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West Africa Meteorology Lecture notes2


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This is an introduction to West Africa Meteorology

This is an introduction to West Africa Meteorology

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  • 2. 2.8 West African Rainfall Distribution• Fig. 16 shows the mean annual rainfall distribution over West Africa (adapted from Nicholson, 2002).• Rainfall distribution over West Africa is dependent on the Inter-Tropical Discontinuity (ITD). Rainfall generally decreases with latitude and with essentially zonal isohyets.• Fig. 17 gives the inter-annual variation of standardized rainfall index S, over the Sahel (top) and Niger south of 15°N (bottom). S = (R – Rm)/σ where Rm is mean rainfall and σ is the standard deviation at the station concerned (after Le barbe et al, 2002).• The figure shows a progressive decline in annual rainfall from the late 1960s.• Fig. 18 shows the composite seasonal rainfall cycles versus ITD latitude – top is long-term average versus wet composite while bottom is long-term average versus dry composite (after Lamb and Abebe, 2008).• Precipitation variability over West Africa is generally studied by comparing a group of wet and dry years.• There is no appreciable rain for about 150km south of the ITD.
  • 3. Fig. 16:
  • 4. Fig. 20:
  • 5. Fig. 18: Composite seasonal rainfall cycles versus ITD latitude (from Lamb andAbebe, 2008)
  • 6. Fig. 19: Location of the AEJ and maximum rainfall beltrelative to the ITD over West Africa for wet years (1958-1962,top) and dry years (1968-1972, bottom) from Grist andNicholson (2001)
  • 7. • Fig. 19 gives the location of the AEJ and maximum rainfall belt relative to the ITD over West Africa for wet years (1958-1962, top) and dry years (1968-1972) from Grist and Nicholson (2001)• AEJ and maximum rainfall belt are closer to the ITD during wet years but farther away from the ITD during dry years. Advance of rainfall maximum and rain event behind the ITD are not a continuous one.• Fig. 20 shows (a) variation of the number of rain events (MCSs) per day and daily rainfall at 5°W (left) and 2°E (right) for wet period 1951-1970 and (b) variation of number of rain events (MCSs) per day and daily rainfall at 5°W (left) and 2°E (right) for dry period 1971-1990 (after LeBarbe et al. 2002).• Bi-modal rainfall maxima prevail south of 9°N, with the April-June the primary one while only one maximum exists in the north i.e. north of 9°N.• Rainfall distribution depends and follows the number of rain events. Note the decrease in both rain events and rainfall amount.The Little Dry Season (LDS)• Fig. 21 gives the variation of rainfall with latutude and month in Nigeria and showing the limits of the little dry season.• One rainfall maximum is noted north of 10°N and bi-modal to the south with a short break known as the Little Dry Season (LDS) which occurs in July and/or August.• Fig. 22 shows areas of west Africa affected by the Little Dry Season. LDS does not usually occur beyond 10°N and is also restricted to within longitudes 10°W and 9°E.
  • 8. (b)Fig. 20:
  • 9. Fig. 21
  • 10. Fig. 22:
  • 11. • LDS does not occur east of Port-Harcourt nor West of Robertsfield, Liberia.• Fig. 23 gives the monthly rainfall pattern over some cities in West Africa (after Hayward and Oguntoyinbo, 1987). Cities affected and not affected by LDS could be seen.
  • 12. Fig. 23:
  • 13. 3. SCALE INTERRACTIONS IN WEST AFRICA• West Africa is a region of unique interactions between several scales of motion. Interactions lead to climate variability and change which include:  Systematic decline in rainfall and inter-annual rainfall variability.  High variability in the onset and cessation dates of the rainy season.  Reduced number of larger and long-lasting Meso-scale Convective Systems (MCSs) and increase in smaller short-lived ones.• All these imply dire consequences for agricultural production, food security and water resources management.• Fig. 24 shows components of scale interaction in West Africa. The interacting scales are:  ITD,  Boundary Layer,  Monsoon Layer,  Northeast and Southwest Trades,  AEJ, TEJ,  Sub-Tropical Anticyclone,  Easterly and Westerly Waves, etc. as shown in the figure.
  • 14. Fig. 24: Components of scale interaction in WestAfrica.
  • 15. 3.1 Summer Situation3.1.1 Monsoon Depth and Strength• Monsoon depth and strength over West Africa exhibit seasonal, decadal, inter-annual and daily variations.• Strong boundary layer θe gradient results in stronger monsoon circulation which implies wetter Sahel while weaker boundary layer θe gradient results in weaker monsoon which implies relatively drier Sahel.• They reach a maximum in July/August. During this time AEJ and thunderstorms/squall lines are absent over coastal areas while little dry season (LDS) prevails.3.1.2 AEJ, TEJ and Deep Convection (See Fig. 8) AEJ• It results from meridional temperature gradient and is at ∼ 3°N in Jan/Feb. It has core speed ∼ 12m/s and maintained through latent heat release.• It is crucial for thunderstorms/squall lines onset and the required vertical shears are: -20≦ ΔUL ≦ -5 and 0 ≦ ΔUm 10, where ΔUL and ΔUm are wind shears between surface – 700hPa and 700 – 400hPa respectively.• Required vertical shear for deep convective systems not possible unless 400hPa winds are easterly and surface wind are southwesterly.
