Chapter 4Aus: World Survey of Climatology

Vol< XT f General Climatology 2

Wind SyStemS

H. FLOHN

Eisevier Publishing Company

Amsterdam 1969

p. 139 - 171

Introduction

In mountains, at large lakes and along oceanic coasts a great variety of local wind Systems,which vary in direction, intensity and time according to local conditions and seasons can beobserved. Some are of extraordinary strength and persistency, such äs the down-slope windSystems at some Antarctic coastal Strips, where the average annual wind speed can reach äsmuch äs 20 m/sec. More frequently such winds blow with surprising regularity during latemorning and afternoon from one direction and then — after a lull — during night and earlymorning from a nearly opposite direction. Such a diurnal cycle of winds indicates atriermal circulation produced by differential heating of cold and heat sources (cf. CONRAD,1936; WAGNER, 1938; TROLL, 1952; BUDYKO, 1956; MILLER, 1965; SELLERS, 1965).In some cases orographical factors — such äs narrowing of valleys, mountain gaps etc. —are responsible for an intensification of winds from selected directions. In other cases themere existence of large-scale slopes causes a downslope motion produced only by gravity(VON FICKER and DE RUDDER, 1948 ; DEFANT, 1951, and WILLETT and SANDERS, 1959).In order to understand the mechanism of such winds, it is quite useful to Start from thewell-known classification of winds derived according to their physical sources. In anabbreviated symbolic form we can write a sum of forces acting on any three-dimensionalrectilinear motion in the atmosphere, based on the well-known Eulerian equation ofmotion:

dt— grad p -l- C — g + F

Here dV/dt is the acceleration of the wind vector V which vanishes under stationaryconditions, i.e., when the opposing forces are balanced. The other three-dimensionalvector quantities are: grad p = pressure gradient; C = deflecting force of the earth'srotation (Coriolis force); g — gravity (usually balanced by the vertical component ofgrad p) ; F = surface friction. For our purposes we may neglect the centrifugal forcesacting on curved motions and the effects of turbulent momentum exchange. Based on theclassical work of H. JefTreys and assuming stationary conditions (9 V/dt = 0, g = con-stant) we may distinguish some special cases:(/) F = 0; geostrophic wind in the free atmosphere, directed parallel to the isobars.(2) C= 0; antitriptic wind, directed towards low pressure.(3) C 0;geotriptic winds (JOHNSON, 1966) i. e., stationary flowwithin the surface friction

layer.

139

Local wind Systems

(4) F — C = 0;Eulerianwind3directedtowardslowpressure(onlymthefreeatmosphere).

(5) Grad;? = 0; gravity wind, directed mainly down-slope.Strictly speaking the cases (2) and (4) can only exist in the immediate vicinity of theequator. Since, however, under quasi-stationary conditions the adaptation between gradp and C takes several hours, short-living wind Systems can be considered, in all latitudes,äs very nearly antitriptic (or Eulerian). The term "katabatic" is used for all down-slope

motions, the term "anabatic" for up-slope motions.In this chapter, we at first discuss examples of simple thermally driven circulations, thesea-breeze System, together with their thermal origin and some simple theoretical con-siderations. Then we proceed to the more complicated Systems of mountain winds with adiurnal component, where motions of different scales act together. The final sub-chapterdeals with local winds driven (or strengthened) primarily by orographical features.

Sea and land breezes

We observe along nearly all coasts, but especially in subtropical and tropical chmaticzones, a regulär diurnal shift of the wind direction. During night and at dawn, a steadyweak breeze blows with a component from land to sea. About two hours after sunrise thisland-breeze ceases and the air is rapidly warmed by the sun. About three hours aftersunrise a markedly cool breeze from the sea develops, at first near the coast, and thenextending in both directions and becoming more and more intense until about 14hOO. At

this time its speed (usually 3-6 m/sec) is much higher than the maximum speed during thenight. Near or after sunset, the sea-breeze lulls, ceases, and is replaced again during thefirst half of the night by the somewhat weaker nocturnal System. This pattern of twomore or less opposite wind Systems following each other during 24 hours at quite regulärtimes is well-known. In the time of sailing craft, fishermen at many continental lakes andat maritime coasts used to start out with their boats early in the morning using the noc-turnal land-breeze towards the sea with firm confidence in the occurrence of the daytimesea-breeze allowing their safe return. The regularity of such a system increases withdecreasing cloudiness and with decreasing intensity of synoptic-scale disturbances or ofthe large-scale flow. Two examples (Table I), of annually persistent diurnal coastalcirculations from the Persian Gulf and Lake Tanganyika illustrate this behaviour.A good example of such a diurnal circulation between land and sea occurs at SantaMonica (near Los Angeles along the coast of California). The coast runs nearly north-northwest to south-southeast, but the pattern is complicated by an east-west oriented

TABLEI

TWO EXAMPLES OF DIURNAL COASTAL CIRCULATIONS FROM THE PERSIAN GULF AND LAKE TANGANYIKA

Sharja25.3°N 55.4DE

Kägoma4.9°S 29.5°E

Time

07hOO16hOO

OShOO14hOO

N

114

32

NE

22

71

E

81

603

SE

321

163

S

331

2

16

SW

56

122

W

533

140

NW

242

J10

Calms

120

113

140

Sea and land breezes

km5-

2-

0-

10 C 1 6 C 22°

Fig.l. Time-height-section of the wind component perpendicular to the coast above Santa Monica,Calif., 29 typical days, July-August. (After EDINGER and KAO, 1959.)

06 12 18 oo h

J

F

M

A

M

J

J

A

5

O

N

D

— •.

\

'J

Vvvy" j

•\

\

' //•/

\

yyy

\

\/

«^y -y *y vy "

, /v\\

r

rr[

s~~

//•

s^„^f

rr

rr/rr

\* /

/•

r4r

/.

r

s*

f\

f4

/ ^ ^ y&

,•$>/

^ /* X" ^

^ J

y ( / y yf \>y v ^

/ " ^\

/> ^ "" ^/ X^ r— I V\^ ,§• Ä

_ „ _ Sea — Breeze — — •-

-i l i i 1 i —

^ .\-

( \

yyy

> y1

^fy

-• \

^, .

jyyj\

«\

_»i

y •y •y "y -

j •i]

NE-Trade

SW-Monsoon

NE-Trade

05 07 09 11 13 15 17 19 21 23

15 m/sec

<Q5 m/sec

Fig.2. Daily and annual course of the resultant surface winds at Berbera (10.4°N 44.5°E). (After BROOKSand DURST, 1934.)

