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The North African Monsoon System
Monsoon traditionally refers to seasonal reversing winds accompanied by changes in precipitation, this seasonal variation of precipitation leads to both a rainy and a dry phase. Indeed, the word monsoon derives from the Arabic mausim which means season, and according to the glossary of the American Me-teorological Society, the word was first applied to the winds over the Arabian Sea, which changed from northeasterly to southwesterly every six months. Then, the term was applied to other regions in the world (Fig. 1.1).
Figure 1.1: Main monsoon systems in the world: North American (NAM), South American (SAM), North African (NAF), South African (SAF), Indian (IND), East Asian (EAS), Western North Pacific (WNP), Australian (AUS). The monsoon domains are defined where the local dif-ference in precipitation rate between boreal summer (May to September) and winter (November to March) exceeds 2.5 mm day−1, and the local summer precipitation represents more than 55% of the annual. Hatched regions correspond to summer precipitation less than 1 mm day−1. Based on rain gauge and satellite data for land and ocean, respectively. From World Climate Research Programme (WCRP).
In this chapter, the main atmospheric circulations at the continental scale dur-ing the north African monsoon season are described. An overview of the spatial and diurnal variability of precipitation are presented and the physical processes driving moist convection are explained. A description of the type of precipitating systems with a focus on mesoscale convective systems is given. The characteris-tics of dust emission and dust effects on the atmosphere, with a special interest on precipitation, are presented. Finally, the objectives of this thesis are detailed.
Large-scale Atmospheric Circulations
Figure 1.2: The key circulations conforming the North African Monsoon system: the low-level monsoon and warm Saharan flows (Harmattan) converge around the ITF/ITD (Intertropical Front/Discontinuity), the Heat Low (depicted by a L over the Sahara), the middle-level AEJ (African Easterly Jet) and AEWs (African Easterly Waves) and the high-level TEJ (Tropical East-erly Jet) and STJ (Subtropical Jet). Reproduced from the Comet program.
The North African Monsoon is a land-ocean-atmosphere coupled system that presents a wide range of spatial and temporal scales during boreal summer. The main large-scale atmospheric elements interacting over Africa are represented in Fig. 1.2.
During the monsoon season, the southern-hemisphere trade winds reverse and pass through the Equator from the ocean to the continent. A strong merid-ional surface temperature gradient between land and ocean is established. This is due to the increased solar forcing over the continent, that creates a low pres-sure region, often enhanced by a sea surface temperature (SST) decrease in the Equatorial Atlantic. The resulting southwesterly monsoon flow can reach 1 or 2-km depth and transports moisture from the ocean to the dry continent, creating favorable conditions for convection there.
When the northern-hemisphere trade winds cross the Sahara towards the south, they take the name of Harmattan flow. The Harmattan wind is hot and dry. It can extend up to the 5-km altitude and transport dust loads from the desert. The convergence zone where both the monsoon and Harmattan flows encounter over the continent is called the Intertropical Discontinuity (ITD).
To the north of the ITD, the surface of the Sahara warms up during daytime due to intense solar heating. As the temperature of the surface increases, en-ergy is transferred from the surface to the atmosphere in the form of outgoing longwave radiation and sensible heat flux by conduction. Almost no latent heat is released because the soil is very dry. The low levels of the atmosphere warm up there and the pressure diminishes considerably. This is the Heat Low region, where dry convection is very intense and helps mix the Saharan boundary layer, that can reach up to 5-km depth (Cuesta et al., 2009). The Heat Low region presents a cyclonic circulation.
The strong meridional temperature gradient over the Sahelian region, which is maximum at the surface and decreases with height, induces a variation of the zonal wind with altitude. This is the thermal wind equilibrium. As temperature increases from the south to the north of the continent, the zonal component of the wind increases with altitude, generating a strong vertical wind shear between the surface and the altitude where the meridional temperature gradient disap-pears. At that altitude, around 650 hPa, this zonal wind has become maximum with a mean speed of 15 m s−1 and located around 12◦N in July. It is known as the African Easterly Jet (AEJ). It presents a diurnal variability.
To the northern and southern sides of the AEJ synoptic disturbances are ob-served over Central and West Africa, where the AEJ is the most intense. These perturbations are called African Easterly Waves (AEWs). They propagate west-ward at a mean speed of 8 m s−1 and with a wavelength of around 3000 km. They present a frequency of 3-5 days. Their initiation is associated with con-vection over the Darfur Plateau and the Ethiopian Highlands (Mekonnen et al., 2006).