  • 16. • AEJ is more crucial and strong factor for MCSs development, sustenance and movement over West Africa. TEJ• It is a semi-global phenomenon. It occurs partly due to reversed thermal gradients. It is also linked to Tibetan highland-Indian Ocean thermal gradient.• It is about twice AEJ speed and is farthest north (10°N) in July/August. It is observed only between June-September.• It is only a modulating effect on MCSs rainfall delivery and has weak influence on their formation and maintenance.• The strength of both Jets, associated vertical wind shear and level of available moisture essentially determine the occurrence, size, intensity and rainfall delivery effectiveness of MCSs. MCSs/Squall Lines• Meso-scale Convective Systems (MCSs)/Squall lines are observed between AEJ AND TEJ positions. TEJ associated divergence at 250hPa over cyclonic side of AEJ enhances deep convection and implies intensified precipitation.• Fig. 25 shows the relative position of region of MCSs/Squall lines to the AEJ and TEJ.• Strong AEJ + weak TEJ (implies reduced Sahel rainfall) | ΔUL|>2 ΔUm : increased rainfall.
  • 17. Fig. 25:
  • 18. • TEJ is very strong and AEJ is weak when thunderstorm is expected. This agrees with observations of Dhonneur (1981) that a rupture of the AEJ is required condition for Cb – cluster development. Once the clusters have developed, the AEJ is re-established and the shear between TEJ and AEJ decreases.• Fig. 26a and b give the RegCM3 moisture build-up (q’-q s) at 700hPa over West Africa for dry year and wet year composites respectively.• For dry year composites (Fig. 26a), there is weak moisture inflow with drier than normal (easterly) air feeding in at mid-levels (700hPa).• For wet year composites (Fig. 26b), there is strong and easterly moisture inflow with moister than normal (westerly) air feeding in at mid-levels (700hPa).• The location of the AEJ and maximum rainfall belt relative to the ITD over West Africa for wet and dry years is shown in Fig. 19).• The striking differences between dry and wet years are: AEJ is weaker and closer to the ITD in wet than in dry years. Moisture inflow begins early and strong in the Boundary Layer (BL) with westerly winds feeding the AEJ during wet years but late and with northeasterly into and out of the AEJ in dry years.
  • 19. Fig. 26a: The RegCM3 moisture build-up (q’-q’ s) at 700hPa over West Africa fordry year composites. Arrows are wind vector anomalies. The thick dotted lineindicates the ITD. Shaded area indicate the position of AEJ – core of maximumeasterly wind (After Omotosho and Abiodun, 2007)
  • 20. Fig. 26b: The RegCM3 moisture build-up (q’-q s) at 700hPa over West Africa for wet yearcomposites. Arrows are wind vector anomalies. The thick dotted line indicates the ITD.Shaded area indicate the position of AEJ – core of maximum easterly wind ((AfterOmotosho and Abiodun, 2007).
  • 21. Squall Lines (SLs) Fig. 27a shows fast moving (≧ 12m/s), cold, line squall system (After Fink, 2008). • SLs are west-south-westward propagating meso-scale cloud clusters with a leading, sometimes convex line of CB’s. • Mean speed is 15m/s and life time is 12-24 hours. • They are responsible for about 80% of the annual precipitation south of 9°N and more than 90% in the Sahel. • There are similar wind system but much less deep westerlies (see vortex-type cross-section)Fig. 27a: Fast moving (≧ 12m/s), cold,line squall system (After Fink, 2008).
  • 22. Meso-scale Convective Systems (MCSs) • Fig. 27b shows numerous unorganized and frequent thunderstorms (Fink, 2008). • These may be caused by any or combination of:  Strong diurnal heating resulting in super- adiabatic lapse rate near the surface. Orographically forced ascent during particularly early and late monsoon season.Fig. 27b: Numerous unorganizedand frequent thunderstorms overWest Africa.
  • 23. Vortex-type Rainfall • Fig.27c shows winds of vortex-type system over Parakou/Benin (9N, 2E) in the Tropical Easterly Jet full monsoon season. • Vortex rainfall last longer and is lighter than rain from squall lines. • They occur in a less sheared (i.e. no AEJ), deep monsoon layer (hence lower CAPE environment) • They occur with a stronger Tropical July/August 2002 Easterly Jet (TEJ) than with MCSs. • Surface winds are westerlies both U-Wind [m/s] before and after the system unlike MCSs with westerlies before but easterliesFig. 27c: Winds of vortex-type system after.over Parakou/Benin (9N, 2E) in the full • TEJ has a modulating effect i.e. eithermonsoon season (After Fink , 2008) increases or decreases rain depending on the position of easterly wave over the area.
