Webster’s New World Dictionary defines the word dynamics as the science dealing with motions produced by given forces. In meteorology, these motions, which are produced by given forces, are called atmospheric dynamics. You learned earlier that flow into the center of a low is convergent and outward flow from the center of a high is divergent in the low levels. In this section, you are introduced to a more complete definition of these terms and the atmospheric motions associated with convergence and divergence. To understand this, you must visualize the motions produced by given forces, or the dynamics within the atmosphere.

We begin by discussing horizontal divergence/convergence and their effects on surface pressure. Then, we examine the horizontal and vertical motions in the atmosphere and how they are produced.

Vorticity, another measure of atmospheric motion, is presented next. Specifically, vorticity is a measure of the spin imparted to air parcels because of differential atmospheric motion (shear and curvature). By learning the basic concepts of vorticity, you’ll gain a greater insight into atmospheric motions.

The last part of the section addresses advection, the process by which atmospheric properties, such as vorticity, temperature, and moisture, move from one place to another. Understanding these basic dynamics of the atmosphere is essential in making a good meteorologically sound forecast.

Evidence shows that surface pressure changes are largely controlled by mass changes in the upper troposphere. The troposphere contains 85 percent of the mass and energy transformations in the atmosphere. It is primarily these energy transformations, caused by divergent and convergent forces, which produce our weather. To understand this, you must visualize the motions within the troposphere.

When we speak of the terms divergence and convergence, you must remember that, in mathematical terms, we are talking only about the measurement of divergent forces. If the resultant is greater than zero, we call it divergence. If it is less than zero, we then call it convergence.

Horizontal divergence

Horizontal divergence refers to the spreading of air. Figure 3–1 illustrates the fact that when horizontal divergence occurs, the air moves away from the center of the column. This results from the air’s pushing downward from the top of the column, which adds mass to the column of air. The original column of air then contracts vertically and expands horizontally. That is, the total volume of the air parcel remains constant.

Horizontal convergence

Horizontal convergence refers to the packing of air. Horizontal convergence of a layer near the surface, keeping the volume constant, is shown in figure 3–2. As the air converges horizontally toward the center of the layer, it creates a flow upward toward the top of the layer of air, which contracts the air layer horizontally and expands it vertically; this is caused by upper-level winds taking mass out of the column. Pressure changes

As indicated, pressure changes at the surface result from changes in mass within the troposphere. The pressure at the surface directly relates to the mass of air in a vertical column above the surface. An increase in this mass increases the surface pressure. A decrease in the mass decreases the surface pressure.

Many separate layers of horizontal divergence or convergence are possible. Therefore, the surface pressure measures the net effect of the convergence and divergence.

As air moves inward toward the center of a surface low-pressure area, it must go somewhere. Since this converging air cannot go into the ground, it must rise. Thus, we have some horizontal and vertical motions in the atmosphere. Now let’s look at some relationships that occur between horizontal and vertical motions in the atmosphere.

As shown in figure 3–3, view A, a surface low pressure develops when the winds aloft forcefully take mass out of a column of air due to upper-level divergence. This causes upward vertical motion and low-level convergence. You have divergence occurring when more air is leaving the atmospheric column than is coming into it (net divergence). When this occurs, the mass decreases within the column. Convergence occurs when more air enters the atmospheric column than leaves it (net convergence). Divergence can be caused by (1) heating, (2) strong upper-level winds, or (3) another force that transports air aloft out of our simplistic column. These outflow/inflow patterns not only produce vertical circulations when they are unbalanced, they produce pressure changes as well. Thus, net divergence can reduce the surface pressure and form lows or destroy highs. Figure 3–3, view A, shows net divergence occurring that results in falling pressure as low-level convergence is generated. The vacancy created by the excess diverging air aloft is filled by air from below. As the low-level air converges, it has no place to go but up, thereby forming the chimney effect depicted in figure 3–3, view A.

You find these divergent wind patterns in association with a developing dynamic low east of an upper-air trough near the 300-mb level. In the middle of the troposphere, about the 600-mb level, a transition zone, commonly called the level of nondivergence (LND), is roughly where upper-air divergence changes to convergence in the low levels. Upward vertical motion is at a maximum at the LND and is zero at the surface of the Earth and at the tropopause (fig. 3–3, view B). Horizontal wind speeds at the LND roughly average 40 knots, whereas the upward vertical motion is about 0.25 knots. As you can see from these numbers, the change taking place within the column is small or slow compared to the distance the column is transported horizontally.

Excess, or net divergence aloft, is the primary cause of deepening a developing dynamic low. Figure 3–4 shows another diagram of the chimney effect with low-level winds converging and excess divergence aloft. Notice how this effect produces clouds and weather due to upward vertical motions and adiabatic cooling. For a dynamic low, a deepening mechanism, does create some minor secondary forces that offset or slow the process to some small degree. Convergence below the LND adds mass to the column of air and, as the air rises, it cools adiabatically. Neither of the effects do anything more than slightly moderate the overall chimney effect in time.

The damper effect is simply the reverse process of the chimney effect. When converging winds aloft forcefully put mass into a column of air, it creates net convergence and thereby increases the mass within the column. As the mass increases, the surface pressure rises and, in turn, generates low-level divergence. You can find these converging wind patterns, in association with a developing dynamic high, east of an upper-air ridge near the 300-mb level. As the winds converge aloft, they can only go downward due to the highly stable tropopause. The cold, dense air at the upper levels tends to sink and the convergent upper-level winds help push the air further down the column. This forms the damper effect, as depicted in figure 3–5, view A.

Downward vertical motion is at a maximum at the LND and is zero at both the surface of the Earth and at the tropopause (shown in fig. 3–5, view B). It is the excess or net convergence aloft that is the primary building mechanism for a dynamic high. Figure 3–6 shows another diagram of the damper effect with low-level winds diverging and excess convergence aloft. Notice how this effect produces fair weather due to downward vertical motions and adiabatic warming. Some minor secondary forces slow the process. Divergence below the LND removes some mass from the column of the air, and as the air sinks, it warms adiabatically. Neither of the effects do anything more than slightly moderate the overall damper effect in time.

Remember that jet maxima are important because of their association with migratory pressure systems. The horizontal divergence and convergence caused by the jet stream is the driving mechanism for creating, dissipating, and maintaining migratory pressure systems.

When you consider convergent and divergent areas associated with jet maximas, you must first consider the flow of the jet axis. The orientation of the jet axis enhances or hinders the magnitude of convergence and divergence occurring. For example, a jet axis flowing in a straight line, as seen in figure 3–7, actually has four separate quadrants associated with the jet maxima. Notice that the predominate vorticity isopleth pattern indicates a shear-jet pattern and that little advection is occurring. This makes sense when you consider that a jet in a high zonal flow pattern is relatively weak and is not associated with significant surface systems. The left-front quadrant is considered to be divergent due to the increasing horizontal shear on the +n side. Conversely, the right-front quadrant is convergent due to the increasing negative shear. The center of the jet max (denoted by the vertical dashed line) is basically a neutral area where the convergence and divergence cease and reverse for the rear quadrants.

As you well know, the chances of the PFJ flowing in a straight line for any appreciable amount of time is rare. Therefore, you must be familiar with the modification of the quadrants in cyclonic and anticyclonic flow. When the flow becomes more cyclonic, as seen in figure 3–8, the two quadrants on the +n side of the jet axis increase in magnitude. This is due to the addition of positive curvature vorticity. On the –n side of the jet axis, the positive curvature has a counter affect on the negative shear. This leaves the two quadrants essentially weak or even neutral in magnitude.

When the PFJ flow becomes more anticyclonic, the effect on the four quadrants is reversed (fig. 3–9). The negative curvature vorticity begins to negate the positive shear on the +n side of the jet axis. This leaves the two quadrants weak to neutral in their relative magnitude.

Because divergence/convergence and their magnitude are difficult to estimate on upper-air products, the vorticity product was developed. The only practical way to evaluate divergence/convergence is by using vorticity advection patterns. Let’s now take a more in-depth look at vorticity.

Vorticity is the measure of spin around the vertical axis of an object. For example, a spinning top has vorticity. In a fluid such as the Earth’s atmosphere, vorticity is the measure of spin of a small quantity, called a parcel, of the fluid. Spin or vorticity results from the application of unequal forces to the fluid parcel. For example, in a river, water that touches the banks is affected by friction and the flow is slower than in the river’s center. This difference in speed causes spinning currents of water (eddies) to be frequently seen near riverbanks. The rate at which an eddy spins is a measure of vorticity.

Spinning eddies in the atmosphere form in much the same way. Adjacent parcels of air moving at different speeds cause spinning or vorticity. Any force affecting the atmosphere causes vorticity. Even if the atmosphere were stationary, parcels of air could spin about a vertical axis because of the rotation of the Earth; this is called planetary vorticity. Planetary vorticity, which is equal to the coriolis parameter, is zero at the Equator and increases to a maximum at the poles. Figure 3–10 illustrates that when an object whose vertical axis is perpendicular to Earth’s surface at the Equator exhibits zero spin about its vertical axis. On the other hand, an object whose vertical axis is perpendicular to Earth’s surface at the North Pole spins once about its vertical axis every 24 hours.

As you know, the Earth’s atmosphere is in perpetual motion so forces other than Earth’s rotation affect a parcel’s spin or vorticity. This spin is caused by wind shear and contour curvature. It remains independent of the Earth’s rotation and is measured relative to the Earth’s surface. This spin is termed relative vorticity. When we couple relative vorticity with planetary vorticity, this becomes a parcel’s total or absolute vorticity.

The spin of a parcel of air may be counterclockwise or clockwise. If the direction of rotation is counterclockwise, we have positive relative vorticity. If the rotation is in a clockwise direction, the relative vorticity is negative. Now let’s see how positive and negative relative vorticity occur.

To examine relative vorticity due to shear, let’s look at small parcels of air in a pattern of upper-level, straight-line flow. In this pattern, differing wind speeds and shear surrounding the parcels result in parcels rotating in different directions and different speeds as seen in figure 3–11. Two unequal forces affecting the parcel on top of the dashed line are the result of stronger wind speeds to its right. Therefore, as the parcel moves along in the flow pattern, it rotates counterclockwise and has positive relative vorticity.

Figure 3–11. Vorticity due to shear.

The bottom parcel is also affected by unequal forces; this time, the stronger flow is to its left (with respect to wind flow). As this parcel moves along the flow pattern, it rotates clockwise and has negative relative vorticity. Notice the likeness between the rotating parcels and the eddies in the earlier example of the river. What would happen if an even distribution of forces with equal wind speeds on either side of a parcel were occurring? Such an occurrence would still cause the parcel to move along in the flow pattern, but it would not rotate or have zero relative vorticity.

Relative vorticity also results when a parcel moves through curved flow. This can be likened to a wood chip floating down a meandering stream. The chip spins or rotates as the water follows the bends and curves in the stream. Consider a parcel moving through a flow pattern that outlines a trough of low pressure and a ridge of high pressure (fig. 3–12).

As the parcel moves into the trough, it begins to rotate counterclockwise as though pivoting to negotiate the curve—it has "positive relative vorticity." Rotation keeps on until the parcel reaches a point midway between the trough and the ridge. Then, it begins to rotate in the opposite direction (clockwise) and has negative relative vorticity until it reaches the midpoint between the ridge and the next trough. The points where spin changes from positive to negative or vice versa are called inflection points. At these points, vorticity from curvature equals zero.

To find the relative vorticity of a given parcel, we must consider both shear and curvature. The two effects can counteract each other. Thus, shear can be positive while curvature is negative or vice versa (fig. 3–13).

