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CHAPTER 3
ATMOSPHERIC CIRCULATION
To understand large-scale motions of theatmosphere, it is essential that the Aerographer’s Matestudy the general circulation of the atmosphere. Thesun’s radiation is the energy that sets the atmosphere inmotion, both horizontally and vertically. The rising andexpanding of the air when it is warmed, or thedescending and contracting of the air when it is cooledcauses the vertical motion. The horizontal motion iscaused by differences of atmospheric pressure; airmoves from areas of high pressure toward areas of lowpressure. Differences of temperature, the cause of thepressure differences, are due to the unequal absorptionof the Sun’s radiation by Earth’s surface. Thedifferences in the type of surface; the differentialheating; the unequal distribution of land and water; therelative position of oceans to land, forests to mountains,lakes to surrounding land, and the like, cause differenttypes of circulation of the air. Due to the relativeposition of Earth with respect to the Sun, much moreradiation is absorbed near the equator than at otherareas, with the least radiation being absorbed at or nearthe poles. Consequently, the principal factor affectingthe atmosphere is incoming solar radiation, and itsdistribution depends on the latitude and the season.
GENERAL CIRCULATION
LEARNING OBJECTIVE: Recognize howtemperature, pressure, winds, and the 3-celltheory affect the general circulation of Earth’satmosphere.
The general circulation theory attempts to explainthe global circulation of the atmosphere with someminor exceptions. Since Earth heats unequally, the heatis carried away from the hot area to a cooler one as a
result of the operation of physical laws. This globalmovement of air, which restores a balance of heat onEarth, is the general circulation.
WORLD TEMPERATURE GRADIENT
Temperature gradient is the rate of change oftemperature with distance in any given direction at anypoint. World temperature gradient refers to the changein temperature that exists in the atmosphere from theequator to the poles. The change in temperature ortemperature differential, which causes atmosphericcirculation can be compared to the temperaturedifferences produced in a pan of water placed over a gasburner. As the water is heated, it expands and its densityis lowered. This reduction in density causes thewarmer, less dense water to rise to the top of the pan. Asit rises, it cools and is forced to the edges of the pan.Here it cools further and then sinks to the bottom,eventually working its way back to the center of the panwhere it started. This process sets up a simplecirculation pattern due to successive heating andcooling.
Ideally, the air within the troposphere may becompared to the water in the pan. The most direct raysof the Sun hit Earth near the equator and cause a netgain of heat. The air at the equator heats, rises, andflows in the upper atmosphere toward both poles. Uponreaching the poles, it cools sufficiently and sinks backtoward Earth, where it tends to flow along the surface ofEarth back to the equator. (See fig. 3-1).
Simple circulation of the atmosphere would occuras described above if it were not for the followingfactors:
3-1
1. Earth rotates, resulting in an apparent forceknown as the Coriolis force (a deflecting force). Thisrotation results in a constant change to the area beingheated.
2. Earth is covered by irregular land and watersurfaces that heat at different rates.
Regions under the direct rays of the Sun absorbmore heat per unit time than those areas receivingoblique rays. The heat produced by the slanting rays ofthe Sun during early morning may be compared withthe heat that is produced by the slanting rays of the Sunduring winter. The heat produced by the more directrays at midday can be compared with the heat resultingfrom the more direct rays of summer. The length ofday, like the angle of the Sun’s rays, influences thetemperature. The length of day varies with the latitudeand the season. Near the equator there are about 12
hours of daylight with the Sun’s rays striking thesurface more directly. Consequently, equatorial regionsnormally do not have pronounced seasonal temperaturevariations.
During the summer in the Northern Hemisphere,all areas north of the equator have more than 12 hoursof daylight. During the winter the situation is reversed;latitudes north of the equator have less than 12 hours ofdaylight. Large seasonal variation in the length of theday and the seasonal difference in the angle at whichthe Sun’s rays reach Earth’s surface cause seasonaltemperature differences in middle and high latitudes.The weak temperature gradient in the subtropical areasand the steeper gradient poleward can be seen in figures3-2A and 3-2B. Note also how much steeper thegradient is poleward in the winter season of eachhemisphere as compared to the summer season.
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NORTH POLE
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POLAR REGIONAREA OF LEAST HEATING
POLAR REGION AREA OFLEAST HEATING
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EQUATORIAL REGION
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Figure 3-1.—Simple circulation.
3-3
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Figure 3-2A.—Mean world temperature for January.
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Figure 3-2B.—Mean world temperature for July.
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Figure 3-3A.—Mean world pressure for January.
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Figure 3-3B.—Mean world pressure for July.
PRESSURE OVER THE GLOBE
The unequal heating of Earth’s surface due to itstilt, rotation, and differential insolation, results in thewide distribution of pressure over Earth’s surface.Study figures 3-3A and 3-3B. Note that a low-pressurearea lies along the intertropical convergence zone(ITCZ) in the equatorial region. This is due to thehigher temperatures maintained throughout the year inthis region. At the poles, permanent high-pressure areasremain near the surface because of the lowtemperatures in this area throughout the entire year.Mainly the "piling up" of air in these regions causes thesubtropical high-pressure areas at 30°N and S latitudes.Relatively high or low pressures also dominate otherareas during certain seasons of the year.
ELEMENTS OF CIRCULATION
Temperature differences cause pressuredifferences, which in turn cause air movements. Thefollowing sections show how air movements work andhow they evolve into the various circulations—primary,secondary, and tertiary.
To explain the observed wind circulation overEarth, three basic steps are used. The first step is toassume Earth does not rotate and is of uniform surface;that is, all land or all water. The second step is to rotateEarth, but still assume a uniform surface. The third stepis to rotate Earth and assume a non-uniform surface.For now, we deal with the first two steps, a non-rotatingEarth of uniform surface and a rotating Earth ofuniform surface.
Static Earth
The circulation on a non-rotating Earth is referredto as the thermal circulation because it is caused by the
difference in heating. The air over the equator is heatedand rises (low pressure); while over the poles the air iscooled and sinks (high pressure). This simplecirculation was shown in figure 3-1.
Rotating Earth
In thermal circulation, the assumption was madethat the Earth did not rotate, but of course this is nottrue. The rotation of Earth causes a force that affectsthermal circulation. This rotation results in thedeflection to the right of movement in the NorthernHemisphere, and to the left of the movement in theSouthern Hemisphere. This force is called the Coriolisforce. The Coriolis force is not a true force. It is anapparent force resulting from the west-to-east rotationof Earth. The effects, however, are real.
Arctic rivers cut faster into their right banks thantheir left ones. On railroads carrying only one-waytraffic, the right hand rails wear out faster than theleft-hand rails. Artillery projectiles must be aimed tothe left of target because they deflect to the right.Pendulum clocks run faster in high latitudes than inlower latitudes. All these phenomena are the result ofthe Coriolis force, which is only an apparent force. Themost important phenomena are that this force alsodeflects winds to the right in the Northern Hemisphere.Therefore, it is important to understand how this forceis produced.
As Earth rotates, points on the surface are movingeastward (from west to east) past a fixed point in spaceat a given speed. Points on the equator are moving atapproximately 1,000 miles per hour, points on the polesare not moving at all, but are merely pivoting, the pointssomewhere between are moving at speeds between1,000 and zero miles per hour depending upon theirrelative position. Refer to view A in figure 3-4.
3-5
A B C
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Figure 3-4.—Coriolis force.
Assume that a missile located at the North Pole islaunched at a target on the equator. The missile does nothave any eastward lateral velocity, but the target has aneastward velocity of 1,000 miles per hour. The result isthat the missile appears to be deflected to the right asthe target moves away from its initial position. Refer toview B in figure 3-4.
A similar condition assumes that a missile locatedon the equator is launched at a target at the North Pole.The missile has an eastward lateral velocity of 1,000miles per hour, while the target on the pole has nolateral velocity at all. Once again the missile appears tobe deflected to the right as a result of its initial eastwardlateral velocity. Refer to view C in figure 3-4.
Due to Earth’s rotation and the Coriolis effect, thesimple circulation now becomes more complex asshown in figure 3-5. The complex on resulting from theinterplay of the Coriolis effect with the flow of air isknown as the theory. (See fig. 3-6.)
3-CELL THEORY
According to the 3-cell theory, Earth is divided intosix circulation belts—three in the NorthernHemisphere and three in the Southern Hemisphere. Thedividing lines are the equator, latitude, and 60°N and Slatitude. The general circulation of the NorthernHemisphere is similar to those of the SouthernHemisphere. (Refer to fig. 3-6 during the followingdiscussion.)