  • 24. Rainfall intensity Zonal wind profiles θe profiles Squall lines Vortex rain Fig. 28: Rainfall intensity, zonal wind profiles for the cases of squall lines and vortex-type rainfall and the θe profiles over West Africa (After Fink , 2008)
  • 25. Extratropically-induced pre-monsoon season convective events (Fink, 2008) SYNOPTIC EVOLUTION upper-level Tropical plume trough LFig. 29a : Infrared satellite imageat 1200 UTC on 17 Jan., 2004 Fig. 29b: 300 hPa geopotential height [gpdm] and MSLP [hPa] at 1800 UTC on17 Jan., 2004
  • 26.  Large subsidence (warm advection) to the west (east) of 2nd upper-trough → dynamically supports further pressure fall  Inland penetration of moist monsoon air allows formation of deep convection and precipitation  Extreme precipitation also in the subtropicsFig. 29c: Extratopically-induced pre-monsoon season convective events(Fink, 2008)
  • 27. Extratropically-induced pre-monsoon season convective events• Mid-latitude upper level trough extends far down into West Africa, “magnetizing” the ITD and pulling it northward.• This causes inland surge of the warm monsoon flow and leads to formation of deep convective clouds and (sometimes) heavy precipitation. These could be seen in Fig. 29c.Precipitation Types in West Africa (Summary)• The Squall Lines (SLs/OCS)• Mesoscale Convective Systems (MCSs)• Vortex-type Rainfall• Extratropically-induced pre- and post-monsoon seasons convective eventsAll are products of Scale Interactions
  • 28. 3.2 Winter Situation3.2.1 Monsoon Depth• The ITD mean position in winter (Jan/Feb) is at 7°N and therefore marks the limit of the south west monsoon over West Africa.• The monsoon depth from coast to ∼7°N is therefore shallow since this area is within the weather zone B having little cloudiness hence frequent early morning fog. Some diurnal convections do occur and could lead to scattered showers/thunderstorm.• North of 7°N there is no cloudiness (clear sky) and the north east trades (monsoon) prevail bringing in Saharan dust. Most of West Africa is in zone A and frequent dust spells occur.3.1.2 LLJ, AEJ, TEJ, STJ (See Fig. 8)• From the coast of West Africa to ∼7°N, shallow south westerly winds (surface-below 900hPa for Kano for instance) prevail and easterly winds are evident up to about 400hPa having a maximum around 700hPa – the African Easterly Jet (AEJ).• Above about 400hPa westerly winds prevail culminating to a maximum at about 200hPa – the Sub-Tropical Jet (STJ) whose core is around 30°N.
  • 29. • North of about 7°N Kano for instance, easterly winds prevail from surface to about 700hPa with a maximum around 900hPa – the Low Level Jet (LLJ) that aids in transporting Saharan dust to West Africa, Nigeria in particular.• Above about 700hPa, the winds become westerlies reaching a maximum around 200hPa – the Sub-Tropical Jet (STJ).• Saharan anticyclone intensifies behind cold fronts (Fig. 30a) due to injection of cold mid-latitude air at higher levels (Fig. 30b). This leads to: Strong wind at the gradient (900hPa) level, as shown in Fig. 30c. Raising of dust in the Bilma – Faya Largeau area, the source region. Advection of dust in a plume-like structure as an informer tongue.• Fig. 31a shows dust tongue associated with strong negative anomalies of equivalent potential temperature at the same level.• In dusty atmosphere, westerly winds aloft attain speed of 40 – 50m/s and descend to as low as 700hPa (Fig. 31b).
  • 30. Fig. 30a: Typical surface synoptic chart showing the passage ofmid-latitude cold front and intensified Sahara high pressure systemover West Africa.
  • 31. Fig. 30b: Injection of cold air from mid-latitude into West Africa upperlevels (500hPa) in association with surface cold fronts causes surfacepressure intensification and consequent raising of dust over desert(Bilma-FayaLargeau area).
  • 32. Fig. 30c: Typical gradient (900hPa) level wind isotach (m/s) over thesource region (top) and wind profile over Faya-Lageau (bottom).
  • 33. Fig. 31a: Vertical cross-section of equivalent potential temperature anomalies (K) overKano, Nigeria in January 1974. Values between arrows denote surface visibility at 1200UTC(After Omotosho, 1989)
  • 34. Fig. 31b: Wind profile over Kano, Nigeria averaged over five (5) dusthaze spells (DS) of period 3-7 days and clear periods of 5-8 days (AfterOmotosho, 1989)
  • 35. 4.0 INSTABILITY PROCESSES IN THE ATMOSPHERE• All atmospheric disturbances (cumulonimbus clouds, thunderstorms, easterly waves, cyclonic waves, hurricanes, Rossby waves, etc.) are as a result of perturbation or instability in the basic flow in the atmosphere.• Different types of the disturbances could occur depending on the processes that lead to those disturbances.• There are disturbances not directly influenced by environmental factors but as a result of perturbation due to obstruction (mountains, etc.) against the basic flow.• The Instability Processes in the atmosphere are: Conditional/Convective Instability (Insitu Development) Conditional Instability of the Second Kind (CISK) Baroclinic Instability and Barotropic Instability4.1 Static Instability• The diagram below (Fig. 32) illustrates static instability.• The sun heats the ground but the air does not absorb heat equally (e.g. sand and grass area.)
