REVIEWS El..SEVIER Earth-Science Reviews 37 (1994) 89-134
Review: the atmospheric boundary layer
J.R. Garratt CSIRO Division of Atmospheric Research, Private Bag No. 1, Mordialloc, Victoria 3195, Australia Received 28 December, 1993; revised and accepted 27 April, 1994
Abstract
An overview is given of the atmospheric boundary layer (ABL) over both continental and ocean surfaces, mainly from observational and modelling perspectives. Much is known about ABL structure over homogeneous land surfaces, but relatively little so far as the following are concerned, (i) the cloud-topped ABL (over the sea predominantly); (ii) the strongly nonhomogeneous and nonstationary ABL; (iii) the ABL over complex terrain. These three categories present exciting challenges so far as improved understanding of ABL behaviour and improved representation of the ABL in numerical models of the atmosphere are concerned.
1. lntroduction factors controlling the erosion and ultimate breakdown of this inversion. 1.1. Role of the atmospheric boundary layer (b) Control and Management of Air Quality is closely associated with the transport and disper The atmospheric (or planetary) boundary layer sal of atmospheric pollutants, including industrial plays an important role in many fields, including plumes. Processes of concern include turbulent air pollution, agricultural meteorology, hydrology, mixing in the ABL, particularly the role of con aeronautical meteorology, mesoscale meteorol vection, photochemistry and dry and wet deposi ogy, weather forecasting and climate. We can tion to the surface. In this general area, research summarise just a few of the problems for which on atmospheric turbulence has a very important boundary-layer knowledge is important, as fol practical application, and local meteorology, in lows cluding the role of mesoscale circulations (sea (a) Urban Meteorology is associated with the breezes, slope winds, valley flows) and the phe low-level urban environment and air pollution, nomenon of decoupling of the low-level flow and including air pollution episodes involving photo the large scale upper flow, is also of major rele chemical smog and accidental releases of danger vance. ous gases. The dispersal of smog and low-level (c) Aeronautical Meteorology is concerned with pollutants depends strongly on meteorological boundary-layer phenomena such as low cloud, conditions. Of particular importance is informa low-level jets and intense wind shear leading to tion on the likely growth of the shallow mixed high intensity turbulence, particularly for aircraft layer resulting from surface heating, and on the landing and take-off. In the case of low clouds
0012-8252/94$26.00 © 1994 Elsevier Science B.V. All rights reserved SSDI 0012-8252(94)00032-R 90 J.R. Garratt / Earth-Science Reviews 37 (1994) 89-134 and low-level jets, factors affecting their forma turbulence is strongly influenced by the diurnal tion, maintenance and dissipation are of great cycle of surface heating and cooling, by the pres importance. ence of clouds and by horizontal variability in (d) Agricultural Meteorology and Hydrology surface properties. Neutra! flow, in which buoy are concerned with processes such as the dry ancy effects are absent, is readily produced in the deposition of natural gases and pollutants to wind tunnel, and may be closely approximated in crops, evaporation, dewfall and frost formation. the atmosphere in windy conditions with a com The last three are intimately associated with the plete cloud cover. The unstably-stratified ABL, state of the ABL, with the intensity of turbulence or convective boundary layer (CBL), occurs when and with the energy balance at the surface. strong surface heating (due to the sun) produces (e) Numerical Weather Prediction (NWP) and thermal instability or convection in the form of Climate Simulation based on dynamical models thermals and plumes, and when upside-down of the atmosphere depend on the realistic repre convection is generated by cloud-top radiative sentation of the Earth's surface and the major cooling. In strongly unstable conditions driven by physical processes occurring in the atmosphere. lt surface heating, the outer region of the boundary has been said (Stewart, 1979) that no general layer in particular is dominated by convective circulation model is conceptually complete with motions and is often referred to as the mixed out the inclusion of boundary-layer effects, and layer. In contrast, the stably-stratified ABL oc that no prediction model can succeed without a curs mostly (though not exclusively) at night, in sufficiently accurate inclusion of the influence of response to surface cooling by lang wave emission the boundary. The boundary layer affects both to space. The unstable ABL is characterized by a the dynamics and thermodynamics of the atmo near-surface superadiabatic layer, and the stable sphere. There are a variety of dynamic effects: ABL by the presence of a surface inversion. more than a half of the atmosphere's kinetic energy loss occurs in the ABL (Palmen and New 1.3. Boundary-layer depth ton, 1969). Boundary-layer friction produces cross-isobar flow in the lower atmosphere, whilst The top of the boundary layer in convective boundary-layer interaction permits air masses to conditions is often weil defined by the existence modify their vorticity. From the thermodynamic of a stable layer (capping inversion) into which perspective, all water vapour entering the atmo turbulent motions from beneath are generally sphere by evaporation from the surface must en unable to penetrate very far, though they may ter through the ABL. Even the oceans are strongly continually erode it, particularly where latent heat influenced by the ABL, since it is through the is released in rising elements of air. The height of boundary layer that they gain most of their mo this elevated stable layer is quite variable, but is mentum so influencing the oceanic circulation. generally below 2 to 3 km. The top of a convec tive boundary layer is weil defined in Fig. 1 by the 1. 2. Turbulence sharp decrease in aerosol concentration at a height of about 1200 m, and coincides with the From both a climate and local weather per base of a deep and intense subsidence invcrsion. spective, the most important ABL processes that Over deserts in mid-summer under strong surface need to be parametrised in numerical models of heating the ABL may be as much as 5 km deep, the atmosphere are vertical mixing and the for and even deeper in conditions of vigorous cumu mation, maintenance and dissipation of clouds. Ionimbus convection. In stable conditions, in con Those Iand-surface properties that are potentially trast to the above, the boundary layer is not so crucial to accurate climate simulation include readily identified, turbulence is much weaker than albedo, roughness, moisture content and vegeta in the unstable case, and consequently the depth tion cover. no more than a few hundred metres at most. At Over land in particular, the structure of ABL night over land, under clear skies and light winds, J.R. Garratt / Earth-Science Reviews 37 (1994) 89-134 91
a capping stratocumulus layer and a depth on the order of 0.5 km. Such a shallow boundary layer is also found in coastal regions when warm air flows from land over a relatively cool sea. In the trop ics, the mean structure is very much dependent on the season, and whether conditions are dis turbed (in the vicinity of the inter-tropical conver (a) 1000 gence zone) or undisturbed. In the former, devel oping cumulus clouds result in poor definition of the ABL top, whilst in undisturbed conditions, the ABL top is well defined by the trade-wind 5!Xl inversion. Under special circumstances, the ABL depth over the ocean can be comparable to that over land in the middle of the day. This can occur during intense cold-air outbreaks over the ocean, 100 200 Aerosol concentration when !arge "jumps" in temperature and humidity that identify the ABL top are particularly notice able, and are the result of cold, dry air flowing out from the continent over relatively warm sea. The purpose of this review paper is to provide 1500 the reader with an overview of boundary-layer structure and behaviour over land and sea (in the
(b) latter case, often cloud-topped), for homoge 1000 neous and heterogeneous surface conditions. In addition, some aspects of boundary-layer parametrisations for use in regional and global climate models are discussed. For more intensive 500 reading, the classical texts should be consulted, e.g. Sutton (1953), Priestley (1959), Lumley and Panofsky (1964), Zilitinkevich (1970), Tennekes 0 L___ __j__:::==~-'-~---'----__J and Lumley (1972), Haugen (1973) and Brown 285 290 295 300 305 Potential temperature (K) (1974). Advanced texts dealing with specialised aspects of the ABL include Wyngaard (1980), Fig. 1. Vertical profiles of (a) aerosol concentration in arbi trary units and (b) potential temperature observed overland in Nieuwstadt and Van Dop (1982), Stull (1988), convective conditions (see Garratt, 1992; fig. 1.2). Sorbjan (1989) and Garratt (1992). For a general overview up to the late seventies consult McBean et al. (1979), and for a comprehensive introduc it may be even smaller, perhaps no more than tion to instruments and techniques for making 50-100 m and strongly influenced by internal measurements in the ABL see Lenschow (1986). wave motions. In recent times, Hess (1992) reviewed aspects of Over the open oceans, where low-level layer observations and scaling of the ABL, with atten cloud (stratus and stratocumulus) is prevalent, tion given to the analysis of observations based the ABL depth may be no more than a few on a framework of similarity scaling. The treat hundred metres and, in extratropical latitudes, ment for inhomogeneous surfaces was particu may have a structure quite similar to that over larly emphasised. Most recently, an up-to-date land. According to results from the north-east review of roll vortices in the ABL has appeared Atlantic (the JASIN experiment - Businger and (Etling and Brown, 1993), following on from that Charnock, 1983), the stability is near-neutral, with by Brown (1980). Etling and Brown (1993) review 92 J.R. Garratt / Earth-Science Reviews 37 (1994) 89-134 recent advances in our understanding of organ lidars {light radars) and a range of Doppler radars. ised large eddies in the ABL and on their role in These mainly involve transmitted acoustic, light the vertical transport of momentum, heat, mois or radio energy, and the detection of the scat ture and chemical trace substances within the tered energy due to (i) natural or artificial atmo lowest part of the atmosphere. spheric targets (dust, salt, rain) as in Doppler radars and lidars; (ii) clear air refractive index fluctuations, as in acoustic sounders and FM-CW 2. Observations radars. All of the above techniques are called active, since they involve transmission of acoustic Much of our knowledge of the ABL can be or electromagnetic radiation to the region of in related to observations of turbulent flows made terest and measuring the backscattered fraction both in the laboratory and in the atmosphere. that is returned to the instrument. In contrast, Such observations are important for understand passive techniques involve the measurement of ing ABL processes, for the validation of numeri radiation naturally emitted from the atmosphere, cal models and for placing suitable constraints on e.g. as in infrared radiometry. the behaviour of models. Remote sensing techniques are attractive Investigation of the structure of the lower part where in situ methods will not work or are uneco of the ABL has mainly utilized sensors located on nomical, and where the atmosphere interacts with tower structures. These towers have ranged in the transmitted or received radiation allowing the size from relatively short masts of a few metres meteorological variable of interest to be mea height, to the 200-300 m tall structures that may sured. Remote sensing utilises spatial averages, support a very comprehensive suite of instrumen usually over volumes, and so is especially suitable tation. For studies of the whole boundary layer, if spatial averages are desired. early approaches had to rely on balloon-borne A brief summary of important field experi instrumentation - free and tethered sounding ments may be of interest to the reader. In the balloons, or constant volume free balloons - early 1950's, the Scilly Isles field experiment of until aircraft techniques and facilities became Sheppard et al. (1952) focussed on vertical mo more widely available. mentum transfer processes over the ocean in the For measurements throughout the ABL, the Northern Hemisphere westerlies. Major land aircraft is now a well-established observational based ABL experiments commenced with the U.S. platform, although balloon platforms have been Great Plains observations in 1953 (Lettau and used with some success. One of the major advan Davidson, 1957), followed by the Australian Wan tages of using aircraft as an atmospheric mea gara (Clarke et al., 1971) and Koorin (Clarke and surement system is their mobility and capacity for Brook, 1979) observations in 1967 and 1974 re making extended line averages of turbulent quan spectively, and by the U.S. Minnesota experiment tities (most aircraft flights occur at speeds of in 1973 (lzumi and Caughey, 1976). Interspersed 1 about 80 m s - ). lt can be used for both vertical with the above were the surface-layer experi profiling well above the tower layer, and for hori ments of Swinbank and co-workers in the 1960's zontal traverses. In many field experiments the at Hay and Kerang in southern Australia (Swin two are combined. The main disadvantage of bank, 1968), and the Kansas experiment in 1968 aircraft measurements is the need to measure (lzumi, 1971). Major international efforts with a accurately the aircraft motion, since corrections strong ABL focus include BOMEX in 1969 to the velocity and temperature measurements (Kuettner and Holland, 1969), GATE in 1974 must be made for this motion. (Kuettner and Parker, 1976), AMTEX in 1974 In the last decade or so, ABL observations and 1975 (Lenschow and Agee, 1976), and JASIN have been enhanced by remote sensing tech in 1978 (Charnock and Pollard, 1983) - all relat niques. The newest instrumentation includes, for ing to the ABL over the ocean - HAPEX in example, a range of sodars (acoustic sounders), 1986 (Andre et al., 1986) and FIFE in 1987 J.R. Ga"att / Earth-Science Reviews 37 (1994) 89-134 93