141

Local wind Systems

ränge of hills \vith a height of about 600 m (EoiNGER and KAO, 1959). During the summerthe maritime cool air is topped by a \vide-spread low-level Inversion; its lower and upperboundaries are situated at a height of 330 m and 820 in, respectively. The temperaturewithin the Inversion layer increases with height upto 8.7°C at I6h00. Oa typical days(Fig.l) the nocturnal land-breeze äs well äs the afternoon sea-breeze reach an altitude of1,500 m. Above that level the direction shifts by nearly 180°, and the reverse branch ofthe thermal circulation (in order to avoid misunderstanding, we use terms like "anti-landbreeze" or "anti-sea breeze") reaches up to about 5km, with its speed maximum at2,500 m. The sea-breeze blows with nearly 3 m/sec, the nocturnal land-breeze with only1.9 rn/sec; both currents decrease with height at and above the Inversion, and the speedof the upper "anti" winds remains below l m/sec. Here we should mention that thisthermal coastal circulation System is reinforced by a valley and mountain breeze Systemin the area of the San Gabriel Mountains in the hinterland, with peaks up to 3,000 m.The depth of the layer of maritime air and the lower boundary of the Inversion rise inareas where the sea-breeze converges (EüiNGER and KAO, 1959) and, in particular, at theslope of the mountains.The behaviour of the sea- and land-breeze in its diurnal and annual cycle may be demon-strated by a remarkable example (Fig.2) from Berbera at the northern coast of the Somalipem'nsula (BROOKS and DURST, 1934). Here the local diurnal circulations are superim-posed on a large-scale annual wind shift between a moderate northeasterly trade windsystem during winter (October-April) and the powerful southwestern monsoon duringsummer (May-September). The recently revealed strong heat-low (with pressures near1,000 mbar) at the hot Danakil Desert reinforces the southwestern monsoon whichreaches, on the open ocean, an average speed of 12-15 m/sec. During the cool season theday-time sea-breeze strengthens the northeasterly trade, blowing nearly in the samedirection, and the nocturnal land-breeze is strong enough to counteract the relativelyweak general flow. During summer, however, the nocturnal land-wind reinforces thesouthwestern monsoon which results in furious gusts in the early morning hours (OöhOO-OShOO). At OShOO, 82% of all winds have a speed of 6 Beaufort and more. After noon,between 13hOO and 15hOO, the sea-breeze from the Gulf of Aden, blowing from north ornorthwest, is strong enough to more or less regularly replace the powerful southwesternmonsoon. On its way inland, it passes the marked escarpment (with an altitude of 1,000-1,800 m) of northern Somalia and converges in a distance of 100-150 km off the coast,with the southwestern monsoon, frequently producing local showers (FLOHN, 1965).Similar conditions exist at the adjacent harbour of Bender Cassim.In the whole region of the Gulf of Aden and the southera Red Sea the general low-tropospheric flow, which is definitely fixed by the topography of the Red Sea Trench, iscombined with either a nocturnal land-breeze system (Fig.SA) or with the stronger day-time sea-breeze system (Fig.3B). As a consequence, a divergence zone develops at seaduring daytime, and in contrast, a convergence zone develops at night (FLOHN, 1965).The sea-breeze system produces afternoon showers at all neighbouring highlands,especially above Ethiopia and Yemen and (äs above-mentioned) along a convergenceline at interior Somalia. During the summer season, however, the nocturnal convergenceat sea is not strong enough to produce an appreciable amount of clouds and/or rain.Generally speaking, the convergence or divergence of the land and sea-breezes depends,äs shown by NEUMANN (1951), on the curvature of the coast line: where the coasts are

142

Sea and land breezes

36° E

12'

10'

143

Fig.3. Streamlines and resultant surface winds at the Red Sea and Gulf of Aden. A. 06hOO-OShOO,July-August. B. 12hOO-15hOO, July-August. Dashed lines = convergence zones. (After FLOHN, 1965.)

Local wind Systems

concave, the sea-breeze diverges in contrast to the converging land-breeze, while at con-vex coastlines the reverse happens.These exarnples are obtained from areas, where semi-arid or arid continents immediatelyborder the ocean; it is here where we should expect the strengest local heating differences.This is certainly not true at the equatorial Lake Victoria covering an area of about68,500 km2. This lake is, along its northern and western coasts, surrounded by swamps orhumid forests and plantations, and there are only weak thermal contrasts (FLOHN andFRAEDRICH, 1966). But even here the sea-breeze Systems usually overpower, during thewhole year, the weak general flow, which blows during northern summer from southerly,during northern winter from northerly directions, in both seasons turning to east withheight. Here we are able to evaluate numerically the daytime divergence (cf. p. 152)(about +2.8 • lO^sec-1 at 14hOO) contrasting with the night-time convergence (at OShOOabout —l.S-lO^sec-1). These results are of the same order of magnitude äs thosederived from synoptic-scale mid-latitude Systems. Based on 12 months of simultaneouspilot-balloon data from four coastal stations and using the continuity equation(p.l52),Fraedrich has evaluated the wind divergences äs a function of height (Fig.4) and thusestimated average vertical wind components (lifting during night, subsidence during theday above the lake). This explains the occurrence of cloud-free areas above the lakeduring the day, while above land in the convergence areas between sea-breeze andgeneral flow, showers and frequent thunderstorms occur. During the night the diurnal

2.0 km above surface

5-8 h 12~18hJun.-Aug. o—Q o—oDec.-Febr. x—x X—x

-3 -2 -lConvergence Divergence

Fig.4. Vertical distribution of convergence (OShOQ-OShOO) and divergence (12hOO-18hOO) above LakeVictoria. (After FRAEDRICH, 1968.)

144

Sea and land breezes

32° 33

Fig,5. Average annual precipitation at Lake Victoria. Rainfall maximum at the lake according to severalstations at Korne and Sesse Islands, with 17-37 years of record. (After FLOHN and FRAEDRICH, 1966.)

circulation reverses its sign, and from niidnight on till early morning large cunrulonimbusclouds with extended anvils and thunderstorms develop above the lake, caused by theall-year convergence of the land-breeze and are gradually displaced to the west coastby the persistent tropospheric easterlies. Here it is possible to verify the existence of amarked rainfall maximum in the central part of the lake (at the Sesse Islands with a long-term annual average of about 230 cm) in contrast to the lake area average of 145 cm/yearand a coastal average of only 126 cm/year (Fig.5). The rainfall maximum at the lake ishigher than that of the highlands situated to the northeast, in spite of their height of upto 3,000 m. The diurnal rainfall pattern, with an afternoon peak at the highlands and anearly morning peak at the lake (including the northern and western coasts) verifies thecorrelation between rainfall amount and diurnal circulations.During day-time, the same phenomenon can be observed above islands and peninsulasof a comparable size. One of the best known examples is Florida, which is, in the absenceof large-scale disturbances, in the afternoon covered by a series of growing and toweringcumuli, which develop into thunderstorms in more than 90% of all summer afternoons. Incontrast to the peninsula the area of Lake Okeechobee remains cloud-free. During thenight the sea-breeze System is replaced by a weak land-breeze System which diverges to-wards the surrounding ocean and is accompanied by subsidence and disappearance ofclouds. At sunrise, huge cumulus towers at sea, some distance off the coast, mark theconvergence of the nocturnal land-breeze with the general flow. BYERS and RODEBUSH(1948) have calculated an average wind convergence in the afternoon, near 2- 10~5sec~1

at the 300 and 600 m-level. This convergence disappears near 1,100 m and is replaced by

145

Local wind Systems

a weak divergence at 1,500 m. The divergence of the night-time land-breeze reaches only450m.Similar examples in Indonesia are well-known. The large islands, Borneo, Sumatra, Java,Celebes, Timor etc. and the Malayan peninsula, together with the adjacent seas promotethe development of still larger Systems. They have been frequently described (BRAAK,1940): in the early morning the observer standing on one of the now cloud-free mountainpeaks looks at the large cunrulonimbus masses appearing in the distance above the seas.In the afternoon, however, thereis a regulär appearance of thunderstorms which developat all mountains and volcanoes of the large islands. Because of the general weakness ofthe upper winds in the equatorial region, these diurnal circulations predominate thesequence of weather with astonishing regularity, varying only in intensity. In the after-noon the all-round convergence of the sea-breeze causes the individual towers above theislands to merge into a massive dark formation with a diameter of several hundreds ofküometres topped by thick grey altostratus covers, which demonstrate, like a giant anvil,the upper outnow. Down-pouring rainfall near the centre and wide-spread turbulence inthe free atmosphere are other characteristics. Along the coasts the daytime sea-breezecauses a marked cloud-free divergence zone. These macro-scale Systems are frequentlyconfirmed by satellite pictures, äs well äs their night-time counterparts above the seas byinfrared pictures. At the Malaysian peninsula and the Strait of Malacca, the diurnal rain-fall pattern coincides well with the formation and decay of the two reversed Systemsduring 24 hours (RAMAGE, 1964).Along the coasts of the Indo-Pakistan subcontinent both sea-breeze Systems are bestdeveloped during the hot pre-monsoon season, when the large-scale wind-systems arerelatively weak and the radiational heat differences reach their highest intensity; now thediurnal pressure difference 05hOO-17hOO (local time) amounts to 6 mbar. During thisseason, the sea-breeze from the Arabian sea is strong enough to surpass the escarpmentsof the Western Ghats and to reach Poona. In the hottest parts of the subcontinent, cityhouses are built with large elevated openings at the roof directed towards the sea-breezein order to catch the cooler air and to ventilate the interior of the house. During thesummer monsoon, however, the strong geostrophic general fiow and high cloudinessminimize the conditions which produce thermal circulations (ANANTHAKRISHNAN andRAO, 1964). In middle latitudes, the horizontal extension of the land-sea breeze circu-lation rarely exceeds 40-50 km. However in subtropical and tropical areas it reaches100-150 km on both sides of the coast with a total diameter of 200-300 km. Thus thesea-breeze from the Indian Ocean can easily be distinguished in the wind statistics of Voi,Kenya, at a distance of 150 km from the coast. In southern Australia, the sea-breeze ininfrequent cases during the evening, can even reach Kalgoorlie at a distance of 300 kmfrom the coast (according to information kindly given by Professor J. Gentilli, Perth).Usually these thermal circulations start along the coast and spread from there in bothdirections. In the presence of a superimposed general flow, the development is eitheraccelerated or delayed. If the local sea-breeze converges with an opposing current, a sortof cold front develops which frequently produces locally organized showers in unstablemoist air. As shown in the above-mentioned examples, the divergences and convergencesbetween different diurnal circulations and with the large-scale prevailing fiow are, to alarge extent, responsible for regional patterns of precipitations and cloudiness, especiallyin low latitudes. In middle and high latitudes the role of radiation is rarely äs significant