  • 36. Tp Tp Tp Tp32
  • 37. • The area that absorbs faster has its density less and rises while others remain.• The parcel will continue to rise and volume expands till the environmental temperature Te equals the parcel temperature Tp and therefore stops rising at p2.• When the parcel is again forced to rise to p 3 it will expand more but if Tp(p3) < Te(p3), the density of the parcel is larger than that of the environment and therefore sinks.• This process does not give us cumulonimbus cloud but only gives cumulus type1 and type2 (ragged cumulus) which cannot develop to cumulonimbus.• Condition for static instability for unsaturated or dry air is mathematically: ∂θ/∂z < 0 i.e. ∂θ/∂p > 0• This is more conveniently represented by the equivalent potential temperature, θe as : ∂θe/∂z < 0 or ∂θe/∂p > 0
  • 38. 4.2.1 Conditional Instability• Conditional Instability is said to exist in the atmosphere when the environmental lapse rate is between the dry adiabatic lapse rate and the moist adiabatic lapse rate, i.e. Γd > γ > Γswhere Γd ≈ 9.8 K/km = constant Γs ∼ ½ Γd ≠ constant (depends on the amount of moisture present)• That is to say that the unsaturated parcel is colder than the environmental air (i.e. stable) while the saturated parcel is warmer than the environmental air (i.e. unstable).• Considering Fig. 27, the unsaturated parcel at temperature T, as it rises it follows the dry adiabat which seems to correspond with the environmental temperature.• As it gets to a point where the dry adiabat intersects with the mixing ratio line through the dew point temperature, we have the Lifting Condensation Level (LCL) and the point is called the Normand Point. This is the cloud base.
  • 39. Fig. 33
  • 40. • However, the environmental lapse rate shows the presence of an inversion from LCL up to some layer above, and the parcel from this point should follow the saturated or moist or pseudo adiabat but it is colder than the environment.• If not forced to rise, it will return to its original position but if forced to rise up to the point where the moist adiabat just intersects the environmental temperature line where the parcel just becomes warmer than the environment and we say that it has attained the Level of Free Convection (LFC).• The negative area bounded by the moist adiabat and the environmental temperature line from LCL to LFC represents Convective Inhibition (CIN).• From Fig. 33, the parcel is quite free to rise with great buoyancy depending on the area enclosed by moist adiabat and the environmental temperature line. This area represents the Convective Available Potential Energy (CAPE).• The point at which the moist adiabat intersects the environmental temperature line up in the atmosphere is the Level of Neutral Buoyancy (LNB) or the Terminal Equilibrium Level (TEL).• At this level the environmental temperature is already increasing again with height – within the stratosphere.• TEL or LNB is the theoretical cloud top. Theoretical in the sense that the cloud top in reality could overshoot this level depending on how large CAPE is.• Figs. 34 & 35 illustrate ways to overcome Convective Inhibition (CIN).
  • 41. 34:
  • 42. 35:
  • 43. • They include: Strong surface heating leading to super-adiabatic lapse rate. Air forced to rise over obstacle – topographic or orographic lifting, Colder Air (Fronts or Coldpools) Convergence associated with a cyclonic vortex or asymptote of convergence.• CAPE can be defined mathematically as:CAPE = ∫F.dl = ∫Bdz ≈ ∫[g(Tcld – Tenv)/Tenv]dz (integral is from cloud base to top) where B is buoyancy given by B = – g(ρ´/ρ) which is approximately equal to - g{ (T´/T) + 0.608q´– ql } for cloudy air to take into account the effects of humidity and condensate. In general all the three terms are important. 1K perturbation in T is equivalent to 5g/kg perturbation in water vapour or 3g/kg in condensate. (T and ρ are mean values)• CAPE represents the amount of potential energy of a parcel lifted to its level of neutral buoyancy.• This energy can potentially be released as kinetic energy in convection.• Conditional instability is sometimes classified into three types: Area associated with CAPE > Area associated with CIN – Latent Type Area associated with CAPE < Area associated with CIN – Pseudo-latent Type
  • 44. • When no Area associated with CAPE appears for parcels of any level – Stable Type4.2.2 Potential or Convective Instability• With the parcel method, the stability properties of the atmosphere have so far been considered when an isolated mass, the parcel is vertically displaced.• This occurs, for example, when the warming of the lower layers causes the ascent of air masses with dimensions of the order of hundreds of meter to 1km.• These masses can eventually become accelerated by a latent instability, as has been seen. They are called thermals or bubbles, and become visible as convective clouds.• When the convection becomes more intense we may have, instead of isolated masses, a continuous jet of ascending air.• It is also important to study the vertical movements of an extended layer of atmosphere, such as may occur, for instance, during forced ascent of an air mass over an orographic obstacle, or due to large-scale vertical motions of appreciable magnitude (as, for example, in frontal situations) see Figs. 28 & 29.• Lifting of a column of moist air until it is saturated throughout also affects the stability. An important difference between dry and moist air is that moist air, initially stable, may be made absolutely unstable or conditionally unstable by lifting but this does not
  • 45. • A column of air which is rendered unstable by lifting to saturation is said to be convectively or potentially unstable.• The criteria for convective stability may be expressed in terms of the lapse rate of equivalent potential temperature as follows: ∂θe/∂z > 0 Convectively Stable ∂θe/∂z = 0 Convectively Neutral ∂θe/∂z < 0 Convectively Unstable• These criteria are best understood with reference to a tephigram.• Convective instability has to do with the lifting of layers and should not be confused with conditional instability, which applies to an undisplaced layer.