(American Meteorological Society, 1990). In Stull For the non-homogeneous ABL, horizontal (1988), a list of major ABL field experiments can variations of properties in the ABL may arise
, be found with associated references. from non-uniform surface features (z0 T0 varia There have commenced or are planned a num tions due to soil moisture, albedo, land-water ber of major international land-surface / ABL ex contrast) or in the presence of orography, and be periments with emphasis on the mesoscale (see associated with thermally forced mesoscale circu Shuttleworth, 1991 for details). Of those com lations (e.g. sea breezes), internal boundary layers pleted, HAPEX-Mobilhy and FIFE are probably (at the coast) and slope flows. Important terms to best known, whilst EFEDA took place in central be included in a mathematical description of the Spain in 1991. There are several other experi ABL are those describing advection, e.g. in the ments affiliated to the International Satellite 2D system, terms such as ua;ax and wa;az. Land Surface Climatology Program (ISLSCP). Of For ease of discussion, we limit ourselves to those planned or recently commenced, HAPEX the 2D case, and start with the equations for the
Sahel in Africa (1992) and BOREAS in Canada set of mean field quantities (u, v, 0v, q, q 1) most (1994) are the most noteworthy (see ISLSCP, relevant to the cloudy, 2D ABL. The five basic 1993). prognostic equations are, in scalar form, au ;at + uau ;ax + wau ;az 1 3. Theoretical framework = -p- ap/ax + fv - a(u'w')/az (1) au ;at + uav ;ax + wav ;az 3.1. Basic equations = -p- 1ap/ay-fu - a(v'w')/az (2) The presence of clouds leads to considerable aevf at + uaev;ax + waev/az complications compared to a dry ABL because of 1 the important role played by radiative fluxes and = (pcp)- aRN/az - a(w'e~);az -AM/pcP phase changes. In a dry ABL, the turbulent struc (3) ture, the mean variables and their evolution in time are controlled by the large-scale external aq/at + uaq;ax + waq/az = -a(w'q');az + M/p conditions and by the surface fluxes. In a cloudy (4)
ABL, the surface fluxes may be important, but aq 1/at + uaq 1/ax + waq 1/az radiative fluxes produce local sources of heating = -a(w'qi)/az + Wq/P -M/p. (5) or cooling within the interior of the ABL and therefore can greatly influence its turbulent Here, M is the mass of water vapour per unit structure and dynamics. The state of equilibrium volume per unit time being created by a phase of a cloud-topped boundary layer (CTBL) is de change from liquid or solid, and Wq is a net liquid termined by competition between radiative cool water source /sink term (e.g. it is negative for net ing, entrainment of warm and dry air from above precipitation leaving the air parcel). The turbu the cloud, large-scale divergence and turbulent lence terms of the form aWs' ;az are of the great buoyancy fluxes. The phase change of water in a est importance in the ABL and represent the cloudy ABL introduces, in addition to 0v and q, effects of small-scale turbulent mixing on the a third variable in the form of the liquid water mean fields. Additional equations required to content q 1 (refer to the Table of symbols for close the set include the gas equation, the conti explanation). In the presence of liquid water nuity equation, and possibly a turbulent kinetic clouds within the ABL, equations for mean tem energy (TKE) equation relevant to the turbulence perature and water variables must take into ac closure, viz. count the possibility of phase changes, and in ae /at + uae ;ax + wae ;az clude a conservation equation for mean liquid = -u'w'au;az -v'w'av;az + (g/8v)w'e~ water content (either as a specific humidity or mixing ratio). -a(W'e+w'p'/p)/az-E (6) 94 J.R. Garratt / Earth-Science Reviews 37 (1994) 89-134
2 2 2 where e is the mean TKE (e = u' + v' + w' ) and temperature, T *, are used where w! = and E is the molecular rate of dissipation of TKE. g(w'6~) 0 h/E>v and T; = E>v(w'@~)~/gh where h Note that on the RHS of Eqs. 1 to 5 a;ax terms is the depth of the CBL. In highly stable condi in turbulent quantities are neglected (this is an tions, weil away from the surface (z > L), local excellent assumption for horizontal length scales scaling is more appropriate and relies on scaling » vertical scales). We can reduce this set so as to turbulence properties using local flux values, not simplify the equations if we confine the problem surface values. to a (i) clear, 2D, (ii) clear, lD (homogeneous) With u * 0 serving as an intrinsic internal veloc ABL. The following simplifications can be made ity scale, suitable scaling of Eqs. (1) and (2) when not dealing with the cloudy, 2D ABL: (i) in simplified to the clear, lD case leads to the logarithmic wind law in neutral conditions (Ten the clear case, q 1 = M = Wq = O; (ii) in the lD case, a;ax = w = 0. nekes, 1982), whilst Monin-Obukhov theory in The above equations provide a sound frame troduces the length scale L to take account of work for study of the physical aspects of the buoyancy effects. Thus, in the surface layer, over ABL, and of its modelling. In solving the set Eqs. surfaces that are not too rough {typically, this (1)-(6), considerable attention must be given to implies values of the aerodynamic roughness the closure problem - this involves parametrisa length z 0 no more than a few tens of centime tion of the covariance or higher-order terms. In tres) and at heights where z/z0 ~ 10-100 (e.g. addition, two Options so far as ABL study is Garratt, 1980, 1983), concerned are generally available (see Clarke, (kz/u* )au;az =
k( (Jva - eo) /0 * 0 =In( z/zT) - 'l'H( 0 (8b) 3. 2. Surface fluxes and profile relations where subscripts "a" and "o" refer to the air and surface respectively, and for humidity Eqs. {Tu), The surface values of the covariances are of (Sb) apply, with 0v or 0 replaced by q, and zT particular importance, both as a basis for scaling replaced by zq. The functions
a set of turbulent scaling parameters used in heights (z/z0 < 10-100) within the roughness 2 Monin-Obukhov theory as follows: u * 0 = sublayer (sometimes called the transition layer), 1 1 2 1 [(u'w')~ + (v'w )~] ! = -(u'w ) 0 if surface-wind the and 'I' profile functions must be modified
coordinates are used, with v = 0. Here, u * 0 is (e.g. Garratt, 1980, 1983; Raupach et al., 1980). the surface friction velocity. Temperature and The details are not pursued here. In general, the
humidity scales, 0 * 0 and q * 0 are defined by roughness lengths z 0 and zT (or zq) are not 1 equal; discussions on their relative values can be 0* 0 = -(w'0~) 0 /U* 0 and q*o = -(W q')0 /U*O' with a stability length scale L (the Obukhov found in Brutsaert (1982; eh. 4), Garratt (1992; Iength) defined by L = u*~0v/(kg0* 0 ) with k eh. 4); also see, Garratt and Hicks (1973), Bel the von Karman constant. The above scaling ap jaars and Holtslag (1991), Betts and Beljaars plies in neutral, moderately unstable and stable (1993), Garratt et al. (1993) and Mahrt and Ek conditions, but in highly convective conditions, (1993). particularly weil away from the surface, mixed The above equations form the basis of the bulk
layer or free convection scales for velocity, w *, transfer relations for surface fluxes ( 'T 0 is the J.R. Garratt / Earth-Science Reviews 37 (1994) 89-134 95
functions of h/z and µ,. Note that these curves surface stress, H0 the surface heat flux and E0 is 0 the surface evaporation), written as follows strongly reflect observations since the functions A and B are directly inferred from these for a (9) range of experiments. 1 Recently, Walmsley (1992) has proposed a new H 0 /pcP = (w 8~) 0 = -u*°8* 0 = CHu(80 - Ova) ABL resistance law for neutrally-stratified flow (10) which, in essence, has strong similarities to that generally accepted. However, comparison against E 0 /p = (w'q') 0 = -u *°q* 0 = CEu( qo - qa) (lla) observations reveals no significant improvement to the geostrophic drag law as represented by
=(qi-qa)/(r5 +ra) (llb) Eqs. (12) and (13). where q* is the saturated humidity, r is a sur 5 3.3. Vertical variation of fluxes and the K approach face resistance, the aerodynamic resistance ra = 1 (CEu)- and the coefficients C , CH, CE are 0 The covariances (vertical fluxes) can be readily deduced as functions of z/z0 and ' from manipulation of Eqs. (7) and (8). By setting the '1' parametrised as functions to zero, the neutral values C 0 N, CHN, Tx = -pu'w' = pKmau;az (16) CEN are also defined. In the above, surface-layer wind coordinates are used, so that u is the sur ,,.Y = -pv'w' = pKmaü ;az (17) face-layer (total) wind. H = pcPw'8~ = -pcpKHaov/az (18) The above transfer relations can be written in terms of ABL-averaged quantities, incorporating E = pw'q' = -pKwaq_;az (19) ABL transfer coefficients (e.g. geostrophic drag where KM is the eddy viscosity, KH the eddy coefficients). The result is (e.g. Garratt, 1992) thermal diffusivity and K w is the eddy diffusivity kug/u* 0 =ku/u* 0 =ln(h/z0 )-A(µ,, Mx, R) for water vapour. In contrast to the molecular (12) case, the eddy diffusivity is not a property of the fluid but may be a function of many quantities, kDs/u * 0 = kD /u * 0 - ku * 0 /fh including position and flow velocity. Here simple forms for K may be used (e.g Garratt, 1992; eh. = -B(µ,, MY, R)sign(f) (13) 8), or calculated by consideration of the TKE 1 2 equation (Eq. 6) with K ex e 1 /, l being a mixing k(Ov - 0 )/8* = ln(h/zT) - C(µ,, M, R) 0 0 length. Altematively for numerical purposes, we (14) can carry equations for the covariances them selves (second-order closure). k( q - q ) /q* =In( h/zq) - D(µ,, M, R) (15) 0 0 For modelling purposes, vertical fluxes of mo which, in analogy with the surface-layer manipu mentum, heat and water vapour and other scalars lation of Eqs. (7) and (8), leads to relations for can be evaluated using either local or non-local the surface fluxes in terms of geostrophic or closure schemes. lt is apparent that convective boundary-layer transfer coefficients (Garratt, mixed-layer turbulence is one of the more diffi 1992; eh. 3). Here, M is a baroclinity parameter cult regimes to model using local closure, because 2 whose magnitude is defined as 1M1 = Ml + M;, the important eddy length scales are comparable
µ, = h/L and R = 1f1h/u* 0 • Variables with the with the thickness of the CBL. Rather, the valid circumflex are vertical averages. Functional forms ity of K-theory and mixing length assumptions for the A, B, C and D can be found in the requires eddies that are much smaller than the literature (Brutsaert, 1982; Garratt, 1992), with scale of interest. In the presence of large eddies Fig. 2 showing the ABL drag coefficient and and intense mixing, the CBL can often be as cross-isobar flow angle (from Eqs. 12 and 13) as sumed to be well-mixed, though on some occa- 96 J.R. Garratt / Earth-Science Reviews 37 (1994) 89-134 sions over land and sea it is distinctly unmixed in The application of K-theory to the 0v profiles one or more of the properties. This lack of mixing in particular has obvious limitations, since the is apparent in the profiles of potential tempera heat tlux may often appear to be countergradient ture and specific humidity shown in Fig. 3. on average, and therefore will require K to be
2
10
'---W.'----L-----L----L----1---...J...--- , (il/~ ) 0 log 11 0 3 4 5 6 7 (a)
40
30
20
10 -1
-100 0 .____.,""'--~---L---__.,------======-10"'------1---- log"' (h/::„l 3 4 5 6 7 (b)
) Fig. 2. Variations of (a) the boundary-Jayer drag coefficient (Cs) and (b) cross-isobar flow angle (a0 as functions of h/z0 and stability parameter µ, (negative µ, corresponds to unstable conditions) - see Eqs. (12) and (13). Here Cs= u * ~/(ui + vi) and tan(a0 ) = ug/us. See Garratt (1992; fig. 3.12). J.R. Garratt / Earth-Science Reviews 37 (1994) 89-134 97 negative, and in parts of the profile (in mid-re from bottom-up (through (w'c')J and top-down gions of the CBL usually) may even be singular. diffusion (through (w'c')h).
