Mesoscale cellular convection over the oceans

Dynamics of Atmospheres and Oceans, 10 (1987) 317-341 Elsevier Science Publishers B.V., Amsterdam - Printed in The Netherlands

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MESOSCAI~K C E L L U L A R C O N V E C T I O N O V E R T H E O C E A N S E.M. AGEE Department of Earth and Atmospheric Sciences, Purdue Unioersity, West Lafayette, IN 47907 (U.S.a.) ABSTRACT Agee, E.M., 1987. Mesoscale cellular convection over the oceans. Dyn. Atmos. Oceans, 10: 317-341. A review of the understanding and progress in the study of mesoscale cellular convection (MCC) in cloud-topped marine boundary layers is provided. A comparison is made between MCC and the classical study of Benard-Rayleigh convection, with noted similarities and differences. External physical forcing mechanisms in the atmosphere are identified and discussed, in view of their effect on convection development and structure. Onset, horizontal planform, circulation direction, aspect radio and scaling phenomena are explained in terms of atmospheric processes. The global climatology of MCC is also updated and briefly discussed in terms of GCM model development.

1. INTRODUCTION Since the inception of the meteorological satellite program in 1960, atmospheric scientists have realized and studied m a n y classical fluid phen o m e n a in the troposphere, unique to the perspective of a space observing platform. The atmospheric manifestation of B e n a r d - R a y l e i g h convection is no exception, it is known as mesoscale cellular convection or M C C (see Agee et al., 1973, Agee, 1984). M C C occurs primarily in the marine b o u n d a r y layer as an extensive organized array of hexagonal convection cells, which can be detected through conventional satellite cloud photography. These cells typically have diameters of 1 0 - 1 0 0 km and can occur as either open or closed circulation systems. Open cells have d o w n w a r d motion and clear skies at cell center, with up motion and clouds in the periphery of the hexagon, while closed cells have the opposite circulation. Convective marine b o u n d a r y layers, which contain MCC, have depths of a b o u t 1 to 2 kin, with a capping inversion layer of 200-600 m. Aspect ratios, defined as the ratio of cell diameter (distance from cell center to cell center) to convective depth, range from values of 5 to 50. A pattern of open M C C presented in Fig. 1 shows the enhancement of convection at the vertices of the individual hexagonal cells as predicted b y b o t h classical linear and nonlinear theories. 0377-0265/87/$03.50

© 1987 Elsevier Science Publishers B.V.

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Fig. 1. Hexagonal open cells north of Cuba, as viewed by Gemini 5 at 1827 GMT, 23 August 1965. Convective enhancement is noted at the vertices of the hexagons, as predicted by theory. In April 1985, the World Meteorological Organization-World Climate Research Program held a workshop at Colorado State University (U.S.A.) on the modeling of boundary layer processes. Participants in this workshop (including the author) recognized and proposed six types of cloud-topped boundary layers (CTBLs), which are briefly summarized below: (a) Type I (CoM-air outbreaks) A convective marine boundary layer in which clouds form as cold air moves to the east of continents over warm ocean currents. The initial cloud formation occurs within 12-18 h after frontal passage and is dominated by strong heating and moisture transport from below. The clouds are initially cumuliform but eventually develop into stratocumulus as they spread out beneath a capping inversion. The CTBL may exist for a few days, typically