146

Heat budget considerations

äs in the tropics and in arid areas; furthermore the general fiow äs well äs the synopticdisturbances are nsually much stronger. Nevertheless, under anticyclonic conditionswell-developed sea-breezes are not infrequent; they penetrate inland behaving like weakshallow cold fronts, usually quite slowly (in southern England with an average speed ofonly 6 km/h). An average uplift of 1-2 m/sec within the "front" has been measured byglider pÜots and also can be niarked by sailing birds and insects on a radar screen(SIMPSON, 1967).

Heat budget considerations

Land and sea-breezes are simple examples of thermally driven circulations. In theatmosphere, the driving agent is locally differentiated heating or cooling of the air.Remarkable artificial examples are the fire storms during catastrophic fires (BUETTNER,1968), such äs the great Hamburg fire during the second World War, Jury 24-27, 1944,at the end of a summer drought. Here an area of 13 km2 was set on fire which produceda heat of 8-12 kiloLangley/h — •! Langley (Ly) = l cal./cm2, i.e., of an energy about 100times larger than the extra-terrestrial solar radiation on a perpendicular plane. The firelasted about 6 hours in its highest intensity; during this time, winds of hurricane force(50-60 m/sec and more) blew concentrically from all sides into the heat centre with acyclonic curvature due to Coriolis forces. In the streets, temperatures up to 750°C havebeen indirectly determined, and above the burning town an enormous brownish firecumnlonimbus developed up to the tropopause.Between land and sea, any thermal circulation is driven by local differences in the heatbudget (BUDYKO, 1956; MILLER, 1965; SELLERS, 1965). Using the same notation äs inChapter 2 we may write the heat budget equation of the earth's surface (neglecting someminor terms) äs :

(JR.-K = net radiation; (?0 = heat flow to ground; HQ — turbulent flux of sensible heatinto air; L = heat of vaporization; EQ = evaporation from the surface). Comparingland and sea, the differences of the net radiation are only of secondary importance: sincethe albedo of the sea is usually lower than that of land (cf. Deacon, p.48), while theatmospheric long-wave radiation is somewhat larger due to the low-level water-vapourcontent, &$ above sea is normally 15-20% higher than above land. The heat flow intothe sea is much greater than that into the soil; its sign varies mainly with season. Themagnitude and the diurnal nuctuations of G0 at land are relatively unimportant since thediurnal temperature wave does not penetrate further than 40-50 cm. The driving factorof diurnal circulation is primarily given by the term ffQ: at the open ocean this term isnearly constant during day and night since the bulk of the net radiation is used forevaporation. In contrast to this at an arid continent (where L-EQ disappears) Jff0 is, atleast during day-time, nearly equal to R^. Only above irrigated rice-fields, swamps, ortropical rain forests isL-E0 much larger than HQ.The local differentiation of the turbulent warming HQ of the air from the ground is theinitial source of energy of all thermal circulations. In low latitudes, HQ at sea is aboutO.IL-EQ. In characteristic cases near the coasts, however, where usually warm air is

147

Local wind Systems

advected by a large-scale flow from the continent to the sea which is persistently cooled byevaporation, HQ can. be negative. In contrast to this, at an arid or semi-arid land surfaceHQ is, in the diurnal cycle, roughly equal to R^ — GQ; the delay of 2-3 honrs between £Nand H is caused by the storage of heat in the ground (G0). In a typical example from thearid interior of California (SELLERS, 1965) (?„ is positive (into the ground) from dawn tillabout 15hOO, then weakly negative. Thus JffQ is positive from about two hours after sun-rise till 1-2 hours before sunset, with a maximum of 40-60 Ly/h shortly after noon.During night-time (including the hours near sunrise and sunset) HQ is weakly negative,in most cases reaching only —3 or —5 Ly/h. A comparison of the heat budget of acontinental desert with that of a large artificial lake in the same climatic zone (SELLBRS,1965) gives an average offfQ at land of 200-250 Ly/day in contrast to —60 to —80 Ly/dayat Lake Mead. Since at the lake, L-Eis large and positive, while H0 is negative, we havehere a typical example of the oasis effect. This effect (cf. Chapter 2, p.74) is character-ized by the energy of evaporation partly derived from the cooling of the not continentalair from below, and the vertical gradients of temperature and specific humidity of differ-ent sign. Unfortunately no direct measurements of the diurnal Variation of the heatbudget at two adjacent stations at land and sea are available äs yet. However, äs anexample, we may show the different heat-budget of an oasis in central Asia and of thesurrounding desertic steppe. BUDYKO (1956) (Fig.6) finds that at the irrigated oasis thebulk of AN is used for evaporation (L- E) äs at sea. Since in such cases additional energyis contributed by advection of sensible heat (—-ff0), £'•£? can be substantially largerthan Rs.At a land-locked mid-latitude lake, Lake of Constance (540 km2), the average temper-

A

Fig.6. Daily Variation of heat budget terms at the jTurkestan semi-desert and at an adjacent oasis. Solidline: oasis (irrigated);dashed line: semi-desert (J50 = 0);Äs =L~E0 4- -So + G0. (After BUDYKO, 1956.)

148

Heat budget considerations

TABLE II

AVERAGE TEMPERATURE DISTRIBUTION (SUMMER) AT LAKE CONSTANCE

Land (observatory)

Lake (centre)

Difference (land — lake)

OöhOO

(°C)12.1

13.8

-1.7

13hOO

(°C)21.4

18.8

+2.6

Difference

CG)9.3

5.0

+4.3

ature distribution during summer (June-August) is äs follows (Table II. After CONRAD,1936, p.228).Since here the lake is not wider than 12 km, thermal advection causes an appreciabledaily Variation of air temperature at the lake centre. At the open ocean, the daily Variationof air temperature is on the order of only 0.2-0.3°C.Taking into account the storage process of GQ and the much weaker horizontal differencesof HQ during night, we understand, in principle, that the onset of land and sea-breezes isdelayed by 2 or 3 hours when. compared with the radiation processes, and that the day-time sea-breeze is usually definitely stronger than the nocturnal land-breeze. This is alsotrue for the fonnation and dissolution of clouds under the innuence of these diurnalcirculations provided that no measurable rain is produced and the cloud fonnationremains a reversible process. In equatorial areas, however, where diurnal circulationsgive rise to enormous cumulonimbus massives with down-pouring rain, the irreversiblerelease of latent heat within the rainclouds acts äs a new powerful source of energy. Infact, during a tropical shower of 10 mm an energy of about 600 Ly is released in a fewminutes, while the long-wave radiation of the water-vapour and of the upper cloudsurface consumes only some 10-30 Ly/h and acts therefore äs a relatively slow heat-sink.These considerations can explain why under super-moist-adiabatic conditions the delaybetween diurnal circulation processes (and related cloud and rainfall phenomena) andsolar radiation amounts to 6 hours and more; the initial heat source of HQ may change itssign, while the much more powerful secondary heat source still maintains the localcirculation. This is the case at Lake Victoria äs well äs above the Indonesian islands andseas and in many other tropical areas with unstable, moist air. In these areas the strengestdevelopment of the two diurnal Systems occurs near (or even after) sunset and sunrise.A typical example of this kindhas been described by Lessmann fromEl Salvador, wherethe rainfall maximum usually occurs between 18hOO and 24hOO (LESSMANN, 1967),Along some subtropical coasts (Peru, northern Chile, Angola, California, Morocco,Somalia), the large-scale wind field produces an orT-shore drift of the surface waters, withupwelling cool water from the oceanic depth below the thermo-cline near the coast.According to the well-known theory of the wind-driven oceanic currents (SVERDRTJP,1938; DEFANT, 1951; HIDAKA, 1954), the surface current is deüected (except at theequator) by 45°in an anticyclonic direction from the wind, i.e., clockwise in the NorthernHemisphere and anticlockwise in the Southern Hemisphere. The difference 7V — TA be"tween the temperature of water (7V) and of air (TA) becomes largely negative, causing