First, note the tropical cell of the NorthernHemisphere that lies between the equator and 30°Nlatitude. Convection at the equator causes the air to heat
and rise, due to convection. When it reaches the upperportions of the troposphere, it tends to flow toward theNorth Pole. By the time the air has reached 30°Nlatitude, the Coriolis effect has deflected it so much thatit is moving eastward instead of northward. This resultsin a piling up of air (convergence) near 30°N latitudeand a descending current of air (subsidence) toward thesurface which forms a belt of high pressure. When thedescending air reaches the surface where it flowsoutward (divergence), part of it flows poleward tobecome part of the mid-latitude cell; the other partflows toward the equator, where it is deflected by theCoriolis effect and forms the northeast trades.
The mid-latitude cell is located between 30° and60°N latitude. The air, which comprises this cell,circulates poleward at the surface and equatorwardaloft with rising currents at 60° (polar front) anddescending currents at 300 (high-pressure belt).However, in general, winds both at the surface and aloftblow from the west. The Coriolis effect easily explainsthis for the surface wind on the poleward-movingsurface air. The west wind aloft is not as easilyexplained. Most authorities agree that this wind isfrictionally driven by the west winds in the two adjacentcells.
The polar cell lies between 60°N latitude and theNorth Pole. The circulation in this cell begins with aflow of air at a high altitude toward the pole. This flowcools and descends at the North Pole and forms ahigh-pressure area in the Polar Regions. After reachingthe surface of Earth, this air usually flows equatorwardand is deflected by the Coriolis effect so that it movesfrom the northeast. This air converges with thepoleward flow from the mid-latitude cell and isdeflected upward with a portion circulating polewardagain and the remainder equatorward. The outflow ofair aloft between the polar and mid-latitude cells causesa semi-permanent low-pressure area at approximately60°N latitude. To complete the picture of the world’sgeneral atmospheric circulation, we must associate thisprevailing wind and pressure belts with some basiccharacteristics.
WORLD WINDS
In the vicinity of the equator is a belt of light andvariable winds known as the doldrums. On thepoleward side of the doldrums are the trade winds; thepredominant wind system of the tropics. These easterlywinds are the most consistent on Earth, especially overthe oceans. Near 30°N and 30°S latitudes lie thesub-tropical high-pressure belts. Winds are light and
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Figure 3-5.—Coriolis effect on windflow.
variable. These areas are referred to as the horselatitudes. The prevailing westerlies, which are on thepoleward side of the subtropical high-pressure belt, arepersistent throughout the mid-latitudes. In the NorthernHemisphere, the direction of the westerlies at thesurface is from the southwest. In the SouthernHemisphere, westerlies are from the northwest. This isdue to the deflection area resulting from the Corioliseffect as the air moves poleward.
Poleward of the prevailing westerlies, near 60°Nand 60°S latitudes, lies the belt of low-pressure basicpressure known as the polar front zone. Here,converging winds result in ascending air currents andconsequent poor weather.
WIND THEORY
Newton’s first two laws of motion indicate thatmotion tends to be in straight lines and only deviatesfrom such lines when acted upon by another force or bya combination of forces. Air tends to move in a straight
line from a high-pressure area to a low-pressure area.However, there are forces that prevent the air frommoving in a straight line.
Wind Forces
There are four basic forces that affect thedirectional movement of air in our atmosphere:pressure gradient force (PGF), the Coriolis effect,centrifugal force, and frictional force. These forces,working together, affect air movement. The forces thatare affecting it at that particular time determine thedirection that the air moves. Also, the different namesgiven to the movement of the air (geostrophic wind,gradient wind, etc.) depends on what forces areaffecting it.
Pressure Gradient
The rate of change in pressure in a directionperpendicular to the isobars is called pressure gradient.Pressure applied to a fluid is exerted equally in all
3-7
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Figure 3-6.—Idealized pattern of the general circulation. (The 3-cell theory).
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Figure 3-7.—Horizontal pressure gradient.
A. CROSS SECTION OF A VERTICAL PRESSURE GRADIENT ALONG LINE AA
HORIZONTAL PRESSURE GRADIENTB.
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Figure 3-8.—Cross section of a vertical pressure gradient along line AA.
directions throughout the fluid; e.g., if a pressure of1013.2 millibars is exerted downward by theatmosphere at the surface, this same pressure is alsoexerted horizontally outward at the surface. Therefore,a pressure gradient exists in the horizontal (along thesurface) as well as the vertical plane (with altitude) inthe atmosphere.
HORIZONTAL PRESSURE GRADIENT.—The horizontal pressure gradient is steep or strongwhen the isobars determining the pressure system (fig.3-7) are close together. It is flat or weak when theisobars are far apart.
VERTICAL PRESSURE GRADIENT.—Ifisobars are considered as depicting atmospherictopography, a high-pressure system represents a hill ofair, and a low-pressure system represents a depressionor valley of air. The vertical pressure gradient alwaysindicates a decrease in pressure with altitude, but therate of pressure decrease (gradient) varies directly withchanges in air density with altitude. Below 10,000 feetaltitude, pressure decreases approximately 1 inch ofmercury per 1,000 feet in the standard atmosphere. Thevertical cross section through a high and low (view A infig. 3-8) depicts the vertical pressure gradient. Asurface weather map view of the horizontal pressuregradient in the same high and low is illustrated in viewB of the figure 3-8.
Pressure Gradient Force
The variation of heating (and consequently thevariations of pressure) from one locality to another isthe initial factor that produces movement of air or wind.The most direct path from high to low pressure is thepath along which the pressure is changing most rapidly.The rate of change is called the pressure gradient.Pressure gradient force is the force that moves air froman area of high pressure to an area of low pressure. Thevelocity of the wind depends upon the pressuregradient. If the pressure gradient is strong, the windspeed is high. If the pressure gradient is weak, the windspeed is light. (See fig. 3-7.)
Figure 3-9 shows that the flow of air is from thearea of high pressure to the area of low pressure, but itdoes not flow straight across the isobars. Instead theflow is circular around the pressure systems. Pressuregradient force (PGF) causes the air to begin movingfrom the high-pressure to the low-pressure system.Coriolis (deflective) force and centrifugal force thenbegin acting on the flow in varying degrees. In thisexample, frictional force is not a factor.
Coriolis Effect
If pressure gradient force were the only forceaffecting windflow, the wind would blow at right anglesacross isobars (lines connecting points of equalbarometric pressure) from high to low pressure. Thewind actually blows parallel to isobars above anyfrictional level. Therefore, other factors must beaffecting the windflow; one of these factors is therotation of Earth. A particle at rest on Earth’s surface isin equilibrium. If the particle starts to move because ofa pressure gradient force, its relative motion is affectedby the rotation of Earth. If a mass of air from theequator moves northward, it is deflected to the right, sothat a south wind tends to become a southwesterlywind.
In the Northern Hemisphere, the result of theCoriolis effect is that moving air is deflected to the rightof its path of motion. This deflection to the right isdirectly proportional to the speed of the wind; the fasterthe wind speed, the greater the deflection to the right,and conversely, the slower the wind speed, the less thedeflection to the right. Finally, this apparent deflectiveforce is stronger at the Polar Regions than at theequator.
Centrifugal Force
According to Newton’s first law of motion, a bodyin motion continues in the same direction in a straight
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Figure 3-9.—Examples of circulation around high and lowpressure systems.
line and with the same speed unless acted upon by someexternal force. Therefore, for a body to move in acurved path, some force must be continually applied.The force restraining bodies that move in a curved pathis called the centripetal force; it is always directedtoward the center of rotation. When a rock is whirledaround on a string, the centripetal force is afforded bythe tension of the string.
Newton’s third law states that for every action thereis an equal and opposite reaction. Centrifugal force isthe reacting force that is equal to and opposite indirection to the centripetal force. Centrifugal force,then, is a force directed outward from the center ofrotation.
As you know, a bucket of water can be swung overyour head at a rate of speed that allows the water toremain in the bucket. This is an example of bothcentrifugal and centripetal force. The water is held inthe bucket by centrifugal force tending to pull itoutward. The centripetal force, the force holding thebucket and water to the center, is your arm swinging thebucket. As soon as you cease swinging the bucket, theforces cease and the water falls out of the bucket. Figure3-10 is a simplified illustration of centripetal andcentrifugal force.
High- and low-pressure systems can be comparedto rotating discs. Centrifugal effect tends to fling air outfrom the center of rotation of these systems. This forceis directly proportional to the wind speeds, the faster thewind, and the stronger the outward force. Therefore,when winds tend to blow in a circular path, centrifugaleffect (in addition to pressure gradient and Corioliseffects) influences these winds.