The nature of the scalar mean profiles in the If we associate eddy diffusivities Kb and K 1 CBL determines whether local closure will be with bottom-up and top-down diffusion respec appropriate. The presence of entrainment pro tively, then it is readily shown under the assump foundly influences the mean temperature and tion of superposition of properties that scalar mixing ratio profiles ( 0v and q, for exam 1 = ) ple) through what is referred to as "top-down ac;az -((w'c 0 (l-z/h)/Kb diffusion". This type of diffusion differs from the + (w'c'h(z/h)/K ) (22) familiar "bottom-up" diffusion because the buoy 1 ant forcing of the convective turbulence occurs which shows that with vertical asymmetry. Numerical simulations (a) If both fluxes are of the same sign, the reveal that the flux-gradient relationship (involv vertical gradient of c will tend to be !arge - this ing an eddy diffusivity) for a passive scalar in a is characteristic of the humidity profile. mixed layer depends on the boundary from which (b) If the fluxes are of opposite sign, the verti the flux originates (e.g. Wyngaard and Brost, cal gradient may be close to zero - this occurs 1984). To see this, and how it relates to the K with the temperature profile. problem, consider a passive, conservative scalar We also have for K, using Eq. (22), whose mean concentration c satisfies K= -w'c'/(ac;az) ac;at = -a(w'c');az (20) where w' c' is the scalar flux. If temporal changes in the scalar gradient ac ;az are assumed negligi ble then the flux profile is linear, and described (23) by Since K depends on the fluxes, this shows that w'c'= (w'c') (l-z/h) + (w'c')h(z/h) (21) 0 the closure is inherently non-linear with 0 < K < so that the flux at any level has contributions oo. The poor behaviour in K can occur even when
(a) (b) 1.5 ::lh
0.5 8 q
0 0 0 -20 0 20 -20 0 20
Fig. 3. Mean 0 and q profiles in unstable conditions in (a) Colorado and (b) south-eastern Australia; in (a) ABL height h = 1710 m and in (b) h = 915 m. Subscript s denotes "screen-level" values. After Mahrt (1976). 98 J.R. Ga"att / Earth-Science Reviews 37 (1994) 89-134
Kb and K 1 are well-behaved, but different (see Holtslag and Moeng, 1991). ://, For the determination of the c profile, Eq. (22) requires knowledge of the eddy diffusivities 0.8 0.6 before integration can be made. Tue form of the diffusivities has been obtained from large-eddy simulations (e.g. Moeng and Wyngaard, 1984;
Holtslag and Moeng, 1991), with K 1 = 1.4 w * z(l -z/h)2 and Kb=w*h(l-z/h)/gb. The top down diffusivity is well-behaved and positive 0.6 throughout the convective boundary layer. In contrast, the mean gradient function gb is posi tive in the lower regions of the mixed layer and near the surface, but changes sign near z/h = 0.6 and is negative above (Moeng and Wyngaard, 1984). This is a clear illustration of the break 0.4 down of the Kb model, with K having a singular ity in the mid-regions of the mixed layer. The result for K shows that the bottom-up process generates a countergradient flux in the upper mixed layer. With suitable bottom boundary conditions, the 0.2 c profile for various values of the ratio R = 1 (w' c')h/(w' c )0 is shown in Fig. 4. Small negative R values give a profile form similar to that for -4 -2 0 potential temperature (e.g. with R = - ß = - 0.2) [c -c(O. I )]w*/(w'c') and illustrate the countergradient heat flux prob 0 Iem in the upper regions of the CBL. Positive R Fig. 4. Scaled concentration profiles for four values of the values yield distinctly "unmixed" profiles, as oc parameter R - see Garratt (1992; fig. 6.10). Here c(O.l) is curs with humidity at times. the concentration at z / h = 0.1, calculated from the inte grated form of Eq. (22). Local closure has problems because it relates the flux of a quantity at a given Ievel to properties at that Ievel. The alternative approach uses non Iocal closure, either in the form of a non-Iocal K (e.g. Holtslag and Moeng, 1991; also see Holtslag velocity, temperature and humidity. This can be and Boville, 1993) as discussed above or in the illustrated for temperature, where, form of integral (e.g. Fiedler, 1984) or transilient methods (Stull, 1984, 1993; Stull and Driedonks, (24) 1987). This, and other analogous equations for slab properties, contain entrainment fluxes at h. These can be formally related to the jumps at z = h, 3.4. Slab approach and rate equation for ABL (e.g. aevh for virtual potential temperature) so depth that for heat transfer,
(25) Where the ABL is treated as a slab, of depth h, vertical integration of Eqs. (1)-(4) in the lD Here the mean vertical velocity at z = h is non-cloudy case leads to equations for the slab represented by wh. J.R. Ga"att / Earth-Science Reviews 37 (1994) 89-134 99
assume wh and y are given, and the surface heat ::/h 8 • flux is specified, closure is achieved by relating the entrainment heat flux to that at the surface. 1.2 Such a relationship may be based on the TKE equation (e.g. see Manins, 1982), and usually ------0 leads to the simple form (~h = -ß~ 0 • 0 0.8 Observations above surfaces of low roughness suggest that in convective conditions ß == - 0.2 ± 0.6 0.1 (Stull, 1988, eh. 11 - see also Manins, 1982), with tentative evidence of values closer to - 0.5 0.4 over large roughness surfaces (e.g. Clarke, 1990). The observed heat-flux profiles in Fig. 5 are 0.2 consistent with the low-roughness value of - 0.2, 0 0 and model results tend to confirm this too (e.g. -0.5 () 0.5 Nieuwstadt et al., 1993 for LES models; Chrobok w'e;;(w'e;.)o et al., 1992 for lD and 2D boundary-layer mod Fig. 5. Vertical variation of the normalised sensible heat flux els). based on experimental data from aircraft - see Garratt (1992; fig. 6.2). Expressions for h range from the simple en croachment-only form, where ah;at = ~)h/
y8 h, to the more detailed descriptions e.g. that of As the slab properties change, so do the re Deardorff (1974). Observed and siinulated ABL spective jumps across the top of the CBL. In the growth under convective conditions are shown in case of temperature, this is given by Fig. 6, revealing the relative behaviour of a range of formulations. a( ii8vh) ;at = ye( ah ;at - wh) - aevm/at, (26) Applications to the nocturnal boundary layer
where y8 is the vertical 0 gradient above h. Eqs. (NBL) are not so well developed, but the ap (24)-(26) need to be solved for h and evm; if we proach is equally as valid, except that h is differ-
1000 h Im l
500
0800 1200 1400 1600 1800 Locol time Fig. 6. Observed (•) and predicted ABL growth in unstable conditions - from Manins (1982). Predictions are from an encroachment model, with (- - -) and without (-) thermal wind; a flux-ratio closure model (- - - ) and from a Froude dynamics model (--). 100 J.R. Garratt / Earth-Science Reviews 37 (1994) 89-134 ent. A suitable rate equation was described by cal expressions in both stable and unstable condi Nieuwstadt and Tennekes (1981), tions can be found in table 4.2 of Sorbjan (1989). In the case of
10· 10-·..._~~~~~---~~~~~~---~~~~~---~~~~~--- 10-· w-• 10• Fig. 7. Dimensionless velocity and temperature gradients ( Spectra and cospectra Tue curves of Kaimal et al. (1972) based on observations made during the Kansas experiment (lzumi, 1971) are still the most comprehensive, covering both unstable and stable conditions. 4.2. ABL and mixed layer over land 10• Mean profiles Validation of ABL similarity theory was pro 10• 10 vided in an analysis of Wangara observations by Yamada (1976), who determined normalised wind Fig. 8. As in Fig. 7, for 0 0 1.0 0.5 oL---1.___.1_.L__L_LJ,dd:~~~~-L..L...w..i..1 10-2 10-1 w'2 / W. Fig. 9. Dimensionless profile of vertical velocity variance in convective conditions for a range of observational data and specific functional forms - over land and sea, and in the laboratory. See Hess (1992) for details. 102 J.R. Garratt / Earth-Science Reviews 37 (1994) 89-134 Brutsaert, 1992). They are useful for estimation profiles, and in Caughey et al. (1979) and Sorb of regional surface fluxes. jan, (1989, figs. 4.23, 4.24) for variances. For the neutral ABL, LES results for the covariance pro Variances files can be found in Mason and Thomson (1987), Variances and higher-order moments have as summarised in Garratt (1992, p. 48). been measured throughout the ABL in both sta ble and unstable conditions, and modelled suc Spectra / cospectra cessfully using LES models in neutral conditions. Suitably nondimensionalised, the observed Systematic behaviour of a range of turbulent spectra and cospectra described by Kaimal et al. statistics is evident if the appropriate scaling is (1976) (also see Sorbjan, 1989; p. 143) for the used (e.g. see Hess, 1992). Thus, in unstable unstable mixed layer, and by Caughey et al. (1979) conditions, this is achieved using convective scal for the stable ABL (also see Garratt, 1992, p. 82) ing; typical results can be found in e.g. Sorbjan represent the most comprehensive set of curves (1989, figs. 4.32, 4.33) and Garratt (1992, p. 77) to date. for profiles of uw (also u and v components) and ur. Composite sets of data for both land and sea have been discussed in Hess (1992) and profiles 5. Marine/ cloudy ABL of uw and ur suitably nondimensionalised are shown in Figs. 9 and 10. Also shown are labora In many situations, clouds exist within or im tory data which, in the lower half of the CBL, mediately above the ABL over the ocean. Put agree well with atmospheric observations. Differ another way, most observations of the cloud ences aloft are probably related to the effects of topped boundary-layer (CTBL) have been ob clouds and vertical wind shear in the observa tained over the ocean where ABL clouds, particu tions, both which affect entrainment dynamics. larly in the form of stratocumulus, are common In stable conditions, local scaling is required, and often persistent over many days at any given and typical results appear in Caughey et al. (1979) locality. Experimental studies have tended to fo and Nieuwstadt (1985) for the covariance (flux) cus on three boundary-layer types 0 0 0 \ D 0 D 0 0 1.0 0 - "11·-·.: „„------D z/z.1 0.5 ° 0 10 2 7fl I T• Fig. 10. As in Fig. 9, for dimensionless temperature variance. See Hess (1992) for details. J.R. Garratt / Earth-Science Reviews 37 (1994) 89-134 103 (i) The tropical and sub-tropical marine ABL, radiative cooling. The cloud forms part of a de usually associated with cumulus (Cu) and stra tached turbulent layer, whose base is separated tocumulus (Sc) clouds. from the true ABL (i.e. the mixed layer adjacent (ii) The mid-latitude marine ABL (open ocean), to the surface). Surface fluxes are usually quite associated with Cu and Sc douds. small; (iii) Tue sub-tropical and mid-latitude marine (iii) a fully mixed single layer, or layer in which ABL generated during cold-air outbreaks to the two or more separated turbulent layers exist, east of the continental land masses, associated which supports multiple cloud layers (not neces with Sc and Cu cloud formation. sarily all being stratocumulus). If more than one The marine ABL during cold-air outbreaks cloud layer exists, the intervening layer of clear can be highly unstable with latent heat fluxes well air may serve as a barrier to vertical mixing. in excess of the sensible heat fluxes (e.g. Francey and Garratt, 1978; Chou, 1993). Absolute values Fully-coupled single layer of the evaporation rates can reach nearly 1000 W This layer may be predominantly surface driven m-2 in extreme conditions. due to heating or friction associated with strong Major boundary-layer cloud types are stratocu winds, or it may be cloud-top driven due to strong mulus, stratus and fair-weather cumulus, and radiative cooling. In the latter case, cloud-top these serve as suitable foci for a discussion of the radiative cooling may be intense enough, in situa cloudy ABL. tions where surface fluxes are small, to support a convectively mixed layer extending from the cloud 5.1. Stratocumulus top to the surface. Examples of mean profiles of thermodynamic Observations reveal the following broad cate variables are shown in Fig. 11 for a thick stra gories of stratocumulus-topped ABL tocumulus layer. The ABL is fully mixed because (i) a fully mixed ABL, with a single elevated of the combination of non-negligible surface cloud layer fully coupled with the sub-cloud layer. fluxes in moderate to strong winds, and the exis This type of CTBL is usually associated with tence of cloud-top radiative cooling. Through the either significant surface heat fluxes or strong cloud layer 0 or ®v increases according to moist radiative cooling at cloud top; adiabatic ascent, whilst ee is approximately con (ii) a single cloud layer decoupled from the stant. Of particular significance are the large surface, in which mixing is the result of cloud-top jumps in properties across the ABL top. p (hPa) 880 920 960 1000 • 280 3CXl 2 8 295 315 -1 O(K) q (g kg ) Oe (K) Fig. 11. Vertical profiles of several thennodynamic variables in the CTBL offshore from California in a situation of thick continuous stratocumulus cover (denoted by pecked lines). See Garratt (1992; fig. 7.3) for further details. 104 J.R. Garratt / Earth-Science Reviews 37 (1994) 89-134 j 1 II 1 j 1 1 1 1 j cause of small surface heat fluxes related to (m) near-neutral stability conditions. The local con densation level coincided with the top of the 1400 shallow ABL and, in the presence of conditional 1200 - ----=-----:;,,- ---- instability above (il0e/ilz < O), a field of cumulus clouds extended into the overlying stratocumulus. 1000 Even though the cumulus was observed to be intermittent and irregularly distributed in the 800 horizontal, it acts to transport mass (water vapour) from the ABL towards the stratocumulus layer. 