  • 46. 4.3 Conditional Instability of the Second Kind (CISK)• In this type of arrangement, strong interactions exist between the smaller and the larger circulations.• Charney and Eliasen (1964) had shown that ordinary conditional or convective instability produces significant growth rate only on the scale of individual cumulus clouds.• The convergence required is on larger scale and is frictionally driven.• With considerable ascent thus ensured, attainment of buoyancy by the rising air parcel is possible and the condensing cloud air then supplies latent heat to the larger scale flow.• The interaction is thus cooperative: the larger scale motion supplies needed moisture for cumulus growth through frictional convergence and Ekman pumping while the cumulus clouds in turn supply the latent heat required to drive the large scale circulation.• When such cooperative interaction leads to the unstable growth of the larger- scale system, the process is referred to as Conditional Instability of the Second Kind (CISK).
  • 47. • The latent heat release is the major source of energy for the maintenance of many tropical circulations unlike the mid-latitude synoptic scale systems which are baroclinic and develop by horizontal temperature advection.• CISK is therefore summarized as the cooperative interaction between small scale cumulus convection and large scale disturbance which results in the unstable growth of the larger scale system.• This is very important in case of tornadoes and hurricane development. 4.4 Baroclinic Instability• The middle latitude atmosphere is baroclinic i.e. the density depends on pressure and temperature i.e. ρ=ρ(p,T) mathematically.• The geostrophic wind has vertical shear related to horizontal temperature gradient.• In this atmosphere, baroclinic waves amplify through the loss of potential energy of the basic flow and develop within the zonal flow.• The axis of the jet stream at any instant tends to coincide with a narrow zone of strong temperature gradients called the polar front.• This is the zone which in general separates the cold air of polar origin from warm tropical air.
  • 48. • The occurrence of an intense jet core above the frontal zone is of course a consequence of thermal wind balance.• It is a common observation in fluid dynamics that jet flows, in which strong velocity shears occur, may be unstable with respect to small perturbations.• This means that any small disturbance introduced into the jet flow will tend to amplify, drawing potential and or kinetic energy from the jet as it grows.• Most synoptic scale systems in the middle latitudes appear to develop as the result of instability in the jet stream flow.• This instability, called baroclinic instability, depends primarily on the vertical shear in the jet stream, and hence tends to occur primarily in the region of the frontal zone.• Horizontal temperature advection plays a crucial role in the development of baroclinic waves.• Baroclinic disturbances may themselves act to intensify pre-existing temperature gradients and hence generate frontal zones.• Baroclinic instability in zonal flows was studied independently at about the same time by Charney (1947) and Eady (1949).• The latter provided our present understanding of the origin of cyclone waves while Charney’s work dealt with longer waves.
  • 49. 4.5 Barotropic Instability• A barotropic atmosphere is one in which density of air depends on pressure only and there is horizontal but no vertical wind shear (thermal wind is zero).• Unlike the mid-latitudes where strong latitudinal temperature gradients provide the primary energy source for synoptic scale disturbances, temperature gradient in the tropical regions are small leading to small available potential energy storage.• Barotropic instability occurs in a barotropic atmosphere which is likely to be approximately non-divergent.• The concept of barotropic waves has been applied at about 600hPa – the so- called ‘level of non-divergence’.• Forecasting of the movement of such waves is based on the conservation of absolute vorticity by individual air particles.• The same principle is applied in respect of long-waves (or Rossby waves) which are essentially barotropic waves in air flow of relatively uniform speed.• Energy for the maintenance of barotropic disturbances comes primarily from latent heat release.• Since most tropical disturbances develop in-situ, their development cannot usually be accounted for by barotropic instability processes.
  • 50. • It is now established that barotropic instability and Conditional Instability of the Second Kind (CISK) are two main mechanisms for the initiation of tropical disturbances.• Now, starting from the barotropic vorticity equation ∂(▽ 2ψ)/∂t = -V ψ·▽(▽ 2ψ + f) and introducing perturbation velocities as for the quasi-geostrophic theory (but here, no assumption of geostrophy) the linearized perturbation stream function equation may be expressed as (∂/∂t + ū∂/∂x)▽ 2ψʹ + (β-d2ū/dy2)∂ψʹ/∂x = 0 where ū is time average of u• Kuo(1949) showed that barotropic instability of zonal flow is possible if β-d2ū/dy2= d/dy(ζ+f) = 0 somewhere• That is, the profile of absolute vorticity must change sign somewhere within the domain of interest. A full treatment of this is in Holton (1979) pp. 352-355.
  • 51. Fig. 36bFig. 36a
  • 52. • Such profiles, taken from Rennick (1976), are shown in Fig. 36a which shows that the condition for barotropic instability is met only in the month of August (i.e. between June and September) at 700hPa at longitude 5°E and within the latitudes of 5 and 15°N as the absolute vorticity is maximum there.• So, only in the month of August does d/dy(ζ+f) change sign. Thus, the season of African wave activity coincides with the season during which the necessary condition for barotropic instability is satisfied.• This implies that the barotropic instability may be an important mechanism by which kinetic energy is transformed from the mean flow to growing wave disturbance.• Rennick (1976) has shown, however, that tropical zonal flows do satisfy the conditions for both barotropic and baroclinic instabilities.• That is, even though horizontal temperature gradients may be small, they are nevertheless sufficient to give some vertical shear of the wind in addition to lateral wind shear.• In fact, strong vertical shear exists over West Africa and Eastern Atlantic. That this is so was shown by Charney (1973) using the internal jet theory of Charney and Stern (1962).• The condition here is that the gradient of potential vorticity q must be negative in the domain i.e.