Frictional Force
The actual drag or slowing of air particles incontact with a solid surface is called friction. Frictiontends to retard air movement. Since Coriolis forcevaries with the speed of the wind, a reduction in thewind speed by friction means a reduction of theCoriolis force. This results in a momentary disruptionof the balance. When the new balance (includingfriction) is reached, the air flows at an angle across theisobars from high pressure to low pressure. (Pressuregradient force is the dominant force at the surface.) Thisangle varies from 10 degrees over the ocean to morethan 45 degrees over rugged terrain. Frictional effectson the air are greatest near the ground, but the effectsare also carried aloft by turbulence. Surface friction iseffective in slowing the wind to an average altitude of2,000 feet (about 600 meters) above the ground. Abovethis level, called the gradient wind level or the secondstandard level the effect of friction decreases rapidlyand may be considered negligible. Air above 2,000 feetnormally flows parallel to the isobars.
WIND TYPES
Since there is a direct relationship betweenpressure gradient and wind speed and direction, wehave a variety of wind types to deal with. We discussbelow the relationship of winds and circulations, theforces involved, and the effect of these factors on thegeneral circulation.
Geostrophic and Gradient Wind
On analyzed surface weather charts, points of equalpressure are connected by drawn lines referred to asisobars, while in upper air analysis, points of equalheights are connected and called isoheights.
The variation of these heights and pressures fromone locality to another is the initial factor that producesmovement of air, or wind. Assume that at three stationsthe pressure is lower at each successive point. Thismeans that there is a horizontal pressure gradient (adecrease in pressure in this case) for each unit distance.With this situation, the air moves from the area ofgreater pressure to the area of lesser pressure.
If the force of the pressure were the only factoracting on the wind, the wind would flow from high tolow pressure, perpendicular to the isobars. Sinceexperience shows the wind does not flow perpendicularto isobars, but at a slight angle across them and towardsthe lower pressure, it is evident that other factors are
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Figure 3-10.—Simplified illustration of centripetal andcentrifugal force.
involved. These other factors are the Coriolis effect,frictional force, and centrifugal effect. When a unit ofair moves with no frictional force involved, themovement of air is parallel to the isobars. This wind iscalled a gradient wind. When the isobars are straight, soonly Coriolis and pressure gradient forces are involved,it is termed a geostrophic wind.
Let’s consider a parcel of air from the time it beginsto move until it develops into a geostrophic wind. Assoon as a parcel of air starts to move due to the pressuregradient force, the Coriolis force begins to deflect itfrom the direction of the pressure gradient force. (Seeviews A and B of fig. 3-11). The Coriolis force is theapparent force exerted upon the parcel of air due to therotation of Earth. This force acts to the right of the pathof motion of the air parcel in the Northern Hemisphere(to the left in the Southern Hemisphere). It always actsat right angles to the direction of motion. In the absenceof friction, the Coriolis force changes the direction ofmotion of the parcel until the Coriolis force and thepressure gradient force are in balance. When the twoforces are equal and opposite, the wind blows parallelto the straight isobars (view C in fig. 3-11). The Coriolisforce only affects the direction, not the speed of themotion of the air. Normally, Coriolis force is not greaterthan the pressure gradient force. In the case ofsuper-gradient winds, Coriolis force may be greaterthan the pressure gradient force. This causes the windto deflect more to the right in the Northern Hemisphere,or toward higher pressure.
Under actual conditions, air moves around high andlow pressure centers toward lower pressure. Turn back
to figure 3-9. Here, the flow of air is from the area ofhigh pressure to the area of low pressure, but, as wementioned previously, it does not flow straight acrossthe isobars (or isoheights). Instead, the flow is circulararound the pressure systems.
The Coriolis force commences deflecting the pathof movement to the right (Northern Hemisphere) or left(Southern Hemisphere) until it reaches a point where abalance exists between the Coriolis and the pressuregradient force. At this point the air is no longerdeflected and moves forward around the systems.
Once circular motion around the systems isestablished, then centrifugal force must be considered.Centrifugal force acts outward from the center of boththe highs and the lows with a force dependent upon thevelocity of the wind and the degree of curvature of theisobars. However, the pressure gradient force is actingtowards the low; therefore, the flow in that directionpersists. When the flow is parallel to the curved portionof the analysis in figure 3-9, it is a gradient wind. Whenit is moving parallel to that portion of the analysisshowing straight flow, it is a geostrophic wind.
We defined pressure gradient as being a change ofpressure with distance. This means that if the isobarsare closely spaced, then the pressure change is greaterover a given distance; it is smaller if they are widelyspaced. Therefore, the closer the isobars, the faster theflow. Geostrophic and gradient winds are alsodependent, to a certain extent, upon the density of theatmosphere and the latitude. If the density and thepressure gradient remain constant and the latitudeincreases, the wind speed decreases. If the latitude
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Figure 3-11.—Development cycle of a geostrophic wind.
decreases, the wind speed increases. If the density andthe latitude remain constant and the pressure gradientdecreases, the wind speed decreases. If the pressuregradient and the latitude remain constant and thedensity decreases, the wind speed increases. If thedensity increases, the wind speed decreases. Truegeostrophic wind is seldom observed in nature, but theconditions are closely approximated on upper-levelcharts.
Cyclostrophic Wind
In some atmospheric conditions, the radius ofrotation becomes so small that the centrifugal forcebecomes quite strong in comparison with the Coriolisforce. This is particularly true in low latitudes where theCoriolis force is quite small to begin with. In this case,the pressure gradient force is nearly balanced by thecentrifugal force alone. When this occurs, the wind issaid to be cyclostrophic. By definition, a cyclostrophicwind exists when the pressure gradient force isbalanced by the centrifugal force alone.
This exact situation rarely exists, but is so nearlyreached in some situations that the small Coriolis effectis neglected and the flow is said to be cyclostrophic.Winds in a hurricane or typhoon and the winds around atornado are considered cyclostrophic.
Movement of Wind around Anticyclones
The movement of gradient winds aroundanticyclones is affected in a certain manner by thepressure gradient force, the centrifugal force, and theCoriolis force. The pressure gradient force acts fromhigh to low pressure, and the Coriolis force actsopposite to the pressure gradient force and at rightangles to the direction of movement of the parcel of air.The centrifugal force acts at right angles to the path ofmotion and outward from the center about which theparcel is moving. (See fig. 3-12.) In the case of ahigh-pressure center, the pressure gradient force andthe centrifugal force balance the Coriolis force. Thisphenomenon may be expressed in the followingmanner:
PG + CF = D
Movement of Wind around Cyclones
As in the case of anticyclones, the pressure gradientforce, the centrifugal force, and the Coriolis force affectgradient winds around cyclones, but the balance of theforces is different. (See fig. 3-12.) In a cyclonicsituation the Coriolis force and the centrifugal forcebalance the pressure gradient force. This balance maybe expressed in the following manner:
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ANTICYCLONE CYCLONE
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Centrifugal force acts with the pressure gradientforce when the circulation is anticyclonic and againstthe pressure gradient force when the circulation iscyclonic. Therefore, wind velocity is greater in ananticyclone than in a cyclone of the same isobaricspacing.
Variations
It has been determined that, given the same density,pressure gradient, and latitude, the wind is weakeraround a low-pressure cell than a high-pressure cell.This is also true for gradient and geostrophic winds.The wind we observe on a synoptic chart is usuallystronger around low cells than high cells because thepressure gradient is usually stronger around thelow-pressure cell.
Geostrophic and Gradient Wind Scales
The geostrophic wind is stronger than the gradientwind around a low and is weaker than a gradient wind
around a high. This is why the isobar spacing andcontour spacing, for a curved flow, differs from thatdetermined by a geostrophic wind scale. If the flowunder consideration is around a high-pressure cell, theisobars are farther apart than indicated by thegeostrophic wind scale. If the flow is around alow-pressure cell, the isobars are closer together thanindicated by the geostrophic wind scale.
Geostrophic and gradient wind scales are used todetermine the magnitude of these winds (based onisobar or contour spacing) and to determine the isobaror contour spacing (based on observed wind speeds).There are a number of scales available for measuringgeostrophic and gradient flow of both surface and upperair charts.
Weather plotting charts used by the NavalOceanography Command has geostrophic wind scalesprinted on them for both isobaric and contour spacing.The most common scales in general use can be used forboth surface and upper air charts. The scales are in 4mband 60m intervals. An example of a geostrophic windscale is shown in figure 3-13. Note that latitude
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accounts for the increases in gradients. In tropicalregions, the geostrophic wind scales become lessreliable because pressure gradients are generally ratherweak.
REVIEW QUESTIONS
Q3-1. Which two factors influence the Earth'stemperature?