149

Local wind Systems

111 YCKJIUH X* _

7777777777777777777777T7/

Fig.7. Sea breeze and coastal fog above cool upwelling ocean waters.

frequentfog and low-level stratus below an Inversion (Fig.7). As a consequence, a strongthermal circulation develops between land (äs heat source) and sea (äs heat sink). Whenthe water temperature drops to 12-15°C In contrast to the adjacent continent, where eventhe night minimum temperature hardly drops below 18°C, the sea-breeze may blowsteadily during the whole night. Then the thermal circulation persists during day andnight, and the low-level stratus is driven far Inland. In spite of these local onshore windsnear the coast, the large-scale wind System maintains the oif-shore component of thesurface waters together with the seasonally permanent upwelling process and the lowwater-temperatures along the coast.

Theoretical experiments on the sea-breeze System

Any thermally driven circulation can be most easily understood by remembering a quitesimple experiment designed by M. Margules (cf. EXNER, 1917, p.139). Let us imaginetwo adjacent rooms of different temperature and let us open the door between them.Then the cool air from one room will start to flow, near the floor, towards the warmroom, while near the ceiling the wann air will flow in the reverse direction towards thecool room. If we open only a slit of the door, we can easily demonstrate the direction ofthe circulation by using a burning candle: at middle heights we observe a neutral level ofno flow (Fig.8). If no heating or cooling takes place (i.e., under adiabatic conditions), thethermally driven circulation ceases when the temperatures in both rooms are equal. Nowthe cool air remains at the bottom and warm air at the top; the centre of gravity of thewhole system lies somewhat lower than at the initial stage. Under non-adiabatic con-ditions, however, a constant circulation will be maintained against friction by a constantheating. According to Carnot's law of thermodynamical processes, such a system canproduce work only if heating occurs under higher pressure than cooling; this is the mostprimitive example of a heat engine in the atmosphere. An application of these generalprinciples to the sea-breeze system äs produced by an isolated island is shown in Fig.9.A numencal treatment could be based on the well-known circulation theorem of BJERK-NES (1898), one of the fundamental theoretical considerations in modern meteorology.The differential form of the theorem, however, is not suitable for a numerical solution.

150

Theoretical experiments on the sea-breeze system

W

Fig.8. Two-chamber-model and thermal circulation.

In a baroclinic atmosphere (where isobaric and isothermic surfaces are cuttmg at a smallangle and thus cause a temperature gradient along an isobaric surface) the circulation C

along a closed material curve accelerates depending on the average isobaric temperaturegradient at the surface F enclosed by the curve. This is completely true within a motionlesscoordinate system: the rotation of the earth produces a deviation of the flow in an aati-cyclonic sense (in the Northern Hemisphere clockwise, in the Southern Hemisphere anti-

clockwise).For a rotating earth, the circulation theorem may be written:

= — (grad a grad p)dF — 2Q sin <p -7—dCr

dt

Here the term "grad" means the increase of a scalar quantity: a = I/Q is the specificvolume (proportional to the temperature); Q = density, p = pressure; 0 = angular

W

W

Fig.9. Vertical section of a thermal circulation between cold air (C) above sea and warm air (W) above aheated Island in the centre; dashed lines px-p3 = isobars.

151

Local wind Systems

velocity of the earth's rotation; (p = latitude; F' is the projection of the closed surface Fon a horizontal plane. The first term represents the thermal circulation in a restingcoordinate System, the second term the influence of the rotating earth. In the case ofdiurnal circulation, like that of land and sea-breezes, which change their sign twice daily,the influence of the second term is relatively small and produces only small deviations.The mutual adaptation of wind and pressure field takes several hours.The horizontal pattern of a diurnal thermally driven circulation depends on the distri-bution of heat sources and sinks, only slightly modified by the secondary role of therotating earth, which decreases towards and vanishes at the equator, Since the differentiaiheating of land and sea changes its sign periodically, twice in 24 hours, we cannot expecta stationary stage. Due to the storage of heat in the ground and especially in the water,

the diurnal circulations reach their greatest intensity at the time when the differentiaiheating is already weakening; the change of sign of the temperature gradient induces thebreak-down of the System. Thus we have to understand the System of diurnal circulationsäs a periodic System. The phase shift between radiation processes and wind usuallyamount to 2-3 hours; if precipitation occurs at a terminal convergence zone of the System,a secondary powerful heat source is introduced causing a delay of 6-8 hours.Any theory of the diurnal thermally driven circulation Starts from such a periodichorizontal temperature gradient äs a forcing function, which initiates horizontal pressureand density gradients and buoyancy and nnally a closed circulation. We also have to takeinto account the turbulent exchange of heat (described by the coefficient K), the transportof heat with the wind V (with the components w, v, w); we can neglect, in a first approxi-mation, the role of the water-vapour and its latent heat. The basic primitive equationsare the equations of motion:

dt

of continuity of mass simplified for an incompressible, homogenous atmosphere:

and of heat transport:

<50 g20dt c dz2

Here k is the Guldberg-Mohn coefficient of surface friction and 0 is the potential

temperature. The horizontal wind divergence div/i V =-~- + •— (u. v beine the x, v-dx dy

components of wind) can be frequently evaluated from surface winds or better from low-level wind data from pilotballoon stations. Since the earth's surface is practically im-permeable, the vertical component w at a certain level z can be determined by verticalIntegration from the surface.

Of the different approaches to a linear solution for the sea-breeze System, we mentionhere only the work of DEFANT (1950), who uses a coordinate System vertical to the coast-line and thus neglects the thermal differences along the coast. By assuming a System ofconvection cells with a widthi = 120 km and putting all quantities periodically in L,

152

Theoretical experiments on the sea-breeze system

TABLE III

RESULTS OF MODEL CALCULATTONS OF THE L AND-SEA-BREEZE CIRCULATION (F. DEFANT)

Latitude 9Coefficient of friction (sec"1)Temperature difference (°C)Wind compoiient u (perpendicular to coast)

near surface, amplitude (m/sec)near surface, phase shift versus temperature (h)height of neutral layer (« = 0) (m)upper return current,, amplitude (m/sec)upper return current, height of rnax. wind (m)

Vertical component w at neutral layer (u = 0) (cm/sec)

Wind component v parallel to coastnear surface, amplitude (m/sec)near surface, phase shift (h)upper return current, amplitude (m/sec)

Model I

0°0Ar = 1°

5.43 AT+3.73201.04 Ar6304.21 Ar

———

Model II

0°2.5 -10-*Ar

1.70 Ar+ 1.15000.26 Ar9201.37AT

—

—• —

Model III

45°2.5- 10-4

Ar

1.84 Ar+1.15000.25 Ar9201.76 Ar

0.72 Ar-9.80.11 Ar

he obtains a complete solution in complex form, in which the values for k and for thehorizontal Coriolis parameter/ = 2Ü sin cp may be varied. Assuming dT/dz — —l • 5°/kmand k — 2.75 • 105 cm2/sec, the results (Table III) depend linearly on the given horizontaltemperature amplitude ZlTbetween land and sea, which is arbitrarily fixed at 1°C.Any linearized model can only represent some general propertles in the atmosphere.Therefore, we cannot expect that such a frequent phenomenon äs the sea-breeze frontprogressing inland will be adequately treated. From the non-linear models we can men-tion here only that of ESTOQUE (1961), who employs a Cartesian coordinate System withthe coast äs j>-axis. Here the vertical turbulent ftuxes of heat and momentum are assumedto decrease with height and to disappear at the upper boundary at the 2 km-level, whichacts, in this way, like an Inversion. Also at this level all deviations of meteorologicalParameters from the average, disappear. The basic equations are nearly the same äs p. 152,if we assume an incompressible atmosphere. While at sea the temperatures remain con-stant, at land they are represented by a simple harmonic wave with an amplitude of± 10°C and extreme phases at 04hOO and 16hOO. Thus Estoque obtains, at day-time, asea-breeze with highest speed (about 10 m/sec) at a height of 250 m, starting at the coast-line and expanding towards both sides. During its way inland a convergence zonedevelops like a cold front, with strong vertical components up to 15 cm/sec. In middlelatitudes the Coriolis term results in a deviation of about 20° towards the right, andspiralling trajectories along the coast. The upper anti-sea breeze is weaker and at sea andabove the coast, subsidence occurs (Fig.10).While these results look quite realistic, they are less so for the night. This may be aconsequence of the model's assumption of a constantlapserate of —8°/km day and night.In a second paper (ESTOQUE, 1962), the local diurnal circulation is combined with a generalflowof 5 m/sec from the four main directions. On-shore winds reduce the horizontalgradients of temperature and density and suppress the sea-breeze circulation. If thegeneral fiow blows parallel to the coast with low pressure at land, friction causes (BRYSON