Vorticity does not cause weather to form, but is related directly to systems that do. Positive relative vorticity, you recall, is counterclockwise spin that occurs because of "cyclonic" shear or curvature. This, of course, means positive relative vorticity is associated with low-pressure systems. Positive vorticity advection indicates upward vertical motions and divergent flow near the tropopause.

Negative relative vorticity (clockwise spin) occurs due to "anticyclonic" shear and/or curvature. Therefore, it is associated with high-pressure systems and indicates downward vertical motions and convergent flow near the tropopause.

To remember these relationships, use this vorticity rule of thumb: "Curl the fingers of your right hand in the direction of spin and your thumb points in the direction of vertical motion."

To summarize, remember vorticity is an instantaneous rotation or spin of very small air parcels. When this rotation results simply from wind shear and/or curvature of the contours, we call it relative vorticity. When twice the angular velocity of the Earth is added (coriolis force) to this rotation, we call it absolute vorticity. Thus, we define absolute vorticity as the sum of the spin of an individual parcel and the spin of the Earth.

The prime advantage of using the vorticity is the detection of vertical motion. It must be emphasized that this vertical motion is indicated by the advection of vorticity and not by the vorticity values themselves. This can be proven mathematically by using the vorticity theorem, but this is beyond the scope of this text. As an element that can be analyzed, vorticity is a very important tool for tracking and predicting the strength of weather systems. The analysis and use of vorticity as a tool is presented in more depth later in this course.

In meteorology, advection is the transport of atmospheric properties by the wind. The atmospheric properties include vorticity, temperature, and moisture. Because moving bodies of air (air masses) tend to be homogeneous, they tend to keep their original temperature, moisture, and other characteristics. As winds move an air mass over the Earth, its characteristics are transported with it (with some modification), causing changes on successive locations. Therefore, the importance of advection lies in the changes it brings about because of horizontal transport of atmospheric characteristics.

Advection of vorticity is determined by the relationship between the 500-mb contours and the vorticity isopleths. We can give a numerical value to this advection; but we more commonly call it weak, moderate, or strong.

Positive vorticity advection (PVA)

PVA occurs when vorticity values increase upstream; i.e., higher vorticity values are moving toward your station. (See figure 3–14.) A strong PVA area usually indicates synoptic conditions creating divergence aloft, convergence in the low levels, and upward vertical motion. As is well known, all this results in poor weather. Several specific results that are not well known are:

• Increasing upward vertical motion (or decreasing downward vertical motion).
• Cooling of the vertical air column (or warming at a slower rate).
• Increasing instability (or becoming more unstable at a slower rate).
• Convergence (or decreasing divergence) at the lower levels.
• Increasing humidity (or becoming drier at a slower rate).
• Beginning of precipitation, if initially saturated.
• A heavier rate of precipitation, if precipitation is already occurring.
• Height/surface pressure falls.

NVA is encountered when vorticity values decrease upstream; i.e., the vorticity values over your location will decrease in the future. (See figure 3–15.) NVA produces downward vertical motion because of convergence aloft and divergence in the lower levels. The weather usually improves with the following specific results:

• Increasing downward vertical motion (or decreasing upward vertical motion).
• Divergence (or decreasing convergence) in the lower levels.
• Warming of the vertical air column (or decreased cooling).
• Decreasing humidity (or increasing at a slower rate).
• Increasing stability (or becoming more stable at a slower rate).
• End of precipitation, if there is only light precipitation initially.
• Lighter precipitation, if there is heavy precipitation initially.
• Height/surface pressure rises.

Neutral vorticity advection means no change in vorticity with time. This implies a parallel relationship (no intersection) between the vorticity isopleths and the 500-mb contour lines or no gradient. The weather usually associated with this is a stagnant or persistent type that is usually dominated by local effects and can be forecast through extrapolation.

Let’s look at a case of temperature advection. For example, the normal process finds surface temperatures decreasing overnight because of the radiation process. Yet, if the temperature warms up overnight, it may be because a warmer body of air moved into the region. Thus, climbing nocturnal temperatures are due to temperature advection. Look at figure 3–16 to see an example of advection of warmer temperatures.

Suppose you are at station A with a temperature of 55°F. Because of westerly winds and warmer temperatures in the direction from which the wind is blowing (upstream), you can expect warming at your station. The amount of warming depends on the strength of the wind and the amount of temperature change versus distance. For example, if the stations in figure 3–16 are 15 nautical miles apart, with 15-knot winds, station A’s temperature increases 5°F per hour because of advection. Since the upstream temperature gradient (change in temperature over distance), wind speed, and wind direction in relation to the temperature gradient stay the same, stations B and C experience the same temperature change per hour. Temperature advection in which temperature warms with time is called warm-air advection. Temperature advection that causes colder temperatures over time is called cold-air advection.

Now, look at figure 3–17 where no temperature advection is taking place because temperatures are the same at all stations. The 72°F temperature stays the same at stations A, B, and C because no warmer or colder temperatures exist upstream from the stations. This figure shows a neutral temperature advection pattern. For advection to occur, there must be changes in air properties upstream from your station.

The gradient level is the top of the boundary layer, which is that part of the atmosphere where the Earth’s surface has an appreciable effect on air motion. The height of the gradient level varies with terrain, wind, and stability. The rougher the terrain, the greater is the vertical wind shear and the more unstable is the atmosphere. These all cause a higher gradient level. The 1000-mb surface and sea-level pressure are found below the gradient level, while 700mb lies above the gradient level when the surface is smooth and located near sea level.

The 850-mb surface may be above or below the gradient level, depending upon the elevation and roughness of the terrain, stability, and wind shear. Typically, in the eastern half of the US, the 850-mb surface lies above the gradient level, while in the west it is below it.

Because temperature advection decreases with height in the atmosphere and wind speed increases with height above the gradient level, temperature advection is typically a maximum somewhere between the gradient level and 700mb.

When cold-air advection occurs, the wind flow is displacing warm, less dense air with cold, dense air. Therefore, it tends to sink (subsidence). Because the cold air is more dense, surface pressures and the height of the 1000-mb surface rise. Heights above the level (gradient level) of maximum CAA fall. Heights and pressures below the level of maximum CAA rise. The resulting effect is a decrease in the thickness. Because the cold air sinks and causes subsidence, the vertical motion is downward.

When warm-air advection occurs, the wind flow is replacing cold dense air with warm, less dense air. Because the warm air is less dense, surface pressures and the height of the 1000-mb surface fall. Heights above the level (gradient level) of maximum WAA rise. Heights and pressures below the level of maximum WAA fall. The result is an increase in thickness. Since the air is warm and less dense, the vertical motion is upward.

Thickness is the vertical distance between two constant-pressure surfaces. The thickness of a column of air is proportional to the mean virtual temperature of the layer. The warmer the column, the higher is its thickness. The colder the column, the lower is its thickness. Although moisture content affects the thickness of a column, we normally associate thickness with temperature. While we can compute thickness using any two constant-pressure surfaces, we usually work with the surface/1000- – 500-mb thickness product.

The surface/1000- – 500-mb product is useful for determining the mean temperature advection in the 1000- – 500-mb layer. We can estimate the mean temperature advection by advecting the 1000- – 500-mb thickness isopleths (lines of equal thickness values) with the surface isobars. WAA occurs in regions where the wind (isobaric flow) advects higher thickness values into a region presently occupied by lower values. CAA occurs in regions where the wind (isobaric flow) advects lower thickness values into a region presently occupied by higher values.

Warm-air advection at the gradient level into a ridge builds it at upper levels (height rises), while cold-air advection into a ridge weakens it at upper levels (height falls). Cold-air advection at the gradient level into a trough deepens it at upper levels (height falls), while warm-air advection into a trough fills it at upper levels (height rises).

We now have an understanding of PVA, NVA, WAA, and CAA and how they individually affect constant-pressure heights and surface pressures. To further your understanding of the dynamics taking place in the troposphere, let’s look at the quasi-geostrophic motions occurring. A quasi-geostrophic system is one in which the motions of the atmosphere are nearly geostrophic. The basic assumptions are that the atmosphere is in hydrostatic balance. That is, the vertical pressure gradient is balanced by the Earth’s gravitational force and the vertical motions in the atmosphere are adiabatic (i.e., no transfer of heat or mass with the environment).

Quasi-geostrophic theory deals with the effects of vorticity advection and thermal advection on the development of developing baroclinic systems. The mathematical underpinnings of quasi-geostrophic theory are complex and are of no direct use to operational weather journeymen/craftsmen. However, this theory is the basis for a number of the approaches we must take to determine where air is moving upward and downward in the atmosphere and where the height of a constant-pressure surface is rising and falling. Vertical motions are important because they let us know where clouds and precipitation are likely to occur. Constant-pressure surface height changes indicate whether troughs are deepening or filling and if ridges are building or weakening.

To use the quasi-geostrophic system to forecast the state of the atmosphere at sometime in the future, we must look at two formulas. The first is the omega equation and the second is the tendency equation. They are both diagnostic equations, but the omega equation forecasts the vertical motion field and the tendency equation forecasts the height field.

We can sum up the omega equation in a nonmathematical sense in the following expressions:

Vertical motions:

1. Upward motions = Rate of increase with height of PVA + WAA on a surface.
1. Downward motions = Rate of increase with height of NVA + CAA on a surface.

The first equation tells us there are two ways to get air to go up. First is positive vorticity advection (PVA) aloft and the second is warm-air advection (WAA). Obviously these two terms can act alone, together, or against each other. Remember that to get clouds and precipitation, we need large-scale ascent; therefore, clouds and precipitation exist in areas of PVA or WAA, especially where the two occur simultaneously.

In the opposite sense (the second equation), we can find clear or relatively cloud-free air in areas of large-scale descent; that is, in areas of negative vorticity advection (NVA) and/or cold-air advection (CAA). Where these two forces are working together, there is clear, subsiding air.

This equation allows for changes in the system’s amplitude and for a qualitative idea of what a height field will look like in time. This is very useful information for us when we try to decide whether a trough will deepen or a ridge will build. We can sum up the tendency equation with the following nonmathematical equations:

1. The rise of the height of a constant-pressure surface = NVA + rate of decrease with height of WAA .
1. The fall of the height of a constant-pressure surface = PVA + rate of decrease with height of CAA.

Using the first equation, the atmosphere can have height rises at a given pressure surface by two means. First, we can have NVA moving into an area and, second, there can be WAA into a region. In figure 3–18, you can see that the WAA causes height rises above the gradient level and falls below it. Of course, if both forces are working together, you’ll have significant height rises occurring at the 500-mb surface in the diagram. The region of WAA can be located on the 850-mb product and on the 1000- – 500-mb thickness analysis.

In the opposite sense (second equation), we can identify areas of height falls by locating regions of PVA or CAA. Each works opposite the terms we explained above. When the two forces work together, significant height falls above the gradient level are likely to occur and significant rises below it (fig. 3–19).

The textbook definition of the second term "thermal advection" as stated above is not particularly clear. Instead of thinking about the rate of decrease with height of the thermal advection (i.e., CAA and WAA), think about the effect that cold air and warm air have on troughs and ridges, respectively. We know that CAA deepens troughs above the gradient level and WAA builds ridges above the gradient level. As cold (more dense, lesser thickness) air advects under a 500-mb trough, the heights fall and the trough deepens.

Conversely, as warm (less dense, greater thickness) air advects under a 500-mb ridge, the heights rise and the ridge builds. The key to understanding this term is that "decrease with height of thermal advection (CAA or WAA)" is really just a fancy way of saying surface/1000- – 500-mb thickness advection. This is why numerical models (discussed earlier) produce a thickness/surface-pressure product.