Q3-2. What are the major factors that result in thewide distribution of pressure over the earth'ssurface?
Q3-3. What effect does Coriolis force have onthermal circulation in the NorthernHemisphere?
Q3-4. According to the 3-cell theory, how manycirculation belts are there?
Q3-5. According to the 3-cell theory, what type ofpressure system would you normally find at30 degrees north latitude?
Q3-6. What is the predominant wind system in thetropics?
Q3-7. Name two types of pressure gradient.
Q3-8. What is the difference between centrifugalforce and centripetal force?
Q3-9. What is the difference between gradient windand geostrophic wind?
Q3-10. What is the relationship between centrifugalforce and pressure gradient force aroundanticyclones?
SECONDARY CIRCULATION
LEARNING OBJECTIVE: Determine howcenters of action, migratory systems, andseasonal variations affect secondary aircirculations.
Now that you have a picture of the generalcirculation of the atmosphere over Earth, the next stepis to see how land and water areas offset the generalcirculation. The circulations caused by the effect ofEarth’s surfaces, its composition and contour, areknown as secondary circulations. These secondarycirculations give rise to winds that often cancel out thenormal effect of the great wind systems.
There are two factors that cause the pressure beltsof the primary circulation to break up into closedcirculations of the secondary circulations. They are the
non-uniform surface of the earth and the differencebetween heating and cooling of land and water. Thesurface temperature of oceans changes very littleduring the year. However, land areas sometimesundergo extreme temperature changes with the seasons.In the winter, large high-pressure areas form over thecold land and the low-pressure areas form over therelatively warm oceans. The reverse is true in summerwhen highs are over water and lows form over thewarm land areas. The result of this difference in heatingand cooling of land and water surfaces is known as thethermal effect.
Circulation systems are also created by theinteraction of wind belts of pressure systems or thevariation in wind in combination with certaindistributions of temperature and/or moisture. This isknown as the dynamic effect. This effect rarely, if ever,operates alone in creating secondary systems, as mostof the systems are both created and maintained by acombination of the thermal and dynamic effects.
CENTERS OF ACTION
The pressure belts of the general circulation arerarely continuous. They are broken up into detachedareas of high and low pressure cells by the secondarycirculation. The breaks correspond with regionsshowing differences in temperature from land to watersurfaces. Turn back to figures 3-2A and 3-2B. Comparethe temperature distribution in views A and B of figures3-2 to the pressure distribution in views A and B offigure 3-3. Note the gradient over the Asian Continentin January. Compare it to the warmer temperature overthe ocean and coastal regions. Now look at view A offigure 3-3 and note the strong region of high-pressurecorresponding to the area. Now look at the same area inJuly. Note the way the temperature gradient flattens outand warms. Look at view B of figure 3-3 and see thelow-pressure area that has replaced the high-pressureregion of winter. These pressure cells tend to persist in aparticular area and are called centers of action; that is,they are found at nearly the same location withsomewhat similar intensity during the same month eachyear.
There is a permanent belt of relatively low pressurealong the equator and another deeper belt oflow-pressure paralleling the coast of the AntarcticContinent. Permanent belts of high pressure largelyencircle Earth, generally over the oceans in both theNorthern and Southern Hemispheres. The number ofcenters of action is at a maximum at about 30 to 35degrees from the equator.
3-14
There are also regions where the pressure ispredominantly low or high at certain seasons, but notthroughout the year. In the vicinity of Iceland, pressureis low most of the time. The water surface is warmer(due to warm ocean currents) than the surface ofIceland or the icecaps of Greenland. The Icelandic lowis most intense in winter, when the greatest temperaturecontrast occurs, but it persists with less intensitythrough the summer. Near Alaska, a similar situationexists with the Aleutian low. The Aleutian low is mostpronounced when the neighboring areas of Alaska andSiberia are snow covered and colder than the adjacentocean.
These lows are not a continuation of one and thesame cyclone. They are, however, regions of lowpressure where lows frequently form or arrive fromother regions. Here they remain stationary or movesluggishly for a time, then the lows move on or die outand are replaced by others. Occasionally these regionsof low pressure are invaded by traveling high-pressuresystems.
Two areas of semi permanent high-pressure alsoexist. There is a semi permanent high-pressure centerover the Pacific westward of California and anotherover the Atlantic, near the Azores and of the coast ofAfrica. Pressure is also high, but less persistently so,west of the Azores to the vicinity of Bermuda. Thesesubtropical highs are more intense and cover a greaterarea in summer than winter. They also extend farthernorthward summer. In winter, these systems move soultoward the equator, following the solar equator.
The largest individual circulation cells in theNorthern Hemisphere are the Asiatic high in winter andthe Asiatic low in summer. In winter, the Asiaticcontinent is a region of strong cooling and therefore isdominated by a large high-pressure cell. In summer,strong heating is present and the high-pressure cellbecomes a large low-pressure cell. (See fig. 3-3A andfig. 3-3B.) This seasonal change in pressure cells givesrise to the monsoon flow over India and Southeast Asia.
Another cell that is often considered to be a centerof action is the polar high. Both Arctic and Antarctichighs have considerable variations in pressure, andthese regions have many traveling disturbances insummer. For example, the Greenland high (due to theGreenland icecap) is a persistent feature, but it is not awell-defined high during all seasons of the year. TheGreenland high often appears to be an extension of thepolar high or vice versa. Other continental regionsshow seasonal variations, but are generally of small size
and their location is variable. Therefore, they are notconsidered to be centers of action.
An average annual pressure distribution chart(figure 3-14) reveals several important characteristics.First, along the equator there is a belt of relatively lowpressure encircling the globe with barometric pressureof about 1,012 millibars. Second, on either side of thisbelt of low pressure is a belt of high pressure. Thishigh-pressure area in the Northern Hemisphere liesmostly between latitudes 30° and 40°N with threewell-defined centers of maximum pressure. One is overthe eastern Pacific, the second over the Azores and thethird over Siberia; all are about 1,020 millibars. Thebelt of high pressure in the Southern Hemisphere isroughly parallel to 30°S. It, too, has three centers ofmaximum pressure. One is in the eastern Pacific, thesecond in the eastern Atlantic, and the third in theIndian Ocean; again, all are about 1,020 millibars. Athird characteristic to be noted from this chart is that,beyond the belt of high pressure in either hemisphere,the pressure diminishes toward the poles. In theSouthern Hemisphere, the decrease in pressure towardthe South Pole is regular and very marked. The pressuredecreases from, an average slightly above 1,016millibars along latitude 35°S to an average of 992millibars along latitude 60°S In the NorthernHemisphere, however, the decrease in pressure towardthe North Pole is less regular and not as great. This islargely due to the distribution of land and water: notethe extensive landmass in the Northern Hemisphere ascompared to those of the Southern Hemisphere.
While the pressure belts that stand out on theaverage annual pressure distribution chart representaverage pressure distribution for the year, these beltsare rarely continuous on any given day. They areusually broken up into detached areas of high or lowpressure by the secondary circulation of theatmosphere. In either hemisphere, the pressure over theland during the winter season is decidedly above theannual average. During the summer season, thepressure is decidedly below the average, with extremevariations occurring such as in the case of continentalAsia. Here the mean monthly pressure ranges fromabout 1,033 millibars during January to about 999millibars during July. Over the northern oceans, on theother hand, conditions are reversed; the summerpressure there is somewhat higher. Thus in January theIcelandic and Aleutian lows intensify to a depth ofabout 999 millibars, while in July these lows fill and arealmost obliterated.
3-15
The polar high in winter is not a cell centereddirectly over the North Pole, but appears to be anextension of the Asiatic high and often appears as awedge extending from the Asiatic continent. The cell isdisplaced toward the area of coldest temperatures—theAsiatic continent. In summer, this high appears as anextension of the Pacific high and is again displacedtoward the area of coolest temperature, which in thiscase is the extensive water area of the Pacific.
In winter over North America, the most significantfeature is the domination by a high-pressure cell. Thiscell is also due to cooling but is not as intense as theAsiatic cell. In summer, the most significant feature isthe so-called heat low over the southwestern part of thecontinent, which is caused by extreme heating in thisregion.
MIGRATORY SYSTEMS
General circulation, based on an average of windconditions, is a more or less quasi-stationarycirculation. Likewise, much of the secondarycirculation depends on more or less static conditionsthat, in turn, depend on permanent and semi permanenthigh and low-pressure areas. Changes in the circulationpatterns discussed so far have been largely seasonal.However, secondary circulation also includes wind
systems that migrate constantly, producing rapidlychanging weather conditions throughout all seasons,especially in the middle latitudes. The migratorycirculation systems are associated with air masses,fronts, cyclones, and anticyclones. These are covered indetail in the next unit.