153

Local wind Systems

T(°C)

50 32 18 18 32 50 72 9 8 k m

-50 -32 -18 -8 -2 18 32 50 72 98 km

Fig.10. Sea breeze ciiculation. A. Vertical component (cm/sec) and isotherms (dashed). B. Horizontalcomponent perpendicular to the coast (m/sec). (After ESTOQUE, 1961.)

and KÜHN, 1961) an on-shore wind component with similar consequences. This is nottrue with a flow parallel to the coast in the reverse direction, i.e., with low pressure at sea.This coastal eftect produces divergence or convergence depending on the local differencesof surface stress and on the angle between coastline and flow direction.

Mountain wind Systems

Mountain wind Systems

Diurnally changing wind Systems in valleys and on slopes are, in principle, also thermallydriven circulations äs described by the circulation theorem. They are caused by theperiodically varying radiation and heat budget of the slopes and depend on azimuth,inclination, and plant cover and other surface conditions. Due to these differences it ismuch more difficult to design sufficiently realistic modeis; many of the ideas discussed inearlier literature are more or less incomplete.A fairly good example of mountain wind Systems (Fig. 11) is known from the Raurisvalley in the Austrian Alps (STEINHAUSER, 1968) just north of the famous meteorologicalobservatory at the Sonnblick. Here the diurnal and seasonal variations of the surfacewind are largely dominated by the day-time valley-breeze from the north and the nocturnalmountain breeze from the south. Only during winter—when the slopes are snow-coveredand the local heating differences are much smaller—is the diurnal System insignificant.In a quite similar Situation in the large Salzach valley near Salzburg, this diurnal windSystem is much less regulär (EKHART, 1944). Since in meteorology the winds are usuallydesignated from their origin (contrary to the use in oceanography), we understand äsvalley breeze (or up-valley breeze) the day-time current blowing upstream and äs moun-tain breeze (down-valley breeze) the reversed nocturnal System blowing downstream.In several cases—such äs the upper Rhone valley in Switzerland—the slope of the trees

l 3 5 7 9 11 13 15 17 19 21 23h

11 13 15 17 19 21 231 3 5 7

Fig.ll. Frequency of down-valley (A. Southwest-southeast) and up-valley (B. Northwest-northeast)winds (in percent of all observations) at Bucheben (Rauris valley, Austria). (After STEINHAUSER, 1966.)

755

a-

3000-

1000

Fig.12. Vegetation profile along La Paz River (Bolivia), simplified. Nevado = glaciers; ptina = high-altitude grassland; montana = mist forest; cabezera devalle = temperate forest; valle templada = mesophytic bushiand; valie calida = hot semi-desert. (After TROLL, 1952.)

Mountain wind Systems

indicates the persistency and strength of the valley breeze (WEISCHET and BARSCH, 1963;YOSHINO, 1964).

The simplest case is that of the katabatic down-slope winds. During the night, the airnear the surface is cooled by the long-wave radiation of the soil; this cold air flows down-slope, driven partly by the density gradient directed towards the relatively warm air offthe slope, but mainly by its own weight, i.e., by gravity (DEFANT, 1933). The combinedeffect of cooling from the surface and katabatic motion along slopes with locally differen-tiated friction causes, in many places, more or less regulär pulsations—avalanches ofair—with periods between 5 and 30 minutes. If such a flow of cold air near the surface ishampered by a forest-strip, an artificial dam, or a narrowing of the valley, it is blockedand forms a cold air lake. At middle and higher latitudes this is a typical Situation forlocal nocturnal frosts. During the day air from the slope surface is heated which causesimmediately a reverse density gradient from the cool and dense free air towards the slopeand thus a buoyancy of the heated air with upslope motion (DEFANT, 1949). TheseSystems can easily be made visible by smoke experiments.In every valley, the sum of the local upslope motions causes nearly simultaneously, atday-time, a valley breeze filling the whole valley (WAGNER, 1938; DEFANT, 1951). Sincesuch upslope motions at both valley flanks cause a divergence of the flow, we have toexpect, normally a weak subsiding motion along the axis of the valley äs a consequenceof the above-mentioned continuity equation. In high mountains of the subtropical andtropical zone these valley breezes frequently reach gale intensity. This occurs in theHimalayas and in the gorges extending downwards frorn the altiplano of Bolivia into theAmazonas plain which are upto 5,000 m deep. In these areas TROLL (1952) and SCHWEIN-FURTH (1956) have demonstrated, that the usually much stronger day-time System fre-quently can be evident in the Vegetation of the valley, reaching from the not semi-desertalong the axis to the humid forest along the slopes, where the ascending and expandingair usually reaches the dewpoint producing a cloud belt, and finally to the cool meadowsabove the timberline (Fig. 12).Similarly, at night time, all downslope motions grow together into a mountain breezewith weak lifting motion along the axis of the valley, due to the convergence of the down-slope winds. Where a valley with a large catchment area of cold air opens into a largervalley or into a plain, these evening mountain breezes are often strongly feit because oftheir remarkable coolness and because they replace the polluted air by fresh air. As ob-served in land-sea circulations, the descending nocturnal System is usually somewhatweaker than the ascending day-time circulation due to the weaker horizontal coolingdifferences; gravitation, however, adds to the driving forces of the circulation. Further-more the thermal stratification of the air is stable during night, but unstable during thedaylight hours. These effects contribute to the fact, that in narrower valleys only the day-time System is reflected in the Vegetation, äs observed by TROLL (1952, 1959). However,in very large tropical valleys—for example the longitudinal rifts of Colombia—bothSystems can grow to the same intensity: then the converging mountain breezes initiate,above the floor of the valley, nocturnal convection with more or less regulär thundershowers (TROJER, 1959). Fig.13 demonstrates the shift of the diurnal course of rainfallwith altitude, from the nocturnal rains at the bottom to the early afternoon showersnear the ridges. The rainfall minimum occurs, at all altitudes, during the morning.Extensive series of pilot balloon measurements of these wind Systems have been under-

757

Local wind Systems

H E I G H T A B O V E

24h

Fig.13. Time-height cross-section of the diurnal Variation of rainfall (in percent of the daily average) inthe Cauca Rift, Colombia, ca. 5°N. (After TROJER, 1959.)

taken by A. Wagner and his colleagues in the Austrian Alps. These measurements showthe wide variety of conditions that exist according to the exposure of the slopes and thewidth of the valley. Buettner and Thyer have reported more recent data from the area ofMount Rainier (Cascade ränge in the northwestern U.S.A.). They measured the iocalizedlifting (up to 2.5m/sec) above the mountain crests where upslope motions from bothsides converge, together with the cross-valley components at the level of the ridges.A characteristic early morning observation was that of a balloon carried at first down-wards along the shaded side of the valley, and then strongly upwards at the oppositesimlit side, thus indicating a simple thermally driven cross-valley circulation.The reversal of the slope circulation depends strongly on the radiation budget. In an Al-pine valley near Davos (URFER-HENNENBERGER, 1967) the down-slope breeze ends, in96% of all cases, within 20 min before or after local sunrise, and the up-slope breezeStarts, in 92% of all cases, between sunrise and 40 min later. The down-valley breeze ends24 min after sunrise; the up-valley wind with only slightly larger deviations Starts 60 minafter sunrise. Since the transition period between the development of the different scalesof local circulation lasts usually less than one hour, it seems unnecessary to distinguishbetween several idealized sub-stages (WAGNER, 1938).