  • 53. ∂q/∂y = ∂(ζ+f)/∂y + ∂{(pf02/Rσ)(∂ū/∂p)}/∂ppot. vort. (barotropic term) (baroclinic term) gradient where ζ = ∂v/∂x-∂u/∂y≈∂u/∂y (since v is small), f02 is the Coriolis parameter at the centre of the region, and σ is the static stability defined by σ =(T/θ) ∂θ/∂p, (q, T and θ are mean values with respect to time and longitude).• The resulting field of ∂q/∂y at 700hPa is shown in Fig. 36b giving firm indication that although both types of instability are important, the baroclinic terms nevertheless makes a smaller contribution.• Charney showed that if surface temperature increases northwards, as is the case in northern Africa, then for instability to occur, ∂q/∂y must be negative somewhere in the domain.• From the results due to Rennick (1976) shown in Fig. 36b, it is obvious that African waves are due to both barotropic and baroclinic instabilities, with the former being the main contributor.
  • 54. Fig. 37:
  • 55. • Furthermore, the observed monthly mean profiles of absolute vorticity at 5°E (from Kuo, 1949) are presented in Fig. 37 at 850, 700 and 500hPa for May, August and October.• It can be clearly seen that there is a local maximum of absolute vorticity, (the meridional gradient of absolute vorticity vanishes) at 700hPa during August and October, at 700hPa during August and October, and at 500hPa during August.• This absolute vorticity maximum occurs only during the months when the wave are observed and near 700hPa where the wind amplitude of the waves is greatest.
  • 56. 5.0 WEST AFRICAN WEATHER SYSTEMS5.1 African Easterly Waves• The model of waves in the easterlies as presented by Riehl (1954) and was aplied successfully for the Caribbean and West Pacific cannot be applied piece-meal to the easterly waves over West Africa and the eastern Atlantic Ocean.• This is because the wind regimes of the areas differ significantly (compare Figs. 10A, B & C for Venezuela, West Africa and West Pacific regions respectively).• The juxtaposition of two tropospheric jets over a rather shallow layer of very moist and warm southerly air is unique and are important over West Africa and the Eastern Atlantic.• Hence, the African Zone of wave activity is unusual when compared to the lower troposphere of the rest of the tropics because temperature changes of 10°C can occur within a latitude band of only 10°; the resulting vertical wind shear leads to the mid-tropospheric African Easterly Jet (AEJ).• The easterly waves develop south of the jet (AEJ) core in a region where the mean zonal flow is barotropically unstable (Burpee,1972).
  • 57. • Carlson, 1969a, b and 1971; Burpee, 1971, 1974 and 1975; and Reed et al. 1977 have shown that the waves typically cross tropical Africa from mid-June until early October with wavelengths ranging from 1500 to 4000 km, periods from 2.2 to 5.5 days and west ward speed of 5 to 10 m/s.• The average period and wavelength are approximately 3.5 days and 2500 km respectively. The waves first appear near 700hPa, the level where they are most easily tracked in the meridional wind field.• Observations indicate that the waves derive their energy from the kinetic energy of the mean flow near 700hPa.• The spectral computations of Burpee(1971 and 1974) showed that the waves transport horizontal momentum away from the low-level easterly jet at latitudes near 10° to 15°N and pressures from 850 to 500hPa, and that the waves transport heat down the temperature gradient during each month of wave activity.• These transports of heat and momentum suggest that both the horizontal and vertical shear of the mean zonal wind may be acting as sources of energy for the African waves.• As the waves propagate westward they strengthen and influence an increasingly greater depth of the atmosphere. Near the west coast of Africa, at latitudes of 10° to 15°N, the wave disturbances reach a maximum amplitude of 5m/s during the latter part of August and September.
  • 58. • Fig. 38 shows a streamline analysis of the 700hPa wind field and corresponding infrared satellite picture for 16 September 1974, during GARP Atlantic Tropical Experiment (GATE).• There are two waves on the map with the troughs at 2° and 18° W and the ridges at 13° and 33° W along 10° N. In this example, both wave and troughs have closed, cyclonic circulations in the wind field, while in June and July most of the disturbances are observed as open waves in the easterlies.• The word “wave” is appropriate regardless of the closed or open appearance of the wind circulation in the streamline analyses because these disturbances are propagating, quasi-periodic phenomena with similar structure.• The satellite picture in Fig. 38 shows considerable convective cloudiness near the trough of the wave along the west coast of Africa, a pattern that was observed frequently during GATE.• A vertical cross section of the average meridional wind pattern associated with the waves near Dakar is shown in Fig. 39 (Burpee, 1975). This figure is an average for 24 waves observed during the three phases of GATE and reveals that the maximum amplitude of 3.5m/s occurs near 700hPa.• Wave features tilt eastward with height from the surface to about 650hPa and then westward with increasing height in the upper troposphere.