Anticyclones
An anticyclone (high) is an area of relatively highpressure with a clockwise flow (wind circulation) in theNorthern Hemisphere and counterclockwise flow in theSouthern Hemisphere. The windflow in an anticycloneis slightly across the isobars and away from the centerof the anticyclone. (See fig. 3-15.) Anticyclones arecommonly called highs or high-pressure areas.
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Figure 3-15.—Anticyclone.
The formation of an anticyclone or theintensification of an existing one is calledanticyclogenesis. Anticyclogenesis refers to thedevelopment of anticyclonic circulation as well as theintensification of an existing anticyclonic flow. When ahigh-pressure center is increasing in pressure, the highis BUILDING or INTENSIFYING. Although a highcan build (or intensify) without an increase inanticyclonic flow, it is rare. Normally, building andanticyclogenesis occur simultaneously. The weakeningof anticyclonic circulation is anticyclolysis. When thepressure of a high is decreasing, we say the high isweakening. Anticyclolysis and weakening can occurseparately, but usually occur together.
The vertical extent of pressure greatly depends onthe air temperature. Since density increases with adecrease in temperature, pressure decreases morerapidly vertically in colder air than in warmer air. In acold anticyclone (such as the Siberian high), thevertical extent is shallow; while in a warm anticyclone(such as the subtropical high), the vertical extentreaches high into the upper atmosphere due to the slowdecrease in temperature with elevation.
Cyclones
A cyclone (low) is a circular or nearly circular areaof low pressure with a counterclockwise flow. The flowis slightly across the isobars toward the center in theNorthern Hemisphere and clockwise in the SouthernHemisphere. (See fig. 3-16.) It is commonly called alow or a depression. This use of the word cyclone
should be distinguished from the colloquial use of theword as applied to the tornado or tropical cyclone(hurricane).
The formation of a new cyclone or theintensification of the cyclonic flow in an existing one iscalled cyclogenesis. When the pressure in the low isfalling, we say the low is deepening. Cyclogenesis anddeepening can also occur separately, but usually occurat the same time. The decrease or eventual dissipationof a cyclonic flow is called cyclolysis. When thepressure in a low is rising, we say the low is filling.Cyclolysis and filling usually occur simultaneously.Cyclones in middle and high latitudes are referred to asextratropical cyclones. The term tropical cyclone refersto hurricanes and typhoons.
VERTICAL STRUCTURE OF SECONDARYCIRCULATIONS (PRESSURE CENTERS)
To better understand the nature of the pressurecenters of the secondary circulation, it is necessary toconsider them from a three-dimensional standpoint.With the aid of surface and upper air charts, you will beable to see the three dimensions of these pressuresystems as well as the circulation patterns of thesecondary circulation as established at higher levels inthe troposphere and lower stratosphere.
In Chapter 2, the study of gas laws showed thatvolume is directly proportional to temperature. Statedanother way, we might say that the thickness of a layerbetween two isobaric surfaces is directly proportionalto the mean virtual temperature of the layer. Becausethe atmosphere is always moist to some degree, virtualtemperature is used. Mean virtual temperature isdefined as the average temperature at which dry airwould have the same pressure and density as moist air.Thus, lines representing thickness are also isotherms ofmean virtual temperature. The higher the mean virtualtemperature, the thicker the layer, or vice versa. Thethickness between layers is expressed in geopotentialmeters. The shift in location, as well as the change ofshape and intensity upward of atmospheric pressuresystems, is dependent on the temperature distribution.
An example of the effects of virtual temperaturecan be demonstrated by placing two columns of air nextto each other. One air column is cold and the other aircolumn is warm. The constant pressure surfaces in the
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Figure 3-16.—Cyclone.
cold column are closer than the ones in the warmcolumn. Figure 3-17 shows an increase in thicknessbetween two pressure surfaces, resulting in an increasein mean virtual temperature. Note the increase in thedistance between the constant pressure surfaces; P, P1,etc., from column A to column B. Using thehypsometric equation can derive the thickness valuebetween two pressure surfaces. Thickness may also bedetermined from tables, graphs, etc.
VERTICAL STRUCTURE OF HIGH PRESSURESYSTEMS
The topographic features that indicate thecirculation patterns at 500 millibars in the atmospherecorrespond in general to those at lower and higher level.However, they may experience a shift in location aswell as a change in intensity and shape. For example, aridge aloft may reflect a closed high on a surfacesynoptic chart. In addition, upper air circulationpatterns may take on a wavelike structure in contrast tothe alternate closed lows, or closed high patterns at thesurface level. The smoothing of the circulation patternaloft is typical of atmospheric flow patterns.
Cold Core Highs
A cold core high is one in which the temperatureson a horizontal plane decrease toward the center.
Because the temperature in the center of a cold corehigh is less than toward the outside of the system, itfollows that the vertical spacing of isobars in the centerof this system is closer together than on the outside.Although the pressure at the center of these systems onthe surface may be high, the pressure decreases rapidlywith height. (See fig. 3-18.) Because these highs areoften quite shallow, it is common for an upper level lowto exist above a cold core high.
NOTE: For the purpose of illustration, figures 3-18through 3-21 are exaggerated with respect to actualatmospheric conditions.
If the cold core high becomes subjected to warmingfrom below and to subsidence from aloft, as it movessouthward from its source and spreads out, itdiminishes rapidly in intensity with time (unless somedynamic effect sets in aloft over the high to compensatefor the warming). Since these highs decrease inintensity with height, thickness is relatively low. In thevertical, cold core highs slope toward colder air aloft.Anticyclones found in Arctic air are always cold cored,while anticyclones in polar air may be warm or coldcore.
Examples of cold core highs are the NorthAmerican High, the Siberian High and the migratoryhighs that originate from these anticyclones.
Warm Core Highs
A warm core high is one in which the temperatureson a horizontal level increase toward the center.Because the temperatures in the center of a warm corehigh are higher than on the outside of the system, itfollows that the vertical spacing of isobars in the centeris farther apart than toward the outside of the high. Forthis reason, a warm core high increases in intensity withaltitude and has an anticyclonic circulation at all levels(see fig. 3-19). From a vertical view, warm core highsslope toward warmer air aloft. A warm core high isaccompanied by a high cold tropopause. Since thepressure surfaces are spaced far apart, the tropopause isreached only at great heights. The temperaturecontinues to decrease with elevation and is cold by the
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Figure 3-18.—Cold core high.
time the tropopause is reached. The subtropical highsare good examples of this type of high. Therefore,anticyclones found in tropical air are always warm core.Examples of warm core highs are the Azores orBermuda High and the Pacific High.
VERTICAL STRUCTURE OF LOW-PRESSURESYSTEMS
Low-pressure systems, like high-pressure systems,are generally a reflection of systems aloft. They, too,experience shifts in location and changes in intensityand shape with height. At times, a surface system maynot be evident aloft and a well-developed system aloftmay not reflect on a surface analysis.
Cold Core Lows
The cold core low contains the coldest air at itscenter throughout the troposphere; that is, goingoutward in any direction at any level in the troposphere,warmer air is encountered. The cold core low (figure3-20) increases intensity with height. Relativeminimums in thickness values, called cold pools, arefound in such cyclones. The temperature distribution isalmost symmetrical, and the axis of the low is nearlyvertical. When they do slope vertically, they slopetoward the coldest temperatures aloft. In the cold low,the lowest temperatures coincide with the lowestpressures.
The cold low has a more intense circulation aloftfrom 850 to 400 millibars than at the surface. Somecold lows show only slight evidence in the surfacepressure field that an intense circulation exists aloft.The cyclonic circulation aloft is usually reflected on thesurface in an abnormally low daily mean temperatureoften accompanied by instability and showeryprecipitation. A cold core low is accompanied by a lowwarm tropopause. Since the pressure surfaces are closetogether, the tropopause is reached at low altitudeswhere the temperature is relatively warm. Goodexamples of cold core lows are the Aleutian andIcelandic lows. Occluded cyclones are generally coldcore in their later stages, because polar or arctic air hasclosed in on them.
At high latitudes the cold pools and their associatedupper air lows show some tendency for location in thenorthern Pacific and Atlantic Oceans where,statistically, they contribute to the formation of theAleutian and Icelandic lows.
Warm Core Lows
A warm core low (figure 3-21) decreases intensitywith height and the temperature increases toward thecenter on a horizontal plane. The warm low isfrequently stationary, such as the heat low over thesouthwestern United States in the summer; this is aresult of strong heating in a region usually insulatedfrom intrusions of cold air that tend to fill it or cause itto move. The warm low is also found in its moving formas a stable wave moving along a frontal surface. Thereis no warm low aloft in the troposphere. The tropicalcyclone, however, is believed to be a warm low becauseits intensity diminishes with height. Because mostwarm lows are shallow, they have little slope. However,intense warm lows like the heat low over the southwestUnited States and hurricanes do slope toward warm airaloft.