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after the passage of the extra-tropical cyclone. The CTBL consists primarily of open cellular convection clouds within a well-mixed layer of 1.5-2.0 km depth, although accompanying closed cells can occur. (b) Type II (Marine stratocumulus) This CTBL forms over colder ocean areas generally to the west of continents. The dominant energy source for vertical mixing is radiative cooling of the cloud tops. The stratocumulus clouds are generally more persistent than those described in Type I and may last several days to weeks. They are predominantly in closed-cellular patterns, and result in only weak heat fluxes at the air-sea interface. (c) Type III (Continental) Several types of CTBL's form over continental areas. Under the influence of high pressure patterns, low cloud layers may form at night and persist through several days: These cloud layers are relatively thin and are dominated by radiative processes. During the day, surface heating may form shallow cumulus which can spread out to form layers of stratocumulus. Other stratocumulus layers may form over and in the lee of large lakes, or in association with orographic influences. (d) Type IV (Polar stratus) Stratus in the boundary layer covers about 70~ of the Arctic polar region in the northern summer. (The incidence of Antarctic stratus and wintertime Arctic stratus remain somewhat uncertain.) Radiative cooling of relatively warm, moist air that is advected over the Arctic Ocean is the dominant formation mechanism. Summertime cloud tops are generally below 1000 m and cloud thickness is about 300 m. The vertical structure of Arctic stratus is frequently complicated by multi-layered clouds that are detached from the surface fluxes and humidity inversions above cloud tops.

(e) Type V (Large-scale weather systems) Extensive low-level cloud layers frequently develop in the PBL in association with large-scale low-pressure weather systems. Although surface heating may be important, the dominant feature is the strong large-scale vertical motion which is upward ahead of the low-pressure center and downward behind. Although extensive in area, these cloud patterns move with the systems and are evolutionary in time.

(f) Type VI (Tradewind) Tradewind clouds form largely in response to the buoyancy caused by evaporation at the ocean surface, and occur mainly within the boundary layer below the tradewind inversion (a region of strong wind shear). They are best developed in the eastern parts of the tropical oceans, where the water temperature is relatively low and the inversion height is less than 1000 m. To the west these heights increase, and the stratiform structure gives way to deeper cumuliform clouds with less area coverage.

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Fig. 2. Cold air outbreak off the east coast of U.S.A., over the Gulf Stream, as viewed by GOES visible imagery at 1430 GMT, 25 December 1983. Open mesoscale cellular convection (MCC) has developed within a Type I CTBL.

As indicated above, the primary focus of this review paper is on the Type I and Type II CTBL's, both of which give rise to patterns of MCC. Fig. 2 shows a typical wintertime example of a cold air outbreak off the east coast of the United States. Cloud streets are first observed just off-shore, with a spacing between bands of 4 - 8 km. Farther downstream (and in time, with additional heating) this two-dimensional band structure gives way to an array of open cells. This pattern of open mesoscale cellular convection extends eastward to the frontal cloud band, trailing southwest from the parent extratropical cyclone (not pictured). These open cells have diameters

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of 20-60 km and depths of 1.5-2.0 km. Heat flux at the air-sea interface beneath these open cells can achieve values up to 1200 W m -2 (see Sheu and Agee, 1977). Such intense heat transfer is instrumental in developing strong baroclinic zones that can fuel the cyclogenesis process as disturbances arise. A representative vertical cross-section of a Type I CTBL is shown schematically in Fig. 3, from near shoreline (or ice edge) to about 1500 km out to sea. The increasing depth of the convective marine layer is due to continued strong heating from the warmer water below. The region of MCC identifies an approximate state of convective equilibrium, but undulations in the height of the inversion base are noted. These include (1) the difference between open (more shallow) and closed (deeper) MCC layers, in part due to the effect of additional latent heat release, and (2) the undulations associated with individual cells, where the inversion base is lowered (and raised) in the respective branches of descending (and rising) motion. Based on observational results from the AMTEX field program over the East China Sea (e.g., see Agee and Lomax, 1978) a more thorough explanation for the difference in convective depth can be seen in Fig. 4. Both the thinner open cells and thicker closed cell regions (for Type I MCC patterns) are seen to occur in a region of large-scale sinking motion. The entrainment process (to be discussed later) is very important to the structure of MCC. A representative example of a Type II CTBL with accompanying MCC is presented in Fig. 5, which largely shows a pattern of closed MCC (flanked by open cells) off the west coast of California. This CTBL is driven from above, largely through radiative-entrainment equilibrium, and occurs over