600 400 5.2. Stratus 200 Observations of stratus cloud layers have tended to concentrate on the Arctic region in 0 summertime, with very few comprehensive data 300 305 310 288 2 4 6 8 (K) (KJ (g kg-1) from elsewhere. Nevertheless, observations to date seem to suggest at least two main categories Fig. 12. Vertical profiles of several thennodynamic variables of stratus-topped ABL in the CTBL over the north-east Atlantic for a decoupled (i) a fully mixed ABL, with a single elevated stratocumulus layer (top pair of pecked lines). Here, q1 = q + q 1• The bottom pecked curve corresponds to the top of the cloud layer fully coupled with the sub-cloud layer. Ekrnan layer. See Garratt (1992; fig. 7.5). This type of CTBL is associated with moderate to strong winds and negligible buoyancy effects. The vertical velocity variance typically shows little en Decoupled single layer hancement near cloud top so that buoyancy ef An example of a decoupled stratocumulus layer fects are probably not significant (Nicholls and is shown in Fig. 12. The cloud layer is barely 150 Leighton, 1986); m thick and located beneath a sharp inversion. (ii) an ABL above which several low-level cloud The dry adiabatic lapse rate and fluctuations in layers exist. the wind components indicate that turbulent mix ing extends beneath cloud base and into the 5.3. Cumulus sub-cloud layer. This whole mixed region has essentially the same values of 0e and total hu Boundary-layer cumuli are cumulus clouds midity ql' and is decoupled from the true ABL or whose vertical extent is limited by the main subsi surface-based Ekman layer (of height "" 0.2 dence or capping inversion. Such clouds are gen u * / 1 f 1 ), in which turbulence is maintained to a erally non-precipitating, and are usually referred height of about 280 m by surface-related pro to as fair-weather cumulus, cumulus humilus, or cesses, as a result of an intervening stably strati trade-wind cumulus. Boundary-layer cumuli are fied Jayer. often composed of an ensemble of forced, active or passive clouds (Stull, 1985) with the deepest of Multiple layers these, the active clouds, ascending no more than Though not shown in diagrammatic form here, a few km above the main inversion. a good example of multiple layers can be found in Forced cumulus clouds form within the top of Nicholls (1985). In this case (from the JASIN mixed-layer thermals while the thermals are over experiment in the north-east Atlantic), coupling shooting into the overlying stable layer. These between a stratocumulus layer of thickness about thermals remain negatively buoyant during the 350 m and the surface-based ABL of depth 400 m overshoot, even though there is warming from was generally weak and intermittent, mainly be- condensation. Thus, the visible thermal rises J.R. Garratt / Earth-Science Reviews 37 (1994) 89-134 105 above the local condensation level, but fails to clouds? lt is apparent from both observations reach its level of free convection (where a parcel (usually involving aircraft) and modelling studies will have positive buoyancy relative to the envi that, where clouds are present and particularly ronment). Active cumulus clouds ascend above fair-weather cumulus, the moisture flux near the level of free convection and thus become cloud base is a significant fraction of the surface positively buoyant. Because of this they tend to evaporation. In contrast, entrainment moisture ascend to greater heights than forced clouds and fluxes near the mixed-layer top in clear skies are develop circulations that are independent of the usually small. This is evident in the observations original triggering thermals. In the forced case, described by LeMone (1980) - see Fig. 13 - little or no transport (venting} of ABL air into the which also reveal small virtual heat fluxes at free troposphere occurs whilst, for the active case, cloud base. These data, in fact, illustrate the transport does take place. Finally, passive clouds slightly unstable conditions often found over trop are the decaying remnants of formerly active ical and sub-tropical open ocean areas, under clouds and, in general, have no interaction with fair-weather conditions, where h/L is typically in the ABL, except that they shade the ground (as the range - 10 to - 1. The surface sensible heat do other non-ABL clouds). flux and virtual heat flux are small, but evapora How is the turbulence structure and the verti tion rates relatively large; the radiative cooling cal fluxes affected by the presence of cumulus near cloud base tends to maintain a sub-cloud 253 600 „ • • .. . ..,.„ ;: (m) 258 600 -WO ~00 () ...___ ...... ___ _ • -0.01 () 0.01 () Ü.Ü5 X 10-' w'T: (m s-1 KJ Fig. 13. Profiles of heat and moisture fluxes for two GATE days, the upper for a cloudy ABL and the Jower for clear skies. Pecked cuives show the approximate level of cloud base and/or ABL top - based on LeMone (1980). 106 J.R. Garratt / Earth-Science Reviews 37 (1994) 89-134 layer cooler than the sea (Betts and Ridgway, and there are additional sources and sinks of 1989). TKE associated with radiative effects and phase We can summarise important observational as changes. lt is unlikely that any single method of pects of the stratus- and stratocumulus-topped scaling data obtained in convective layers contain ABL as follows ing stratocumulus will yield results which have (i) As with the clear ABL, strong jumps in E>v universal applicability since the TKE balance is and q occur across cloud top (defined as the potentially much too complicated. At other times, ABL top), with 0e often showing both positive particularly under strong wind conditions, the and negative jumps. The change in 0e across the CTBL is indistinguishable in a dynamical sense cloud top is an important parameter characteris from the neutral ABL without clouds. ing entrainment, and is closely associated with strong cloud-top radiative cooling and cloud-layer stability. 6. Non-bomogeneous abl over land (ii} In the presence of strong winds, large sur 6.1. Introduction face fluxes or large entrainment effects (related to cloud-top radiative cooling}, the CTBL is well Significant horizontal variations in ABL struc mixed, with the cloud layer coupled to the sub ture and properties usually arise from hetero cloud layer and to the surface. The top of the geneity in surface properties, though they may ABL usually corresponds with cloud top. also be induced by horizontal variations in cloudi (iii} In light to moderate winds, with negligible ness. Even at the microscale (order of tens of surface fluxes of sensible heat and momentum, metres), surface variations can induce an organ the cloud layer is decoupled from the surface. In ised response in the form of internal boundary this case, the CTBL is not well mixed. layers, though these are usually confined to the (iv) The CTBL may consist of more than one lowest few metres or tens of metres. lndeed, for cloud layer. surfaces whose heterogeneities are disorganised (v) Precipitation in the form of drizzle may at length scales somewhat less than 10 km, there occur, with a tendency towards stabilization of is no apparent organised mesoscale response in the CTBL. The stability is affected because the the lower atmosphere. For surfaces that have cloud layer is associated with the latent heat of organised heterogeneities at scales greater than condensation, whilst in the sub-cloud layer evapo 10 km, there may be an organised mesoscale rative cooling occurs. response throughout the ABL and above. Fig. 14 (vi) In some instances, the cloudy stratiform illustrates this schematically. The response of ABL is convective in nature and has a number of area-averaged evaporation at these two scales to similarities with the clear CBL heated from be the imposed available energy, surface and atmo low. Generally though, the main source of buoy spheric conditions may also be different (Shut ancy for this CTBL is at the entraining boundary tleworth, 1991). TYPEA TYPEB ,------..... ------m•oet+m 22.A f :4~000bMOOM~~~--:~~::~~:J ABL L..--1__ ™_1okm_I 1 10km 1 Fig. 14. Classification of land surface types (two) based on Shuttleworth (1991). In type A, the surface structure is disorganised at length scales of 10 km or less; in type B, there is organisation at length scales greater than 10 km. J.R. Ga"att / Earth-Science Reviews 37 (1994) 89-134 107 The problem of surface heterogeneity can be work of the World Climate Research Programme considered on several scales. On the smallest (WCRP) of WMO. The focus was on a program scale, the response is confined to the surface of observations and modelling aimed at studying layer and is dominated by local advection effects. the water budget and evaporation at the scale of The ABL as a whole can then be taken as hori a GCM grid. One main objective was to provide a zontally homogeneous. At the next larger scale, data base against which parametrisation schemes the effects of the surface disturbance can be for the land surface water budget can be tested followed downwind until the whole of the ABL is and developed. Several aspects of the experiment affected. This is the case of mesoscale advection. and results have been reported, including an Finally, at this scale and larger, the ABL adjusts overview (Andre et al., 1986), initial results to the new surface conditions and, under opti (Andre et al., 1988); regional estimates of fluxes mum large-scale conditions, mesoscale circula and comparisons of aircraft fluxes with surface tions develop (sea breezes, lake breezes, inland measurements (Andre et al., 1990); application of breezes) which interact further with the horizon a mesoscale model using an advanced surface tally varying ABL. Surface heterogeneity can also parametrisation (Bougeault et al., 199la, b; Noil be identified with changes in surface elevation, as han et al., 1991); canopy processes and applica occurs in regions of hills, ridges and valleys. Un tions (Mehrez et al., 1992a; Pinty et al., 1992); der these conditions, large scale changes in the calibration of a bare soil energy model (Mehrez pressure field are induced producing a complex et al., 1992b). In addition, the complete data set response in the ABL flow. is readily available from the French group (see e.g. Andre et al., 1990). 6.2. Major field experiments There is no doubt that the HAPEX-Mobilhy experiment has resulted in a high quality compre HAPEX-Mobilhy took place in south-west hensive data set focussing on aspects of ABL France for 3 months in 1986 under the frame- structure on selected days, and surface energy 1.5 1.26 ,r--MS 1989 a H 0.5 o..._~~~~~~~-----~~~~~~~_._~~~~~~~~~ 10 100 1000 10000 rc (sm-1) Fig. 15. Dependence of the Priestley-Taylor coefficient a upon canopy stomatal resistance rc; a value of 1.26 is the accepted value for homogeneous saturated surfaces. Of particular interest is the mean values for the FIFE (F) and ARME (A) experiments - both type A surfaces - and for the HAPEX (H) experiment - a type B surface. Other observational data, and the curve, are from McNaughton and Spriggs (1989). After Shuttleworth (1991). 108 J.R. Garratt / Earth-Science Reviews 37 (1994) 89-134 and water budgets at several locations over a significant mesoscale advection and circulations period of 3 months. The data have proved invalu reducing the regional evaporation relative to the able for validation and calibration of new surface net available energy. This has important implica and ABL schemes used in mesoscale and climate tions for the parametrisation of regional energy models. fluxes. FIFE - the first ISLSCP field experiment (see e.g. Sellers et al., 1988) took place in east 6.3. Internat boundary layers Kansas in 1987 and 1989, with major observing periods focussed within a 15 X 15 km region of General Konza prairie. One of its major objectives con Internal boundary layers (IBL's) in the atmo cerned the use of satellite observations to infer sphere are associated with the horizontal advec climatologically significant land-surface parame tion of air across a discontinuity in some property ters. of the surface. Studies usually specify the surface The first set of major results has been de forcing in terms of a step change in surface scribed recently in a special issue of Journal of roughness, temperature or humidity, or in the Geophysical Research (Vol. 97, No. Dl 7, 1992). surface flux of heat or moisture. Several classes Studies have focussed on six areas, as follows: (i) of IBL problem can be readily identified in the ABL, with particular emphasis on airborne fluxes literature, concerning both laboratory and atmo and measurement techniques; (ii) Surface Fluxes, spheric flows, and related both to the thermal with emphasis on heat, moisture and C02 fluxes stability characteristics and horizontal scale of the and sensor performance; (iii) Correction and cali flow. bration of satellite sensor data; (iv) Surface radia Earlier work (from the late fifties to the mid tion and biology; (v) Soil moisture; (vi) Integrative seventies) was concerned mainly with the prob science. The results to date have led to improved lem of neutral flow across a step change in sur understanding of the exchanges between the land face roughness. Later in this period, attention surface and atmosphere at the local scale (1-100 turned to the effects of thermal stratification m) and at the intermediate scale (100 m to 15 upon the flow across a roughness change, and to km). In addition, much is being learnt on the the growth and structure of the IBL related to application of remote sensing science at these step changes in surface heat flux and tempera two scales, and the use of remote sensing and ture. Throughout this period the main emphasis modelling at the intermediate scale. The applica was on the small-scale aspects of the flow i.e. on tion of FIFE data to validation of model simula relatively small downwind fetches and where, in tions has recently received some attention, with the atmosphere for example, the IBL was con emphasis on aspects of surface and boundary fined to the atmospheric surface layer of the layer behaviour (Betts et al., 1993). advected planetary boundary layer. Such a con Both HAPEX and FIFE observations of evap straint allowed for several simplifying assump oration have been used by Shuttleworth (1991) to tions in analytical and numerical treatments, and illustrate the impact of land heterogeneity upon confined the downstream fetch to maximum val regional evaporation, and its relation with avail ues of about 1 km. able (net radiant) energy and bulk canopy resis In more recent times (the mid-seventies to the tance (rJ. Fig. 15 shows the dependence of the present), the emphasis moved from the microme Priestley-Taylor parameter a upon rc, based on teorological (and local advection) problem associ observations and predictions of a simple lD ABL ated with relatively small fetches to mesoscale model with a range of initial conditions (note that advection, and in particular to the effects of a is directly proportional to AE/(RN - G); the buoyancy and the development of the thermal reader is referred to Priestley and Taylor, 1972). IBL towards an equilibrium boundary layer far lt is thought that the HAPEX anomaly reflects downstream of the leading edge (relatively large the impact of larger-scale heterogeneities, with fetches). The main topic has been on the growth J.R. Garratt / Earth-Science Reviews 37 (1994) 89-134 109 of the convective thermal intemal boundary layer pollution problems in the coastal region. Al (TIBL) at the coast, mainly because of practical though the advection of air across the coastline concern on the influence of the IBL on coastal relates to both surface roughness and tempera pollution from industrial sites located in the ture changes, the primary consideration is often coastal region. Less attention has been given to the response and growth of an IBL to a marked the parallel problem of a stably stratified IBL, step-change in surface temperature. lt is recog whose full development may require fetches of nised that in real world situations roughness several hundreds of km. For a recent review of changes also are present. The emphasis generally the topic, and many relevant references, see Gar has been on the growth equation for the IBL ratt (1990). We limit the present discussion to the depth in both convective and stable conditions, problem of mesoscale flow and IBL growth as and the factors affecting the depth. Definition of schematically shown in Fig. 16. the thermal IBL is usually in terms of the marked Most mesoscale studies focus on the structure changes in temperature and humidity gradient and growth of the TIBL, and stem from the found near its top, and this seems to be most perceived relevance of the IBL to diffusion and appropriate for the stable case. Observations of (b) Land Fig. 16. (a) Schematic Elv profiles under conditions of onshore flow and formation of a convective IBL of depth hb(x ); (b) As in (a) but for offshore flow and formation of a stable IBL. 110 J.R. Garratt / Earth-Science Reviews 37 (1994) 89-134 convective TIBL structure (usually for fetches tram, 1977). For practical application, the heat less than about 50 km) during sea-breeze events flux may be parametrised in terms of a bulk and for general onshore flow conditions have relation using suitable transfer relations. found that the top defined from