The AEWs present perturbations both at 850 and 700 hPa, that consist on a single dynamic mode propagating westward simultaneously over the African continent (Pytharoulis and Thorncroft, 1999). AEWs have an amplitude maxi-mum is near 700 hPa, close to the AEJ level, and a low-level vorticity maximum near 850 hPa. In general, the composite AEW has a northeast-southwest oriented axis, the trough (Reed et al., 1977). The trough region presents generally two cyclonic vortices at 850 hPa (Fink and Reiner, 2003). The maximum of 850 hPa vorticity, corresponding to the southern vortex, occurs near the 700 hPa center in the region of moist convection. The secondary peak of 850 hPa vorticity, the northern vortex, is associated with the ITD. Convergence (divergence) at 850 hPa occurs ahead (behind) of the trough. Inversely, divergence (convergence) at 700 and 200 hPa occurs ahead (behind) of the trough (Reed et al., 1977). The AEWs are the most intense near the altitude of the AEJ, to its southern flank. Energy transfers between the jet and the waves allow the growth of the latter. Also to the southern flank of the AEJ, the humidity transported by the monsoon flow over the continent is large, favoring the development of convective systems. The latter appear to propagate within the AEJ and to be modulated by the AEWs (Fig. 1.3). An AEW can be viewed as alterning northerly and southerly wind perturbations at 700 hPa, immersed in the strong zonal AEJ flow. The souther-lies transport cooler and moister air whereas dry, hot and dust ladden air masses are transported by the northerlies. The northerlies and the trough regions of the AEW are found to be favorable for convection.
Figure 1.3: Most relevant synoptic and mesoscale systems of the northern African monsoon.
Reproduced from the Comet program.
The intense divergence in the higher troposphere due to convection feeds the Tropical Easterly Jet (TEJ), located at 200 hPa around 10◦N, and extending from the east of the Asian continent to the west coast of Africa.
Precipitation
Spatial Distribution
Precipitation over northern tropical Africa takes place mostly during the North African monsoon season. Since the development of the monsoon in May over the Guinean Coast, that reaches its maturity in August in the Sahel, to its retreat in October, the monsoon season concentrates more than 70% of the annual pre-cipitation in West Africa (CLIVAR, 2015). Over Ethiopia, the precipitation during the summer contributes to between 50 and 95% of the annual totals (e.g., Segele and Lamb, 2005; Korecha and Barnston, 2007).
The average precipitation in the Earth for June 1998-2011 is presented in Fig. 1.4. Over Africa, most of the precipitation lies between the Equator and the Sahara, from Ethiopia to the Atlantic.
Figure 1.4: Mean rainfall for June 1998-2011, from product 3B43 of the Tropical Rain-fall Measuring Mission (TRMM). Reproduced from https://pmm.nasa.gov/trmm/trmm-based-climatology.
The rain belt is colocated with the Intertropical Convergence Zone (ITCZ) over the northern tropical Africa. The beginning of the rainy season corresponds to the ITCZ migration from the Southern Hemisphere in winter to the Northern Hemisphere in summer. Its location is at around 5◦N in May-June and 10◦N in July-August (Sultan and Janicot, 2000). Sultan and Janicot (2003) established the 14 May as the mean climatological date for the monsoon preonset, when the first increase of precipitation in West Africa occurs. This is called the Guinean regime. Guinean coastal rainfall is mainly controlled by SST in the Atlantic ocean (Nguyen et al., 2011). The monsoon onset is set to the climatological value of 24 June (Sultan and Janicot, 2003). It corresponds to a sudden shift of the ITCZ from 5◦N to 10◦N and a second increase in precipitation. These dates are consistent with the 11 May and 26 June for the mean Guinean coastal rainfall onset and retrait, respectively, found by Nguyen et al. (2011). From infrared (IR) satellite imagery, the ITCZ manifests as a band of clouds, usually groups of thunderstorms moving westward within the AEJ. Coincident with the southern flank of the AEJ, the ITCZ presents a line of vorticity around the 850 hPa level, enhancing convergence and the development of convection.
AEWs have a major role on precipitation at intraseasonal scale. Maxima in precipitation is observed ahead of the AEW troughs, coincident with low-level convergence (Reed et al., 1977). Northward bursts of the monsoon every 3-5 days can lead to enhanced moisture available for convection (Couvreux et al., 2010) at the synoptic scale. For instance, a rainfall surplus over Algeria was observed to be associated with the northward penetrations of the monsoon flow forced by the south sector of AEWs (Cuesta et al., 2010).
Diurnal Cycle
During the daytime, the solar radiation warms up the Earth surface and oceans, which heat, in turn, the atmosphere through infrared radiation and conduction at the surface. The air masses close to the surface expand and become less dense than the air aloft, which leads to conditionnaly unstable conditions. When con-vection triggers, it redistributes the excess of energy to higher altitudes through vertical mass transport. In the early morning convection is shallow and in the absence of clouds it is called dry convection. Thermals mix the atmosphere in a layer of around 1-km depth, the atmospheric boundary layer. When water vapor condenses we speak of moist convection. In moist convection, the vapor conden-sation releases latent heat, resulting in extra warming of the air masses, that will pursue their uplift vertically and can give birth to precipitating clouds if other factors are met. For instance, the mechanical uplifting forced by orography, cold pool or gravity wave is usually required, or at least beneficial, for the formation of precipitating systems. The processes of moist convection will be addressed in section 1.2.3.