Both equations, omega and tendency, depict vorticity and thermal advection as important forces to consider in producing a good forecast. Give the two items equal consideration. When you consider these factors together, some combinations tend to reinforce each other, while other combinations offset each other. For example, in an area where PVA and WAA are working together, you can expect to find significant weather. Furthermore, scientific research has shown that low-level WAA plays a key role in the formation of large storm cells.

Having a thorough knowledge of these dynamic interactions is a key in understanding how to forecast the weather. These dynamic processes give us some valuable forecasting rules. Remember, cold-air advection into an upper-level trough deepens that trough/low, while warm-air advection into an upper-level ridge/high builds it. These are important system factors to consider as they indicate whether a system is intensifying or not. This intensification can and will ultimately affect the weather you forecast.

Moisture commonly originates in one region and moves over another. For example, if an air mass forms or stagnates over an ocean region, the air mass becomes moister. If the air mass moves on shore, it carries or advects moisture with it. Frequently, clouds, fog, and precipitation characterize this type of advection. Moisture increasing because of advection is called moisture advection. Conversely, drying because of advection is called dry-air advection. Moist and dry-air advection and temperature advection may occur at any or all levels of the troposphere. Still, most moisture advection occurs in the lower tropospheric layers.

Another situation involving moisture advection occurs when moisture associated with weather systems in a geographic site moves over another site. For example, this commonly occurs when a warm front with many clouds and precipitation ahead of it moves toward your station. Figure 3–20 shows warm front weather and its associated moisture advection. Station A is experiencing generally scattered cirrus clouds and unrestricted visibility. As you look upstream from station A, you see increased cloudiness, lowering ceilings, and eventual precipitation at station D. As warm-air advection drives the warm front toward you, the changes are advected over station A. Increasing dew points characterize surface moisture advection while clouds and precipitation indicate advection of upper-level moisture. Further examples of the advection of atmospheric properties are covered later.

Working largely from surface observations, a group of scientists in Bergen, Norway, developed a model explaining the life cycle of an extratropical storm; that is, a storm that forms at middle and high latitudes outside the tropics. This extraordinary group of meteorologists included Vilhelm Bjerknes, his son Jakob, Halvor Solberg, and Tor Bergeron. They published their theory shortly after World War I. It was widely acclaimed and became known as the polar-front theory of a developing wave cyclone or, simply, the polar-front theory. What these meteorologists gave to the world was a working model of how a mid-latitude cyclone progresses through the stages of birth, growth, and decay. An important part of the model involved the development of weather along the polar front. As new information became available, the original work was modified, so that, today, it serves as a convenient way to describe the structure and weather associated with a migratory storm system.

Since weather systems are somewhat intricate, we need a thorough understanding of the terms associated with systems. This familiarization is intended to assist you in understanding how systems develop.

We use the term "cyclone" to describe any area of closed counterclockwise circulation in the Northern Hemisphere. When this closed circulation occurs on a frontal surface, it is called a wave cyclone.

A favorable place for wave cyclone development is on the polar front. This is where cold air from the high latitudes or poles converges with warm air from the low latitudes or tropics. In theory, the polar front is stationary; often, a series of wave cyclones develop along it because of the enormous difference in the air masses. As they develop, the cyclones distort the polar front and follow a distinctive life cycle.

Cyclogenesis describes the formation of a cyclone or the deepening of an existing cyclone. Thus, when cyclogenesis occurs, the central pressure falls more rapidly than surrounding areas. There is also evidence of an increase in the intensity of the counterclockwise circulation.

Cyclogenesis may begin at the surface and work upward, or begin aloft and work downward to the surface. Waves that work upward are usually associated with surface fronts; waves working downward usually are not.

When the central pressure of a low decreases, it is said to be deepening. Deepening usually coincides with cyclogenesis, but the two terms are not synonymous.

Cyclolysis is the decrease in circulation or dissipation of a low-pressure system. When the counterclockwise circulation area decreases or disappears, the low is undergoing cyclolysis.

Filling refers to an increase in the central pressure of a low, which may or may not coincide with cyclolysis. Remember, cyclolysis refers to an decrease of the cyclonic circulation.

An anticyclone is a closed, clockwise wind circulation in the Northern Hemisphere with relatively high pressure. An anticyclone is commonly called a high. Semipermanent and migratory anticyclones, as wind systems, never reach the intensity that cyclones do. Even anticyclones of great vertical extent are less organized in formation, shape, and development.

Anticyclogenesis describes the formation or intensification of an existing anticyclone. Therefore, if the circulation intensity or area of closed circulation increases, the high is undergoing anticyclogenesis. The definition of the term "anticyclogenesis" requires that intensity or the area of circulation increase.

When the central pressure of the high increases, it is said to be building. Building usually coincides with anticyclogenesis, but the two terms are not synonymous. Remember, there must be a change in circulation to have anticyclogenesis.

Anticyclolysis is the decrease in circulation or dissipation of a high-pressure system. When the clockwise circulation area decreases or disappears, the high is undergoing anticyclolysis.

Weakening refers to a decrease in the central pressure of the high, which may or may not coincide with anticyclolysis. Remember anticyclolysis refers to a decrease or weakening of the anticyclonic circulation.

Short waves are small atmospheric waves associated with temperature advection (increased baroclinicity). Short waves form as a result of an imbalance in the atmosphere as reflected in thermal instability associated with the jet stream. The long waves move cold air Equatorward and warm air poleward in an attempt to balance the energy and heat in the atmosphere. Short waves develop within the long-wave pattern as a result of smaller scale instability that is discussed in detail later.

Short waves are associated with temperature advection and are, therefore, baroclinic. The height contours and isotherms are out-of-phase and cross, resulting in temperature advection. Short-wave wavelengths are generally less than 60° of longitude. Their speeds average about 20 knots per day in the summer while in the winter, the average speed increases to 30 knots per day due to the increased speeds associated with the polar-front jet (PFJ). If a short wave supports a baroclinic low, we call it a short-wave trough. If a short wave supports a baroclinic high, we call it a short-wave ridge.

A major short wave is a short wave that is large enough to distort the long-wave pattern. They also reflect through a large depth of the atmosphere (visible on more than one constant-pressure level) and support surface features such as baroclinic lows or baroclinic highs.

A minor short wave is a small short wave that generally moves faster than a major short wave and causes very little distortion to the long wave. They may not be detectable at more than one level.

The following terms relate to low-pressure systems:

• Baroclinic low.
• Barotropic low.
• Stable waves.
• Unstable waves.
• Mature waves.
• Decaying waves.

When contours and isotherms are out-of-phase, advection is occurring and the atmosphere is described as baroclinic. In other words, a baroclinic low is a low that has temperature advection associated with it. In a baroclinic system the axis tilts with height. The baroclinic low resides in a region of strong thermal contrast. As a result, strong thermal advection exists around the low. The system tilts with height to a short-wave trough. The system stacks down toward the warm air or up to the cold air. The tilted stack of the system causes the wind to change direction with height, thus suggesting thermal advection.

When contours and isotherms are in phase, the atmosphere is said to be equivalent barotropic¾ or simply barotropic. There is no temperature advection. Lows located in pockets of cold air are also called barotropic. Barotropic lows are vertically, or nearly vertically, stacked. This means that the upper tropospheric reflection (upper-level low) of the surface system is located directly above the surface system. These systems are not associated with a contrast in air masses and no fronts accompany them.

You can identify the baroclinic and barotropic lows by noting the relationship between the flow and the temperatures around the system (fig. 3–21). Notice the baroclinic situation has isotherms that cross the flow pattern (temperature advection is occurring). In the barotropic situation, the isotherms and flow pattern are parallel to each other (no temperature advection is occurring).

This low-pressure system is usually shallow, seldom seen above the 850-mb level, and lacks upper-air support. It has a closed-cyclonic circulation and is baroclinic. This system does not intensify nor does it occlude. As it begins to form, you’ll find parallel windflow in the opposite direction on either side of the stationary polar front.

Local convection or heating increases the thickness in the local area, causing pressure falls that induce cyclonic flow and vertical motion. The winds on the cold side turn northerly and the winds on the warm side become southerly. Due to the low-level warm-air advection ahead of the low center and cold-air advection behind it, you’ll find height and pressure falls and rises, respectively. Figure 3–22 shows how this stable wave looks, both horizontally and vertically.

Stable waves tend to propagate along fronts but without changing intensity. Since there is no upper-air support, the movement of this wave depends on the low-level warm-air advection. The amplitude of the wave is small and the pressure rises and falls are of a comparatively small size. The wave can produce significant weather over a small area and often evolve into unstable waves.

Unstable waves deepen and undergo cyclogenesis. This low is also baroclinic, but it is much deeper than the stable wave because it extends higher into the troposphere. This system has upper-air support (i.e., major short-wave trough), which helps intensify it, and may progress into an occlusion (mature wave).

The initial development of an unstable wave is very similar to the stable wave. The upper-air support (an upper-air trough) is the mechanism that intensifies the cyclone. The major short-wave trough and the wave on the frontal system amplify with time as self-development occurs.

The unstable wave is slower moving than the stable wave and the amplitude is much larger, as depicted in figure 3–23. As the cyclone intensifies, the cold air mass pushes further south and the warm air mass is pushed to the north. It also has a large magnitude of pressure rises and falls compared to the stable wave. It is an extensive weather producer, especially ahead of the warm front portion of the system.

As an unstable wave continues to amplify, the low deepens back into the cold air and migrates to the +n side of the jet. The term "+n" is from the three dimensional natural coordinate system that moves with the wind flow. It’s very useful in understanding the processes involved with atmospheric motions. The n axis is oriented perpendicular to the flow, with +n oriented to the left of the flow (cold side of the jet) and –n to the right of the flow (warm side of the jet).

At this point the unstable wave evolves into a mature wave. The surface low center becomes removed from the surface temperature contrast and separated from the warm and cold fronts.

As the cyclone develops to this stage, you must have three different air masses: cold, warm, and cool. For the wave to occlude, the cold front rotates around the cyclone faster than the warm front and eventually overtakes it. As the cold front overtakes the warm front, one front is forced up over the other; this forms an upper front. An occlusion extends from the surface fronts into the cold air mass. The location of the coldest air mass compared with the wave determines the type of occlusion that forms. The point of intersection of the surface warm front, cold front, and occlusion is called the triple point. Unfavorable weather conditions occur here. It is a favorable area for the development of a new low since cyclonic turning is already occurring at this point.

Decaying waves

Eventually, the warm air in the warm sector associated with the mature wave is pushed aloft. In being pushed aloft, the warm air is cut off from the cyclone. Then the cyclone becomes cold-core or barotropic and the low-pressure center begins to fill. This cold-core system is known as a decaying wave. The thermal gradient across the occlusion is gradually lost, washing out the occluded portion of the front. Figure 3–24 shows the placement of the air masses as the system washes out or dissipates. The wave then disappears and the stationary polar front reestablishes.

Baroclinic instability (Holton, 1979) is a process by which short waves amplify (increase amplitude with time) by extracting energy from the north/south temperature gradient. We bypass a formal investigation of baroclinic instability theory and examine two implications. First, long waves are relatively stable. Second, major short waves are most likely to amplify due to baroclinic instability. The optimum growth rate occurs for wavelengths of 3000 to 4000 km.

Recall that differential heating leads to a strong north/south temperature gradient in the mid-latitudes. It also results in the long-wave pattern discussed previously. This gradient is the source of potential energy which baroclinic systems tap as they strengthen. In order for a major short wave to amplify significantly, there must be a large energy transfer from the temperature gradient to the wave. The potential energy of the north/south temperature gradient is transferred to the major short wave by thermal advection throughout the troposphere.

Aloft, the thermal wave and contour wave are out of phase. Cold-air advection occurs into the contour trough; warm-air advection occurs into the contour ridge. This thermal advection, coupled with advection around the well-developed baroclinic system on the surface, causes the upper-level wave to amplify. The greater the amplitude of a short wave, the more energy is associated with it. The well-developed baroclinic system on the surface must be present to ensure thermal advection throughout the troposphere.