In general, the temperature field is quiteasymmetrical around a warm core cyclone. Usually thesouthward moving air in the rear of the depression is
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Figure 3-19.—Warm core high.
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Figure 3-20.—Cold core low.
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Figure 3-21.—Warm core low.
not as warm as that moving northward in advance of it.A warm core low decreases intensity with height orcompletely disappears and are often replaced byanticyclones aloft. The heat lows of the southwesternUnited States, Asia, and Africa are good examples ofwarm core lows. Newly formed waves are generallywarm core because of the wide-open warm sector.
DYNAMIC LOW
Systems that retain their closed circulations toappreciable altitudes and are migratory are calleddynamic lows or highs. A dynamic low is acombination of a warm surface low and a cold upperlow or trough, or a warm surface low in combinationwith a dynamic mechanism aloft for producing a coldupper low or trough. It has an axis that slopes towardthe coldest tropospheric air. (See figure 3-22.) In thefinal stage, after occlusion of the surface warm low iscomplete, the dynamic low becomes a cold low with theaxis of the low becoming practically vertical.
DYNAMIC HIGH
The dynamic high is a combination of a surfacecold high and an upper-level warm high orwell-developed ridge, or a combination of a surfacecold high with a dynamic mechanism aloft forproducing high-level anticyclogenesis. Dynamic highshave axes that slope toward the warmest troposphericair. (See fig. 3-22.) In the final stages of warming thecold surface high, the dynamic high becomes a warmhigh with its axis practically vertical.
REVIEW QUESTIONS
Q3-11. What is the term that defines the formation ofan anticyclone or the intensification of anexisting anticyclone?
Q3-12. What is the direction of the windflow around acyclone?
Q3-13. How do temperatures change within a coldcore low?
Q3-14. Low pressure due to intense heating over thesouthwestern United States is an example ofwhich type of low-pressure system?
TERTIARY CIRCULATION
LEARNING OBJECTIVE: Define tertiarycirculation and describe how tertiarycirculations affect local weather and winddirection and speed.
Tertiary (third order) circulations are localizedcirculations directly attributable to one of the followingcauses or a combination of them: local cooling, localheating, adjacent heating or cooling, and induction(dynamics).
Many regions have local weather phenomenacaused by temperature differences between land andwater surfaces or by local topographical features. Theseweather phenomena show up as circulations. Thesetertiary circulations can result in dramatic local weatherconditions and wind flows. The most common tertiarycirculations are discussed in this lesson. However, thereare numerous other circulations and related phenomenain existence around the world.
MONSOON WINDS
The term monsoon is of Arabic origin and meansseason. The monsoon wind is a seasonal wind thatblows from continental interiors (or large land areas) tothe ocean in the winter; they blow in the oppositedirection during the summer. The monsoon wind ismost pronounced over India, although there are otherregions with noticeable monsoon winds.
Monsoon winds are a result of unequal heating andcooling of land and water surfaces. During winter amassive area of cold high pressure develops over theextensive Asiatic continent. This high pressure is dueprimarily to cold arctic air and long-term radiationcooling. To the south, the warm equatorial waters existand, in contrast, the area has relatively lower surfacepressures. The combination of high pressure over Asia
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Figure 3-22.—Vertical slope of low -pressure systems.
and low pressure over the Equatorial Belt sets up apressure gradient directed from north to south. Becauseof the flow around the massive Siberian high, northeastwinds begin to dominate the regions from India to thePhilippines. (See fig. 3-23).
During the winter months, clear skies predominateover most of the region. This is caused by the massmotion of air from a high-pressure area over land to anarea of lower pressure over the ocean. As the air leavesthe high-pressure area over land, it is cold and dry. As ittravels over land toward the ocean, there is no source ofmoisture to induce precipitation. The air is also
traveling from a higher altitude to a lower altitude;consequently, this downslope motion causes the air tobe warmed at the adiabatic lapse rate. This warmingprocess has a still further clearing effect on the skies.
During the summer the airflow over the region iscompletely reversed. The large interior of Asia isheated to the point where the continent is much warmerthan the ocean areas to the south. This inducesrelatively low pressure over Asia and higher pressureover the equatorial region. This situation produces asouthwesterly flow as shown in figure 3-24.
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The weather associated with the summer monsoonwinds is thunderstorms, almost constant heavy rain,rain showers, and gusty surface winds. This condition iscaused by mass motion of air from the relativelyhigh-pressure area over the ocean to a low-pressurearea over land. When the air leaves the ocean, it is warmand moist. As the air travels over land toward thelow-pressure area, it is also traveling from a loweraltitude to a higher altitude. The air is lifted by amechanical force and cooled to its condensation pointby this upslope motion (pseudo adiabatic process).
LAND AND SEA BREEZES
There is a diurnal (daily) contrast in the heating oflocal water and land areas similar to the seasonalvariation of the monsoon. During the day, the land iswarmer than the water area; at night the land area iscooler than the water area. A slight variation in pressureis caused by this temperature contrast. At night thewind blows from land to sea and is called a land breeze.During the day, the wind blows from water areas to landareas and is called a sea breeze.
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Figure 3-25.—Circulation of land and sea breezes.
The sea breeze usually begins during midmorning(0900-1100 local time) when the land areas becomewarmer than adjacent ocean waters (see fig. 3-25). Thistemperature difference creates an area of slightly lowersurface pressures over land compared to the now coolerwaters. The result is a wind flow from water to land.The sea breeze starts with a shallow flow along thesurface; however, as maximum heating occurs, the flowincreases with height. The height varies from anaverage of 3,000 feet in moderately warm climates to4,500 (or more) in tropical regions. The effects of thesea breeze can be felt as far as 30 miles both onshoreand offshore. In extreme cases, the sea breeze is felt 100miles inland depending upon terrain. By mid afternoon(maximum heating) the sea breeze will reach itsmaximum speed and may be strong enough to beinfluenced by the Coriolis force, which causes it to flowat an angle to the shore. The sea breeze is mostpronounced in late spring, summer, and early fall whenmaximum temperature differences occur between landand water surfaces. A decrease in temperature and anincrease in humidity and wind speed mark the start of asea breeze.
The sea breeze continues until the land area coolssufficiently to dissipate the weak low pressure. Aftersunset, the land cools at a faster rate than the adjacentwaters and eventually produces a complete reversal ofthe winds. As the land continues to cool through theevening hours, a weak area of high pressure forms overthe land. The water area, with its warmer temperatures,has slightly lower pressure and again a flow isestablished; however, the flow is now from land towater (offshore). (See fig. 3-25.)
The land breezes, when compared to the seabreezes, are less extensive and not as strong (usually
less than 10 knots and less than 10 miles offshore). Thisis because there is less temperature contrast at nightbetween land and water surfaces as compared to thetemperature contrast during daytime heating. Landbreezes are at maximum development late at night, inlate fall and early winter. In the tropical land regions,the land and sea breezes are repeated day after day withgreat regularity. In high latitudes the land and seabreezes are often masked by winds of synoptic features.
WINDS DUE TO LOCAL COOLING ANDHEATING
In the next sections we discuss tertiary circulationsdue to local cooling and heating effects. Under normalcircumstances, these winds attain only light tomoderate wind speeds; however, winds often occur inand near mountain areas that have undergone dramaticchanges in normal character. At times, mountain areastend to funnel winds through valleys and mountainpasses. This funneling effect produces extremelydangerous wind speeds.
FUNNEL EFFECT
Winds blowing against mountain barriers tend toflatten out and go around or over them. If a pass or avalley breaks the barrier, the air is forced through thebreak at considerable speed. When wind is forcedthrough narrow valleys it is known as the funnel effectand is explained by Bernoulli’s theorem. According toBernoulli’s theorem, pressures are least wherevelocities are greatest; likewise, pressures are greatestwhere velocities are least. This observation is true forboth liquids and gases. (See fig. 3-26.)
3-23
HIGHPRESSURE
LOWPRESSURE
MO
UN
TA
IN
AG5f0326
STRONGWINDS
PA
SS
RA
NG
E
CO
LD
AIR
LIGHT
WINDS
Figure 3-26.—Strong wind produced by funneling.
