection also determined CBL conditions. In SOP 1, theCBL reached the height of the monsoon layer top. Verti-cal exchange between the SAL and CBL occurred whichled to a strong increase of moisture in the lower SAL

0

250

500

750

1000

1250

1500

1750

2000

2250

2500

CA

PE

in J

kg−1

330 335 340 345 350 355 360

Θe (K) Θe (K)

330 335 340 345 350 355 360−25

−20

−15

−10

−5

0

5

10

KO

in K

(a) (b)

Figure 10. (a) CAPE and (b) KO as functions of the equivalent potential temperature, θe. θe is averaged over the lower part of the boundarylayer (25 hPa).



and decrease of moisture in the CBL. This is in goodagreement with observations in the area of Niamey dur-ing pre-monsoon onset, reported by Canut et al.(2009). InSOP 2, by contrast, the CBL did not exceed the monsoonlayer top and the SAL was not involved in entrainmentprocesses. Calculation of the heat and moisture budget ofthe CBL allowed for an estimation of the extent to whichthe CBL state was determined by turbulent flux conver-gence. This estimation showed that the heat budget wasdominated by sensible heat flux convergence and that themoisture budget was determined equally by latent heatflux convergence and advection processes. A comparisonof the observed convection indices CAPE and KO withthe values expected on the basis of heat and moisturechange in the CBL indicated that the diurnal evolutionsof the conditional and the convective instability weredetermined mainly by processes in the CBL. Analysisof radiosoundings at Dano revealed a linear correlationbetween the near-surface equivalent potential temperatureand convective and conditional instability, respectively.However, the technical expenditure associated with thelaunch of radiosondes to determine convective indicesis rather high. Future work should therefore focus onthe above correlation, because the near-surface equiva-lent potential temperature is easy to obtain by surfacemeasurements or aeroplanes and makes it easy to deter-mine the spatial distribution of convective indices. TheKO index which is normally used for midlatitudes wasalso found to be a good indicator of the energetic imbal-ance in the Tropics. For this purpose, however, the mid-latitude thresholds for the likelihood of the developmentof thunderstorms have to be adapted to the tropical envi-ronment. Finally, further work will focus on the use of thedataset for the validation of soil–vegetation–atmospheretransfer (SVAT) models. Only a few of such compre-hensive datasets are available which include soil, energybalance and boundary-layer data from different phases ofthe African monsoon.

Acknowledgements

This work is part of the AMMA project. Based on aFrench initiative, AMMA was established by an inter-national scientific group and is currently being fundedby a large number of agencies, especially from France,the UK, the USA and Africa. AMMA was the benefi-ciary of a major financial contribution under the EuropeanCommunity’s Sixth Framework Research Programme.Detailed information on the scientific co-ordination andfunding is available on the AMMA International website http://www.amma-international.org Figure 2 of thepresent paper was acquired from the GES-DISC Inter-active Online Visualization ANd aNalysis Infrastructure(Giovanni) as part of NASA’s Goddard Earth Sciences(GES) Data and Information Services Center (DISC).

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