the 0 profiles Most observational studies have been confined can be weil below the level of minimum turbulent to inland fetches in the range 5 to 50 km, since kinetic energy. Such observations of the mean this is the typical distance required to approach and turbulence structure reveal the IBL as a the full depth of the ABL. Growth of the TIBL fairly well-mixed layer, with large horizontal gra over this fetch range has generally supported the 1 2 dients in 0v and turbulent kinetic energy for relation hb = ax 1 , with hb and x both in me several tens of km inland from the coast. In tres, and with a in the range 2 to 5 approxi contrast, turbulence closure-model studies of the mately. These numerical values can readily be TIBL structure and growth (e.g. Durand et al., shown to be consistent with that implied in Eq. 1 1989) reveal that the TIBL top defined in terms (29a). If we take ß = 0.2, and y 8 = 0.01 K m- (a of the 0v inversion layer tends to coincide with rather !arge value), then the level of near-zero heat flux and turbulent (29b) kinetic energy, and with maximum temperature 1 = variance. Realistic values of um 5 m s- and H 0 = 100 W m- 2 give a::::: 2, whilst um= 2 m s- 1 and 2 = The conuective thermal IBL H 0 400 W m- give a::::: 7. Decreasing y 8 to a The convective IBL (depth hb) usually forms low value of 0.001 K m - l increases these a 1 2 over land under clear skies during the daytime values by a factor 10 1 • Thus, a values lying in when there is an onshore flow of air. Simple the range 2 to 22 can be expected and are compa models of convective IBL growth can be devel rable with observed values (Hsu, 1986). oped from a slab-model approach, based on This x 112 behaviour is found because station mixed-layer dynamics. This can be used to derive arity is assumed as well as horizontal uniformity a set of governing equations from which hb(x) of the heat flux. The effect of a time varying heat can be determined, where the fetch is designated flux can be readily deduced if the heat flux is as distance x from the coastline (Fig. 16a). Equa assumed to vary sinusoidally through the day (a good approximation). Let H = H max tions for mixed-layer wind components and tem 0 0 perature, the continuity equation and an equa sin(27Tt' /T), where T = 24 hours and t' is the tion for hb permit numerical solution, with ap time elapsed since sunrise and also the time at propriate boundary conditions. The hb equation which air crosses the coastline. lf t is the travel required for closure results from applying the time from the coast to a point x inland, the TIBL TKE equation (Eq. 6) at the inversion level and height at time t can be found from ( Garratt, utilising a suitable entrainment assumption. 1992; p. 188) For practical application, it is possible to de 1 = 2y - (l + 2ß)jt+t' ( H max /pcp) rive a suitable model, based on the steady state, h~ 8 0 pure advective case, with equations identical to t' those for the growing CBL over a homogeneous Xsin(27Tt"/T) dt" surface. The reader is referred to Garratt (1990) 1 :::::: 2y8- (1+2ß)(H:;ax/pcp)(T/21T) for details. If we assume that the Japse rate y 8 is constant and that the surface heat flux is inde X {cos(2'1Tt'/T) - cos[27T( t + t')/T]}. pendent of x, and use the initial condition that (30) hb = 0 at x = 0, then, with um the slab velocity, we have The depth of the TIBL after a travel time t ( =x/um) depends on t' since this sets the aver (29a) age level of the heat flux during the passage of so revealing an hb a x 112 behaviour (e.g. Venka- the air to a point x inland from the coastline. For J.R. Ga"att / Earth-Science Reviews 37 (1994) 89-134 111 t' = 0, the air crosses the coast at sunrise when historical data and dimensional analysis the heat flux is zero over the land surface whilst if (Mulhearn, 1981; Hsu, 1983), by use of numerical t' = T / 4 the air crosses at midday when the heat and simple physical models (Garratt, 1987), and flux is maximum. Growth of the TIBL will be by analysis of detailed aircraft data ( Garratt and greater near midday than near sunrise. What Ryan, 1989). Growth rates are found tobe small, typical travel times are appropriate? For the TIBL with fetches of several hundreds of km required depth to approach the "equilibrium" CBL depth to develop an IBL several hundred metres deep. expected "far inland", t of order 1 hour suffices, The nature of the 0v profiles within the IBL over noting that for t' = 0, the CBL will be shallow the sea is found to be quite different to that (""' 100 m) and for t' = T /4, the CBL will be found in the stable boundary layer over land. approaching its maximum daytime depth ( ""' 1-2 Over the sea, the 0v profiles are found to have km). Two solutions of Eq. (30) serve to illustrate large positive curvature with vertical gradients the problem, increasing with height, interpreted as reflecting (i) With t' = 0 and t =x/um ( « T), the dominance of turbulent cooling within the layer. The behaviour contrasts with the behaviour h~ 2y 1(1+2ß)( H max /pcpum)(2'7TX 2 /umT) = 0 0 in the nocturnal boundary layer over the land, (31a) where curvature is negative (vertical gradients decreasing with height) and where radiative cool (ii) With t' = T / 4, ing is dominant (Garratt, 1990). = 1 A suitable model of IBL growth can be formu h~ 2y0 (1+2ß)( H0max/pcpum)x. (31b) lated based on the basic structure shown in Fig. Thus, after one hour's travel time inland from 16b. We assume for simplicity that the advected the coast (x/um = 1 hour) the midday IBL will be continental mixed-layer air of virtual potential greater than the early morning IBL by a factor temperature 0vc remains constant. Above the (umT /21Tx)112 z 2. The analysis also shows that IBL we take y8 = 0. The starting point is the 2D, the stationarity assumption is likely to be much steady-state 0v equation (Eq. 3) integrated be better around the time of maximum heat flux tween the surface and hb, with an assumed linear when, for an assumed sinusoidal variation, the flux profile and assumed self-preserving forms for heat flux undergoes its minimum rate of change the profiles. The result for hb(x) is (Garratt, with time. 1990) At very large fetches, boundary-layer heights 2 must tend towards an equilibrium value. That is, h~ = aiLJ ( gß.(J /8v) -l X (32) the square root dependence must ultimately be where invalid at large enough x where hb tends to a and roughness Jength (z ). Data from Garratt (1982) 6. 4. Flow over orography 0 Site 'Ye Observed Calculated Theory and other considerations Koorin -15 0.1 0.002 0.13 0.14 Minnesota 45 0.01 0.001 0.37 0.38 Sloping terrain. In the presence of uniform slop Wangara -30 0.001 0 0.39 0.37 ing terrain many aspects of ABL structure remain Cabauw 45 0.1 0 0.42 0.39 unchanged from that over horizontal surfaces. At nighttime, however, the presence of slopes can affect significantly the depth of the noctumal a significant slope angle and a cross-slope boundary layer (NBL) and induce shallow geostrophic flow. drainage or katabatic flows almost independent of the overlying NBL. By day, upslope flows tend Flow over hills. The study of flow over ridges and to be deep and the CBL little affected by the hills represents an extra degree of complexity underlying slope. compared with flow over horizontal or sloping The presence of sloping terrain is represented terrain, or even flow across a simple change in in Eqs. (1) and (2) by an additional buoyancy surface conditions. The pattern of flow around a term; thus, with the x-axis taken along and down hill is determined not only by the hill shape but the slope, the RHS of Eq. (1) contains a new also by its size. If the hill is large enough (usually term, of the form -(g/8)0'sina where the tem of height > 100-500 m) and occupies a signifi perature deviation 0', which is negative in stable cant fraction of the daytime ABL depth then flow conditions, produces a height dependent local disturbances throughout the ABL and into the pressure gradient. The slope of the terrain is stable region above will result (both day and represented by the angle a. night). For smaller hills, flow disturbances within In the presence of sloping terrain, and hence and above the NBL only can be expected. of slope flows and interfacial mixing, NBL char Ta a greater degree than in many other areas, acteristics are modified compared with flow over our understanding of hill flow has been shaped by a horizontal surface. Both numerical models and mathematical modelling. This is not surprising observational studies have been used to study the when instrumentation used in lD micrometero effects of slope. For example, Brast and Wyn logical experiments must be located in a multi gaard (1978) used a simplified second-order ABL tude of locations to capture the complex (3D) model to show how slope inclination and the flow field around even simple topography (see relative direction of the Iarge-scale wind affected the review by Taylor et al., 1987). At Askervein NBL depth and internal structure. Garratt (1982) Hill, the site of the most complete field experi summarized observations and model results of ment to date, 50 towers were deployed with 27 of the nondimensional NBL depth for a range of them equipped with three-component turbulence slope angles. Table 1 summarises these results sensors (Taylor and Tuenissen, 1987). The most and shows how the non-dimensional height pa influential theoretical work to emerge has relied rameter is sensitive to both the slope magnitude upon techniques of asymptotic matching applied and the angle between the wind direction and the to flow over low hills, where the equations of slope fall vector. The impact of katabatic flow in motion (Eqs. 1, 2) can be linearized (Jackson and the Koorin case is clearly evident, with the re Hunt, 1975; Hunt et al., 1988a, b). There is a duced value of 'Ye (see Eq. 28) being the result of great deal more structure to be explored in the J.R. Garratt / Earth-Science Reviews 37 (1994) 89-134 113 mean flow above an arbitrary hill than in the and described by Kustas and Brutsaert (1986) and horizontally homogeneous and advective situa Brutsaert and Kustas (1985, 1987). The Pre-Alps tions discussed above. In the lowest layer, the region is characterised by hills of the order of 100 inner layer, whose thickness depends on the size m above the valley elevations and by distances and roughness of the hill, the Monin-Obukhov between ridges of the order of 1 km. The empha similarity theory can be used in modified form, sis in the first study was on the near-neutral wind , but farther from the surface the presence of profile and the relation between z 0 the zero advection does not allow the condition of local plane displacement and the surface geometry. equilibrium in the turbulence that such scaling This profile was found to be logarithmic over a requires (see Kaimal and Finnigan, 1993; eh. 5). substantial part of the ABL depth, leading to consistent results in estimating z 0 (about 3.5 m), Field experiments and observational studies in cam d (about 45 m) and the friction velocity. In the p/ex terrain second study, neutral humidity profiles were We summarise below several relevant experi analysed, yielding values of zq some 10 to 15 • ments and results of observational analyses that times smaller than z 0 In the third study, surface provide some insight into ABL structure and fluxes in unstable conditions were evaluated from behaviour over regions of complex terrain. measured vertical profiles using surface-layer re (i) SESAME (1979) - this included, inter alia, lations (Eq. 8, for wind and humidity) applied studies of the low-level jet over the Great Plains throughout the lower half of the ABL. However, of the USA (e.g. Lenschow et al., 1988a, b). best results were found with the z / L dependence (ii) ALPEX (1982) - the emphasis, inter alia, ignored and the '11 functions replaced by a non was on ABL aspects of flow around and within zero constant. the European Alps (e.g. see Smith, 1986 for re (iii) ASCOT (1984) - the overall objective of search in the US; see volume 36, pp. 1-296 of this major US dispersion-focussed experiment Meteorol. Atmos. Phys., 1987). The main ALPEX seemed to be a better understanding of pollutant experiment was held during March and April in transport and diffusion associated with valley 1982. flows, with particular emphasis on the transport Studies related to the ABL were mainly con of pollutants from energy related facilities to pop cemed with diumal heating effects, particularly ulation or agricultural centres set within valleys. on valley and slope winds, and on the local wind The original programme was developed in 1978, systems themselves. For example, Hafner et al. with major studies held in 1979, 1980 and 1981 in (1987) studied the diumal variation in ABL struc the Geysers geothermal area in northem Califor ture on both sides of the Alps in the context of nia, USA (Gudiksen and Dickerson, 1983) andin the regularly observed elevated heat low. 1984 in the Brush Creek valley in westem Col Vergeiner and Dreiseitl (1987) summarised val orado, USA (see Journal of Applied Meteorology, ley-wind research, emphasising the importance of Vol. 28, 1989). quasiperiodic thermal forcing in driving local wind (iv) MESOGERS (1984) - this ABL experi circulations, and the role of slope winds and ment involved studies of flow properties over topographical relief. They found that the diumal inhomogeneous terrain and regional-scale fluxes variation in the horizontal thermal contrast be (e.g. Durand et al., 1987; Weill et al., 1988). tween plains and valley induces a like pressure (v) AUTAN (1984) - this ABL experiment gradient variation, thus forcing both up- and studied the orographic influence on large-scale down-valley winds. Finally, Hennemuth (1987) flow and its consequences on the dynamics of the discussed the role of small valleys in the interplay local ABL (e.g. see Benech et al., 1987). of thermal circulation systems, and how these (vi) FIFE (1987, 1989) - see above. The es related to heating of the mountain ABL. sential requirements for FIFE were the need for ABL observations from ALPEX in the Pre simultaneous data acquisition and multiscale ob Alps region of Switzerland have been analysed servations and modelling. Main ABL data con- 114 J.R. Garratt / Earth-Science Reviews 37 (1994) 89-134 sisted of measurements of surface and near found that during daytime the thermal effects of surface fluxes from ground-based and aircraft surface heating and entrainment on form drag systems, and of thermal and biophysical proper were more significant than dynamic effects in an ties of the vegetation, and soil moisture. A strat adiabatic, non-entraining ABL. Typical form-drag egy for optimum station distribution within the coefficients, for the example chosen of intermedi experimental area, and a discussion of estimating ate terrain variations, reached 0.005 in the middle regional-scale surface fluxes from local values, of the day, greater than the frictional coefficient can be found in Mahrt (1987). relative to a reference height of about 100 m. The (vii) HAPEX (1986) - see earlier discussion. work suggests how form drag might be parametrised in GCM's, a problem under fairly Theoretical / numerical studies in comple.x terrain intensive study at present. These authors cer A number of these in the past few years have tainly question the use of effective roughness given new insight into ABL evolution over sloping lengths in the standard ABL formulations, since or mountainous terrain, into ABL structure