As the potential energy is transferred from the temperature gradient to the short wave, the short wave uses the potential energy to develop the low-level circulation of the low or high. The low-level circulation (wind) is kinetic energy. Energy is made available to the system when warm air rises and cold air sinks, as is true in mid-latitude systems. Warm air moves northward ahead of a baroclinic low; behind a baroclinic high, it ascends. Cold air moves southward ahead of a baroclinic high; behind a baroclinic low, it descends.

Because the process proceeds without outside influences if conditions are right, we call it instability. As the cyclogenesis/anticyclogenesis proceeds, the thermal advection increases and more energy is transferred to the wave. The energy then strengthens the low-level circulation. The process continues, repeating itself.

We can draw an analogy to an absolutely unstable layer on a Skew-T. Once a parcel is forced to rise, it continues to rise on its own, drawing energy from the unstable atmosphere. Once a short wave begins to amplify, it continues to amplify by drawing energy from the north/south temperature gradient.

Baroclinic instability is the primary mechanism responsible for the development of mid-latitude synoptic-scale systems (baroclinic lows and highs). These systems help maintain the global heat balance by transporting warm air northward and cold air southward as the atmosphere attempts to gain equilibrium. Amazingly, only 0.5 percent of the total potential energy of the atmosphere is available for conversion to kinetic energy. Additionally, only 10 percent of that 0.5 percent converts to kinetic energy. Imagine the consequences if the atmosphere was more efficient!

Cyclogenesis is favored by certain large-scale conditions. In the long-wave pattern, cyclogenesis typically occurs at, and just downstream from, a long-wave trough axis. Intense cyclogenesis normally occurs in troughs with a negative tilt (Sanders and Gyakum, 1980). Negatively-tilted troughs have an axis that is oriented northwest/southeast, cause stronger divergence, and, therefore, can support stronger cyclogenesis. Positively-tilted troughs have an axis that is oriented northeast/southwest and are more often associated with non-developing systems.

The relationship between the jet stream and cyclogenesis is also very important. Cyclogenesis typically occurs under difluent flow aloft. Supergradient winds deepen troughs and are associated with difluent flow. Upper-level troughs are likely to intensify when the wind speeds upstream from the trough axis are greater than the wind speeds downstream from the axis. The determination of whether the wind speeds are decreasing downstream or not is based on the contour spacing across the trough axis.

Short-wave troughs strengthen as they move into long-wave trough positions and become negatively-tilted downstream from the long-wave trough axis. Short-wave troughs intensify under difluent upper flow, which is common in negatively-tilted troughs. The negatively-tilted difluent major short-wave trough (six words to describe one trough!) stacks 1 to 3° latitude per standard level to a baroclinic low on the surface. This trough must be present to support the baroclinic low. The baroclinic low must be there to ensure temperature advection occurs throughout the troposphere, in order for the wave to amplify.

Baroclinic instability is the link between the upper-and lower-level features. You know that baroclinic instability converts the potential energy of the north-south temperature gradient to the kinetic energy of the disturbance. The strongest north-south temperature gradient occurs at the polar front. Consequently, baroclinic lows form along the polar front. Petterssen’s rule describes the initial interplay between the frontal system and the divergence associated with the major short-wave trough. Thus, all the mechanisms causing cyclogenesis are linked together.

Petterssen’s rule

Petterssen’s rule states that cyclogenesis occurs when and where an area of upper-level divergence becomes superimposed over a low-level frontal zone across which the thermal advection is weak. Figure 3–25 shows the jet (implied by the close spacing of thickness lines), short wave, thermal field, and PVA all must interact for rapid cyclogenesis to occur. If only some criteria are met, a weak surface low may form, but development is slow.

We can use the application of Petterssen’s rule in other synoptic situations. Short waves are constantly moving through the upper levels but only a few lead to the development of strong baroclinic lows. Short-wave troughs approach cold fronts from upstream. If divergence ahead of a short-wave trough occurs behind a surface cold front, in a region of cold-air advection, the tendency is for pressure rises to occur. Conversely, thermal advection offsets the tendency for pressure falls resulting from the divergence aloft. A surface low will not form.

If the divergence ahead of the short-wave trough occurs well behind the surface front, in a cold air mass, intensification at the surface cannot occur. Even if a weak surface low begins to form, no thermal advection occurs. Thermal advection throughout the troposphere is required for amplification of the short wave.

The thermal advection required to amplify the upper-level short wave must be supplemented with surface pressure falls. This is initiated by divergence aloft for the entire system to intensify. This combination is most common along the slow moving portion of a surface front. Since winds already turn cyclonically across surface fronts, cyclogenesis more likely occurs along the front than away from it.

We can illustrate the process of baroclinic instability further by focusing on the developing system rather than on energy transformations. This is known as self-development. (Palmén and Newton, 1969).

• Baroclinic lows tilt with height and stack from upper-level short-wave troughs. Because of the tilted stack, the divergence aloft ahead of the short-wave trough is located above the surface low. Divergence over the surface low causes the low to deepen. This coincides with cyclogenesis as the cyclonic circulation strengthens.

• The cyclonic circulation strengthens the temperature advection occurring around the low. Stronger warm-air advection occurs into the short-wave ridge located ahead of the surface low. Stronger cold-air advection occurs into the short-wave trough supporting the low. This pattern causes the short-wave ridge to build and the short-wave trough to deepen (fig. 3–26).
• The ridge becomes more anticyclonically curved so the vorticity minimum in the ridge becomes smaller. (For example, a "10" vort min becomes a "06.") The trough becomes more cyclonically curved so the vorticity maximum within the trough becomes larger. (For example, a "14" vort max becomes a "20," fig. 3–27).
• The divergence ahead of the short-wave trough strengthens. The positive vorticity advection between the trough and ridge is stronger. The stronger the positive vorticity advection, the stronger is the divergence. This divergence causes the surface low to deepen further and further cyclogenesis to occur (fig. 3–28).

The above pattern continues as a positive feed-back loop by which the upper-level dynamics deepen the surface low and the temperature advection intensifies the upper-level dynamics. This process is called the self-development process because once development starts, it continues until another influence acts to offset or stop it. Self-development, another way of looking at baroclinic instability, was discussed by Sutcliffe and Forsdyke in 1950.

Although self-development causes baroclinic lows to strengthen rapidly, the intensification immediately initiates other processes known as braking mechanisms. These mechanisms slow, and ultimately stop, self-development and prevent the system from deepening indefinitely.

As a baroclinic low deepens due to divergence aloft, the low-level cyclonic circulation increases. Cyclonically-curved flow in the boundary layer causes low-level convergence (a consequence of friction). The addition of mass through low-level convergence offsets the mass removed by divergence aloft. This process, which can act as a braking mechanism, is known as boundary-layer convergence (BLC) or Ekman pumping.

The vertical motion field around the low intensifies as the low strengthens. Stronger upward vertical motion cools the air adiabatically. The adiabatic temperature change within the developing system acts as a braking mechanism.

Physical processes can affect the development of lows as the ascent of dry, stable air opposes development and extracts significant amounts of energy from the low. Forced ascent of moist stable air opposes the development of the low, although not as strongly as dry stable air. This occurs because some energy is returned to the low by the release of latent heat. As conditionally unstable air is forced to ascend, it can actually enhance the development of the low—if the energy returned by latent heat release is greater than that extracted from the low in lifting the air to its level of free convection (LFC).

Baroclinic lows may dissipate in one of two ways. First, the low can lose its upper-level support at any point in its life cycle (the life cycle of lows are covered in the next section). Without upper-level support, the low fills.

Second, the low can proceed through its entire life cycle and become a decaying wave. The system slowly dissipates from the bottom up as boundary-layer convergence adds mass to the column of air and surface pressure rises. Adiabatic cooling of the rising air causes thickness to decrease; the upper low remains, and is situated in a cold pocket. The closed low at 500mb often remains well after the surface low fills.

Now that we’ve examined the development of baroclinic lows and highs through Petterssen’s rule and self-development, we turn our attention to their classic development patterns. By familiarizing yourself with the symptoms evident on the 500-mb heights/vorticity and surface/thickness products, you can determine how far a particular system has evolved.

When several cyclones are present along a front, we call them a family. These wave cyclones develop in similar patterns and progress through a maturing process, or life cycle, unless the system loses its dynamic support (upper-level trough/low). They tend to develop in a series, with each succeeding wave cyclone developing similarly to the preceding one. Figure 3–29 illustrates how a typical wave cyclone family appears on the surface product.

A baroclinic low developing along a front is called a frontal wave cyclone. Frontal wave cyclones can be divided into two types: those that amplify and those that do not. A baroclinic low in which the amplitude of the wave along the frontal system does not change with time is a stable wave. An unstable wave is a baroclinic low in which the amplitude of the wave in the frontal system increases with time.

We now explore the evolution of the classic baroclinic low. Its life cycle can be arranged in five stages:

• Initial conditions.
• Wave initiation stage.
• Wave intensification stage.
• Mature wave stage.
• Dissipation stage.

The initial conditions stage consists of a baroclinic zone located downstream from a long-wave trough. The baroclinic zone may be either a slow-moving front or a thermal gradient organizing into a front. On the 500-mb heights/vorticity product you’ll notice a short-wave trough approaching the baroclinic zone. Development does not occur until the divergence aloft ahead of the trough is over the baroclinic zone. On the surface/thickness product, the baroclinic zone/surface front separates cold from warm air. This front is usually a stationary front or a slow-moving cold front.

During wave initiation, the second stage, the upper-level short-wave trough, causes a wave to form along the front. On the 500-mb heights/vorticity, the short-wave trough nears the surface front. Divergence aloft ahead of the short-wave trough starts to occur over the surface front. On the surface/thickness product, a wave develops along the surface front. Indications of wave formation include winds on the cold air side of the front turning in toward the front and precipitation in the region of the developing wave increasing. Widening of the cloud band at the center of the wave, which is evident on satellite imagery, occurs as the upper-level short-wave trough approaches.

The third stage, wave intensification, is reached as the frontal wave and the upper-level short wave continue to interact and self-development occurs. At 500mb, thermal advection associated with the developing baroclinic low causes the upper-level short-wave trough to deepen. The vorticity maximum in the base of the trough strengthens. On the surface, the frontal wave is unstable and amplifies with time. The surface low is rapidly deepening.

Mature wave stage

As self-development continues to occur, the wave reaches its maximum intensity during the mature wave stage. During this fourth stage, the short-wave trough begins to "close off" at 500mb. In other words, a closed contour forms around a 500-mb low. The vorticity maximum climaxes during this stage. PVA, and the corresponding divergence aloft, are intense initially, but weaken rapidly as the vorticity maximum becomes coincident with the closing low at 500mb. At the surface, the low has occluded and is near peak intensity (lowest central pressure). It reaches its zenith just after it occludes. The surface low is removed from the warm air mass at low levels and continues to develop back into the cold air. The vertical tilt between the 500-mb low and the surface low decreases, but the system is not yet vertically stacked.

The fifth and final stage is called dissipation. Here, the system evolves into a cold barotropic low. The 500-mb low develops more closed contours and becomes more circular and less wave-shaped. The vorticity center is still strong, but is centered in the 500-mb low. Weak PVA suggests the lack of upper-level divergence. The surface low is well back in the cold air mass, beneath the upper low. The system is vertically stacked. The surface low is separating from the occlusion and fills and vanishes long before the upper low. The low has now evolved into a decaying wave (fig. 3–30).