322

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cool ocean currents to the west of continents where air-sea temperatures are nearly in equilibrium. The Type II MCC patterns can sometime develop the rare actinia cloud species (see Fig. 6) that evolves between regions of open and closed cells. Studies of the Type II CTBL (comparable to the field investigations of Type I marine boundary layers such as AMTEX, STREX and KONTUR) have not been conducted, however, the marine boundary layer segment of F I R E (see Randall et al., 1984) could go far in providing additional understanding of Type II regions. A clear distinction between Type I and Type II physical processes is the important role played by radiation in the latter. The eventual breakup of stratocumulus cloud decks, as well as diurnal variations in fractional cloud amounts, is highly dependent on combined radiation and entrainment effects. The climatology of MCC events is very distinctive, with preferred times and location of occurrence. Generally, the Type I pattern is a cold season phenomenon and the Type II pattern is a warm season event. Type I MCC occurs to the east of continents over warm ocean currents, and south (or north) of the respective polar ice caps. These patterns consist mainly of open cells, but as noted earlier, closed cells can also occur. Similarly, for Type II MCC, patterns are preferred to the west of continents over cool ocean currents. Circulation direction is largely of the closed cell variety, but open cells can also occur. In studying the MCC phenomena for nearly two decades (including the viewing of a few thousand satellite images), the author can state comfortably that MCC are occurring somewhere on the planet all of the time. An update of the author's original climatological map for MCC is given in Fig. 7, which shows two additional regions for Type I open MCC. These are in the Norwegian and North Sea areas, and in the Gulf of Alaska. These regions of interest were brought into focus by the U.S.A.-Canada STREX experiment (see Fleagle et al., 1982) and the F.R.G.

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Fig. 5. A Type II CTBL off the west coast of the U.S.A. as viewed by GOES visible imagery at 1745 GMT, 1 August 1984, with primarily closed mesoscale cellular convection.

K O N T U R experiment (see Bakan, 1982). It is noteworthy that the M C C p h e n o m e n o n provides an interesting climatological structure across mid-latitude continents and adjacent ocean current. The North America continent for instance is flanked by a warm Gulf Stream to the east, where large energy fluxes occur during the winter from the ocean to the atmosphere. However, to the west during the summer the cool California current is not

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Fig. 6. Nimbus I visible imagery off the coast of Peru at 1813 GMT, 15 September 1964. This photo, centered near 10°S and 95°W, shows actinia (spoke-pattern convection) between regions of open and closed MCC within a Type II CTBL. an energy source, and the cloud-topped boundary layer radiates away large amounts of energy. This feature has earned Type II regions the label 'ocean desert' (Tom Vonder Haar, personal communication, 1985). The author further proposes that the Type I and Type II CTBLs, which flank mid-latitude continents, constitute a 'climatic thermal torque' of systematic heating and cooling of the marine b o u n d a r y layer. 2. MCC AND CLASSICAL THERMAL CONVECTION A reasonable question to pose in geophysical fluid dynamics is that of the relationship, if any, between the classical study of thermal convection and

325

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mesoscale cellular convection. The following analogous considerations can be made in the study of thermal convection and Type I and II CTBL's (with MCC): (a) criteria for the onset of convection; (b) wavelength and geometric planform of the initial convective mode; (c) circulation direction (i.e., open cells versus dosed cells); (d) transitional convective geometry and patterns with increased thermal stress; (e) aspect ratio of convection cells; (f) heat transfer; and (g) convective scaling at large supercritical Rayleigh Number (Ra). 2.1. Convection: onset, planform and circulation direction

An early attempt to write a meteorological version of the Boussinesq equations for thermal convection was made by Agee and Chen (1973), with particular emphasis on the study of MCC patterns in the PBL. Their model introduced an atmospheric Rayleigh number for dry convection, where thermal conductivity and molecular viscosity were replaced by their eddy counterparts. Krishnamurti (1975) extended this type of formulation for