In the southern Red Sea Trench, during the cool season, shallow orographically chan-nelled wind Systems usually converge somewhere between 14°-19°N (Red Sea conver-gence zone). In this area, where the layer of moist maritime air rises usually up to about2,000m, the diurnal circulations along the escarpments at both flanks produce, near andafter noon, regulär orographic clouds at 1,200-1,800 m. These clouds produce precipi-tations of not less than 1,000 mm/year and a humid mist-forest, which contrasts marked-ly to this otherwise arid latitude. During the night, the down-valley winds converge abovethe Red Sea, forming nocturnal stratus clouds and occasional rains, e.g., at Massawa.This is a typical example of a combination of sea and valley-breezes and of iand andmountain-breezes, äs it often occurs at mountainous coasts (cf. Fig.3).

158

Mountain wind Systems

For reasons of continuity we have to expect that the diurnal Systems of valley andmountain-breezes are complemented above the crests and peaks of the mountain rangesby a reverse systern of "anti-winds". We find, in fact, that these anti-winds embrace amuch larger area than the narrow valleys and are, therefore, much weaker. Above theAlps an anti-wind in the order of not more than 30 cm/sec has been verified throughsystematic difference calculations (BURGER and EKHART, 1937). However, if the generalflow above the mountains is sufficiently weak, direct evidence can be found, e.g., at MountRainier (BUETTNER and THYER, 1965). This wind System is maintained by the horizontalintegration of all vertical components in the mountain area. BLEEKER and ANDRE (1951)have demonstrated that such a System produced by the large Rocky Mountain System inwestern North America is responsible for the wide-spread occurrence of nocturnalthunderstorms in the Great Plains (cf. also LETTAU, 1968). During summer night-time thelarge-scale katabatic flow from the Rockies converges with the prevailing southerlywinds at the Great Plains, while during day time the ascending flow causes large-scaledivergence onthe order of 10~6 sec^1. Similar processes are very likely at the Argentinepampas. Such orographically induced and thermally driven wind Systems of varyingscales—from the local scale of a few metres upto a regional scale of several hundreds ofkilometres—interact and eventually merge into a giant respiration System of the largemountains.Of the many local peculiarities only one will be mentioned here: the local infringement ofascending valley-breezes across asymmetric watersheds. The best-known example is theMaloya wind: the valley-breeze ascending from the steep Bergell valley and extendingacross Maloya Pass downwards into the open Engadin valley (Switzerland) to St. Moritzand still further. Here the strengest heating of the air takes place at the flanks of the wideEngadin valley, where the peaks are higher and the slopes more extended than in thenarrow Bergell. Similar cases are not rare (e.g., at Davos). We have already mentionedthe combined sea-valley-breeze crossing the asymmetric escarpments of the WesternGhats (India) between Bombay and Poona and of the Somali peninsula south of Berbera.It is the locally differentiated intensity of heating, the term HQ, which is responsible forsuch local anomalies.Naturally, the intensity and duration of all mountain wind Systems depend on localradiation and heat budget—again primarily on the term 770. In this case azimuth and

TABLE IV

HEAT BUDGET DATA FROM THE TURKESTAN MOUNTAINS

Lat. 41°N, height 3150 m. Monthly averages for September, above bare debris.(According to AISENSHTAT, 1966)

Albedo

Horizontal area 0.14

North-facing slope (33°) 0.20

South-facing slope (31°) 0.15

Daily sums (Ly/day) Noon values (Ly/h)

RN H0 L-E0 G0 RN H0 L-Eo Go

348 248 50 50 52 38 11 3

145 85 41 19 22 12 7 3

426 304 75 47 61 45 12 4

159

Local wind Systems

sw

NW-

360'

-180'

-270'

360°

E x t r a t e r r e s t r i s c h e Sonnenstrahlung, 40° N, Hänge von 31 ° (60%) Neigung,in Ly /d

Fig.14. Daily sums of direct solar radiation (neglecting the influence of atmosphere and clouds) at steepslopes (60% = 31°), Lat. 40°, äs a function of azimuth and season. (After LEE, 1963.)

inclination of the individual slopes which dominate the radiation pattern are subject togreat variety and the local heat budget and friction are controlled by the Vegetation. Wequote some data of net radiation T?N and heat budget terms (Table IV) from the highmountains of central Asia with their powerful radiation.

Here the net radiation RN and the flux of sensible heat HQ at the south-facing slope areabout three times larger than at the opposite side. Because of the semi-arid conditions ofthat area, H0 uses 55-75% of the available radiation energy.

As shown in Chapter 3, p. 109, the direct solar radiation (here neglecting atmosphericdiffusion and cloudiness, i.e., under extra-terrestrial conditions) can be easily computedfor each slope äs a function of five parameters: latitude, declination and azimuth of thesun (equivalent to season and hour), azimuth (exposure) and inclination of the slope.Fig. 14 shows the above-mentioned results of LEE, 1963, äs daily radiation sums valid forlatitude 40°N and a steep slope (inclination 60% or 31°) äs a function of exposure andseason. The effective global radiation consists of the direct radiation of the sun afterpassing through the atmosphere and the diffuse radiation of the sky. Since the latter partincreases with decreasing elevation of the sun, the atmosphere causes some levelling ofthe contrasts. This is even more true if we consider the effective outgoing radiation whenwe calculate the net radiation RN. Under simplified assumptions and taking into accountdiurnal variations and the role of orographically induced variations in day length, modelcomputations of these quantities could be undertaken without serious complications.Unfortunately no numerical results of this kind are available äs yet.At isolated tropical mountains the normal diurnal course of the Insolation can be sub-stantially varied by the clouds induced by the local diurnal circulations (TROLL, 1952).

160

Mountain wind Systems

During the morning—at least until about lOhOO—the sky is mostly clear with only a fewpatches of high or medium clouds. During the afternoon, the ascending circulationsproduce regularly strong convection with cumulus clouds covering the slopes and im-peding direct Insolation (Fig.15). As a consequence of this more of less regulär oocurrence, the eastward facing slopes obtain, before noon, mnch more radiation than theopposite western slopes, which in the afternoon receive only diffuse sky radiation andlong-wave counter-radiation (Geiger, p.116). This local differentiation can also bereflected in the Vegetation which is usually more humid at the western slopes.If the catchment area of a valley is partly covered either by an extended area of perma-nent snow or by a glacier, the day time Insolation is ineöective due to the high albedo(0.7-0.9) of snow and ice. In this case the cooling effect of ice leads to a permanent temper-ature Inversion, together with negative values of H0. Under the influence of gravity, thecold air near the surface flows persistently downwards: this glacier-wind is usually quiteshallow, with a vertical extension of a few decametres, but extremely persistent and, inthe vicinity of the ice margin, has great influence on the local climate (HoiNKES, 1954).Due to the increased heating of the surrounding area, the intensity of such thermalcirculation reaches its highest value near noon. At the snout of the great Pasterze Glacierat the Großglockner (Austria), the relative frequency of such glacier winds at 14hOOreaches, during July and August, 90%. They spread underneath the ascending valley-

07h

W

16h

Fig.15. Effect of Insolation and diurnal convection on an isolatedtropicalmountain.^ = snowboundary.(After TROLL, 1952.)

161

Local wind Systems

breeze, with striking coolness. From the view-point of energy, it should be stressed thatonly in the case of snow and ice surfaces, is the heat sink situated near the surface, i.e.,lower than the heat source.Most important and quite similar are the nearly persistent katabatic winds at the marginsof the great ice-caps of Greenland and Antarctica. At Port Martin (Adelieland) they blowwith furious intensity and remarkable persistency; they have a vertical thickness of about300 m and an average annual velocity of 19 m/sec. The average speed of individual dayscan be äs high äs 45 m/sec (LoEWE, 1960). This is one of the strongest thermally drivencirculations with nearly constant direction; it is only shortly interrupted, even by intensetravelling cylones.

