As the day progresses, warm air rises and draws the cool air up from the valley, creating a valley breeze. At night the mountain slopes cool more quickly than the nearby valley, which causes a mountain breeze to flow downhill. Katabatic winds move up and down slopes, but they are stronger mountain and valley breezes.
Katabatic winds form over a high land area, like a high plateau. The plateau is usually surrounded on almost all sides by mountains. In winter, the plateau grows cold. The air above the plateau grows cold and sinks down from the plateau through gaps in the mountains. Wind speeds depend on the difference in air pressure over the plateau and over the surroundings. Katabatic winds form over many continental areas.
Extremely cold katabatic winds blow over Antarctica and Greenland. Chinook winds , also called Foehn winds , develop when air is forced up over a mountain range. This takes place, for example, when the westerly winds bring air from the Pacific Ocean over the Sierra Nevada Mountains in California. As the relatively warm, moist air rises over the windward side of the mountains, it cools and contracts.
If the air is humid, it may form clouds and drop rain or snow. When the air sinks on the leeward side of the mountains, it forms a high pressure zone.
The windward side of a mountain range is the side that receives the wind; the leeward side is the side where air sinks. The descending air warms and creates strong, dry winds. Chinook winds can raise temperatures more than 20oC 36oF in an hour and they rapidly decrease humidity. Snow on the leeward side of the mountain disappears melts quickly. If precipitation falls as the air rises over the mountains, the air will be dry as it sinks on the leeward size. Santa Ana winds are created in the late fall and winter when the Great Basin east of the Sierra Nevada cools, creating a high pressure zone.
The high pressure forces winds downhill and in a clockwise direction because of Coriolis. The air pressure rises, so temperature rises and humidity falls. The winds blow across the Southwestern deserts and then race downhill and westward toward the ocean.
Air is forced through canyons cutting the San Gabriel and San Bernardino mountains. The hot, dry winds dry out the landscape even more. If a fire starts, it can spread quickly, causing large-scale devastation. High summer temperatures on the desert create high winds, which are often associated with monsoon storms.
While that is true there is a more fundamental process that takes place for the cause of rising warm air. Warm air rises primarily due its lower density as compared to cooler air. As the temperature increases, the density of the air decreases. But even air that is of a lower density will not begin to rise by itself.
Isaac Newton's first law of physics is that the velocity of an object will remain constant unless another force is exerted on that object. The more common way of saying this is 'an object at rest tends to stay at rest and an object in motion tends to stay in motion'.
This is why decreasing the density of air alone is not sufficient enough to cause air to rise. There must be another force exerting on the less dense air for it to begin its upward motion. That force is 'gravity'. Gravity's role is its pull of cooler, denser air toward the earth's surface.
As the denser air reaches the earth's surface it spreads and undercuts the less dense air which, in turn, forces the less dense air into motion causing it to rise. If the rising air is warmer and less dense than the surrounding air, it will continue to rise until it reaches some new equilibrium where its temperature matches the environmental temperature.
In this case, because an initial change is amplified, the air parcel is unstable. In order to figure out if the air parcel is unstable or not we must know the temperature of both the rising air and the environment at different altitudes.
One way this is done in practice is with a weather balloon. We can get a vertical profile of the environmental lapse rate by releasing a radiosonde attached to a weather balloon.
A radiosonde sends back data on temperature, humidity, wind, and position, which are plotted on a thermodynamic diagram. This vertical plot of temperature and other variables is known as a sounding. If an air parcel is dry, meaning unsaturated, stability is relatively straightforward. An atmosphere where the environmental lapse rate is the same as the dry adiabatic lapse rate, meaning that the temperature in the environment also drops by 9.
After some initial vertical displacement, the temperature of the air parcel will always be the same as the environment so no further change in position is expected.
If the environmental lapse rate is less than the dry adiabatic lapse rate, some initial vertical displacement of the air parcel will result in the air parcel either being colder than the environment if lifted , or warmer than the environment if pushed downward.
This is because if lifted, the temperature of the air parcel would drop more than the temperature of the environment. This is a stable situation for a dry air parcel and a typical scenario in the atmosphere.