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moist convection, and determined the regime-stability diagram for both linear and nonlinear model equations under conditions of weak supercritical Ra. Her findings are shown schematically in Fig. 8, where the anisotropy parameter 3' represents vertical motion. As seen in Fig. 8, if the fluid is initially at rest (3' = 0) or moving with constant horizontal velocity, the convective shape at onset is two-dimensional stable rolls. These cloud street patterns in convective marine boundary layers have at onset a typical aspect ratio (wavelength-to-depth) of about 3 to 1. If there is large-scale vertical motion the onset shape becomes three-dimensional and occurs as a stable array of hexagonal cellular convection. It should be noted that the presence of strong vertical shear in the horizontal wind in the PBL can externally force the 2-d structure, which is often the case in convective boundary layers over the land. Also, the aspect ratio of the 3-d convection at onset is up to an order of magnitude smaller than that often seen in patterns of MCC. The issue of circulation direction in the onset of 3-d convection has been addressed in both of the above mentioned studies. Agee and Chen showed that the reversal of circulation direction could be accomplished in their model by changing the sign of the vertical gradient of eddy viscosity (a feature often present in atmospheric PBL's, but not in laboratory convection tanks). This is analogous to the classical result by Palm (1960), which showed a circulation direction of open (closed) cells if molecular viscosity increases (decreases) with temperature as seen in the behavior of gases (liquids). Krishnamurti also showed that the sign of the large-scale vertical

327 velocity controlled the circulation direction, with up motion ('t > 0) producing closed cells and down motion (y < 0) producing open cells. As discussed in a later section, several additional physical processes must be incorporated into the study of convective marine boundary layers (which can offset an apparent dominant effect such as large-scale vertical motion). AMTEX studies (see Fig. 4) show both open and closed cells in the presence of large-scale sinking motion. 2.2. Increased thermal forcing The steady 2-d and 3-d onset convective modes become time dependent as the Rayleigh Number (or thermal forcing) increases. A schematic of the Rayleigh Number-Prandfl Number regime stability diagram (in the -/= 0 plane) is presented in Fig. 9, which shows the tendency of larger values of Ra to produce turbulent convection. Krishnamurti discussed the concept of the atmospheric Prandtl Number, which can range from 1 to as high as 50 (but typically is 10 or less). Although not indicated in Fig. 9, a recent nonlinear spectral model by Rothermel and Agee (1986) indicates that organized convection can occur in the turbulent regime at large supercritical Ra (from 500 to 1000 × Ra¢). 2.3. Convective scaling The concept of convective scaling pertains to the ability of a convection pattern (e.g., the wavelength or aspect ratio of the onset mode) to resolve

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itself into a much larger horizontal length scale (even though the convective depth remains fixed). This is particularly true for both 2-d and 3-d patterns in Type I convective boundary layers over water. The evolution of these larger aspect ratios can produce asymmetric structures in the process of undergoing transition from one length scale to another. Generally, it is viewed that the onset thermal mode represents the smallest length scale, which is called the basic convective mode (BCM). A convective layer can conceivably form an initial pattern of BCM's that can subsequently resolve