In considering cyclogenesis, you must consider the relationship between cyclones and fronts. Remember, wave formation is more likely on slowly moving or stationary fronts than on rapidly moving fronts. These waves usually move rapidly along the front after their formation.

Cyclogenesis is also likely when an upper trough overtakes a surface front or trough. This method is especially important when the short waves move out of the long-wave trough. Another situation, sometimes associated with short waves, is the movement of a jet maximum over a surface front. When this occurs, there is a steep gradient of temperature and wind speed and the circulation develops a strong cyclonic shear. This shear leads to a cyclonic circulation aloft, which builds downward to the surface.

When the winds aloft flow parallel to the surface front, cyclogenesis is likely. When the winds up to 700-mb level flow parallel to the surface front, divergence aloft usually leads to the development of waves along the front; thus, we have cyclogenesis.

The shape and curvature of the isobars suggest possible cyclogenesis. When the isobars behind a cold front are curved anticyclonically, it shows the front is probably slow moving and, therefore, exposed to cyclogenesis. Cyclonically curved isobars in the cold air behind the cold front show a fast-moving front and little chance of cyclogenesis. The reverse is true ahead of a warm front.

When cyclogenesis occurs, you can usually find the following parameters near the center of the low:

• Marked increase in falling pressure in the cold air.
• Marked local changes in wind direction or speed.
• Increasing cloudiness north of the frontal area.
• Start of precipitation on the cold air side of the front.

As stated previously, anticyclones are closed, clockwise wind circulations in the Northern Hemisphere. Let’s now look at some characteristics associated with anticyclones.

Anticyclones are usually larger than cyclones. An anticyclone may span 2,000 miles or more. Anticyclone size is determined by the size of the air mass with which the anticyclone is associated. Normally, semipermanent anticyclones have a greater horizontal extent than their migratory relatives.

The vertical extent of anticyclones depends on the characteristics acquired in the air-mass source region and later modification. Subtropical (semipermanent) systems have great vertical extent and are persistent. They are stationary or move very slowly. Still, small parts of these semipermanent features may become separated from the main system and move as a migratory high along the edge of the semipermanent high. The smaller highs of lesser vertical extent provide areas of clearing weather between frontal systems.

The more migratory anticyclones, such as those associated with continental polar air masses, are of far less vertical extent than subtropical highs. They stack to short-wave ridges; the ridge is the upper system. The high itself rarely extends above 10,000 feet. These very cold and dry air masses are forced out of their source regions and cold fronts mark their leading edge. Although slow to change, after 2 or 3 days over warm water, they begin to take on the characteristics of subtropical anticyclones.

As with cyclogenesis, anticyclogenesis is favored by certain large-scale conditions. Anticyclogenesis typically occurs at, and just downstream from, long-wave ridge axes in confluent flow aloft. Upper-level ridges are likely to intensify when the wind speeds upstream from the ridge axis are weaker than the winds downstream, as is common with confluent flow. This describes a subgradient condition; subgradient winds build ridges.

A major short-wave ridge and surface baroclinic high must be present for amplification of the wave to occur. As with baroclinic lows, baroclinic highs derive their energy from the north-south temperature gradient. Baroclinic highs do not form along a front, however, but on the cold air side of the front. The surface divergence associated with baroclinic highs precludes the formation of a front in the region of anticyclonic flow.

Both Petterssen’s rule and self-development are analogous for anticyclones. A difference in the application of Petterssen’s rule is that the high doesn’t develop on the front, but on the cold air side of the front.

Baroclinic highs tilt with height and stack from upper-level short-wave ridges.

• Because of the tilted stack, the convergence aloft ahead of the short-wave ridge is located above the surface high. Convergence over the surface high causes the high to build. This coincides with anticyclogenesis as the anticyclonic circulation strengthens.
• The anticyclonic circulation strengthens the temperature advection occurring around the high. Stronger warm-air advection builds the short-wave ridge supporting the high. Stronger cold-air advection deepens the short-wave trough. This pattern causes the short-wave ridge to build and the short-wave trough to deepen.
• The ridge becomes more anticyclonically curved so the vorticity minimum in the ridge becomes smaller. (For example, a "08" vort min becomes a "06.") The trough becomes more cyclonically curved so the vorticity maximum within the trough becomes even larger.
• The convergence ahead of the short-wave ridge strengthens. The negative vorticity advection between the ridge and the downstream trough is stronger. The stronger the negative vorticity advection, the stronger is the convergence. This convergence causes the surface high to build further and additional anticyclogenesis to occur.

There is a limit on the strength of the anticyclonic gradient wind and self-development does not occur as vigorously as for baroclinic lows. Braking mechanisms for highs are much more efficient than for lows. As the high builds, these braking mechanisms begin to strongly oppose further development. As a result, highs never "wind-up" as much as lows. Additionally, the same braking mechanisms that affected the low retard the development of the baroclinic high.

As a baroclinic high builds due to convergence aloft, the low-level circulation increases. Within the boundary layer, anticyclonically curved flow causes low-level divergence which "partially" offsets the mass being added to the system aloft. Friction is primarily responsible for this effect. This phenomenon removes mass from the surface high center.

The vertical motion field around the high intensifies as the high strengthens. Since highs are associated with subsidence, the energy used to force the subsidence is taken from the developing high, thus reducing the amount of energy available for anticyclonic development. This is true no matter what the characteristics of the air mass are (dry/stable, moist/stable, conditionally unstable). This factor limits the strength the baroclinic high can attain. Because the air is forced to subside, there is no latent heat energy returned to the system.

The baroclinic high initially forms under convergence aloft, on the cold air side of the polar front within a polar air mass. Occasionally, the high well north of the polar-front jet may move southward because of a change in the long-wave pattern and evolve into a baroclinic high. As the high moves southeast, where cold-air advection is causing pressure rises, subsidence and diabatic heating cause the air at the center of the high to warm. During its life-cycle, the high moves from the +n (cold) side of the polar-front jet to the –n (warm) side. The air mass warms, the contrast across the front ahead of the high diminishes, and eventually the high is absorbed into the subtropical ridge.

Baroclinic highs may dissipate in one of two ways. First, the high can lose its upper-level support. This can happen at any point in the life cycle of the high. Without upper-level support, the high weakens.

Second, the high can proceed through its entire life and become absorbed into the subtropical ridge. As the high proceeds through its life cycle and moves southeast, the air in the center warms. The front ahead of the advancing high frontolysizes as the temperature across the front diminishes. The high eventually becomes a warm high and is absorbed into the subtropical ridge.

You already know there are at least two types of pressure systems—highs and lows. Highs and lows are categorized further according to individual characteristics, such as vertical extent, vertical structure, and horizontal temperature distribution, which is the key to classification and identification of the systems. We can divide horizontal temperature distribution into two classes: baroclinic and barotropic.

Remember, the term "baroclinic" describes a horizontal temperature distribution that is non-uniform within the system. Thus, temperature advection is associated with the system. We also stated that the term "barotropic" describes temperatures not changing within the system; there is neutral advection.

Knowledge of the types of pressure systems and their structure aids your understanding of weather and eventual analysis of these pressure systems. Our study centers on the basic types of pressure systems:

A cold barotropic high is one in which temperatures on a horizontal level decrease toward the center. Thus, the temperature in the center of the high is lower than it is toward the outside of the system. The pressure at the center of these systems on the surface is very high, but the pressure decreases rapidly with height, as does the anticyclonic circulation. A cold barotropic high rarely extends more than 10,000 feet above the surface and is characterized by low-pressure or cyclonic circulation above it. This type of high is a migratory pressure system. For example, cP air masses are cold barotropic highs (fig. 3–31).

A warm barotropic high is one in which temperatures on a horizontal level increase toward the center. Temperatures in the center of a warm barotropic high are higher than they are on the outside of a system. The pressure at the center is high. Also, the high and anticyclonic circulation increases in intensity with altitude. This type of high has great vertical extent and is usually found over water areas. Semipermanent subtropical high-pressure systems, such as the North Pacific High and the North Atlantic High, are classified as warm barotropic highs (fig. 3–32).

A baroclinic high is a high-pressure system with an axis that tilts toward warm air with increasing height above the surface. At upper levels, it exists as a short-wave ridge. Because they are baroclinic, these systems have temperature advection associated with them. Figure 3–33 shows the basic structure of a baroclinic high.

A warm barotropic low has horizontal temperatures that increase toward the center of the pressure system. A warm barotropic low is like the cold barotropic high in that it exhibits decreasing intensity and cyclonic circulation with height. A warm barotropic low rarely extends more than 10,000 feet vertically and is often found in warm regions under the subtropical ridge. An example is the "heat" low that forms over the desert southwest of the United States during the summer months. Warm barotropic lows of this nature are considered to be semipermanent in nature (fig. 3–34).

A cold barotropic low has horizontal temperatures that decrease toward the center of the low (fig. 3–35). A cold low is like the warm barotropic high in that its intensity and cyclonic circulation increases with height. It also has great vertical extent. Cold barotropic lows may be associated with surface occluded frontal systems (as decaying waves).

As stated previously, a baroclinic low has an unstable frontal wave and associated upper-air short-wave trough. Also, colder temperatures behind the system and warmer temperatures ahead characterize it. In other words, cold-air advection is occurring to the rear of the low. This type of low is normally in the process of deepening. These lows have a vertical axis that tilts toward the cold air associated with the upper-level short-wave trough (fig. 3–36).

The basic feature of synoptic meteorology is the air mass. A front cannot exist unless two air masses of differing characteristics or properties are brought next to each other.

To identify a front, you need to know the elements expected in the frontal zone. These characteristics tell you a lot about what weather to forecast for your station. A front or frontal zone marks the boundary between adjacent air masses. Whether you classify a front as warm or cold depends on the direction of movement of the air mass and the relative temperatures of the air masses involved. The boundary associated with a cold air mass displacing a warm air mass is a cold front. When a cold air mass retreats and is replaced by warmer air, the boundary between the air masses is a warm front. For example, cP air moving out of Canada into the central United States is preceded by a cold front. When this cold air mass moves eastward, a warm front marks its trailing edge because warmer air is replacing the retreating colder air mass. Remember, because cold air is denser, it can displace or move warm air out of the way. Warm cannot, however, displace cold air; it can only replace retreating cold air.

A significant feature in frontal identification is the lapse rate through a front. The lapse rate in the frontal zone shows a strong stabilization or warming. Consequently, the frontal zone resists vertical exchanges of heat and moisture between the air masses on either side of the front. Figure 3–37 illustrates an idealized lapse rate through a frontal zone. An actual sounding might show a less definite frontal zone and many minor variations in the lapse rate. The sharpness and the extent of the frontal inversion shows the height, strength, and degree of mixing between the air masses.

While frontal zones are depicted as blue or red lines on a weather product, there is much more to understanding them. The frontal zone is an area of varying width that acts as a buffer or transition zone between the two types of air masses. Fronts are also zones of low pressure; as such, they are associated with convergence and cyclonic circulation patterns at the surface. Frontal zones have certain other characteristics that aid in their identification and help to explain the type of weather that forms because of air mass movement.

Figure 3–38 illustrates the classical humidity pattern through a frontal zone. The dew point (Td) usually increases sharply with the inversion in the temperature curve. Figure 3–38 also shows that, because of the many variations in the temperature curve, the dew-point curve better indicates the frontal zone.

Figure 3–38. Dew-point temperature curves through a frontal zone.

In association with a subsidence inversion, the dew-point curve shows an entirely different trend. The dynamic warming of air as it subsides reduces the relative humidity. The dew point normally drops off to an undetectable value through a strong subsidence inversion. Figure 3–39 illustrates the characteristics of the dew-point curve through a subsidence inversion and a frontal inversion as they might appear in association following a cold frontal passage. The usual subsidence within southward moving cold air often creates one or more inversions below the frontal inversion. These lower discontinuities can be mistaken for frontal zones, or may be so close to the frontal zone that only one deep inversion is evident. On the sounding shown in figure 3–39, the dew-point curve identifies the inversion between 870 and 820mb as one caused by subsidence. If precipitation should fall from above, however, the raising of the dew point would leave little or no clear indication of the subsidence origin of the layer.