and they do not reflect the influence of ABL height terrain-induced mesoscale motions, and into the and mixed-layer temperature variations at the evolution of thermally-induced valley circulations. mesoscale. For example, Banta (1984, 1986) was con Nevertheless, much work has been reported on cerned with evolution of the daytime ABL over the problem of effective roughness lengths. For a the Roclcies, McNider and Pielke (1984) with mixed-cover landscape, the vertical flux of mo general slope and mountain flows, and Segal et mentum to the surface is proportional to 2 al. (1988) with the effects of vegetation on ups [ln(/b/z0 e)J- , where z 0 e is the effective rough lope flows. The ALPEX and ASCOT experi ness length and /b is a blending height above ments have stimulated many theoretical and nu which the effects of local IBL's are not felt. The merical studies related to ABL behaviour on the following summary is taken from Hess (1992). mesoscale - for example the work of Egger The effective roughness length for momentum (1987a, b) and papers in the 1989 Journal of can be estimated from a weighted average rela Applied Meteorology (Vol. 28) related to ASCOT. tion Egger's work involved the modelling of thermally-induced valley-plain circulations, with emphasis on clarifying the dependence of these where fi is the fractional cover of local roughness upon the valley's geometry; stratification; advec length zoi· Consideration of Eq. (33) for typical tive processes and the diurnal cycle of heating. values of lb in the range 10-100 m shows that the Model simulations which were highly resolving in areal roughness length, and hence stress, is domi the horizontal were able to produce the gross nated by the roughest sub-areas and tallest features of observed flows near the mouths of roughness elements. Where orographic influences valleys. are present, terrain can play a significant role in Of particular relevance are the numerical stud determining surface roughness. Recently, Wood ies of Han et al. (1982) and Deardorff et al. and Mason (1993) - see also the earlier study by (1984) on the heated mixed layer over a mesoscale Grant and Mason (1990) - parametrised the region of complex terrain. In part, they were total drag by a sum of the shear-stress contribu interested in the problem of topographically-in tions associated with the small-scale roughness duced form drag, and evaluation of the related elements and form drag created by the terrain: drag coefficients, over irregular terrain in the 2 2 2 ln- ( zm/z e) = C /k + ln- ( Zm/z i) (34) presence of a well-defined mixed layer. Their 0 3 0 results are highly relevant to the problems of Where Ca is a form-drag coefficient and Zm is a effective roughness lengths or drag coefficients scale height approximately equal to the height of for use in GCM's, and of sub-grid variability and the obstacles (trough to peak for a ridge-valley how this should be represented in models. They system). A comparison of the predictions from J.R. Garratt / Earth-Science Reviews 37 (1994) 89-134 115 numerical modelling. The sea breeze is basically a 10.0 •... ···················"··" boundary-layer phenomenon and is a response to .. thermal forcing; it is driven by the differential 100 z0 e/h 0 0 •••• •• surface heating between land and sea that occurs 0 i;i.·· ·· ...... ····· during the daytime. This induces a sea-land tem perature difference of several Kelvins throughout the ABL, with the corresponding horizontal pres 1.0 sure gradient producing the sea-breeze circula tion. lts intensity, longevity and inland penetra tion is significantly dependent upon several major factors including, dryness of the land surface, large-scale flow, terrain slope, latitude, insolation over land (e.g. Pielke, 1984 (eh. 13) and Arritt, 0. 1 L..__L--=-':-'.::--'---::--':-=----'---::-':-=-~--::-= 0.00 0.05 0.10 0.15 0.20 1989 for general overviews; Physick, 1980; Kondo, A/Sd 1990; Bechtold et al., 1991; Arritt, 1993; for spe cific studies). Fig. 17. The variation of z 0 e / h (h here is the terrain height) as a function of a parameter A / Sd closely related to hill A recent review by Abbs and Physick (1992) slope (here, A is the frontal or silhouette area of the hill, Sd discusses several aspects of sea breezes, including is a base area). The squares are observed values from Grant fine-scale structure, behaviour in complex terrain, and Mason (1990), triangles from Hopwood (1991). The solid line shows the prediction of Eq. (34) for two-dimensional hills inland penetration, role in internal-bore genera tion and numerical modelling. We summarise with z 0 = 0.3 m, ,\ = 1000 m; the dashed line is for hills with z0 = 0.5 m and ,\ = 1000 m. From Wood and Mason (1993). these issues here. The sea-breeze front over land has strong simi larities with a density current structure, the dense Eqs. (33) and (34) with observations (Fig. 17) sea air that flows towards the front at low levels shows general agreement. Note that values of z 0 e being swept up and backwards at the leading of several metres are implied under some circum edge in the region referred to as the "head" stances, and that the ALPEX derived z 0 of 3.5 m (Simpson 1969, 1972). Mixing between cool moist (second part of Section 6.3), with the terrain sea air and warm dry land air takes place at the height h = 100 m and a terrain horizontal wave rear of the head through breaking billows (see length A = 1 km, is entirely consistent with the Nakane and Sasano, 1986 for observations of results of Fig. 17. these), generated at the front of the head by Kelvin-Helmholz shear instability. The mixed air 6.5. Mesoscale circulations flows back towards the coastline above the inflow and constitutes the return flow. A number of Sea breezes studies are relevant to the relation between the The diurnal cycle of onshore winds during the depths of the boundary layer, the inflow and the day and offshore winds at night is a characteristic raised head (see e.g. Simpson and Britter, 1979). weather feature in many coastal regions when Typically, the inflow of depth several hundred synoptic-scale winds are light. This wind regime is metres flows into a mixed layer ahead of the sea generally referred to as the sea- and land-breeze breeze of depth 1 to 2 km, with the head about 2 circulation (Mak and Walsh, 1976). to 3 times the depth of the inflow and extending Much has been learnt in the past three decades back several kilometres only behind the leading of the behaviour of sea breezes, and the ABL edge (see e.g. Simpson et al., 1977; Nakane and structure pertaining to the sea and land compo Sasano, 1986; Kraus et al., 1990). Along the front nents of the complete sea-breeze circulation. In during the day, the leading edge consists of a that time considerable research has been devoted continually changing pattern of lobes and clefts to theoretical and observational aspects, and to about 1 km across. These arise from convective 116 J.R. Garratt / Earth-Science Reviews 37 (1994) 89-134 instability caused by the overrunning of warm the cold air flattens and spreads inland propagat land air at the ground by denser sea air (Britter ing as an unsteady gravity current. and Simpson, 1978), and disappear at night when Abbs and Physick (1992) have reviewed the radiative effects lead to cooling of the near history of sea-breeze modelling much of which surface air. The first fine-resolution modelling comprises numerical simulations in both two and study of the sea breeze to appear in the literature three space dimensions. Many of the numerical was published by Sha et al. (1991). They were studies utilise hydrostatic mesoscale models on able to simulate all the fine-scale features found horizontal grids ranging from 2 to 20 km spacing. in laboratory experiments and verify the empiri A number of authors (Pielke, 1972; Martin and cal relations involving the inflow and head depths, Pielke, 1983; Song et al., 1985; Yang, 1991) have and billow amplitudes and wavelengths. concluded that the hydrostatic assumption is valid The topography of the coastline has an impor for grid spacings as small as 1 km, with non-hy tant controlling effect on the sea breeze. As weil drostatic effects tending to weaken the sea breeze as influencing the inland penetration it affects when compared to a hydrostatic simulation. Most convergence / divergence and consequently the recently, Sha et al. (1991) used a non-hydrostatic sea-breeze intensity and associated weather (see model with a horizontal grid spacing of only 100 Abbs, 1986; Abbs and Physick, 1992). Coastal m to examine the mixing processes in the head orography may also act as a barrier to the inland region. This clearly showed the entraining of penetration of the sea breeze (Abbs and Physick, warmer ambient air into the cooler low-level sea 1992). There is much evidence, both from obser air by the breaking of Kelvin-Heimholz billows. vations and numerical experiments, that sea breezes do indeed travel large distances inland Land breezes from the coast, the size of the penetration de The land-breeze phenomenon needs to be dis pending upon several factors including time of tinguished from other nighttime flow or circula year (hence surface heating and moisture con tion phenomema. For example, katabatic or tent) and latitude. Sea breezes have been ob drainage flows usually occur at night under light served as far as 280 km inland ( Garratt and wind, clear sky conditions wherever there exists Physick, 1985), with many studies reporting pene sloping terrain, inland or at the coast. In addi trations of about 200 km (see references in Abbs tion, Coriolis effects acting upon the sea-breeze and Physick, 1992). Movement of the sea breeze circulation may induce an offshore component inland is not uniform, but is dependent in part during the night. upon the pre-frontal ABL structure. Three dis As with the sea breeze, the land breeze is a tinct phases appear to exist (e.g. Simpson et al., shallow boundary-layer phenomenon resulting 1977; Sha et al., 1991) - initial uniform move from the differential surface heating between land ment followed by deceleration in the early after and sea that occurs during the night. The induced noon; acceleration about sunset; deceleration and horizontal pressure gradient produces the land decay late at night. The first phase corresponds breeze circulation. This tends to be somewhat to Clarke's (1984) immature stage when the rela weaker and shallower than the sea breeze be tive flow is through the isentropes and is far from cause there exists no lower heat source to enable steady state. The second phase is Clarke's early penetration of the circulation to heights found and late mature stages, when the sea-breeze surge with the sea breeze (e.g. see table 23 and fig. 69 has not reached steady state but, with increased in Atkinson, 1981 for observations and theoreti density contrast across the front, acceleration has cal predictions relating to the depths of both sea commenced. During early evening the supply of and land breezes). cool air from .the sea ceases, decoupling of the In contrast with studies of the sea breeze, the frictional link with the surface occurs and accel land breeze has not been studied extensively. eration continues. In the last phase, correspond Indeed, the studies that do exist are often carried ing with Clarke's early and late degenerate stages, out jointly with those of the sea breeze (e.g. J.R. Garratt / Earth-Science Reviews 37 (1994) 89-134 117 Neumann and Mahrer, 1971; Mizuma and time history of the surface sensible heat flux: Kakuta, 1974; Mak and Walsh, 1976; Mizuma, across the region. 1985). For general discussion relating to the land In analogy to the above, significant spatial breeze, the reader is referred to chapter 5 in variations in the daytime surface sensible heat Atkinson (1981) and chapter 13 in Pielke (1984). flux: (H) are probably quite common over sub stantial areas of land. Thermally induced circula Non-classical mesoscale circulations tions (termed NCMC's to distinguish them from the classical sea-breeze circulation) can be ex General. Sea breezes (and the analogous lake pected under such conditions (e.g. Segal and Ar breezes) are associated with significant and spa ritt, 1992). Following the terminology of these tially coherent horizontal differences in surface authors, we refer to areas in which H values are temperature, hence surface sensible heat flux:. suppressed or enhanced relative to the surround Most importantly they are associated with hori ing region as Perturbed Areas (PA's). These are zontal pressure differences related directly with most likely to occur due to spatial variations in temperature differences between distinctive air surface evaporation or evapotranspiration, solar masses of significant depth. The marine air mass irradiance absorption or reflection, thermal stor comprises a shallow cool, moist weakly stratified age by the subsurface. Perturbed areas inducing ABL and the terrestrial air mass a deeper, dry significant NCMC's are usually clearly distin convective ABL by day, and (with the sea-breeze guishable from the surroundings, though they are front further inland by nightfall) a shallow, dry not likely to be as uniform as water surfaces in but stably stratified ABL by night. The air mass the sea-breeze case. When they are !arge enough, density differences are largely dependent on the the induced circulations may be as intense as a ßo PA Fig. 18. Illustrative vertical cross section of an NCMC, showing vertical temperature profiles and surface heat fluxes at various locations across the circulation. From Segal and Arritt (1992). 118 J.R. Garratt / Earth-Science Reviews 37 (1994) 89-134 sea breeze. Fig. 18 schematically illustrates a typi cal NCMC vertical cross section for land situa tions (from Segal and Arritt, 1992). In most cases, the PA is associated with reduced H to the ~25~ • • : • : : : : : • l:I l atmosphere (positive) compared with the region around the PA. In some situations, the sensible heat flux H over the PA may be towards the surface (negative). In both cases, a suppressed CBL or a stable boundary layer will develop over ~]::=s;;J the PA. Figs. 19 and 20 illustrate clearly the differences in temperature above adjacent snow Fig. 20. Horizontal variations of surface heat flux and low-level covered and clear surfaces, and the horizontal 0 across a snow- to bare-surface transition - from Segal et variation of H (inferred from a low-flying air al. (1991). craft) between the two regions (see Segal et al., 1991). and Arritt (1992). Development of an NCMC Effects of PA size on generation of NCMC's. The requires that the induced pressure-gradient force following arguments are largely based on Segal be sufficient to counteract the imposed ambient wind. We consider an initial ® stratification ß 0 = ae ;az which evolves over the surrounding area 2000r-r-r---r.....,....-.--r-..,...... r-r--.~--,---r-r-, towards a mixed layer (ß =O) of depth h. Using a} the Exner function, with Il = cPT /®, we can SNOW scale the pressure perturbation Il' over the PA as 1500 E (35) ~1000 C> where 8 is the average potential temperature w :c: within the ABL. Taking a sinusoidal variation in H through the day, and using typical growth models for the CBL (e.g. Tennekes, 1973), we can deduce an expression for the time varying scaled pressure perturbation, viz. 286 289 292 295 298 8(Kl 2 Il' = -1.2grH0 [l - cos( 1Tt/r)] /pcP1Te . (36) Here, r is the duration of sunshine hours, H (b) 0 NO SNOW is the heat flux over the PA surroundings at noon 1500 (t = 0 at sunrise). To simplify matters, we assume E that the time scale for the development of the pressure perturbation is limited to the advective ~1000 time scale L/Ug, where L is the minimum width C>w :c: of the PA that produces an NCMC countering the background flow, Ug. We also assume the ideal case of unheated air over the PA (H = 0). The pressure perturbation that can be produced OL...... __..