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329 or coalesce into a pattern of convection with a much larger aspect ratio. The cause for this coalescence or scaling may be intrinsic to the nonlinear partial differential equations that govern the convection, or else the scaling is externally forced by a host of possible physical mechanisms that are discussed below. The latter is particularly evident in Type I and Type II convective marine boundary layers. The MCC mode is viewed as a larger horizontal length scale that emerges from the BCM or onset mode. The spectral model results by Rothermel and Agee are particularly relevant to the scaling phenomenon, in that no external 'esoteric' physics is present to force the flatter convection cells. Figure 10 shows numerically simulated power spectra of vertical velocity at mid-level for equilibrium convective states at values of 4, 30 and 600 times Rayleigh critical. The weak supercritical case ( R a / R a c = 4) shows the BCM or onset thermal mode (at a wavelength of 2V~- for unit depth). For a computational domain of aspect ratio 28.28, 10 cells appeared initially. The role of nonlinear effects in the governing equations, however, became increasingly important for the higher values of Ra. At R a / R a c = 600, the MCC type of mode is seen to appear (corresponding to an aspect ratio of 14.14, or two cells in the computational domain). 3. FORCING MECHANISMS IN CONVECTIVEMARINE PBL'S In addition to the strong thermal forcing due to heating at the surface (Type I CTBL) and cooling at cloud top (Type II CTBL), a host of thermodynamic and dynamical mechanisms can be identified that influence the development and structure of convective marine boundary layers and the formation of MCC. A list of these mechanisms and influences is given below, many of which represent an external asymmetry feature that alters the convection pattern (as seen above in the consideration of vertical motion and eddy viscosity effects): (a) thermal forcing at the bottom boundary due to nearly uniform heating by the warmer ocean surface (resulting in a superadiabatic surface layer); (b) radiative forcing due to the combined effects of strong infrared cooling at the top of the CTBL ( - 30 m depth) and solar heating ( - 100 m depth), resulting in cloud-top destabilization; (c) latent heat release and added convective buoyancy due to cloud formation; (d) large-scale vertical motion; (e) temperature and moisture properties of the capping inversion layer above the CTBL (the source region of air that is entrained into the convective layer);

330 (f) cooling through the evaporation of cloud material, especially as a result of drier air being entrained into the cloudy layer (and the possible formation of upside-down convection, see Randall, 1980); (g) vertical shear of the horizontal wind within the CTBL and inversion layer; (h) superposition of a horizontal temperature gradient at the air-sea interface (T~ea-Ta~r) and the subsequent modulating effect of a Hadley type circulation on Benard-Rayleigh cells (see Ookouchi et al., 1977); (i) dynamic instability through the interaction of the length scales for the various physical processes that can occur, such as mesoscale entrainment instability (see Fiedler, 1984); (j) surface friction (e.g., land versus water surface); (k) small-scale turbulence within the CTBL that results from or is effected by most of the above processes, which in turn largely controls the entrainment process; and finally (1) entrainment (discussed below).

3.1. Asymmetry In the above list of physical mechanisms, both horizontal (fl) and vertical (T) asymmetry influences can be noted. The T-type of influences include large-scale vertical motion, viscosity variations as a function of height, and any combination of the various heating and cooling mechanisms that produce a curvilinear temperature profile (and thus a departure from the Rayleigh-type of profile or constant lapse rate). Roberts (1967) was the first to address the vertical asymmetry effects represented by heating a n d / o r cooling departures from a Rayleigh profile. Notable horizontal asymmetries (fl effect) include: (1) anisotropic turbulence, where the horizontal mixing (compared with the vertical mixing) becomes excessively greater; and (2) the asymmetry due to a horizontal temperature gradient at the bottom of the convective layer (arising from horizontal gradients in air-sea temperature differences). An example of the first type of horizontal asymmetry can be seen in Fig. 11, where cold air is moving off the coast into the Gulf of Mexico. This GOES imagery at 2031 GMT 2 April 1985 shows a convective boundary layer with cloud streets embedded in northwesterly flow over southeast Alabama and the adjacent waters of the Gulf. The development of a visible boundary of clouds along a northeast to southwest axis can be noted, with streeting along the northwest to southeast direction. Interestingly, the portion of the cloud streets over the land shows an aspect ratio of about 3 to 1. Over the water, however, scaling has occurred, apparently due to the more aerodynamically smooth surface (compared to the land surface). This results in less mechanical eddies or

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Fig. 11. GOES visible imagery at 2031 GMT 2 April 1985 showing different length scales (between land and water) in the same cloud street pattern along the Gulf coast.