Precipitation over tropical Africa exhibits thus a marked diurnal cycle mainly driven by solar heating and yielding the generation of precipitating systems usu-ally in the afternoon. Over eastern Africa, convection is triggered in the after-noon along the western slopes of the Ethiopian Highlands (e.g., Laing et al., 2008, 2012). Over West Africa, precipitation typically shows a single diurnal peak (Zhang et al., 2016). The time distribution of the peak of precipitation varies between land and ocean (Fig. 1.5). Over land, the maximum of precipita-tion tends to occur between the late afternoon and night, from 16 to 04 LT (local time). Afternoon peaks are generally associated with topographic features such as the Jos Plateau, Cameroon and Ennedi Mountains. Nocturnal peaks are due to precipitating systems downstream of orography, over Niamey and the northeast-ern part of Benin which is devoid of mountainous regions (e.g., Mathon et al., 2002; Fink et al., 2006; Janiga and Thorncroft, 2014; Zhang et al., 2016). In-deed, the average precipitation diurnal maxima result from the superposition of local maximum, coincident with the triggering of convection, and the delayed arrival of systems propagating from the east. The diurnal peak of oceanic pre-cipitation is attained in night or early morning, between 04 and 10 LT (Fig. 1.5).
Figure 1.5: Time of the maximum in the diurnal cycle of precipitation for (a) TRMM product 3B42 and (c) TRMM PR (Precipitation Radar) July-September (JAS) 1998-2012. Reproduced from Janiga and Thorncroft (2014).
Precipitating Clouds
Precipitation mainly arises from moist convective clouds. Precipitating sys-tems originate as individual convective cells triggered by the development of the boundary layer during daytime. Moist convection is a response to atmospheric instability, which is evaluated with respect to the vertical displacement of moist air parcels. Every air parcel is subject to buoyancy, a force resulting from the density difference between the air parcel and its environment. If a parcel is less dense than the ambient air, it experiences an upward acceleration, it is positive buoyant. If the parcel is denser, it suffers a downward acceleration, it is nega-tively buoyant. For moist convection to happen, the air parcel needs to ascend further up than the altitude where it becomes saturated. At this level, its wa-ter vapor condenses leading to the formation of the first hydrometeors, cloud droplets. A cloud then begins to form.
An emagram like in Fig. 1.6 is very convenient to determine the vertical extent of the cloud and to represent its convective development as an air parcel ascends in the atmosphere.
Figure 1.6: Sounding taken at Savé (Benin) at 1800 UTC on 1 July 2016, during the DACCIWA (Dynamics-Aerosols-Chemistry-Cloud Interactions in West Africa) field campaign. The temper-ature is represented by the black line and the dew point temperature by the dark blue line. The theoretical trajectory of an ascending parcel is indicated by red line. It follows a dry adia-bat (solid) up to the LCL and a moist adiabat from the LCL until the EL. The area between the sounding profile and the parcel trajectory below (above) the LFC corresponds to the CIN (CAPE).
They are divided into cumulus humilis, cumulus congestus and cumulonim-bus, depending on their vertical development, that is given by the instability of the warm moist parcels. The altitude of the cloud top for the cumulus humilis is constraint by the stable layer at the top of the boundary layer. Cumulus humilis are the shallowest, and cloud droplets do not have time to produce enough large rain drops. Cumulus congestus and cumulonibus exhibit a much greater verti-cal extent, and are precipitating. The top of the cumulus congestus is placed at around 5 to 6 km, the altitude of the freezing level. No solid phase is present, and buoyancy comes from the latent heat release of condensation processes. Precipitation is produced by coalescence of liquid droplets. Cumulonimbus or thunderstorms correspond to deep convection. As they develop above the 0◦C isotherm, extra latent heat is released through vapor deposition on ice crystals and the freezing of liquid water. This energy supply is much greater than in con-gestus, allowing the cumulonimbus to grow past the tropopause, at around 17 km. Well-developed cumulonimbus are characterized by an anvil-like top, caused by wind shear or inversion near the tropopause. The anvil is the visible conse-quence of divergence close to the tropopause. Cumulonimbus are mixed-phase clouds (with liquid and iced hydrometeors), and capable of producing lightning. The lifetime of an individual cumulonimbus is around 30 to 50 minutes. These isolated convective clouds can evolve, if the environmental conditions are fa-vorable, into mesoscale convective systems (MCSs), a highly organized type of moist convection. These systems present longer lifetime and a huge potential for intense and lasting precipitation.