The global average tropospheric lapse rate is 6. Finally, if the environmental lapse rate is greater than the dry adiabatic lapse rate, some initial vertical displacement of the air parcel will result in the air parcel either being warmer than the environment if lifted , or colder than the environment if pushed downward. This is because if lifted, the temperature of the air parcel would drop less than the temperature of the environment. This is an unstable situation for a dry air parcel. When moisture is added, everything gets more complicated.
In Chapter 4 we learned that whether or not an air parcel is saturated depends primarily on its temperature and, of course, its moisture content. The graph of the Clausius-Clapeyron relationship shows us that given the same amount of moisture, air is more likely to be saturated at a lower temperature.
We know that as an air parcel is lifted, its temperature drops according to the dry adiabatic lapse rate. So what happens when the air parcel is cold enough that the air becomes saturated with respect to water vapor? The short answer is that if it continues to cool, water vapor will condense to liquid water to form a cloud.
When water vapor condenses, it goes from a higher energy state to a lower energy state. Energy is never created nor destroyed, especially in phase changes, so what happens to all that excess energy? The energy gets released in the form of latent heat. The latent heat of condensation is approximately equal to 2. This has large consequences for the lapse rate of an air parcel and distinguishes the dry adiabatic lapse rate from the moist adiabatic lapse rate.
As latent heat is added from the process of condensation, it offsets some of the adiabatic cooling from expansion.
Because of this, the air parcel will no longer cool at the dry adiabatic lapse rate, but will cool as a slower rate, known as the moist adiabatic lapse rate. The effects of moisture change the lapse rate of the air parcel and, therefore, affects stability. However, the concepts are still the same and we still compare the air parcel temperature to the environmental temperature. We have just one added complication to worry about—we need to know whether the air parcel is dry or moist.
Some definitions are included below, which take into account both dry and moist adiabatic lapse rates. The atmosphere is said to be absolutely stable if the environmental lapse rate is less than the moist adiabatic lapse rate.
This means that a rising air parcel will always cool at a faster rate than the environment, even after it reaches saturation. If an air parcel is cooler at all levels, then it will not be able to rise, even after it becomes saturated when latent heating will counteract some cooling. The atmosphere is said to be absolutely unstable if the environmental lapse rate is greater than the dry adiabatic lapse rate.
This means that a rising air parcel will always cool at a slower rate than the environment, even when it is unsaturated. This means that it will be warmer and less dense than the environment, and allowed to rise. The atmosphere is said to be conditionally unstable if the environmental lapse rate is between the moist and dry adiabatic lapse rates. This means that the buoyancy the ability of an air parcel to rise of an air parcel depends on whether or not it is saturated.
In a conditionally unstable atmosphere, an air parcel will resist vertical motion when it is unsaturated, because it will cool faster than the environment at the dry adiabatic lapse rate. If it is forced to rise and is able to become saturated, however, it will cool at the moist adiabatic lapse rate. In this case, it will cool slower than the environment, become warmer than the environment, and will rise.
Around Hawaii, the atmosphere is almost always conditionally unstable, meaning that the environmental lapse rate lies somewhere between the dry and moist adiabatic lapse rates. For this reason, Hawaii almost always has convective clouds. Convective clouds are clouds where the edges are bumpy and cumuliform, like cauliflower. The clouds are convective because the atmosphere is stable to dry lifting and unstable to moist lifting. Once the air is saturated, instability sets in and vertical motion takes off.
This is especially common as air is lifted over our mountainous islands. The forced lifting from the terrain creates clouds and rain right over the mountains! In scientific terms, the initial lifting of the stable low level dry air by the terrain causes the air to adiabatically expand and reach saturation, at which point the environment is unstable to moist lifting and convection is the result.
There are many different types of thermodynamic diagrams, but the main one we will discuss are Skew-T Log-P diagrams, so-named because the isotherms lines of equal temperature, T on the diagram are slanted skewed and the isobars lines of equal pressure, P on the diagram are in log space. Here we will focus on how to read and utilize Skew-T Log-P diagrams often shortened to Skew-T diagram to determine parcel buoyancy and atmospheric stability.
You can see the vertical environmental temperature profile T plotted as the black jagged line on the right. The dew point temperature T d with height is plotted with the black jagged line on the left.
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