turbulence, and effectively allows horizontal mixing to be even more dominant over the vertical mixing. The aspect ratio of the cloud streets over water is at least an order of magnitude larger ( > 30 to 1). A scale analysis of the horizontal mixing (Kh) to vertical mixing (Kv) ratio, after Priestley (1962), gives a change of KhL = 9KvL (3 : 1) to Khw = 900K w (30 : 1) primarily due to the reduced value of Kv, even though convection is strong over both land and water. The next example of asymmetry and resultant scaling is presented in Fig. 12, in response to a Type I cold air outbreak over the Gulf of Mexico. GOES visible imagery at 1430 GMT 29 February 1984 shows a larger horizontal length scale or bandwidth evolving (compared with the bandwidth of the cloud streets that initially developed, i.e., the BCM mode). One can also note in Fig. 12 that the cloud free path (distance from shoreline to the start up of visible clouds; see Chou and Atlas, 1982) is very short in the northeast sector of the Gulf of Mexico, but much longer as one moves southwest. Such a pattern is indicative of a horizontal gradient in air-sea temperature difference (strong AT in the northeast sector to weak in the southwest). This is a different shoreline configuration than that for most cold air outbreaks off the east coast of the U.S.A., where cloud free paths are essentially the same up and down the coast. It is now proposed that the

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Fig. 12. GOES visible imagery of a Type I cold air outbreak over the Gulf of Mexico. Convective cloud streets with two length scales are present consisting of the BCM and a much larger horizontal band width (see text for proposed explanation). asymmetry and subsequent scaling seen in the Gulf of Mexico is due to the imposed Hadley-type of circulation (see Ookouchi et al., 1977) on a B e n a r d - R a y l e i g h regime. The favorable horizontal heating gradient at the b o t t o m b o u n d a r y has introduced another length scale (namely one large circulation cell) which interacts with the basic convective mode. The resulting large aspect ratio and thick cloud bands occur frequently in the Gulf of Mexico, but are seldom seen off the east coast of the U.S.A. in Type I CTBL's. 3.2. An A M T E X case of convective scaling The Air Mass Transformation Experiment (AMTEX) held over the East China Sea in 1974 and 1975 afforded several opportunities to study M C C and Type I CTBL's. Aircraft investigation of one of these events (see Rothermel and Agee, 1980) is noted at this time, because it demonstrates the coexistence of both the M C C and B C M convective modes. Figure 13 shows a pattern of open and closed cellular convection over the A M T E X measurement network with 1 - 2 km reported depths of the convective marine layer. Figure 14 shows some of these individual open and closed cells as photographed b y Japan Airlines. Figure 15 is a simultaneous power spectrum of

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Fig. 13. DMSP satellite image of Type I MCC patterns over the East China Sea at 1154 JST 16 February 1975. The AMTEX observational network is noted, along with the depth of the convective layer (equivalent to the height of the inversion base) from 1-2 kin.

t~lg. t4. v h o t o ~ t l e ~ lry Sapan Air~dnes of individual open and ctosed cells ~tween Naha, ~ a Isimultaneous with Fig~ 13).

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specific humidity q for the flight leg by the N C A R Electra aircraft from Naha, Okinawa, to the Keifu ocean vessel. This shows spectral peaks, as indicated for both the MCC and BCM modes at the 970 m flight level (just beneath cloud base). It is interesting to conjecture that the larger aspect ratio for the MCC (about 30 to 1) is sustained by a coalescence of the smaller BCM circulations (about 3 to 1). Such energy cascading and redistribution has been numerically simulated by Rothermel and Agee (1986) as discussed in a previous section. 3.3. Turbulence and entrainment The small-scale turbulence and entrainment processes within CTBL's are inseparable, and work collectively to affect convective formations and changes within the marine boundary layer. Entrainment is viewed as a kinematic process representing the movement or 'pulling' of air down into