Bernoulli’s theorem is frequently used to forecasttertiary winds in the mountainous western UnitedStates. The famous Santa Ana winds of southernCalifornia are a prime example. Winds associated withhigh pressure situated over Utah are funneled throughthe valley leading into the town of Santa Ana near theCalifornia coast. Low pressure develops at the mouth ofthe valley and the end result is hot, dry, gusty andextremely dangerous winds. When the Santa Ana isstrong enough, the effects are felt in virtually everyvalley located along the coast of southern California.Visibility is often restricted due to blowing sand. It iscommon to see campers, trailers, and trucks turned overby the force of these winds. When funneled windsreach this magnitude, they are called jet-effect winds,canyon winds, or mountain-gap winds.
Winds Due to Local Cooling and Heating
There are two types of tertiary circulationsproduced by local cooling—glacier winds and drainagewinds.
GLACIER WINDS.—Glacier winds, or fallwinds (as they are sometimes called) occur in manyvarieties in all parts of the world where there areglaciers or elevated land masses that become coveredby snow and ice during winter. During winter, the areaof snow cover becomes most extensive. Weak pressureresults in a maximum of radiation cooling.Consequently the air coming in contact with the coldsnow cools. The cooling effect makes the overlying airmore dense, therefore, heavier than the surrounding air.When set in motion, the cold dense air flows down thesides of the glacier or plateau. If it is funneled through apass or valley, it may become very strong. This type ofwind may form during the day or night due to radiationcooling. The glacier wind is most common during thewinter when more snow and ice are present.
When a changing pressure gradient moves a largecold air mass over the edge of a plateau, this action setsin motion the strongest, most persistent, and mostextensive of the glacier or fall winds. When thishappens, the fall velocity is added to the pressuregradient force causing the cold air to rush down to sealevel along a front that may extend for hundreds ofmiles. This condition occurs in winter on a large scalealong the edge of the Greenland icecap. In some placesalong the icecap, the wind attains a velocity in excess of90 knots for days at a time and reaches more than 150nautical miles out to sea.
Glacier winds are cold katabatic (downhill) winds.Since all katabatic winds are heated adiabatically intheir descent, they are predominantly dry. Occasionally,the glacier winds pick up moisture from fallingprecipitation when they underride warm air. Even withthe adiabatic heating they undergo, all glacier or fallwinds are essentially cold winds because of the extremecoldness of the air in their source region. Contrary to allother descending winds that are warm and dry, theglacier wind is cold and dry. It is colder, level for level,than the air mass it is displacing. In the NorthernHemisphere, the glacier winds descend frequently fromthe snow-covered plateaus and glaciers of Alaska,Canada, Greenland, and Norway.
DRAINAGE WINDS.—Drainage winds (alsocalled mountain or gravity winds) are caused by thecooling air along the slopes of a mountain.Consequently, the air becomes heavy and flowsdownhill, producing the MOUNTAIN BREEZE.
Drainage winds are katabatic winds and like glacierwinds; a weak or nonexistent pressure gradient isrequired to start the downward flow. As the air near thetop of a mountain cools through radiation or contactwith colder surfaces, it becomes heavier than thesurrounding air and gradually flows downward (fig.3-27). Initially this flow is light (2 to 4 knots) and only afew feet thick. As cooling continues, the flow increasesachieving speeds up to 15 knots at the base of themountain and a depth of 200 feet or more. Winds inexcess of 15 knots are rare and only occur when themountain breeze is severely funneled.
Drainage winds are cold and dry. Adiabatic heatingdoes not sufficiently heat the descending air because ofthe relative coldness of the initial air and because thedistance traveled by the air is normally short. Drainagewinds have a very localized circulation. As the cold airenters the valley below, it displaces the warm air.Temperatures continue to fall. If the flow achievesspeeds of 8 knots or more, mixing results between thewarm valley air and the cold descending air that resultsin a slight temperature increase. Campers often preferto make summer camps at the base of mountains to takeadvantage of the cooling effect of the mountain breeze.
Funnel Effects
VALLEY BREEZES.—The valley breeze is theanabatic (uphill) counterpart of the mountain breeze.When the sun heats the valley walls and mountainslopes during the morning hours, the air next to theground is heated until it rises along the slopes. Rocky or
3-24
sandy slopes devoid of vegetation are the most effectiveheating surfaces. If the slopes are steep, the ascendingbreeze tends to move up the valley walls. Theexpansion of the heated air next to the surface producesa slight local pressure gradient against the groundsurface. As the heating becomes stronger, convectivecurrents begin to rise vertically from the valleys (figure3-28). The updrafts along the valley walls continue tobe active, particularly at the head of the valley. Thevalley breeze usually reaches its maximum strength inthe early afternoon. It is a stronger and deeper wind
than the mountain breeze. It is difficult to isolate thevalley breeze effect because of the prevailing gradientwinds. Consequently, the valley breeze is much morelikely to be superposed as a prevailing wind than is themountain breeze, which by its very nature can developonly in the absence of any appreciable gradient wind.The valley breezes are generally restricted to slopesfacing south or the more direct rays of the sun, and theyare more pronounced in southern latitudes. They arediurnally strongest in the late afternoon and areseasonally strongest in summer.
3-25
UP-DRAFTWARM
RADIATION
AG5f0327
COLD
DRAIN
AGE
OF
COOLAIR
Figure 3-27.—Mountain breeze or katabatic wind. During the night outgoing radiation cools air along hillsides below free airtemperature. The cooled air drains to lowest point of the terrain.
COOLDOWN-DRAFT
AG5f0328
COLD
WARM WARM
Figure 3-28.—Valley breeze or anabatic wind. During the daytime hillsides heat quickly. This heating effect causes updrafts alongslopes—downdrafts in the center.
THERMALS.—Thermals are vertical convectivecurrents that result from local heating. They stop shortof the condensation level. Thermal convection is theusual result of strong heating of the lower atmosphereby the ground surface. A superadiabatic lapse rateimmediately above the ground is necessary to thedevelopment of strong thermals. They form mostreadily over areas of bare rock or sand and in particularover sand dunes or bare rocky hills. In the presence of amoderate or fresh breeze, especially in a hilly terrain, itis impossible to distinguish between turbulent andthermal convection currents. Pure thermal convectionnormally occurs on clear summer days with very lightprevailing wind. In the eastern United States, drythermals are usually of only moderate intensity, seldomreaching an elevation in excess of 5,000 feet above thesurface. The high moisture content of the air masses inthis section in summer reduces the intensity of surfaceheating to some extent. This moisture content usuallykeeps the condensation level of the surface air near oreven below a height of 5,000 feet above the ground. Inthe dry southwestern part of the country, where groundheating during clear summer days is extreme, drythermal convection may extend to a height of 10,000feet or more. Under these conditions, extremelyturbulent air conditions can occur locally up towhatever heights the thermals extend, frequentlywithout a cloud in the sky.
One variation of the dry thermal is seen in the dustor sand whirls, sometimes called dust devils. They areformed over heated surfaces when the winds are verylight. Dust whirls are seldom more than two or threehundred feet high and they last only a few minutes atmost. Over the desert on clear hot days as many as adozen columns of whirling sand may be visible at once.The large desert sand whirls can become severalhundred feet in diameter, extend to heights of 4,000 feetor higher, and in some cases last for an hour or more.They have been observed to rotate both anticyclonicallyand cyclonically, the same as tornadoes.
An almost identical phenomenon is observed overwater in the form of the waterspout. Waterspouts occurfrequently in groups and form in relatively cool humidair over a warm water surface when the wind is light.The waterspout is visible due to the condensed watervapor, or cloud formation, within the vortex. Thecondensation is the result of dynamic cooling byexpansion within the vortex. In this respect it differsfrom the sand whirl, which is always dry. Both the sand
whirl and the waterspout represent simple thermalconvection of an extreme type. They are not to beconfused with the more violent tornado.
When dry thermal convection extends to anelevation where the dry thermals reach thecondensation level, then cumulus convection takes theplace of the dry convection. A cumulus cloud, whosebase is at the condensation level of the rising air, topseach individual thermal current. Beneath every buildingcumulus cloud a vigorous rising current or updraft isobserved. Thus the local thermal convection patternbecomes visible in the cumulus cloud pattern. Thecumulus clouds form first over the hills where thestrongest thermals develop. Under stable atmosphericconditions, little convective cloud development occurs.However, under unstable conditions these thermalsmay develop cumulonimbus clouds.
INDUCED OR DYNAMIC TERTIARYCIRCULATIONS
There are four types of induced or dynamic tertiarycirculations. They are eddies, turbulence, large-scalevertical waves, and Foehn winds.
Eddies
An eddy is a circulation that develops when thewind flows over or adjacent to rough terrain, buildings,mountains or other obstructions. They generally formon the lee (downwind or sheltered) side of theseobstructions. The size of the eddy is directlyproportional to the size of the obstruction and speed ofthe wind. Eddies may have horizontal or verticalcirculations that can be either cyclonic or anticyclonic.