  • 59. Fig. 38:
  • 60. Fig. 39:
  • 61. • Reed et al., 1977 have used the preliminary data from GATE to determine the average three-dimensional structure of eight waves observed during the last half of August and the first half of September in western Africa and the eastern Atlantic.• They used a compositing technique similar to that used by Reed and Recker, 1971; and Burpee, 1975, which averages the observations according to eight categories defined by the wind field associated with the waves at 700hPa.• The composited wind fields of the waves are shown with the mean wind included in Fig. 40 and with the mean wind removed in Fig. 41 for surface, 850, 700, 500 and 200hPa.• The average wavelength of these waves is slightly less than the 2500km found by Burpee for 24 waves observed during the three phases of GATE.• Wave structure in Fig. 40 shows a surface convergence zone only, a cyclonic vortex at 850hPa and a distinct wave pattern at 700hPa. Wave intensity decreases upwards thereafter.• With the mean wind removed (Fig. 41), the composited wind fields have closed, cyclonic and anticyclonic circulation centres at the lowest three levels.• These were done on the waves over West Africa divided into categories as was the case for Figs. 40 and 42. Categories 4 and 5 are areas of the wave ahead of the trough. The waves are very much evident at 850 and 700hPa with the latter more pronounced.
  • 62. Fig. 40:
  • 63. Fig. 41:
  • 64. Fig. 42:
  • 65. • The circulation pattern in the zonal perturbation shows a cyclonic one ahead of the wave trough and anticyclonic behind the wave trough.• South of these circulation centres, the trough and ridge lines have a well defined slope from southwest to northeast, while to the north of the circulation centres very little slope is observed. The wind field at 200hPa is dominated by a region of strong inflow and a second region of strong outflow.• The distribution of relative humidity and vertical motion from the study of Reed et al, 1977 are illustrated in Fig. 42 a and b for an east-west line through the waves at a latitude of approximately 11°N.• The variations of relative humidity are closely related to the meridional wind with relatively dry air in the northerlies and relatively moist air in the southerlies following the trough.• Vertical motion has been computed kinematically from the wind field presented in Fig. 41. The maximum upward motion is about 3hPa/hr and occurs at 750hPa between the trough and the maximum northerly winds.• There is strongest convergence and upward motion located ahead of the wave trough where cyclonic vorticity is also strongest.• At the trough line categories 4 and 5, there is deeper layer of relative humidity > 80%.
  • 66. • Highest rain rainfall is ahead of the trough – categories 2 and 3 (Fig. 36c).• Weather and precipitation distribution are therefore a reverse of the Caribbean situation.• The waves are divided into categories as shown below (Fig. 43a). Cloud types are shown in the various categories in Fig. 43b. The tall cumulonimbus clouds could be seen in categories 3-5 ahead of the trough axis while the others have none or shallow clouds.
  • 67. Fig. 43a:Fig. 43b:
  • 68. 5.2 MESO-SCALE CONVECTIVE SYSTEMS (MCS) AND SQUALL LINES• To a large extent, West African meso-scale convective systems and squall lines are dependent on or controlled by the summertime monsoon oscillations.• They are absent when the monsoon retreats from most parts of West Africa (November – March), suppressed if the monsoon is too deep ( giving way to ‘monsoon rain’ over coastal regions in June/July and very violent when the monsoon depth is about 1 – 2 km especially north of 11°N.• Deep convective systems and storms over West Africa are consequences of one or a combonation of the following: Differential surface heating – convection is essentially an insolation phenomenon. Destabilization due to upper-level or cloud-top radiative cooling which leads to convective overturning (see, for example, Omotosho, 1981). The inherent convective or conditional instability present in the basic flow up to 800hPa (this is true for the entire tropical atmosphere). The availability of abundant moisture supply at low levels, with a relatively drier middle troposphere. The presence of the so-called African Easterly Jet (AEJ), and the associated vertical wind shear (Moncrieff and Miller, 1976).
  • 69. • However, despite all the above, the instability has to be released: A release mechanism for fourth condition in the form of low-level convergence associated with vortices and asymptotes of convergence (Okulaja, 1970; Obasi, 1974). Another release mechanism is the Richardson number which indicate the role of boundary layer energetics in storm development (Omotosho, 1987).Meso-Scale Convective Systems (MCSs)• They are numerous and unorganized frequent thunderstorms and may be caused by any or combination of: Strong diurnal heating resulting in super-adiabatic lapse rate near the surface. Orographic forced ascent particularly during early and late monsoon season.Squall Lines (SLs)• Squall lines (SLs) are west-south-westward propagating meso-scale cloud clusters with a leading, sometimes convex line of cumulonimbus clouds.• Mean speed is 15m/s and life time 12 – 24 hours.• The are responsible for about 80% of the annual precipitation south of 9°N and more than 90% in the Sahel.