336 the CTBL from the overlying inversion layer. The entrainment rate is directly proportional to the small-scale turbulence, which is largely due to various forcing mechanisms discussed in section 3. Also, the moisture and temperature (or density) properties of the inversion layer play an important role, as well as large-scale vertical motion. A summary of the physical processes that affect turbulence production and dissipation are given below: (a) surface heating--generates convective buoyancy and thermal turbulence; (b) latent heat release--adds to the buoyancy and convective turbulence; (c) radiative destabilization at cloud top--generates convective overturning and turbulence; (d) wind shear--generates mechanical turbulence; (e) cloud top cooling--entrains stable drier air that consumes turbulent kinetic energy (TKE); (f) evaporation within cloudy layer--produces cooling that decreases convective turbulence (except for CIFKU, discussed below); and (g) turbulence--entrains stable air from the inversion layer that consumes TKE. It is generally recognized that sinking motion near the top of CTBL (for whatever reason) consumes TKE. Lilly (1968) pointed out that sinking in association with turbulent entrainment can result in the evaporative cooling of unsaturated air, which may (under certain conditions) cause a convective downdraft. Such an event could result in the production of TKE, rather than the usual loss. Randall (1980) termed such an instability as C I F K U (Conditional Instability of the First Kind Upside-Down), a condition that occurs in his model when the virtual temperature j u m p at cloud top is sufficiently weak. Deardorff (1980) examined the cloud top entrainment instability and stratocumulus break up in a numerical model. The entrainment rate was found to increase decisively when AOe dropped below the critical value. As AOe becomes negative, the final breakup of the cloud deck occurs whereby C I F K U operates and the fractional cloud cover becomes small. Fiedler (1984) proposed the concept of Mesoscale Entrainment Instability (MEI), in which mesoscale fluctuations of buoyancy and humidity were reinforced in phase (if certain conditions were met). MEI represents yet another mechanism by which large aspect ratios can be obtained in cellular convection (such as MCC). Controversy, however, does surround this mechanism for cell flattening (see Feidler, 1985; Van Delden, 1985). An explanation of the processes leading to MEI is now offered. The maximum mesoscale-buoyancy region in a convective CTBL gives rise to maximum turbulence (and thus maximum entrainment). As noted in the discussion of CIFKU, entrainment usually results in the reduction of turbulence. How-

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ever, MEI can occur and convective turbulence increase, if there is a sufficiently large virtual potential temperature gradient across the inversion layer. Under such conditions the entrainment of warm air enhances the convective turbulence, and this eventually increases the density jump across the inversion layer. LiUy's (1968) hypothesis states that entrainment into the mixed layer is inversely proportional to the density jump, and this subsequently produces a condition of reduced entrainment where the mesoscale buoyancy is maximum. Weaker entrainment in this region allows for the moisture and buoyancy fluctuations to be in phase and the MEI instability can occur. 4. C O N C L U S I O N S

An effort has been made to review the progress in the study and understanding of mesoscale cellular convection (MCC), as well as the cloud-topped marine boundary layers within which this convection develops. Several interesting comparisons have b e e n made between MCC in marine PBL's and classical Benard-Rayleigh convection. Similarities and differences have been duly noted, but atmospheric convection is subjected to a myriad of physical processes that are external to the thermally-driven convective overturning. For example, the onset mode and circulation direction can be controlled by several factors (such as vertical motion, vertical gradient of eddy diffusivity, and mechanisms due to local heating and cooling such as radiation or latent heat release). The aspect ratio or degree of cell flatness can be induced by nonlinear interactions of a thermally-driven system, an imposed horizontal temperature gradient at the bottom boundary or entrainment (particularly MEI). Atmospheric convectionists should never fall into the trap of categorically stating that a particular convective structure is always caused by a single external forcing mechanism. Over and over this is seen not to be the case (e.g., the occurrence of both open and closed MCC patterns within a region of large-scale sinking motion). A final example is offered to show further the seemingly never ending challenge one faces in the study of mesoscale cellular convection. This example is presented around the following proposed question: Does MCC always occur over the ocean? The author and colleagues who study MCC have always viewed these convective patterns as occurring in marine boundary layers. More specifically, as discussed, these patterns occur in a systematic and climatological fashion within the Type I and Type II CTBL's. Reasons to expect MCC over the oceans include the following: (a) near uniform heating and moisture supply at the bottom boundary; (b) minimal diurnal forcing (allowing a natural convective time scale); (c) no surface topography variations;