__.___.__._ ...... _.__._....._.._.._.._...... during the advective time scale over the PA can 283 286 289 292 295 298 8(K) be estimated by evaluating Eq. (36) for t = L/Ug. Fig. 19. Vertical profiles of 0 over adjacent snow-covered and lt is then assumed that the NCMC will exist if the bare surfaces - from Segal et al. (1991). perturbation pressure gradient at distance L is at J.R. Garratt / Earth-Science Reviews 37 (1994) 89-134 119 50 lations for surface irregularities of scale > 10 // km). 30 „// /...... ::: 10 Recent studies on NCMC's 100 .r / · 200 I The generation of significant horizontal heat E 8 /300 flux differences will lead to differential ~ I /400 6 boundary-layer evolution, and to horizontal tem _J /; perature gradients throughout the lower atmo 4 /j sphere, resulting in a sea-breeze-like circulation. 2 l1, Sharp contrasts in surface temperature are read ily observed from satellite imagery (Segal et al., 0 0 2 3 4 5 6 7 8 1989). The intensity of this circulation will be Ug(m 5-I) critically dependent upon the synoptic flow and on the presence of topographically-induced Fig. 21. The minimum PA width, L, needed in order to develop an NCMC flow for a range of opposing synoptic mesoscale flows. winds, Ug, and several values of the surface heat flux - Theoretical and conceptual evaluations of the 2 shown for values from 100 to 400 W m- • Note the change in kinematic and thermodynamic processes associ ordinate scale at L = 10 km. From Segal and Arritt (1992). ated with NCMC's have been provided, in recent years, by Smith and Mahrt (1981), Anthes (1984), Segal et al. (1984), Pielke and Segal (1986). For example, Anthes (1984) hypothesised that bands least equal to the pressure gradient associated of vegetation in semi-arid areas, of optimum with the opposing background flow Ug. Thus, spacing - 100 km, could, under favourable large-scale conditions, result in enhanced convec fUg = - Ball' /aL tive precipitation. The basis of this related to (i) increase in low-level moist static energy, (ii) in = l.2gHO sin(7rL/UgT)/pcp8ug, (37) crease in ABL water vapour due to increased where L provides the minimum PA width for evaporation and decreased runoff and (iii) gener development of an NCMC. Note that the above ation of NCMC's associated with the surface in analysis assumes that the heat flux H over the homogeneities created at this scale by the vegeta PA is zero. Where H > 0, or the PA has a num tion. These ideas were broadly confirmed in a ber of small-scale patches with H > 0, the effec numerical study (Han and Anthes, 1988) which tive H 0 in the above equation will have to be showed that strips of dry and wet surface with reduced. Fig. 21 provides values of L for various width and separation - 100-200 km can, in a • combinations of Ug and H 0 lt suggests that for convectively unstable environment with weak syn Ug = 3 ms- 1 and H = 100 Wm-2 the impact of optic flow and plentiful moisture, initiate convec the opposing background flow on the suppression tive rainfall. of the NCMC is noticeable for L = 10 km. Thus, Numerical model studies by Mahfouf et al. P As having L > 10 km are needed in this case for (1987) and Segal et al. (1988) suggested that for the NCMC to counter the background flow. For prescribed dense, well-watered and extended crop Ug = 4 ms- 1 and the same value of H, suppres areas, mesoscale circulations of an intensity close sion of the NCMC occurs for L < 26 km. For the to that of a sea breeze may be produced under ideal extreme case where H = 400 wm- 2 (typical optimum conditions. In Segal et al. (1988) for of mid-summer, inland, dry surrounding condi example, the presence of vegetation on extensive tions with H = 0 over the PA), NCMC's would be sloping terrain significantly reduced the upslope noticeable with similar background flows for PA's flow existing under ideal conditions. Where dense as small as several km (otherwise these results are vegetation was introduced into dry-soil areas, and consistent with Fig. 14, suggesting coherent circu- covered an extensive area under optimum envi- 120 J.R. Garratt / Earth-Science Reviews 37 (1994) 89-134 ronmental conditions, circulations close to sea Integral (slab) models breeze intensity were developed. The require The integral approach predicts the vertically ment for weak synoptic and upslope flows was averaged properties of the ABL, so that details emphasised in the observational study of Segal et on the vertical profile structure of any property al. (1989) related to the irrigated crop areas of are unavailable. The approach is particularly north east Colorado in the USA. The expected suited to cases where vertical gradients are small thermally-induced NCMC's were never unam throughout much of the ABL, or where vertically biguously identified, though wind-field changes averaged quantities are specifically required. In across the contrast region were observed, and the former case, the daytime entraining CBL is coincident with the thermal and moisture hori the best example whilst GCM schemes that have zontal gradients. lt was concluded that the synop limited vertical resolution are examples of the tic flow and the daytime, terrain-induced upslope latter. Closure of the slab equations for both the flow combined to mask the NCMC. clear and cloudy ABL requires knowledge of sur face and entrainment fluxes and the ABL depth. 7. Numerical modelling and parametrisation schemes High-resolution models "High resolution" as used here implies multi 7.1. Models ple levels in the vertical, so allowing the internal structure of the ABL to be evaluated. To solve Numerical solutions of the averaged conserva for the mean fields, the Reynolds flux-divergence tion equations require both a suitable closure term must be approximated at each level. There scheme to account for the turbulent flux terms, are two main categories of this type of model: (i) and usually a set of physical parametrisations of those utilising the ensemble-averaged equations, the Earth's surface, clouds and radiation. Both or volume-averaged equations that approximate one-dimensional ABL models, and two and ensemble averages because the averaging scale is three-dimensional atmospheric models with hori much greater than the ABL scale and (ii) LES zontal grid scales much greater than the typical models with volume averaging and explicit repre ABL length scale ( - 1 km), utilise ensemble aver sentation of the !arge eddy structure. The ensem aged equations. In contrast, in !arge eddy simula ble-averaged models may be structured either as tion (LES) models (which potentially represent a specific ABL model, or as an interactive com the best approach to calculating the 3-D, time ponent of a mesoscale or general circulation dependent structure of the ABL), true volume model. averaging is incorporated. In LES models, sub grid turbulence parametrisation in the limit of small grid scale (relative to the turbulence length Large-eddy simulation models scale) is likely to be far simpler, and universally In these, the averaging volume is sufficiently more applicable, than in the ensemble-averaged small that the largest energy containing eddies case. This is so, since large eddies are explicitly are resolved explicitly, at least weil away from the resolved and small eddies are likely to take on a lower boundary and the overlying inversion layer. more universal character. Unfortunately, in many This is of paramount importance, since turbulent applications, LES is impracticable - in mesoscale flows tend to differ from one another mainly in and large-scale models, for example - and em their large-eddy structure whereas the small scales phasis here will be given to the details of closure in all turbulent flows tend to be statistically simi schemes used with the ensemble averaged equa lar. The !arge eddies are very sensitive to the tions. environment (geometry and stratification), and in ABL models are usually structured in integral particular to the buoyancy forcing. Thus, only the (slab) form, or in high (vertical) resolution form less sensitive small scales need to be parametrised. utilising either first-order or higher-order closure. LES models allow the use of relatively simple J.R. Garratt / Earth-Science Reviews 37 (1994) 89-134 121 sub-grid closure schemes, including the first- and Alternatively, it may be computed using a full second-order schemes used in ensemble average treatment of soil heat diffusion in a multi-level models. soil model, with the SEB equation solved itera Solution of the volume-averaged equations tively using the Newton-Raphson method (e.g. provides variables that are partly random. Thus, Jacobs and Brown, 1973; Pielke, 1984 - eh. 11). to compute mean (in the ensemble sense) vertical In the prognostic case, soil heat flux may be profiles of any variable, there is a need to average crudely parametrised or even set equal to zero. over a series of horizontal planes and/ or over a One method which is widely used is the force-re number of time steps, and/ or several runs or store method, where the surface temperature is simulations. For area-averaged turbulent statis approximated by the temperature of a thin upper tics the total value of the property will be the sum layer (e.g. Deardorff, 1978; Dickinson, 1988). of the resolved and unresolved (parametrised sub-grid) parts. Before one can determine the surface fluxes, 7.3. Surface humidity (soil moisture) and fluxes throughout the ABL (e.g. through Eqs. 1 to 4 ), in any model it is often necessary to In determining the surface humidity for a bare evaluate a number of surface characteristics. soil surface, two quite different approaches need Same properties, including roughness length and to be recognised - interactive and non-interac albedo, are most likely specified, but this is not tive. The non-interactive approach means that generally the case with surface temperature and the surface humidity or soil wetness does not humidity. If we are dealing with the sea surface, respond to atmospheric forcing in a realistic way, the surface temperature may weil be prescribed if at all. Several examples can be found in Carson and the surface humidity taken as the saturated (1982) - see also Garratt (1992 - eh. 8). value at this temperature. The interactive approach allows for feedback between the surface and the prevailing atmo spheric conditions. Two common approaches are 7.2. Surface temperature the "bucket" approach, where surface moisture is identified with that of a single thick slab. Its main Numerical models of the atmosphere usually shortcoming is that the evaporation does not re compute the ground surface temperature (T) spond to short-period occurrences of precipita from either a diagnostic form of the surface en tion, which in fact change the moisture content, ergy balance (SEB) equation or from a prognostic and hence evaporation, only gradually (Deardorff, • form, i.e. a rate equation for T0 The diagnostic 1978). The thick layer is analogous to a bucket form is usually written, which holds, say, 15 cm of water at saturation (it overflows if more water is added from rainfall). (38a) The other approach is force restore, where near-surface soil moisture can be treated in an where RN is the net radiation and the prognostic form, analogous way to that of surface temperature using a two-layer soil model. As with tempera (38b) ture, the model must represent the rapid re sponse of the near-surface moisture to forcing by where Cg is a volumetric heat capacity per unit precipitation or evaporation and must also in , area of ground, and G0 1 are soil heat fluxes at clude a source of moisture from the deep soil to the surface and some small depth respectively. the surface when there is no precipitation (e.g. In the diagnostic case, the ground heat flux Deardorff, 1977). The evaporation may be evalu may be parametrised very crudely - e.g. as a ated as a fraction of the potential rate, where the constant fraction of RN or by assuming the heat fraction is dependent on the near-surface soil = capacity of the ground is zero and setting G0 0. moisture content (see Noilhan and Planton, 1989). 122 J.R. Garratt / Earth-Science Reviews 37 (1994) 89-134 7.4. Canopy parametrisation canopy cover, either as a sparse uniformly dis tributed cover or as full cover occupying only a Simple canopy models fraction of the grid area, more complexity is in The presence of vegetation over an area of volved. Discussions on multi-level canopy models, ground modulates the evaporation from the soil, where the variation of fluxes through the canopy and contributes further to the vertical flux of is evaluated, can be found in e.g. Finnigan and water vapour into the ABL through transpiration. Raupach (1987) and Raupach (1988). A realistic canopy formulation must ultimately For a complete vegetation cover, the simplest represent the effects of vegetation (averaged over canopy model uses a constant rs in Eq. (llb) for the grid square in a 30 numerical model) upon evaporation, with 's values consistent with known evaporation, energy partitioning, rainfall inter bulk stomatal resistances, together with the speci ception and soil moisture, as weil as albedo and fied albedo and z 0 • In contrast, a complex aerodynamic roughness. lnclusion of canopy ef single-level canopy model contains many parame fects allows the deep soil moisture (in the root ters with which to evaluate fluxes from the soil zone) to act as a source for evapotranspiration. beneath the canopy, from open areas between Except when completely wet, the canopy foliage the canopy elements as weil as from the foliage exerts some degree of physiological control upon itself (e.g. the SiB model of Seilers et al., 1986; the evaporation rate, and the surface humidity the BATS model of Dickinson et al., 1986). In becomes indeterminate. Under these conditions, addition the component fluxes are averaged over a canopy or surface resistance (conductance) is the grid area in some realistic way. With this introduced into the evaporation formulation, and approach, quite sophisticated treatments for the the resistance (conductance) concept is at the surface resistance can be used. heart of most canopy models. Simple, but realistic, canopy models can be Single-level, canopy formulations are the most formulated for both isothermal and non-isother appropriate for use in mesoscale and general mal surfaces. In the isothermal case, both canopy circulation models, and these will be emphasised and surface-soil layers are assumed to have the here. GCM's, for example, have the option of full same temperature, but in the non-isothermal ap canopy or bare soil grid coverage. For partial proach, the canopy and soil temperatures are p l _j- ). Fig. 22. Schematic representation of an isothermal soiljcanopy model (near-surface soil layer and canopy are at temperature T01 P is precipitation rate. J.R. Garratt / Earth-Science Reviews 37 (1994) 89-134 123 allowed to differ. These models are essentially fluxes at the top of the canopy, the canopy / soil 10 formulations applied to a grid area that might surface temperature must be evaluated. This is comprise either uniform