Horizontal eddies form in sheltered areasdownwind of rough coastlines or mountain chains. Anexample of a horizontal eddy is the weak cycloniccirculation that develops in the channel off the coast ofSanta Barbara, California. The winds frequently blowparallel to the northern California coastline during thewinter fog and stratus season. The Santa Barbarachannel often remains fog-free because the waters areprotected from winds that transport the fog inland.However, when the winds are sufficiently strong,friction along the tough coastal range produces a weakcyclonic eddy over the channel. This cyclonic flow,though weak, is sufficient to advect fog into the region.
3-26
Vertical eddies are generally found on the lee sideof mountains, but with low wind speeds, stationaryeddies or rotating pockets of air are produced andremain on both the windward and leeward sides ofobstructions. (See figure 3-29.) When wind speedsexceed about 20 knots, the flow may be broken up intoirregular eddies that are carried along with a wind somedistance downstream from the obstruction. Theseeddies may cause extreme and irregular variations inthe wind and may disturb aircraft landing areassufficiently to be a hazard.
A similar and much disturbed wind conditionoccurs when the wind blows over large obstructionssuch as mountain ridges. In such cases the windblowing up the slope on the windward side is usuallyrelatively smooth. On the leeward side the wind spillsrapidly down the slope, setting up strong downdraftsand causing the air to be very turbulent. This conditionis illustrated in figure 3-30. These downdrafts can bevery violent. Aircraft caught in these eddies could beforced to collide with the mountain peaks. This effect isalso noticeable in the case of hills and bluffs, but is notas pronounced.
Turbulence
Turbulence is the irregular motion of theatmosphere caused by the air flowing over an uneven
surface or by two currents of air flowing past each otherin different directions or at different speeds. The mainsource of turbulence is the friction along the surface ofEarth. This is called mechanical turbulence. Turbulenceis also caused by irregular temperature distribution.The warmer air rises and the colder air descends,causing an irregular vertical motion of air; this is calledthermal turbulence.
Mechanical turbulence is intensified in unstable airand is weakened in stable air. These influences causefluctuations in the wind with periods ranging from afew minutes to more than an hour. If these windvariations are strong, they are called wind squalls andare usually associated with convective clouds. They arean indication of approaching towering cumulus orcumulonimbus clouds.
Gustiness and turbulence are more or lesssynonymous. Gustiness is an irregularity in the windspeed that creates eddy currents disrupting the smoothairflow. Thus, the term gust is usually used inconjunction with sudden intermittent increases in thewind speed near the surface levels. Turbulence, on theother hand, is used with reference to levels above thesurface. Gustiness can be measured; turbulence,however, unless encountered by aircraft equipped witha gust probe or an accelerometer, is usually estimated.
Large-Scale Vertical Waves (Mountain Waves)
Mountain waves occur on the lee side oftopographical barriers and occur when the wind-flow isstrong, 25 knots or more, and the flow is roughlyperpendicular to the mountain range. The structure ofthe barrier and the strength of the wind determines the
3-27
AG5f0329
LOW WIND SPEED(BELOW 20 MPH)
HIGH WIND SPEED(ABOVE 20 MPH)
Figure 3-29.—Eddy currents formed when wind flows overuneven ground or obstructions.
AG5f0330
WINDWARD LEEWARD
WIND
Figure 3-30.—Effect of windflow over mountains.
amplitude and the type of the wave. The characteristicsof a typical mountain wave are shown in figure 3-31.
Figure 3-31 shows the cloud formations normallyfound with wave development and illustratesschematically the airflow in a similar situation. Theillustration shows that the air flows fairly smoothlywith a lifting component as it moves along thewindward side of the mountain. The wind speedgradually increases, reaching a maximum near thesummit. On passing the crest, the flow breaks down intoa much more complicated pattern with downdraftspredominating. An indication of the possible intensitiescan be gained from verified records of sustaineddowndrafts (and also updrafts) of at least 5,000 feet perminute with other reports showing drafts well in excessof this figure. Turbulence in varying degrees can beexpected and is particularly severe in the lower levels;however, it can extend to the tropopause to a lesserdegree. Proceeding downwind, some 5 to 10 miles fromthe summit, the airflow begins to ascend in a definitewave pattern. Additional waves, generally less intensethan the primary wave, may form downwind (in somecases six or more have been reported). These are similarto the series of ripples that form downstream from asubmerged rock in a swiftly flowing river. The distancebetween successive waves usually ranges from 2 to 10miles, depending largely on the existing wind speedand the atmospheric stability. However, wavelengths upto 20 miles have been reported.
It is important to know how to identify a wavesituation. Pilots must be briefed on this condition sothey can avoid the wave hazards. Characteristic cloudforms peculiar to wave action provide the best means ofvisual identification. The lenticular (lens shaped)clouds in the upper right of figure 3-31 are smooth incontour. These clouds may occur singly or in layers atheights usually above 20,000 feet, and may be quiteragged when the airflow at that level is turbulent. Theroll cloud (also called rotor cloud) forms at a lowerlevel, generally near the height of the mountain ridge,and can be seen extending across the center of thefigure. The cap cloud, shown partially covering themountain slope, must always be avoided in flightbecause of turbulence, concealed mountain peaks, andstrong downdrafts on the lee side. The lenticular, likethe roll clouds and cap clouds, are stationary, constantlyforming on the windward side and dissipating on the leeside of the wave. The actual cloud forms can be a guideto the degree of turbulence. Smooth clouds generallyshow smoother airflow in or near them with lightturbulence. Clouds appearing ragged or irregularindicate more turbulence.
While clouds are generally present to forewarn thepresence of wave activity, it is possible for wave actionto take place when the air is too dry to form clouds. Thismakes the problem of identifying and forecasting moredifficult.
3-28
AG5f0331
STRONG WINDS
TURBULENT LAYER
LENTICULARWINDS
CAP CLOUD
TURBULENT
ROLLCLOUD
Figure 3-31.—Schematic diagram showing airflow and clouds in a mountain wave.
Foehn Winds
When air flows downhill from a high elevation, itstemperature is raised by adiabatic compression. Foehnwinds are katabatic winds caused by adiabatic heatingof air as it descends on the lee sides of mountains.Foehn winds occur frequently in our western mountainstates and in Europe in the late fall and winter. InMontana and Wyoming, the Chinook is a well-knownphenomenon; in southern California, the Santa Ana isknown particularly for its high-speed winds that easilyexceed 50 knots. For the purpose of illustrating a Foehnwind, the Santa Ana is used.
The condition producing the Foehn wind is ahigh-pressure area with a strong pressure gradientsituated near Salt Lake City, Utah. This gradient directsthe wind flow into a valley leading to the town of SantaAna near the coast of California. As the wind enters thevalley, its flow is sharply restricted by the funnelingeffect of the mountainsides. This restriction causes thewind speed to increase, bringing about a drop inpressure in and near the valley. The Bernoulli effectcauses this pressure drop in and near a valley.
Generally speaking, when the Santa Ana blowsthrough the Santa Ana Canyon, a similar windsimultaneously affects the entire southern Californiaarea. Thus, when meteorological conditions arefavorable, this dry northeast wind blows through themany passes and canyons, over all the mountainousarea, including the highest peaks, and quite often atexposed places along the entire coast from SantaBarbara to San Diego. Therefore, the term Santa Anarefers to the general condition of a dry northeast windover southern California.
In the Rocky Mountain states, the onset of Foehnwinds have accounted for temperature rises of 50°F ormore in only a few minutes. In southern California, thetemperature, though less dramatically, also rises rapidlyand is accompanied by a rapid decrease in humidity (to20 percent or less) and a strong shift and increase inwind speeds. Although these winds may on occasionreach destructive velocities, one beneficial aspect isthat these winds quickly disperse the severe airpollutants that plague the Los Angeles Basin.
REVIEW QUESTIONS
Q3-15. What is the cause of monsoon winds?
Q3-16. What causes land and sea breezes?
Q3-17. Describe Bernoulli's theorem.
Q3-18. When does a valley breeze usually reach itsmaximum?
Q3-19. What causes eddies?
Q3-20. What causes Foehn winds?
SUMMARY
In this chapter, we studied the primary, secondaryand tertiary circulation of the atmosphere. We learnedabout large-scale circulations, worldwide locations ofmajor pressure systems, horizontal and verticalpressure systems. We studied how pressure systems,temperature, and world winds relate to each other, andfinally we studied small-scale effects, due to localfeatures. A good understanding of atmosphericcirculation is essential in order to understand thecharacteristics of air masses and fronts.
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