Figure 1.8: Mean fraction of (a) shallow convective, (b) deep convective and (c) stratiform precipitation from the TRMM PR during JAS 1998-2012. Reproduced from Janiga and Thorncroft (2014).
In a general way, precipitation is considered to be of two types, stratiform or convective. These types are defined in terms of their vertical velocity scales. Con-vective (stratiform) precipitation is a precipitation process in which the vertical air motion is large (small) as compared to the terminal fall velocity of ice crystals and snow (Houze, 1994). Deep convection, which is very frequent in continen-tal Africa, generates between 50 and 70% of precipitation in the ITCZ region (Fig. 1.8b). There, from the Atlantic to Ethiopia, a significant contribution of 30 to 50% to total precipitation is stratiform-like (Fig. 1.8c). Offshore over the ITCZ, the contribution of deep convective and stratiform precipitation decreases and increases, respectively. Note that stratiform precipitation is closely linked to convective processes. Stratiform clouds are primarily associated with massive clouds such as MCSs, that exhibit a mixture of convective and stratiform struc-tures. Note as well that the fraction of rain attributed to deep convection is due to both isolated clouds or MCSs. Finally, the contribution of shallow convective clouds to precipitation is negligible. It only appears over the coastal regions of the Gulf of Guinea (Fig. 1.8a). These clouds do not produce much precipitation because they lack of vertical extent for rain drops growth.
Mesoscale Convective Systems
Moist convection can be organized into large complexes denominated mesoscale convective systems (MCSs). Houze (1994) defines an MCS as a cloud system that occurs in connection with an ensemble of thunderstorms and produces a contiguous precipitation area ∼100 km or more in horizontal scale in at least one direction. In a MCS, this large precipitation is both of convective and stratiform type, which present horizontal scales of about 10 and 100 km, respectively. It is associ-ated with circulations of different scales as well, such as convective updrafts and downdrafts and mesoscale circulations. Mesoscale convective organization can happen in several forms: squall line (Zipser, 1977), mesoscale convective com-plex (MCC, Maddox (1980)), organized convective system (OCS). Squall lines, which are very typical in West Africa, have been more intensively studied than other types of mesoscale convection.
A squall line is characterized by a front line of convective cells perpendicular to its motion, and a trailing stratiform region (Fig. 1.9).
Figure 1.9: Structure of a tropical squall line represented by a cross section parallel to its motion. The squall line is in its maturity stage of its lifecycle (see text for details). It presents convective cells in the front: new developping ones, mature ones with the strongest updrafts generating overshoots at the tropopause and old ones that eventually become embedded in the trailing stratiform region. The large stratiform region presents mesoscale circulations: an ascending flow from the convective parts and a descending inflow fed by the surrounding air. Both the convective and stratiform regions produce precipitation. Reproduced from Houze et al. (1989).
Fast updrafts and downdrafts are located in the convective region, where the convective cells present different stages of development. New shallow cells (cu-mulus humili and congestus) may appear in the very front of the MCS, followed by older cells, mature ones. The updrafts of a mature cell transport moist air parcels and hydrometeors up to the equilibrium level that corresponds to the cloud top, and sometimes overpass it generating overshoots until their upward velocity is zero. The convective downdrafts contain the negatively buoyant air cooled by evaporation of rain, and when arriving at the surface these cool and dense air masses create cold pools. In their cloud-resolving model study, Diongue et al. (2002) simulated a squall line with an associated temperature drop of 9 K at the surface due to cold pools. Cold pools are found to maintain convection in two ways. On the one hand, they favor the growth of the convective system (Tompkins, 2001), whose updrafts have, in turn, the potential to remain more undiluted, reaching higher altitudes and becoming deeper (Khairoutdinov and Randall, 2006). Thus, cold pools help deepen the front cumuli into strong cu-mulonimbus. On the other hand, the turbulent density currents in the borders of the cold pools act like a cold front forcing the uplift of warm moist air in the near-surface. This mechanism may trigger new convective cells.
Table of contents :
Introduction
Introduction (en français)
1 The North African Monsoon System
1.1 Large-scale Atmospheric Circulations
1.2 Precipitation
1.3 Dust
1.4 Objectives of the Thesis
2 A Dust Outbreak over Northern Africa
2.1 The 9-14 June 2006 AMMA Case Study
2.2 Simulations with the Méso-NH Model
2.3 Observations
2.4 MCS tracking
3 The Westward Increase of Deep Convection Organization in Northern Africa
3.1 Introduction
3.2 Article
4 Radiative Impact of Dust on Mesoscale Convective Systems and their Organization in Northern Africa
4.1 Introduction
4.2 Article
Conclusions and perspectives
Conclusions et perspectives (en français)
References