Theoretical considerations on mountain winds

Starting from simple assumptions, DEFANT (1933) has estimated the speed u of a coldcurrent gliding down along a slope with an elevation angle ff, assuming the low-levclatmosphere consisting of two homogeneous layers with potential temperatures of 00

(below) and 0, (above). In this case he obtains the equation:

U2 = &. Q ' -©o sin aK 0

with the friction coefficient K derived from G. Taylor's skin-friction theory (K = 0.0025)and the thickness z of the downwards moving cold air. Such thickness values have beenfound to be 20-50 m above Alpine or Scandinavian glaciers, but 200-400 m above thecontinental ice-caps of Greenland and Antarctica. Initially this cold air drainage Startsquite weakly in a shallow surface layer of some 20 cm; even in strong katabatic currentsthe wind maximum lies only a few decametres above the snow. Since most glacier sur-faces are convex, the local divergence of glacier wind causes subsidence, persistentlycarrying fresh air near the snow surface where it will be cooled by turbulent flux ofsensible heat directed downwards. Considering the driving forces it can be shown thatgravity, the interior (local) pressure gradient produced by the cold air itself, and theexterior (synoptic-scale) pressure gradient are of equal magnitude; thus the latter canparalyze—at least in strong-wind situations—the development of the glacier wind.The physical mechanism of all kinds of mountain winds has caused much discussion for along time; this discussion has been reviewed by DEFANT (1951). Most of these consider-ations deal only with one or another partial aspect of the highly complex phenomenon.This is also true for the very important work of the Innsbruck school of WAGNER (1932)where the upper anti-winds are not always sufficiently stressed. In fact, valley andmountain-breezes are only branches of a thermally driven circulation. We can, therefore,neglect the authoritative, but somewhat misleading ideas of J. Hann, who treated theseSystems äs being periodic air-mass displacements (cf. WAGNER, 1932 and MÖLLER, 1951);this is inconsistent with most aerological observations.In fact, the fundamental role of Bjerknes' circulation theorem cannot be neglected, evenif the upper anti-winds are always much weaker than the lower branches, which arelocally enforced by mesoscale orographical effects. Considering a cross-section through aslope in a vertical plane, it is obvious that the air is heated during day-time and cooled

162

Theoretical considerations on mountain winds

00-06.H

Fig.16. Slope-wind circulation in a valley during day (A) and (B) night (clouds only in wide valleys).

during the night from the underlying slope. This causes an. inclination of the isothermallevels compared with the isobaric levels and thus a baroclinic circulation between thesurface layer of air immediately above the slope and the air columa at the central axis ofthe valley. This thermal circulation induces along the slope a shallow but intense flowupwards during the heating period and downwards during the cooling period. As aneffect of mass contlnuity, a closed circulation develops with cross-valley and verticalwind components (Fig.16). Heating of the ascending air near the slopes from the surface(by the term ffQ~) and subsidence above the valley centre lead to a warming of the wholevalley atmosphere compared with the free atmosphere above the adjacent plain at thesame level. This differential heating causes and maintains a larger-scale thermal circu-lation between the whole mountain area and the plain. The effect of these large-scalecirculations has been demonstrated, at the Rocky Mountains, by BLEEKER and ANDRE(1951). Similar Systems exist above the Altiplano of Bolivia and Peru, äs shown byGUTMANN and SCHWERDTFEGER (1965) above the huge Tibetan highlands (FLOHN, 1968)and most presumably also over Ethiopia. In these cases, the latent heat liberated by theprecipitations induced by the ascending branches of the circulation adds much to thetotal energy of the large-scale System, at least in a super-moist adiabatic atmosphere. Inthe case of Tibet this effect intensifies and prolongs the day time circulation to suchan extent that the local night time circulations are partly subdued. A fairly completenumerical theory of mountain wind Systems has been designed by THYER (1966), basedon the same fundamental equations äs used by Estoque in his theory of land and sea-breezes. In his model he assumes a V-shaped Symmetrie valley and surrounding ridgeswith constant altitude, which is closed at. the upper end and stopping abruptly at its

163

Local wind Systems

D

Fig.17. Components of a valley breeze System in its initial stage (cm/sec). Slope wind circulation;A. horizontal component (i>); B. vertical component (H>); C. longitudinal component (u); D. observedlongltudinal component («). Negative signs = up-valley; positive signs = down-valley. (After THVER,1966.)

lower ("distal") end, when opening into the plain. He also assumes constant heatingof the air from the surface. The Coriolis term is neglected äs is any general flow. Thecomputation is carried out only for the first short time-steps; it can therefore yieldonly some hints of the future developments. These hints, however, seem quite realistic.The cross-valley circulation is, äs can be expected from the assumptions made, sym-metric in cases of both day time heating and night-time cooling, but with reversedflow directions (Fig.17). If the complete three-dimensional model with the valley axis äsabscissa is used, a valley-breeze and (during night) a mountain-breeze along the valleyappears, filling the bulk of the valley, äs actually observed. Above the ridge the anti-breezes occur much weaker than the lower currents. At the exit of the valley into theplain the Systems weaken and spread laterally. In a longitudinal section the highest speedoccurs at this exit since heating in the plain remains of minor importance. With steeperslopes the currents are intensified. The height of the boundary between valley-breeze and

164

Mechanical effect ofmountains

anti-wind is lowest above the valley axis; this is not necessarily a mere consequence of thesubsidingverticalbranch. Obviouslythesimilaritiesbetweenmodeland reality are ratherconvincing; this indicates the validity of the model assumptions and the usefulness ofsuch numerical computations.In actual cases we have to consider the periodic behaviour of all meteorological para-meters. Whüe in a dry atmosphere only a minor time-lag occurs, the introduction of thehydrological cycle into the model should lead to the observed time-lag of about onefourth of the total phase angle. These few considerations of numerical approaches tothe theory of such thermally driven circulations show how far we can understand the basicprinciples of these Systems. Under such circumstances it will be to some extent possibleto predict the climatological consequences of the construction of a large-scale artificiallake and other experimental climatic modifications. Such local effects act persistently,depending on radiation conditions, i.e., on advective changes of cloudiness, water-vapour content etc. In subtropical and tropical areas at least their scale and their energyare of the same order of magnitude äs those of travelling synoptic Systems; both Systemsinteract in a quite individual manner. In the tropics their influence on the local v/eatherand climate can hardly be overlooked since horizontal pattern and diurnal course ofprecipitation are largely controlled by such geographically fixed Systems.A new type of mountain circulation named "thermo-tidal wind" has been describedrecently by LETTAU (1968), starting from evidence at the western slope of the PeruvianAndes. This flow type results from the horizontal temperature gradient reversing fromday to night—above a large-scale terrain slope such äs the Great Plains east of the RockyMountains: it creates a geotriptic low-level flow, where, under special conditions, thefrictionally induced cross-isobaric component may counterbalance the direct thermalcirculation äs described above. The interaction between this geotriptic flow and thethermal circulations can be very complex, due to the time-dependent turbulent difiusioncoefficient K and to the role of a term/2 — Q2 = H2 (4 sin2 9? — 2)—the horizontal Corio-lis parameter/— 2£> sin cp—which changes its sign at lat. 30° (2 sin 9? = ± 1).

Mechanical effect of mountains

In addition to these thermally driven local wind Systems, we frequently observe climato-logically relevant features of the wind field field which are mechanically produced byisolated hills, mountain ranges, escarpments and jet-like wind-gaps. Due to their largenumber (SCHAMP, 1964) we have to restrict our considerations to a few generalized andsimple cases:(1) As predicted by a theory first applied to the slow currents at the ocean noor (EKMAN3

1932), air tends to deviate in an anticyclonic direction when forced to flow over ascendingground. Thus a westerly airflow above a large plateau or a mountain ränge deviatesmainly by curving around its poleward flank. When flowing over descending ground, acyclonic curvature is observed. Due to this effect, large mountains modify even globalpatterns of atmospheric circulation, äs shown by BOLIN (1952).When large-scale currents approach the mountains, the accumulation of mass in the areaof speed decrease (convergence) results in pressure rise, lifting of air and anticyclonicdeformation of the flow. At the lee-side the speed of the current increases again which

165

Local wind Systems

V g e o s t r.