The best sources for identifying frontal inversions are a combination of temperature and dew-point curves coupled with vertical wind distribution. Clues are provided by vertical wind direction changes as you view the Skew-T from the lower levels up through the upper levels. The wind direction backs (changes counterclockwise) with height through inversions associated with cold fronts and veers (changes clockwise) with height through inversions associated with warm fronts. In figure 3–40, look at the wind changes from the 800- to the 650-mb levels. The winds veer (southwesterly changing to northwesterly) between the two levels; this identifies this as a warm frontal zone.

The boundary zone between two sharply differing air masses is usually accompanied by clear discontinuities in temperature, dew point, and vertical wind distribution. You must consider each of these parameters for the greatest degree of certainty in frontal zone identification. Any parameter by itself might prove inconclusive. For example, a backing wind in the vertical might suggest cold-air advection not associated with a cold front. Conversely, a veering wind in the vertical might suggest warm-air advection not associated with a warm front.

The vertical discontinuities of a frontal zone are also represented in the horizontal. Besides temperature and dew-point discontinuities, an easily identifiable wind shift usually accompanies a frontal passage at the surface (particularly a cold frontal passage). The horizontal shift in wind direction is in a clockwise or veering sense (in the Northern Hemisphere) because fronts lie in a trough of low pressure. Usually, wind speed increases when a wind shift occurs. Greater horizontal discontinuities accompany cold frontal passages. Direction and speed discontinuities associated with a warm frontal passage are usually not as sharp, but are normally identifiable.

A streamline is a line showing the direction all air parcels are moving at any instant in time. A baroclinic low has three such streamlines we call conveyor belts. Figure 3–41 illustrates the three different airflow patterns that flow around and through a developing baroclinic low. These different patterns are composed of the warm, cold, and dry-air conveyor belts.

A set of streamlines which originates at low levels in the moist tropical air mass Equatorward of the surface low center we call the warm conveyor belt. The warm conveyor belt flows northward, ascends, and turns anticyclonically at the jet-stream level. The strongest ascent occurs near the surface low. Baroclinic zone cirrus (baroclinic zone cirrus is discussed elsewhere) forms within the northern portion of the conveyor belt, where the warm moist air is being lifted.

The warm conveyor belt flows northward ahead of the cold front in one of two configurations. If the front is an active cold front (anafront¾ discussed in the next section), the warm conveyor belt ascends the frontal surface. This is called rearward sloping ascent. If the front is an inactive cold front (katafront¾ discussed in the next section), the warm conveyor belt will not rise over the frontal surface. The strongest lift is over the warm frontal surface. This is called forward sloping ascent.

The cold conveyor belt originates in the low levels in the cold air east of a low center. The cold conveyor belt flows westward beneath the warm conveyor belt and is associated with subsidence well ahead of the low center. As it nears the low center, it ascends rapidly and emerges from beneath the warm conveyor belt west of the low center. This strong ascent is associated with widespread cloudiness and precipitation. These clouds compose the "head" of the comma-shaped cloud system we see on satellite imagery (comma-shaped cloud systems are discussed in the unit on satellite interpretation).

Once west of the low center, the cold conveyor belt may take one of two paths. It may continue to rise, turn sharply anticyclonically, and join the warm conveyor belt in the upper levels, or it may turn cyclonically and descend well west of the low center.

A third set of streamlines, which originate at upper levels, upstream from the major short-wave trough supporting the low, we call the dry-air conveyor belt. As the dry-air conveyor belt approaches the short-wave trough, it undergoes strong subsidence and drying. As it nears the low, it splits into two branches. One of these branches turns cyclonically and flows north and west of the warm conveyor belt. The boundary between the two conveyor belts is visible as the smooth high cloud border north of the low. The other branch turns anticyclonically and descends to low levels well behind the low.

Fronts are classified by the relative motions of the warm and cold air. The four types of fronts are cold, warm, stationary, and occluded.

A warm front occurs when warmer air replaces cold air. Cold air has a horizontal component of flow away from the front. The warm air flows up over the cold air, forming a frontal slope of to . Remember from the resident course that the frontal slope is the ratio of the height of the upper front at some location (rise), to the distance between that location and the surface front (run). In other words:

A cold front exists when the cold air displaces the warm air. The windflow in the cold air would be toward the front. The cold air forces the warm air aloft, forming a frontal slope of to . Cold fronts move faster than warm fronts.

The stationary front shows no significant change between the warm and cold air. Here, the flow of cold air is parallel to the front. This means the front should have no movement.

Occluded

Remember, an occluded front occurs when the cold front overtakes the warm front and forces the warm air aloft. There are two types of occluded fronts: warm-frontal type and cold-frontal type. The warm-frontal type occurs when the cool air behind the cold front overrides the colder air ahead of the warm front. This results in a cold front aloft. When the cold air behind the cold front lifts the warm front, it is called a cold-frontal type. Figure 3–42 illustrates these two types of occlusions. The line A-B shows the cold-frontal type and line C-D shows the warm-frontal type.

The type and intensity of frontal weather is determined largely by such factors as the slope of the front, the water vapor content and stability of air masses, the speed of the frontal movement, and the relative motion of air masses at the front. Because of the variability of these factors, frontal weather may range from a minor wind shift with no clouds or other visible weather activity to severe thunderstorms accompanied by low clouds, poor visibility, hail, icing, and severe turbulence.

Weather associated with a frontal passage at one place is frequently very different from that associated with a frontal passage at another place along the same front. This is because when an air mass modifies from below as it moves over a surface, the mixing creates variations in characteristics within the air mass. Therefore, frontal zone characteristics vary along the length of the same frontal boundary.

Because air masses have differing temperature and moisture characteristics, they also have different densities; cold dry air masses are the densest.

When air masses meet, the denser air mass displaces the less dense air mass and pushes it aloft. Thus, less dense warm air slides up over the denser cold air, resulting in a sloping frontal surface. The measure of steepness of the slope is determined by the angle measured between the frontal surface and Earth’s surface (fig. 3–43). The varying slopes of frontal surfaces are due to the dynamics involved with the movement of the associated air mass.

We use differences in air-mass properties, such as temperature, moisture, wind, cloud types, and pressure, to find, identify, and track fronts. One of the most easily recognized discontinuities is temperature. At the Earth’s surface, the passage of a front is most often recognized by a noticeable temperature change. The rate and amount of change are partial signs of frontal intensity. Abrupt and sizable temperature changes accompany strong fronts while gradual or minor decreases in temperature characterize weak or diffuse fronts. With passage of a cold front, the cold air mass displaces a warm air mass at the surface. Cold fronts usually move faster and have a steeper slope than warm fronts.

At the surface, decreases in temperature and dew point, a wind shift, and, occasionally, gusty winds mark a cold frontal passage. Cold frontal weather is found in a somewhat narrow band as opposed to warm fronts, which have widespread weather patterns.

Surface cold fronts are classified as active or inactive, both of which are discussed below.

An active cold front moves at average speeds between 10 to 15 knots. Its slope is about miles, as indicated in figure 3–44. The speed and slope are directly related to the perpendicular wind component at the front, which decreases with height through the front. Thus, the net flow is upward throughout the front. This type of front is also called an anafront because most of the clouds and weather caused by the upward flow are at and behind the front. The term "anafront" comes from the word anabatic, which means "moving upwards." When the warm air is stable, altostratus and nimbostratus extend several hundred miles behind the front. If the warm air is unstable or conditionally unstable, thunderstorms and cumulonimbus form up to 50 miles or more behind the front.

The ceiling is normally at the height of the frontal surface, since most of the clouds occur in the warm air. The exception is stratus, which may form in precipitation areas near or immediately behind the front in the cold air.

Visibility may be poor near the front, due to fog or precipitation, but usually improves rapidly as the front moves away. Pressures fall steadily before frontal passage and rise weakly behind it. The temperature and dew point drop rapidly behind the front, due to the sharp temperature gradient created by the upward airflow. The wind veers with frontal passage. Maximum velocity occurs at passage of the front and decreases afterward.

Figure 3–45 illustrates an inactive type of front. Frontal slope is much steeper than the active front, averaging about to . The wind component perpendicular to the front increases with height, which results in a net downward flow of air along the frontal surface. This downward flow helps the front average a movement of about 25 knots. This type of front is called a katafront. The term "katafront" comes from the word katabatic, which means "moving downwards."

The subsidence associated with the downward flow prevents the formation of clouds at or behind the front and visibility increases with frontal passage. However, the downslope air from the front converges with the warm air ahead of the front and produces weather and storms ahead of it. This will be discussed more in the section on squall lines. Squall lines usually occur 50 to 200 miles ahead of the front.

The subsidence in the front weakens the temperature gradient across the front. Temperature falls, due to precipitation in the squall line, further weakening the temperature gradient across the front. Maximum temperature drops usually occur far behind the front. Dew point and wind direction are better indicators of frontal passage than temperature. The wind veers with frontal passage and is strong, gusty, and turbulent for a considerable length of time. Dew points decrease sharply with passage. Pressure falls ahead of the front and rises strongly after its passage. Frequently, the squall line is mistaken for the inactive cold front.

051. Warm and stationary frontal weather effects

Warm fronts have certain characteristics that distinguish them from other types of fronts. Because of dynamics associated with retreating colder air, warm fronts have shallow slopes (fig. 3–46). Warm fronts move at an average speed of 10 to 15 knots. A warm front differs from a cold front in that the associated weather pattern may be more extensive. Clouds are usually more stratified and precipitation is continuous.

Clouds associated with a warm front usually occur in this sequence: cirrus, cirrostratus, altostratus, nimbostratus, and stratus. Cirrus and cirrostratus form several hundred miles ahead of the surface warm front. Gradually, the cirrostratus thickens to form altostratus 300 to 500 miles ahead of the front. Closer to the frontal boundary, the altostratus thickens, becoming nimbostratus with its associated precipitation. The amount and type of clouds and precipitation with a warm front vary with the characteristics of the air masses involved.

When the overrunning warm air is moist and stable, nimbostratus clouds with continuous light to moderate precipitation are found up to 300 miles ahead of the front. The base of the clouds lowers rapidly as more clouds form in the cooler air under the frontal surface. The clouds form by evaporation of the falling rain. When the cooler air mass is stable, the clouds are stratus with fog reducing visibilities. The clouds in the cooler air are stratocumulus when the cooler air mass is unstable.

The pressure ahead of the warm front usually makes a rapid or unsteady fall with a leveling off after frontal passage. As the cold front approaches, the pressure begins to fall again. Wind speed may increase gradually in advance of warm fronts. The temperature rises ahead of the front and levels off after frontal passage.

When overrunning air is moist and unstable, nimbostratus and altostratus clouds frequently have cumulus and cumulonimbus clouds embedded in them. In such cases, heavy rain showers or thunderstorms, which are intense and intermittent, are combined with the continuous precipitation.

When the overrunning warm air is dry, it must ascend to high altitudes before condensation can occur. Therefore, only high and middle clouds are observed.

Visibility, except where the precipitation occurs, is usually good. When the cold air mass is stable, fog may restrict visibility.

Pressures ahead of the front usually make a rapid or unsteady fall, with a leveling off after frontal passage. Wind velocity increases ahead of warm fronts, due to the increased pressure gradient. Velocity reaches a maximum just before frontal passage. The wind also veers with frontal passage. Temperature and dew point are usually constant or slowly rising ahead of the front and increase markedly behind it.