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Fig. 16. NOAA-7 visible image at 1423 GMT on 27 March 1984 showing open cells over Europe. (Photo provided by Professor Hans Oerlemans, Utrecht University, Netherlands.)

339 (d) favorable proximity of continents and ice sheets to ocean currents; and (e) sufficiently uniform conditions in large-scale meteorological state (both space and time) within which the convective patterns develop. During the author's visit to the Netherlands in 1984 a most interesting satellite photograph was obtained from Professor Hans Oerlemans at Utrecht University (see Fig. 16). This NOAA-7 visible image at 1423 G M T on 27 March 1984 shows a pattern of open cells over the land covering much of Europe from the Alps to the North Sea. Cold air boundary layer instability coupled with daytime heating was responsible for this convection, which had penetrated the inversion layer and grown to the precipitation stage. This remarkable pattern of open cells was most surprising to the author, and seemingly the entrainment process was somehow responsible. To heighten

Fig. 17. GOES visible image at 2031 GMT on 8 April 1985 showing open cells over the midwest region of the U.S.A. MCC-type of convection is rare and unexpected over the land.

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the interest in this phenomenon, another example of open cells over the land was observed by the author on 8 April 1985 (as shown in Fig. 17). This GOES visible image at 2031 G M T shows the development of large open cells over the midwest region of the U.S.A. in response to a cold air outbreak from Canada. Cloud streets are noted, aligned with the northwesterly flow, which correspond to the size of the BCM mode. Interestingly, within this pattern of 2-d convection, a pattern of large 3-d cells evolved as shown. This convection was penetrative, and developed intense convective snow showers. The open cell noted in Fig. 17 passed directly over the author's location at West Lafayette, Indiana. This passage gave an appearance that the convective snow showers had ended, as the sky became (temporarily) completely clear. Entrainment instability at the cloud top level may have been responsible for the development of the array of large open cells. The 2-d mode was probably still weakly present in the boundary layer beneath the open cells, as moderately strong vertical shear in the horizontal wind was present. An observational study of this case of 'MCC' over the land is in progress. In closing, it is worth noting the importance of MCC in the Type I and Type II CTBL's from the standpoint of G C M - N W P models. The need to parameterize the marine boundary layer processes is essential to improve model performance. Increased emphasis on the turbulent transport of heat. m o m e n t u m and water vapor and the initiation of stronger convective PBL's is required, along with the proper treatment of radiative effects. Model approaches to CTBL turbulence include: (a) mixed layer model, (b) firstorder (K theory) and higher-order closure, and (c) large eddy simulation. Parameterization techniques in G C M models should focus on the prediction of: (a) heat exchange at the air-sea interface, (b) cloud amounts, extent and persistence, and (c) radiative fluxes. As noted at the begirming, MCC occur somewhere on the planet daily and with climatological regularity. These cloud-topped boundary layers play an important part in the daily and seasonal patterns of weather and climate. ACKNOWLEDGMENTS

The author is grateful to Jotm Wagner for drafting the figures, and Carla Breiner for typing the manuscript. The research was sponsored by the National Science Foundation, Atmospheric Science Division, through NSF Grants ATM-8410560 and ATM-8505477, a n d b y the D O D Office of Naval Research under contract N ~ 1 4 - 8 6 - K - 0 1 7 9 awarded to Purdue University. REFERENCES Agee, E.M., 1984. Observations from space and thermal convection--a historical perspective. Bull. Am. Meteorol. Soc., 65: 938-949.

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