cover (soil or vegetation) done through a surface energy balance equation, or an assumed distribution of patches of bare soil with the soil/ canopy system treated as a two-layer and full canopy. structure, and canopy temperature (T0 r) calcu The isothermal canopy / soil model represents lated from the force-restore method. a combined overstorey and soil layer, and as In contrast, the major features of the non-iso sumes that the soil surface and the foliage are at thermal canopy / soil model are shown schemati the same temperature T0 f (Noilhan and Planton, cally in Fig. 23. In this, the soil surface tempera 1989). Fig. 22 sketches the main elements of this ture (Tg) is allowed to differ from the foliage approach. The main task is to compute the turbu temperature (Tr), resulting in fluxes from both lent fluxes of heat and water vapour, H0 and A.E0 underlying ground (Eg, Hg) and foliage (Er, Hr) respectively, from the canopy to the air. The heat to the atmosphere (Deardorff, 1978). The total , flux is given by the surface-layer relation Eq. (10), fluxes, H0 and E 0 are given by H0 = Hr +Hg and = with 0 0 r the canopy (surface) potential tempera E 0 Er+ Eg. In order to solve for the component ture, and for evaporation, a distinction is made fluxes, the temperature T0 must be evaluated between dry and wet canopies. Thus, for a wet from a solution of the SEB equations applied at canopy, evaporation is at the potential rate, and the canopy top (for Tr) and at ground level (for for a dry canopy, with evapotranspiration under Tg). physiological control, evaporation is given by Eq. (llb). The relative amounts of dry and wet evapo Models of partial canopy cover ration are made to depend on the amount of The canopy formulations discussed above are liquid water residing on the foliage, due either to best applied to vertical exchange from an area of rainfall or dewfall. In addition, the soil moisture uniform, dense vegetation cover. Their extension depends on precipitation reaching the ground, to sparse or partial canopy cover is not straight and so the effects of canopy interception of rain forward, although progress has been made in fall should be included. recent times (Shuttleworth and Wallace, 1985; , In order to solve for H 0 and E 0 the turbulent Shuttleworth and Gurney, 1990). In the real E p l 1 j- Fig. 23. As in Fig. 22, but for the non-isothermal case - the canopy and upper soil layer are at temperatures Tr and Tg respectively. 124 J.R. Garratt / Earth-Science Reviews 37 (1994) 89-134 world, partial canopy cover is structured in an equations and volume-averaged equations used in infinite number of ways, but two extreme struc LES modelling. Their principle aim is to allow tures suffice to illustrate the problem. In the first, calculation of vertical fluxes throughout the tur a low-density uniform cover exists across the grid bulent ABL, covering the whole range of stabili area; in the second, the area is comprised of ties. patches of dense canopy and bare soil. From a grid area perspective, the fractional canopy cover First-order closure (ensemble models) = E0 In many respects, the second-order closure and the associated constants relating it to sub problem reduces to the approximations for the sidiary length scales is at the heart of the prob molecular destruction, turbulent transport, and lem. Although length-scale prescriptions cause pressure covariance terms in the second-moment fewer problems in modelling, they often oversim equations. The pressure transport term in the plify the physics by constraining the modelled TKE equation is typically either neglected or turbulence to an expected state. When multiple simply absorbed into the turbulent transport term. scales are present (as in any complex flow, e.g. in Closure for the pressure covariance in the flux the ABL with shear and buoyant forcing; in hori equations is usually an extended version of Rotta's zontal flow over a heterogeneous surface) there return-to-isotropy hypothesis. are advantages in using a differential equation for The turbulent transport (triple correlation) l. Physically, the alternative approach of using a terms are not significant in the neutral and stable differential equation for the dissipation rate E ABL, but in the CBL they may dominate the flow has theoretical problems and may not improve dynamics, and the entrainment at the ABL top. the physical realism over the equation for /. But The most commonly used closure for the trans there is evidence of improved simulations of the port terms is the downgradient diffusion model, turbulence in complex flows when using a prog involving a suitable turbulent length scale (Moeng nostic equation for l or E rather than the diag and Wyngaard, 1989). Finally, many second-order nostic (algebraic) approach (e.g. Mellor, 1985; closure models parametrise molecular dissipation Holt and Raman, 1988). rates in terms of velocity and length scales based Second-order modelling of the ABL is still on standard scaling arguments, thus avoiding the evolving, though the increasing use of volume need to carry a rate equation for the dissipation averaged models which restJlve the large-eddy itself. structure in the ABL (LES models) suggests that Closures summarised above were originally de the refinements to second-order schemes will not veloped to model turbulent neutral shear flows, be so crucial in the future. The approach has and the associated adjustable constants were been called "turbulence engineering" (Wyngaard, mostly obtained from laboratory data. Although 1982), and a wide variety of models can be found transport modelling contains some fundamental in the literature, some based entirely on shear problems, these are nowhere as critical as the flow closures and some tuned specifically for problems associated with modelling the pressure buoyancy-dominated turbulence. A recent inter and dissipation terms. In the case of the transport comparison of four LES models based an CBL closure, the downgradient diffusion assumption is simulations can be found in Nieuwstadt et al. clearly inadequate in the CBL where K theory (1993). fails. For most atmospheric model applications Finally, a number of studies have been re this is not important because the mean-flow char ported involving comparisons of simulations of acteristics are not much affected. In fact, LES mean and turbulence fields in the ABL which results have verified the poor performance of the give some insight into the relative performances downgradient assumption in the CBL, particu of first-order (K), one-and-a-half order (TKE) larly in the prediction of some turbulent statistics and second-order closure schemes. Thus, the (Moeng and Wyngaard, 1989). The simulation of reader is referred to Holt and Raman (1988) - turbulent second-order statistics using a TKE-dis K vs. TKE; Chrobok et al. (1992) - local K vs. sipation closure approach has also been de nonlocal closure in a cold air outbreak; Smith scribed by Andren (1991) who compared the re and Hess (1993) - K vs. second-order closure in sults with observations for a range of stabilities. the oceans. Another major deficiency lies in the use of 7.6. ABL c/oud parametrisation length scales, and the fact that all process scales are ultimately related back to one master turbu In many ABL and larger-scale models, ABL lent scale (/). Determination of this length scale layered cloud is assumed to exist if the ABL top J.R. Ga"att / Earth-Science Reviews 37 (1994) 89-134 127 lies above the local condensation level, i.e. if the marine case, the clear ABL far more than the relative humidity at z = h is 100%. In many cloudy case and the homogeneous or weakly non GCM's, low-level stratiform clouds are assumed homogeneous ABL far more than the strongly to be present when the grid-point relative humid nonhomogeneous case, particularly over complex ity at a specified level in the model exceeds some terrain. critical value (about 85-95%). Between this criti We have attempted to summarise current cal value and saturation, fractional cloudiness is knowledge on a number of these issues in order allowed to increase according to some empirical that the reader might appreciate the range of relation (e.g. a quadratic relation), becoming full problems for which knowledge of boundary-layer cover at 100% relative humidity. Some schemes processes is an important prerequisite for solu then introduce a dependence of cloudiness upon tions of these problems. other factors e.g. (i) low-level stability; (ii) the Much effort has been, and is likely to be, cloud-top entrainment instability. expended on the incorporation of surface and In both slab and high resolution ABL models, ABL processes in numerical models of the atmo incorporation of moist thermodynamic equations sphere. In this context, and possibly for other allows the presence of cloud to interact both with applications, several areas of research are likely the radiation and turbulence fields, and to affect to be given emphasis over the coming years: both the surface energy balance for example, and (i) area averaging of ABL properties in hetero the depth of the ABL. geneous terrain; In many GCM's in use today, the depth of the (ii) the structure and behaviour of the strongly ABL is not governed by a suitable rate equation time-varying and nonhomogeneous ABL; and may not even be computed (although it may (iii) the application of ABL theory over com be set equal to a fixed value). Hence, the pres plex terrain, and the influence of sub-grid orogra ence of stratiform cloud is unaffected by bound phy on regional transport; ary-layer dynamics. Such a non-interactive system (iv) the structure, behaviour and parametrisa can never aspire to simulate the real-world be tion of the cloudy ABL. haviour and distribution of boundary-layer cloud. These four areas extend investigation beyond Because the ABL controls the evaporation and the steady, homogeneous ABL for which exten turbulent redistribution of water substance into sive theoretical and observational treatments ex the atmosphere, it strongly determines the global ist. Even so, significant problems still exist in distribution of both cumuliform and stratiform relation to the ideal ABL that have resisted com clouds. A comprehensive parametrisation of ABL pletely acceptable solutions. A good example is processes for a numerical model of large-scale the turbulence closure problem, recently dis atmospheric circulations must take into account cussed by Hasse (1993) who emphasised the lack the interaction of the ABL with clouds. This of a commonly accepted ABL theory analogous aspect of the ABL parametrisation problem has to the Monin-Obukhov similarity theory of the an importance comparable to that of determining surface layer. He noted that the lack of conver the turbulent fluxes. gence of ABL theory towards a consensus is of practical relevance, e.g. in application of theory within larger-scale atmospheric and oceanic nu 8. Final comments merical models. His paper was intended to stimu Recent texts provide the reader with copious late discussion on whether the ABL community information on the atmospheric boundary layer needs to search for an established ABL theory, or (e.g. Stull, 1988; Sorbjan, 1989; Garratt, 1992), whether different theories should be adapted ac but with the emphasis predominantly on the sta cording to the problem to be solved. tionary and time-varying horizontally homoge All of the above suggests that much still re neous cases. For practical reasons, the continen mains to be learnt about the atmospheric bound tal ABL has received far more attention than the ary layer. 128 J.R. Garratt / Earth-Science Reviews 37 (1994) 89-134 List of symbols HAPEX Hydrologie Atmospheric Pilot Experiment u, v, w components of velocity along x, y, z ISLSCP International Satellite Land Sur axes face Climatology Project x, y, z Cartesian coordinates ( z is the vertical JAS IN Joint Air Sea Interaction (Pro axis) ject) p density of air LES Large-eddy simulation p pressure MESOGERS Mesoscale experiment in the T absolute temperature Gers region 0 potential temperature NCMC Non-classical mesoscale circula ev, ee virtual and equivalent potential temper- tion atures SESAME Severe Environmental Storms q specific humidity and Mesoscale Experiment liquid water content WMO World Meteorological Organisa surface values of 0 and q tion Coriolis parameter (N.B. sign(f) = ± 1) s generic quantity for u, v, w, T, q standard deviation of s fluctuations References e mean turbulent kinetic energy aerodynamic roughness length Abbs, D.J„ 1986. Sea-breeze interactions along a concave roughness lengths for temperature and coastline in Southern Australia; Observations and numeri humidity cal modeling study. Mon. Weather Rev., 114: 831-848. ß negative of ratio of entrainment to sur Abbs, D.J. and Physick, W.L., 1992. Sea-breeze C'bservations face heat fluxes and modelling: A review. Aust. Meteorol. Mag., 41: 7-19. Albrecht, B.A., Penc, R.S. and Schubert, W.H., 1985. An observational study of cloud-topped mixed layers. J. At mos. Sei., 42: 800-822. List of acronyms American Meteorological Society, 1990. Symposium an FIFE. Am. Meteorol. Soc., Mass., USA, 180 pp. ALPEX Alpine Experiment Andre, J.C., Goutorbe, J.P. and A. Perrier, 1986. HAPEX AMTE X Air Mass Transformation Experi MOBILHY: A hydrologic atmospheric experiment for the study of water budget and evaporation flux at the climatic ment scale. Bull. Am. Meteorol. Soc., 67: 138-144. ARME Amazon Region Micrometeorol Andre, J.C. et al., 1988. 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Garratt / Earth-Science Reviews 37 (1994) 89-134 up diffusion of a sealar in the eonveetive boundary layer. J. John Garratt, ARCS, BSe, DIC, PhD, Atmos. Sei., 41: 102-112. spent seven years at Imperial College, Yamada, T., 1976. On the similarity funetions A, B and C of University of London in the Depart the planetary boundary layer. J. Atmos. Sei., 33: 781-793. ments of Physics and Meteorology be Yan, H. and Anthes, R.A., 1988. The effeet of variations in fore emigrating to Australia to take surface moisture on mesoseale eireulations. Mon. Weather up a 3-year fellowship with CSIRO. Rev„ 116: 192-208. He has been a researeh seientist at Yang, X., 1991. A study of nonhydrostatie effeets in idealized the CSIRO Division of Atmospheric sea breeze systems. Boundary Layer Meteorol., 54: 183- Research since October 1970, and is 208. currently leader of the Atmospheric Zeman, 0., 1981. Progress in the modeling of planetary Proeesses Programme. His career has boundary layers. Annu. Rev. Fluid Mech., 13: 253-272. included visiting appointments to the Zilitinkevich, S.S., 1970. Dynamics of the Atmospheric National Center for Atmospherie Research (1980-1981), Boundary Layer. Hydrometeorol., Leningrad, translated Boulder, Colorado and to Colorado State University (1987- by C. Long, 239 pp. 1988) in Fort Collins. Zilitinkevich, S.S., 1972. On the determination of the height His current research interests include the representation of of the Ekman boundary layer. Boundary Layer Meteorol., soil-canopy and boundary-layer cloud processes in numerieal 3: 141-145. models, and the validation of numerieal simulations (in gen eral eireulation models) of the surfaee energy and surface radiation budgets. Each year he gives a full leeture eourse on the atmospherie boundary layer to honours and postgraduate students in the Department of Mathematies, Monash Univer sity. Cambridge University Press published bis book The At mopsheric Boundary Layer in 1992 (softback in 1994).