On the upper-air products, isotherms are parallel to the front and packed ahead of it. The stronger the packing, the more active is the front.

A front that becomes stationary, or nearly so, is called a stationary front. Normally, a front moving less than 5 knots or one that does not move steadily in one direction is considered stationary.

A front becomes stationary when the cooler air mass does not move toward or away from the front. This means the wind blows generally parallel to the front rather than toward or away from it.

Though the front is stationary, there is still a net inflow of warm air toward the front that causes slow, shallow advancement of air over the frontal surface. As air is lifted and condensed, clouds and precipitation form in the warm air on and above the front. The width of the precipitation band and low ceilings varies from 50 to about 200 miles.

Weather patterns associated with stationary fronts are much like the weather pattern of the warm front. If warm air associated with the stationary front is stable, stratiform clouds predominate with possible drizzle, light rain, or snow. If warm air is unstable, clouds are a mixture of cumuliform and stratiform varieties; precipitation is showery; thunderstorms may occur.

Extensive fog and low ceilings may result within the cold air mass when the warm rain or drizzle falling through the cold air from the warm air mass above saturates the cold air.

The examples just discussed involve frontal boundaries with only two air masses. With a third air mass, we must consider other frontal configurations. Remember, these configurations are variations on cold and warm frontal structures we call occluded fronts. Occlusions are the result of an interaction between cold, cool, and warm air masses. They combine weather of the warm and cold front into one extensive system. Specifically, convective activity associated with cold fronts merges with low ceilings and visibilities associated with the warm front.

Occluded frontal systems are climatologically more common in the northern United States than in the South. The North favors the development of frontal occlusions. They occur frequently in the winter months in the northwest and northeast sections of the country.

There are two types of occluded fronts—cold and warm occlusions. The type of occlusion is dictated by the relative densities of the air masses. Remember the moisture content of the air mass also affects the air-mass density. If two air masses have equal temperatures, the one with the higher moisture content is the less dense of the two.

Cold occlusions

As we stated previously, a cold-type occlusion occurs when three air masses are situated so the densest, coldest air is behind the cold front with cool, less dense air ahead of the warm front (fig. 3–47). Then, as the three air masses come together, the advancing cold front remains on the surface. Thus, it forces both the less dense cool and warm air masses up and over the cold air mass. As the process continues, the surface warm front becomes a warm front aloft. Figure 3–47 also shows a cross section of a cold frontal occlusion that depicts the typical weather and associated cloud patterns. The embedded thunderstorms with the cold occlusion usually occur with passage of the surface occluded front. This is the most common type of occlusion. The name cold occlusion was derived from the front that stays on the surface—here, the cold front.

Warm occlusions

In a warm-type occlusion, the densest (usually the coldest) air mass is ahead of the warm front with a less dense (cool) air mass behind the cold front (fig. 3–48).

When the three air masses come together, the densest cold air ahead of the warm front stays at the surface. This forces the advancing less dense, cool air mass aloft. Being the least dense, the warm air is wedged between the cool and cold air masses and is forced aloft, just as it was with the cold frontal occlusion. Since the cool air mass is overrunning the cold, a warm front stays at the surface, forcing the cold front aloft. As with the cold occlusion, the warm occlusion is named for the front remaining at the surface.

The cloud pattern with this system is much wider than with the cold-type occlusion. This is because of extensive overrunning of the cool air over the cold air. This overrunning situation is like that associated with a warm front. A line of thunderstorms with the warm occlusion is often embedded within the stratiform overcast layer and may precede the occlusion and surface warm front by 200 to 300 miles.

In the initial stages of a cold occlusion, the weather and cloud sequence ahead of the occlusion resembles a warm front pattern. Near the surface position of the front, the cloud and weather patterns resemble a cold front. As the occlusion continues to lift the warm air higher, the warm front and prefrontal clouds disappear. Now the cold frontal pattern prevails, and most of the precipitation occurs just ahead of the occlusion. Clearing occurs rapidly following passage of the upper warm front or the surface occlusion, depending on the stage of development. Surface pressure drops rapidly ahead of the occlusion and rises rapidly following its passage.

Figure 3–49 shows the typical lapse rate changes in a cold front type occlusion. In the three curves, station X is first in the cool air overrun by the warm air (curve A), then in the warm sector (curve B), and last in the cold, subsiding air (curve C). Station Y is deep within the circulation of the low and the frontal discontinuities are weak due to the convergent lift of the air. Curve A shows station Y in the cool air overrun by the warm air. Curve B shows station Y in cold air overrun by cool air. Curve C indicates the station is in the cold air.

The occlusion usually disappears before it reaches the 700-mb level due to the loss of the warm front portion. As this occurs, the thickness gradient or mean isotherms associated with the surface front weakens. This is depicted in figure 3–48. In
figure 3–48, view A, the upper warm front still exists and the gradient between X and Z is strong because of the warm air aloft. In figure 3–48, view B, the upper warm front has disappeared and the gradient has weakened between X and Z because the temperature contrast has decreased with the loss of the warm air. The vertical wind distribution resembles the individual frontal systems wind distribution.

Figure 3–49 illustrates the typical lapse rate changes with the passage of a warm-type occlusion. Station X is first in the cold air overrun by the warm air (curve A), then in the warm sector (curve B), and last in the cool air (curve C). Station Y is first in the cold air overrun by the warm air, next in the cold air overrun by cool air but also with the warm air above this (curve B), and then in the cool air (curve C).

The weather pattern of a warm front occlusion has the characteristics of both warm and cold fronts. Ahead of the occlusion, the sequence of clouds is similar to that of a warm front. The upper cold front may be accompanied by showers and thunderstorms if the warm air is moist and unstable. As the warm air is lifted higher, the weather activity diminishes. Normally, there is clearing weather after passage of the upper cold front.

The surface pressure shows a steady fall ahead of the upper cold front, leveling off after passage, and a steady rise following passage of the surface occlusion. The temperature rises slightly following passage of the occlusion. Winds in the horizontal and vertical are similar to the individual fronts.

The term "front" refers to a frontal zone, frontal surface, or surface front, depending on the user. For instance, a frontal zone is a transition zone between two air masses of different properties. A frontal surface is a surface of separation of the two different air masses. A surface front is the line of intersection of either a frontal surface or frontal zone with the Earth’s surface.

Frontogenesis is the formation of a new front or the regeneration of an old front. This occurs only when two conditions are met: (1) two air masses of different densities exist next to one another, and (2) a prevailing wind field brings the two air masses together.

Frontolysis, simply put, is the dissipation of a front. This occurs when either the temperature difference between the two air masses disappears or the wind carries the particles of the air masses away from each other. Frontolytical processes are most effective in the lower layers of the atmosphere. Surface heating and turbulent mixing often quickly modify the air in a frontal zone, which eliminates the temperature contrast. Therefore, fronts are often more easily identified above the friction layer or with a thickness analysis.

Frontogenesis and frontolysis occur under different conditions. Let’s examine ways to identify them.

The distribution and concentration of isotherms within a deformation field consist of a col area of flat pressure between two opposing highs and two opposing lows.

Cols are favored areas for frontogenesis because the horizontal motions of the atmosphere in these areas contribute to sharp horizontal temperature gradients. Consider, for example, the circulation pattern shown in figure 3–50 to represent a col between opposing highs and lows. The flow along line B-B tends to move the isotherms toward the neutral point at the center of the col. This flow of maximum contraction along a common axis is known as the axis of contraction. This shrinking (or contraction) is complemented by a stretching (or dilatation) along line A-A, which is known as the axis of dilatation. The isotherms tend to become concentrated and parallel to line A-A. Petterssen showed that the angle between the axis of dilatation and the isotherm packing is important in determining whether frontogenesis will develop. If the angle between the axis of dilatation and the isotherm ribbon is less than 45° , like in figure 3–50, frontogenesis will likely occur. A warm or cold front can develop along the axis of dilatation if the conditions for frontogenesis are satisfied.

The axis of contraction is easily seen behind a moving cold front. You just simply look for the delineation between the cyclonically and anticyclonically curving isobars as seen in figure 3–51. What isn’t apparent is the axis of contraction on the warm side of the front. Because the system is moving, the direction of the inbound parcels on the warm side of the front is masked. The axis of dilatation is parallel to the front and, therefore, the isotherms. This type of deformation zone is responsible for the constant reinforcement of the cold frontal boundary. Without it, the cold front would undergo rapid frontolysis due to diabatic effects (heated from below).

Another deformation zone pattern is the frontolytic pattern. In this pattern (fig. 3–52), the axis of contraction (line B-B) and the isotherms are parallel and the axis of dilatation (line A-A) is perpendicular. This would result in the isothermal packing pattern being pulled apart by the axis of dilatation.

The above two examples are extreme. A vast majority of the deformation zones identified on products posses some degree of angle between the axes and the isotherms. The easiest way to remember this is to reference this angle difference to either the axis of dilatation or the axis of contraction. For ease of memorization (and this author’s preference), we stick with the axis of dilatation. When the angle between the axis of dilatation and the isotherms are greater then 45°, frontolysis is favored. Conversely, if the angle is less than 45° , then frontogenesis is favored. Angles of 45° show doubtful frontogenesis, or even frontolysis. Figure 3–52 illustrates the isotherm and windflow pattern most favorable for frontolysis.

The next factors we look at are convergence and divergence. Low-level convergence (divergence) always supports frontogenesis (frontolysis). It does not, however, guarantee frontogenesis (frontolysis). Upper-level dynamics associated with the polar-front jet are critically important for frontogenesis and frontolysis. Divergence (convergence) aloft, as shown by PVA (NVA) on the vorticity product, is associated with low-level convergence (divergence) and supports frontogenesis (frontolysis).

Cyclonically (anticyclonically) curved flow combined with the friction found in the boundary layer leads to low-level convergence (divergence). Since surface fronts are found in regions of cyclonic turning of the wind, this effect helps maintain the front. A sharpening of the cyclonic turning strengthens the front. A front moving into a region where the low-level wind turns anticyclonically undergoes frontolysis.

The final phenomena affecting the frontal intensity are diabatic processes. A diabatic process is one that involves the exchange of energy (heat) between air parcels and their environment, the opposite of an adiabatic process. Recall from the resident course that in an adiabatic process, no heat or mass is exchanged between the parcel and its environment. The main diabatic effect we are concerned with is the transfer of energy between the Earth’s surface and the low levels of the atmosphere.

Frontogenesis is supported if some mechanism heats the warm air or cools the cold air. This is the primary mechanism by which air masses form; this effect is common in air mass source regions. Outside air-mass source regions, this situation is uncommon. One example of this is when the warm air ahead of a cold front is heated further by the warm waters off the East Coast.

Frontolysis is supported if some mechanism cools the warm air or heats the cold air. This is the typical case outside air-mass source regions. The warm air behind a surface warm front moves over a relatively cool land mass (previously occupied by cold air) and cools. The cold air behind a surface cold front moves over a relatively warm land mass (previously occupied by warm air) and warms. This effect is further enhanced because as cold air moves Equatorward, it becomes shallow. The shrinking in the vertical (subsidence) causes adiabatic warming. A shallow layer of cool air is modified rapidly over a warm surface.

Any combination of these factors (low-level deformation zones, convergence/divergence, boundary-layer processes, and diabatic effects) may be present. If all of the factors are supporting frontogenesis (frontolysis), then a significant change in frontal intensity or the development (dissipation) of a front occurs. If some factors oppose one another, then the change in frontal intensity is small and difficult to estimate. Continuity is essential for tracking the intensity of fronts. Monitoring the lapse rate in the stable layer and the thickness/temperature contrast across the transition zone reveals important changes in the strength of the front. The strength of the front plays a large role in determining the conditions around the surface front. The stronger the front, the more likely significant weather is to occur along it.