V *

Fig.IS. Deviation of atmospheric fiow above mountains. A. Mechanical uplift and thermally inducedcirculation cause a divergent flow component V*. B. Induced by an additional Coriolis force C*t

superposed to the general fiow, the general flow deviates anticyclonically around the mountain ridge andproduces downstream a cyclonic curvature (trough).

results in pressure fall, subsidence and cyclonic flow. If we assume that the large-scalepressure gradient above the mountain, which maintains more or less zonal flow, is, at thewindward side, smaller than the average, then the Coriolis component/causes a pole-ward deviation of the flow. Vice-versa at the lee-side: the increasing speed of the upperflow is faster than the pressure gradient, and is therefore accompanied by an equatorwarddeviation. (Fig.lS).By applying general lee-wave theory (QuENEY et al.3 1960) to the case of a large-scalemountaia (with a diameter of about 100 or 1,000 km), we find for reasonable values of theaverage zonal fiow u (^ 10 m/sec), a train of damped waves, starting with an anticyclonic

166

Mechanical effect of mountains

displacement of the flow immediately before reaching the obstacle. After crossing themountain, a first cyclonic lee-wave trough develops in a distance of about 2,000 km,while the crest of the next anticyclonic wave occurs after some 6,000 km and a secondcyclonic trough after some 9,000 km. Rossby's simplified longwave theory has beenfrequently used to describe the large-scale average flow of the westerlies in middle lati-tudes. In a horizontal non-divergent flow, where the Rossby parameter ß = df/By(f= 2Q sin (p) is constant, the speed C of propagation of a Rossby wave is given by:

where L is the wave length. The stationary case is identified by C = 0; in this case thewave-length changes slov/ly with the square root of u. When investigating the areas oflong wave formation and decay, it could be shown (ESSENWANGER, 1953), that the for-mation of troughs occurs most frequently at the lee-side of the largest mountains, and theformation of ridges at their windward side. Furthennore, the theory predicts the oc-currence of weak secondary quasi-stationary waves in the time-averaged high-tropo-sphericflowe.g.jintheEuropeanarea, whicharelocallyinitiated in the Rocky Mountainsin western North America, and that similar mountain-induced phenomena occur in thePacific area and within the southern westerlies.(2) In the vertical direction each mechanical obstacle tends to produce wave-like motions(QUENEY et al, 1960) extending to heights many times higher than the mountain itself.In these cases the flow is intensified and forced into a series of waves (Fig. 1 9); theirascending branches are indicated by elongated clouds shaped like a fish or cigar, with alens-like cross-section. Such "lenticularis" or wave clouds are frequently found in the leeof an obstacle, especially in the middle and upper troposphere and sometimes at severalsuperimposed levels. Their wavelength varies between about 5 and 30-40 km, with avertical amplitude of a few kilometres. They are characterized by large organized up-draughts and downdraughts which have vertical components up to 6 m/sec and even

Lee Waves (Lenticularis)

30 km

Fig,19. Lee waves and lenticularis clouds behind an obstacle. (Simplified after KUETTNER, cf. QUENEY et al..1960.)

167

Local wind Systems

more, which is enough to lift gliding planes äs if by a large elevator. Usually they areaccompanied by moderate or severe turbulence. Satellite pictures have revealed, thatsuch elongated lee-wave Systems may extend over hundreds of kilometres, with up to20-25 waves (ANDERSON et al., 1966). In the lowest layers behind the obstacle, turbulentrotor clouds frequently develop. These have been observed at the Alps and at the Scandi-navian mountains, äs well äs at the Rocky Mountains, the South American Andes, atCaucasus, Hindukush, at the South Island of New Zealand and at many other places.(J) In the troposphere, ascending cloud air condenses water-vapour; due to the liberationof condensation heat, the cooling rate is only about 5-7°/km. In contrast to this, subsidingair warms at the usual dry-adiabatic lapse-rate, i.e., 10°/km (Fig.20). Since much of thewater-vapour has been precipitated out at the ascending part of the motion, the de-scending air is desiccated and its relative humidity drops to desert-like values. Along manyvalleys of the northern side of the Alps, this unusually warm and dry wind, accompaniedby excellent visibility and lenticularis clouds, has been known for centuries äs "Föhn"(cf. VON FICKER and DE RUDDER, 1948). During the winter half-year, its mainclimaticeffect is rapid melting and evaporation of the snow-cover; only after the large-scalepressure gradient causes the drainage of the cold air at the bottom of the valley, is thedescending motion of the föhn instigated. Similar descending warm winds are frequent-ly observed at the lee-side of mountain ridges: äs examples we mention here only theChinook at the eastern flank of the Rocky Mountains in Canada and U.S.A. and theZonda (GEORGII, 1954) at the Argentine Andes, äs well äs dry winds in the eastern low-lands of southern New Zealand. When the prevailing atmospheric flow is forced to crossmountain ranges, the ascending branch at the windward side of the mountain producesa belt of heavy precipitation; nearly all regional and local rainfall maxima are caused bythis eifect. As a further consequence the descending branch at the leeward side ofthe mountains is characterized by dry conditions; this is especially pronounced ininterior basins and valleys. A striking example occurs in Norway: at the western side ofthe mountains (with an average height of 1,600-2,000 m) the annual precipitation risesto 250-400 cm and in some areas (according to run-off data and hydrological budget)even to 600 cm. Across the mountains in the Otta valley (61.5°N) only 25-30 cm fall

900

100O

Fig.20. Temperaturevariations in a Föhn Situation, a = dry-adiabatic lifting,-£ = moist-adiabatic liftingof cloud air above condensation level (900 mbar) up to mountain height (700 mbar ~ 3,100 m); c = dry-adiabatic subsidence.

168

References

annually; agriculture needs extended irrigation, and arid salt-soils are observed. Simi-larly the rainfall at the San Gabriel Mountains in the hinterland of Los Angeles, can beäs much äs 150 cm/year, while at the Mohave Desert it is only about 12 cm/year in a hori-zontal distance of a few tens of kilometres.(4) If the descending motion is imposed on very cold air-masses, adiabatic warming is insome areas insufncient to mask the polar origin of the air. This is especially true at thenorthern Adriatic Sea where the "Bora" is a very typical northeasterly air current of galeforce, partly accornpanied by clear anticyclonic weather, partly caused by local cyclo-genesis and then accornpanied by high cloudiness, snow and rain. Quite similar conditionsexist on the Caucasus coast of the Black Sea. When cold Continental air flows above anunfrozen sea, the heating of the air by the warm water (our term #„) together with theevaporation (L -E0) reach extremely high values and cause very unstable lapse-rates, andby vertical exchange of mornentum, heavy gusts. Generally speaking at a stable lapse-rate the warm subsiding air needs special synoptic conditions before it can replace thecool and dense air below the Inversion, while unstable lapse-rate favours downwardmotion and vertical exchange of momentum, potential temperature and absolute humidi-ty. At the southern Alps, föhn winds from northerly directions penetrate into the valleysmore frequently than the southerly föhn winds at the northern flank. In each case theformation of a baroclinic field (i.e., temperature differences along isobaric surfaces)favours the downward penetration of the föhn (FREY, 1953).(5) Where the large-scale flow runs nearly parallel to a valley or a mountain gap, the low-level winds are locally strengthened and canalized. Sometimes the direction of suchorographically induced jet-streams with a wind maximum in lower layers is demon-strated by the shape of trees (WEISCHET and BARSCH, 1963; YOSHINO, 1964), äs in theRhone valley in Switzerland and southern France. The "Mistral" is such an example of anortherly flow, intensified in the gap between the Alps and the Cevennes and quitefrequently occurring during winter and early spring. A similar northeasterly flow causesthe "Bise" (beeze) at the Lake of Geneva, between the Alps and the French Jura Moun-tains. Similar strongjet-like winds are observed at Gibraltar (see Chapter 3, p.119), andat the large mountain gaps and passes of central Asia.Other local winds are of synoptic origin, i.e., caused by travelling cyclones and anti-cyclones; here we mention only a few examples in the Mediterranean. At the approach ofcyclonic disturbances from the west, southerly winds cause, in southern Italy, the"Sirocco", in Lybia the "Ghibli", in Egypt and Middle East the "Khamsin" and "Sa-mum". These winds carry air of African origin which is often laden with dust and sandparticles, and are characterized by extended altostratus-altocumulus sheets from theadvancing upper trough.

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