  • 70. • There are two squall line/thunderstorm seasons for regions up to 11°N but only one to the north of 11°N.• The monsoon precipitation in June and later, the so-called ‘little dry season’ (LDS) of West Africa which affect only coastal areas south of 10°N during July/August both occur between the two storm seasons.• Thus, most of the precipitation is from these deep convective systems while the entire rainfall of the Sahel comes from the storms.• In fact, it has been shown (e.g. Glantz, 1976 and Omotosho, 1985) that squall lines and thunderstorms deliver more than 60% of the total annual precipitation of areas south of 10°N and as much as 80% or more north of this latitude as shown in Fig. 44.• Also, their destructive capacity is magnified in the loss of millions of naira worth of food and cash crops and property by accompanying strong winds (gusts).• These convective systems are therefore vey important to the socio-economic politics of the region as a whole.
  • 71. Fig. 44b:Fig. 44a
  • 72. Mechanism of Storm Development and Sustenance• Several researchers have attributed squall line formation to different mechanisms.• It was initially believed that orography played a major role in their initiation and alignment e.g. Barrefors (1964), until Omotosho (1984) showed that although these physical features do influence the intensity of the storms, the storms do not, in general, owe their existence to orography.• Later and even at present, the propagation storms are regarded as the products of low-level convergence associated with the synoptic-scale easterly waves in areas where the atmosphere is convectively and conditionally unstable (see for example, Reed et. al. (1979) and Adefolalu (1985)) because they occur during the period (May – September) of wave activity.• Fig. 38b shows that squall lines are found almost any where on the easterly wave although mostly ahead of the wave trough.• Also Omotosho (1981), Mcbride & Gray (1980) have demonstrated that the waves only have a modulating but not a forcing effect on deep convection over the GATE area. (GATE is GARP Atlantic Tropical Experiment and GARP is Global Atmospheric Research Programme).
  • 73. • Notwithstanding all these, it is to be emphasized that the release mechanism is a crucial stage in thunderstorm development.• The required low level convergence usually manifests itself as cyclonic vortices or convergence lines (often as asymptote of convergence) at 900 or 850hPa level.• When the vortices are accompanied by easterly waves at the 700hPa level, the depth of convergence increases and hence upward motion is generally enhanced.• Vigorous thunderclouds results especially if the convective available potential energy (CAPE), proportional to the buoyancy of the in-cloud rising air parcel, is large. This is the perceived role of the easterly perturbation.• Now, regarding the forcing mechanism for thunderstorms and squall line, it is now evident that the existence of a mid-tropospheric wind maxima, the AEJ, with reversed shear in the lower levels is important to their development (Moncrief and Miller, 1976; Bolton, 1981).• The systems are in quasi-state with the density current (gust front) feeding the cloud cells with warm and moist (high θe) air of the monsoon as shown in Fig. 45.
  • 74. • It is also suggested that the simultaneous presence of positive boundary layer vertical wind shear, moist layer of at least 1km depth and a distant and organized AEJ are all necessary conditions for wide spread thunderstorms to develop and be aligned into a line (see e.g. Fig. 46).• Most recently, Omotosho (1987) found that the storms occur most frequently when boundary layer Richardson Number R i is within the range - 2≤ Ri ≤ 0 and also 1≤ Ri ≤4 in the middle troposphere. Ri = - (g/θ)(∂θ/∂z)/∣∂V/∂z∣ 2 = - (g/θ)(∂θ/∂z)/∣(∂u/∂z)2 + (∂v/∂z)2∣ (This is the rate of destruction of thermal stratification by kinetic energy)• The interaction between the clouds, gust front and the boundary layer is essentially that of the Conditional Instability of the Second Kind (CISK).• Such interaction is possible only because of the peculiar regime over West Africa which causes the relative flow to enter the storm clouds from the front.• The storm cloud sucks in cool, dry (low θe) air from the front in the layer 800- 600hPa and becomes the evaporative downdraft due to the drag of the falling precipitation.• Some of the precipitation is evaporated into the downdraft which is cooled and then reaches the surface as strong gust of wind.
  • 75. Fig. 45: Fig, 46:
  • 76. • Because of the large temperature contrast and wind shifts across the boundary of this cold inward rushing air and the very warm, moist ambient (high θe) air at and close to the ground, a frontal zone (gust front) is formed.• The warm, moist boundary layer air is thus forced upward into the cloud. This way the storms are assured of much needed moisture and hence the accompanying latent heat release.• The precipitation into the large scale flow then enhances the moisture convergence at the gust front and the continued sustainance of the storms until the front travels too far ahead of the storm cloud.• At this stage, the moisture supply (and hence latent heat) is drastically reduced or cut-off and the storm begins to decay while a new one develops ahead of the decaying one. Comparison with Squall line of other Region• West African squall lines are similar in many respects to their counterparts over Venezuela although they occur within dissimilar flow regime.• They all propagate through, and in general travel faster than the basic flow at all levels in the cloud layer. This property causes inflow of air into these systems from the front.
  • 77. • This is in contrast to mid-latitude squall lines which are advected since they travel at nearly the mean wind speed in the cumulonimbus layer.• Again, unlike temperate squall lines, tropical ones are associated with larger values of the Richardson number, Ri.• While tropical storms are characterized by values of R i greater than 2.0 in the convective layer, temperate latitude storms are associated with lower Richardson number (Ri ≤ 1.0, high wind shear).• The transfer of momentum in tropical storms of Venezuela and West Africa are distinctive and of large magnitude.• Also, their propagation speed cs is almost constant over a wide range of Ri.