Frontal intensity is a measure of the strength of the density contrast across the transition zone. This density contrast can be seen in the thickness gradient on a thickness product, temperature gradient of a front, and lapse rate across the front. The thickness gradient is the first property we investigate.

The thickness gradient is the difference between the thickness on the warm and cold sides of the front. It is also a measure of the density contrast across the front. A weak front has a weak thickness gradient (widely spaced thickness lines) in the transition zone. A strong front has a strong thickness gradient (closely spaced thickness lines) in the transition zone.

Examining the difference between the temperature on the warm and cold sides of the front (temperature gradient) is the second way to measure the density contrast across the front. The temperature gradient is evident at various levels (surface, 850mb, etc.). A weak front has a weak temperature gradient (widely spaced isotherms) in the transition zone. A strong front has a strong temperature gradient (closely spaced isotherms) in the transition zone.

This is the parameter we most frequently use in finding the intensity of fronts. In practice, the temperature gradient for determining frontal intensity is the difference between the representative warm air immediately next to the front and the representative surface temperature 100 miles from the front on the cold air side. By convection, the transition zone is taken to be part of the cold air mass.

A suggested set of criteria based on horizontal temperature defines a weak front as having a gradient less than 10°F per 100 miles, a moderate front as when the gradient is 10 to 20°F per 100 miles, and a strong front as where the gradient is greater than 20°F per 100 miles.

The third property used to evaluate the density contrast across a front is the lapse rate through the stable layer (frontal inversion) on a Skew-T. Remember that the lapse rate is positive if the temperature decreases with height. A weak front has a small positive lapse rate. The temperature decreases slowly with height. A strong front has an inversion (negative lapse rate). The temperature increases with height.

Not only are density differences across a front indicative of frontal intensity, you can also use turbulence and wind shear as an indicator of it.

A front may be intense with discontinuity across it, but may have no weather except strong winds and a change in temperature. A front, normally classified as weak, is considered moderate if turbulence and gustiness occur along it. The term "gustiness" includes thunderstorms and strong winds despite the amount of wind shear.

Wind shear may be either the vector difference between the surface-geostrophic wind components parallel to and immediately on either side of the front or the 1,000- – 500-mb thermal wind shear. The greater the geostrophic wind shear, the more intense is the front. Convert thermal wind shear into frontal intensity by using the following relationships. If the thermal wind shear is:

• Equal to or less than 25 knots, no front exists or frontolysis is in progress.
• Greater than 25 knots but equal to or less than 50 knots, it is a weak front or frontogenesis is in progress.
• Greater than 50 knots but less than or equal to 75 knots, the front is of moderate intensity.
• More than 75 knots, the front is strong.

Certain regions of the world have a high frequency of frontogenesis. These regions usually have the greatest temperature contrasts. Two of the most important areas are those over the North Pacific and the North Atlantic Oceans.

The polar front, usually present all year, is a boundary between either tropical and polar air (the most frequent case), between maritime polar and continental polar air, or between modified polar air and a fresh outbreak of polar air. This front is common over North America in winter and near 50°N latitude.

Coastal areas are most favorable for formation and intensification of the polar front in winter because of the large temperature contrast between the ocean and continent.

To explain the geographical distribution of the principal zones of frontogenesis and the principal frontal zones, the approach has been through a consideration of wind systems and source regions for air masses. Generally, the source of temperature difference is partly the normal temperature drop with increasing latitude and partly the temperature contrasts between oceans and continents. In areas of convergence where air masses with different properties are brought together, moisture and temperature gradients are steep. It is this steepness of gradient that defines a front.

There are four principal types of air masses in each hemisphere. The four are named according to their source regions: arctic, polar, tropical, and Equatorial. Arctic air is separated from polar air by the Arctic front. The most important frontal system is the polar front, which exists between tropical and polar air masses.

Besides the principal fronts, secondary fronts may form temporarily within the air masses when the distribution of temperature and wind is favorable for convergence. They form most frequently in polar air masses, particularly between "old" and "fresh" polar air or between continental polar and maritime polar. The secondary fronts are usually neither extensive nor long lived, while the principal fronts are extensive and more permanent.

The polar front does not make up a single, continuous, band of convergence and bad weather. It appears as a series of active, convergent zones, which move in a general direction from west to east. These convergent zones are the migratory low or traveling cyclones of the middle latitudes. It has been shown that in middle and high latitudes it is impossible for air masses with different densities (temperatures) to lie side by side in continuous equilibrium. Waves form along the discontinuity or front that separates the two. If the wavelengths of the disturbances are about 300 to 1,800 miles, the waves are unstable and their amplitude increases with time. Here, the cold air in the rear of the waves is pushed toward the Equator, where it tends to underrun the warm air, thus forming the cold front of the cyclone. A compensating poleward flow of warm air occurs in the forward portion of the advancing wave, thus forming the warm front of the cyclone. The warm light air behind the warm front tends to ride up and over the denser cold air ahead of the front. The waves develop into occlusions; a typical extratropical cyclone is the result.

The formation of a baroclinic low along the polar front is first noted as a slight indentation of the front. This indentation, or wave, deepens and moves along the front, rapidly at first, and then becomes progressively slower as it becomes associated with the self-development process.

The waves nearly always move from west to east for two reasons. In the first place, this is the prevailing direction of the windstream in which they form. Secondly, according to wave theory, a wave-type cyclone should move so that the warmer air is on the right in the Northern Hemisphere and on the left in the Southern Hemisphere. The warmer air is usually to the south of the polar front in the Northern Hemisphere and north of the polar front in the Southern Hemisphere. Thus, the zones of active convergence along the polar fronts in each hemisphere appear as migratory lows that travel from west to east, generally more than 20° from the Equator.

There are several major differences between the two types of cyclones. The most important is the lack of symmetry (of temperature, wind, pressure, cloudiness, and precipitation) in the extratropical type. This lack of symmetry is a result of the presence of two or more different air masses. Other important points of variance arise because the extratropical cyclone is not so intense as the tropical. This is because it covers a broader area and has less torrential rain and because the calm center, or "eye," is rarely observed outside the tropics. The extratropical cyclones are most vigorous in winter, when the air mass differences and intensity of the polar-front jet are greatest. Tropical cyclones are characteristic of the warm season.

The weather pattern around baroclinic lows is typified by the conventional model of the extratropical cyclone. Note that the extratropical cyclone of the Southern Hemisphere is a mirror image of the conventional model of the Northern Hemisphere.

The extent of the warm front precipitation area depends largely on the structure of the warm air mass. With comparatively warm, dry air on the Equatorial side of the polar front, the precipitation and cloud area may be small and uneven. If the overrunning warm air possesses a large amount of moisture but is fairly stable, the cloud and precipitation area will be large, with uniform precipitation intensity throughout. If the warm air was initially unstable, the precipitation area may be large but irregularly distributed (showery).

Similarly, the intensity of weather along the cold front is determined by the temperature difference across the front and by the air-mass properties of the warm air. If the warm air is moist but stable, the precipitation zone will be narrow and of moderate intensity.

If the warm air was initially unstable, the bad weather zone will be more extensive and of greater intensity. Although the average rate of movement of extratropical cyclones is about 25 knots, the speed of individual storms varies widely about this mean value.

In winter, the most favorable conditions for vigorous frontal activity are concentrated along the east coasts of North America and Asia. Cold air masses from continental sources meet warm, moist air from the oceans. The warm ocean currents along these coasts greatly accentuate the frontal activity. The great difference in temperature between the air masses, caused by the contrasting characteristics and proximity of their sources and the great amount of moisture that feeds into the air from the warm ocean currents, accounts for the intensity and persistence of these frontal zones off east coasts in winter.

In the Atlantic, the polar front of winter swings back and forth between the West Indies and approximately 45°N with a maximum of intensity when the front coincides with the coastline. Waves with cold and warm fronts form along this polar front and move northeastward along the front. Like all cyclonic waves, they develop low-pressure centers along the frontal trough. They may grow into strong disturbances called bombs (lows that deepen by more than 1-mb/hour) and go through the usual stages of development: wave initiation, intensification, mature wave, and dissipation.

The cyclonic waves occur in families. Each family of waves is associated with a southward surge, or outbreak, of cold polar air. The winter polar front commonly extends from 30°N in the western Atlantic to near 45° N in the Eastern Atlantic. As the polar air advances, it pushes the front southward. The outbreak occurs and polar air spills Equatorward.

No regular interval exists for the large outbreaks of polar air. Under average conditions, there are from three to six cyclonic waves on the polar front between each outbreak of polar air. The first of these usually travels along part of the front that lies farthest to the north. As the polar air accumulates north of the front, the front is pushed southward. The last wave, therefore, follows a path that starts farther south than the path followed by the first wave. These families of polar-front cyclones appear most regularly over the North Atlantic and North Pacific in winter.

During the summer months, the polar front of the Atlantic recedes to a location near 50°N with the average summer storm track extending from the St. Lawrence Valley, across Newfoundland, and on toward Iceland. Polar outbreaks, with their accompanying family groupings of cyclones, are very irregular in summer and often do not exist at all. Frontal activity is more vigorous in winter than in summer because the polar and tropical air masses have greater temperature contrasts in winter and polar highs reach maximum development in winter. Both factors increase the speed of winds pouring into fronts.

Over the oceans of the middle latitudes, a third factor helps to make winter fronts more vigorous than summer fronts. In winter, continental air becomes very unstable when it moves over the comparatively warm ocean surface. In summer, it remains relatively stable over the comparatively cool ocean. Summer frontal activity (in middle latitudes) is therefore weak over oceans and over land. The high-moisture content of maritime air causes much cloudiness, but this moisture adds little energy to frontal activity in the relatively stable summer air. In winter, frontal activity is most vigorous over oceans off the east coast of continents and moderate over oceans off west coasts.

The coldness of cP air masses that move eastward over the oceans, combined with the warmth of ocean currents over which they move, accounts for the vigorous activity off east coasts. Modification of the air masses as they sweep eastward across the ocean cause the modified frontal activity off west coasts. The polar-front activity of the Pacific is similar to that of the Atlantic, except that in winter there are usually two fronts.

When one high dominates the subtropical Pacific in the winter season, the Pacific polar front forms near the Asiatic coast. This front gets its energy from the temperature contrast between cold northerly monsoon winds and the tropical maritime air masses they meet and from the warm, moist Kuroshio Current. Storms usually occlude when they move along this polar front before they reach the Aleutian Islands or the Gulf of Alaska. This frontal movement is limited to the southern side of the Aleutian low because mountains and the North American winter high-pressure center prevent fronts from passing northward through Alaska without considerable modification.

When cold season cyclones reach the Aleutians and the Gulf of Alaska, cold continental winds from the Arctic feed them from the north and cool mP winds from the ocean feed them from the south. Here, where Arctic air meets maritime air over comparatively warm water, is the Pacific Arctic front of winter. Although many occluded storms dissipate in the Gulf of Alaska (decaying waves), others become strongly regenerated with waves developing at the triple points of the occluded systems.

When the Pacific subtropical high divides into two cells or segments, a front forms near Hawaii. Along this front, the "Kona" storms develop and move northeastward. When this second polar front exists, two systems of cyclonic disturbances move across the Pacific. Because of their greater sources of energy, however, storms that originate over the Kuroshio Current and move toward the Aleutians are always more severe. In the Atlantic, a second polar front, similar in nature and cause to the second polar front of the pacific, sometimes, though rarely, develops. During the summer months, the Pacific polar front lies to the north of Kamchatka and the Aleutians and shows no rhythmic polar outbreaks.

Classify the frontal type for each of the following characteristics: