Determination of the height of the levels of condensation and sublimation. Vertical structure of the atmosphere How much temperature increases with height
In the first sections, we met in general terms with the vertical structure of the atmosphere and with changes in temperature with altitude.
Here we consider some interesting features temperature regime in the troposphere and in the overlying spheres.
Temperature and humidity in the troposphere. The troposphere is the most interesting area, since rock-forming processes are formed here. In the troposphere, as already mentioned in chapter I, the air temperature decreases with height by an average of 6° per kilometer rise, or by 0.6° per 100 m. This value of the vertical temperature gradient is observed most often and is defined as the average of many measurements. In fact, the vertical temperature gradient in temperate latitudes The earth is changeable. It depends on the seasons of the year, time of day, nature atmospheric processes, and in the lower layers of the troposphere - mainly on the temperature of the underlying surface.
IN warm time year, when the layer of air adjacent to the surface of the earth is sufficiently heated, a decrease in temperature with height is characteristic. With a strong heating of the surface layer of air, the value of the vertical temperature gradient exceeds even 1 ° for every 100 m uplift.
In winter, with a strong cooling of the surface of the earth and the surface layer of air, instead of lowering, an increase in temperature is observed with height, i.e., a temperature inversion occurs. The strongest and most powerful inversions are observed in Siberia, especially in Yakutia in winter, where clear and calm weather prevails, which contributes to the radiation and subsequent cooling of the surface air layer. Very often, the temperature inversion here extends to a height of 2-3 km, and the difference between the air temperature at the earth's surface and upper bound inversion is often 20-25°. Inversions are also characteristic of central regions Antarctica. In winter, they are in Europe, especially in its eastern part, Canada and other areas. The magnitude of the change in temperature with height (vertical temperature gradient) largely determines the weather conditions and types of air movement in the vertical direction.
Stable and unstable atmosphere. The air in the troposphere is heated by the underlying surface. Air temperature changes with height and with atmospheric pressure. When this happens without heat exchange with environment, then such a process is called adiabatic. Rising air does work at the expense of internal energy, which is spent on overcoming external resistance. Therefore, when it rises, the air cools, and when it descends, it heats up.
Adiabatic temperature changes occur according to dry adiabatic And wet adiabatic laws. Accordingly, vertical gradients of temperature change with height are also distinguished. Dry adiabatic gradient is the change in temperature of dry or moist unsaturated air for every 100 m raise and lower it by 1 °, A wet adiabatic gradient is the decrease in temperature of moist saturated air for every 100 m elevation less than 1°.
When dry, or unsaturated, air rises or falls, its temperature changes according to the dry adiabatic law, i.e., respectively, falls or rises by 1 ° every 100 m. This value does not change until the air, when rising, reaches a state of saturation, i.e. condensation level water vapor. Above this level, due to condensation, the latent heat of vaporization begins to be released, which is used to heat the air. This additional heat reduces the amount of air cooling as it rises. A further rise in saturated air occurs already according to the humid adiabatic law, and its temperature does not decrease by 1 ° per 100 m, but less. Since the moisture content of air depends on its temperature, the higher the air temperature, the more heat is released during condensation, and the lower the temperature, the less heat. Therefore, the humid adiabatic gradient in warm air is smaller than in cold air. For example, at a temperature of rising saturated air near the earth's surface of +20°, the humid adiabatic gradient in the lower troposphere is 0.33-0.43° per 100 m, and at a temperature of minus 20° its values range from 0.78° to 0.87° per 100m.
The wet adiabatic gradient also depends on the air pressure: the lower the air pressure, the smaller the wet adiabatic gradient at the same initial temperature. This is due to the fact that at low pressure, the air density is also less, therefore, the released heat of condensation is used to heat a smaller mass of air.
Table 15 shows the average values of the wet adiabatic gradient at various temperatures and values
pressure 1000, 750 and 500 mb, which approximately corresponds to the surface of the earth and heights of 2.5-5.5km.
In the warm season, the vertical temperature gradient averages 0.6-0.7° per 100 m uplift. Knowing the temperature at the surface of the earth, it is possible to calculate the approximate values of the temperature at various heights. If, for example, the air temperature at the earth's surface is 28°, then, assuming that the vertical temperature gradient is on average 0.7° per 100 m or 7° per kilometer, we get that at a height of 4 km the temperature is 0°. The temperature gradient in winter in the middle latitudes over land rarely exceeds 0.4-0.5 ° per 100 m: There are frequent cases when in separate layers of air the temperature almost does not change with height, i.e., isothermia takes place.
By the magnitude of the vertical air temperature gradient, one can judge the nature of the equilibrium of the atmosphere - stable or unstable.
At stable equilibrium atmospheric masses of air do not tend to move vertically. In this case, if a certain volume of air is shifted upwards, it will return to its original position.
Stable equilibrium occurs when the vertical temperature gradient of unsaturated air is less than the dry adiabatic gradient, and the vertical temperature gradient of saturated air is less than the wet adiabatic one. If, under this condition, a small volume of unsaturated air is raised by an external influence to a certain height, then as soon as the action stops external force, this volume of air will return to former position. This happens because the raised volume of air, having spent internal energy on its expansion, was cooled by 1 ° for every 100 m(according to the dry adiabatic law). But since the vertical temperature gradient of the ambient air was less than the dry adiabatic one, it turned out that the volume of air raised at a given height had a lower temperature than the ambient air. Having a greater density than the surrounding air, it must sink until it reaches its original state. Let's show this with an example.
Suppose that the air temperature near the earth's surface is 20°, and the vertical temperature gradient in the layer under consideration is 0.7° per 100 m. With this value of the gradient, the air temperature at a height of 2 km will be equal to 6° (Fig. 19, A). Under the influence of an external force, a volume of unsaturated or dry air raised from the surface of the earth to this height, cooling according to the dry adiabatic law, i.e., by 1 ° per 100 m, will cool by 20 ° and take a temperature equal to 0 °. This volume of air will turn out to be 6 ° colder than the surrounding air, and therefore heavier due to greater density. So he starts
descend, trying to reach the initial level, i.e., the surface of the earth.
A similar result will be obtained in the case of rising saturated air, if the vertical gradient of the ambient temperature is less than the humid adiabatic one. Therefore, when steady state atmosphere in a homogeneous mass of air, there is no rapid formation of cumulus and cumulonimbus clouds.
The most stable state of the atmosphere is observed at small values of the vertical temperature gradient, and especially during inversions, since in this case warmer and lighter air is located above the lower cold, and therefore heavy, air.
At unstable equilibrium of the atmosphere the volume of air raised from the earth's surface does not return to its original position, but retains its upward movement to a level at which the temperatures of the rising and surrounding air are equalized. The unstable state of the atmosphere is characterized by large vertical temperature gradients, which is caused by heating of the lower layers of air. At the same time, the air masses warmed up below, as lighter ones, rush upwards.
Suppose, for example, that unsaturated air in the lower layers up to a height of 2 km stratified unstable, i.e. its temperature
decreases with altitude by 1.2° for every 100 m, and above, the air, having become saturated, has a stable stratification, i.e., its temperature drops already by 0.6 ° for every 100 m uplifts (Fig. 19, b). Once in such an environment, the volume of dry unsaturated air will begin to rise according to the dry adiabatic law, i.e., it will cool by 1 ° per 100 m. Then, if its temperature near the earth's surface is 20°, then at a height of 1 km it will become 10°, while the ambient temperature is 8°. Being 2° warmer and therefore lighter, this volume will rush higher. At height 2 km it will be already 4° warmer than the environment, since its temperature will reach 0°, and the ambient temperature is -4°. Being lighter again, the considered volume of air will continue its rise to a height of 3 km, where its temperature will be equal to the temperature environment (-10°). After that, the free rise of the allocated air volume will stop.
To determine the state of the atmosphere are used aerological charts. These are diagrams with rectangular coordinate axes, along which the characteristics of the state of the air are plotted. Families are plotted on upper-air diagrams dry And wet adiabats, i.e., curves graphically representing the change in the state of air during dry adiabatic and wet adiabatic processes.
Figure 20 shows such a diagram. Here, isobars are shown vertically, isotherms (lines of equal air pressure) horizontally, inclined solid lines are dry adiabats, inclined dashed lines are wet adiabats, dashed lines are specific humidity. The above diagram shows curves of air temperature changes with a height of two points at the same observation period - 15:00 on May 3, 1965. On the left - the temperature curve according to the data of a radiosonde launched in Leningrad, on the right - in Tashkent. It follows from the shape of the left curve of temperature change with height that the air in Leningrad is stable. In this case, up to the isobaric surface of 500 mb the vertical temperature gradient averages 0.55° per 100 m. In two small layers (on surfaces 900 and 700 mb) isotherm was recorded. This indicates that over Leningrad at heights of 1.5-4.5 km located atmospheric front, separating the cold air masses in the lower one and a half kilometers from the thermal air located above. The height of the condensation level, determined by the position of the temperature curve with respect to the wet adiabat, is about 1 km(900 mb).
In Tashkent, the air had an unstable stratification. Up to height 4 km vertical temperature gradient was close to adiabatic, i.e., for every 100 m rise, the temperature decreased by 1 °, and higher, up to 12 km- more adiabatic. Due to the dryness of the air, cloud formation did not occur.
Over Leningrad, the transition to the stratosphere took place at an altitude of 9 km(300 mb), and over Tashkent it is much higher - about 12 km(200 mb).
With a stable state of the atmosphere and sufficient humidity, stratus clouds and fogs can form, and with an unstable state and a high moisture content of the atmosphere, thermal convection, leading to the formation of cumulus and cumulonimbus clouds. The state of instability is associated with the formation of showers, thunderstorms, hail, small whirlwinds, squalls, etc. The so-called "bumpiness" of the aircraft, i.e., the aircraft throws during flight, is also caused by the unstable state of the atmosphere.
In summer, the instability of the atmosphere is common in the afternoon, when the layers of air close to the earth's surface are heated. Therefore, heavy rains, squalls and similar dangerous weather phenomena are more often observed in the afternoon, when strong vertical currents arise due to breaking instability - ascending And descending air movement. For this reason, aircraft flying during the day at an altitude of 2-5 km above the surface of the earth, they are more subject to "chatter" than during night flight, when, due to the cooling of the surface layer of air, its stability increases.
Humidity also decreases with altitude. Almost half of all humidity is concentrated in the first one and a half kilometers of the atmosphere, and the first five kilometers contain almost 9/10 of all water vapor.
To illustrate the daily observed nature of the change in temperature with height in the troposphere and lower stratosphere in different regions of the Earth, Figure 21 shows three stratification curves up to a height of 22-25 km. These curves were built based on radiosonde observations at 3 pm: two in January - Olekminsk (Yakutia) and Leningrad, and the third in July - Takhta-Bazar (Central Asia). The first curve (Olekminsk) is characterized by the presence of a surface inversion, characterized by an increase in temperature from -48° at the earth's surface to -25° at a height of about 1 km. During this period, the tropopause over Olekminsk was at a height of 9 km(temperature -62°). In the stratosphere, an increase in temperature with height was observed, the value of which is at the level of 22 km approached -50°. The second curve, representing the change in temperature with height in Leningrad, indicates the presence of a small surface inversion, then an isotherm in a large layer and a decrease in temperature in the stratosphere. At level 25 km the temperature is -75°. The third curve (Takhta-Bazar) is very different from the northern point - Olekminsk. The temperature at the earth's surface is above 30°. The tropopause is at 16 km, and above 18 km there is an increase in temperature with altitude, which is usual for a southern summer.
- Source-
Pogosyan, Kh.P. Atmosphere of the Earth / Kh.P. Poghosyan [and d.b.]. - M .: Education, 1970. - 318 p.
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The rays of the Sun, when passing through transparent substances, heat them very weakly. This is due to the fact that direct sunlight practically does not heat the atmospheric air, but strongly heats the earth's surface, capable of transmitting thermal energy adjacent layers of air. As it warms, the air becomes lighter and rises higher. In the upper layers, warm air mixes with cold air, giving it some of the heat energy.
The higher the heated air rises, the more it cools. The air temperature at an altitude of 10 km is constant and is -40-45 °C.
A characteristic feature of the Earth's atmosphere is a decrease in air temperature with height. Sometimes there is an increase in temperature as altitude increases. The name of this phenomenon is temperature inversion(temperature change).
Temperature change
The appearance of inversions may be due to the cooling of the earth's surface and the adjacent air layer in a short period of time. This is also possible when dense cold air moves from mountain slopes to valleys. During the day, the air temperature changes continuously. During the daytime, the earth's surface heats up and heats the lower layer of air. At night, along with the cooling of the earth, the air cools. It is coolest at dawn and warmest in the afternoon.
IN equatorial belt there is no diurnal temperature fluctuation. Night and day temperatures are the same. Diurnal amplitudes on the coasts of the seas, oceans and above their surface are insignificant. But in the desert zone, the difference between night and day temperatures can reach 50-60 ° C.
In the temperate zone maximum amount Solar radiation on Earth falls on the days of the summer solstices. But the hottest month is July in the Northern Hemisphere and January in the Southern. This is because, despite the fact that solar radiation less intense during these months, great amount thermal energy is given off by a very hot earth's surface.
The annual temperature amplitude is determined by the latitude of a certain area. For example, at the equator it is constant and is 22-23 ° C. The highest annual amplitudes are observed in the regions of middle latitudes and in the depths of the continents.
Absolute and average temperatures are also characteristic of any area. Absolute temperatures determined by long-term observations at weather stations. The hottest area on Earth is the Libyan Desert (+58°C), and the coldest is Vostok Station in Antarctica (-89.2°C).
Average temperatures are set when calculating the arithmetic mean of several thermometer readings. This is how average daily, average monthly and average annual temperatures are determined.
In order to find out how heat is distributed on Earth, temperature values \u200b\u200bare plotted on a map and connect points with the same values. The resulting lines are called isotherms. This method allows you to identify certain patterns in the distribution of temperatures. Yes, most high temperatures are recorded not at the equator, but in tropical and subtropical deserts. A decrease in temperatures from the tropics to the poles in two hemispheres is characteristic. Given that in the Southern Hemisphere, water bodies occupy a larger area than land, the temperature amplitudes between the hottest and coldest months are less pronounced there than in the Northern Hemisphere.
According to the location of the isotherms, seven thermal zones are distinguished: 1 hot, 2 moderate, 2 cold, 2 permafrost areas.
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In the troposphere, the air temperature decreases with height, as noted, by an average of 0.6 ° C for every 100 m of altitude. However, in the surface layer, the temperature distribution can be different: it can decrease or increase, and remain constant. temperature with height gives the vertical temperature gradient (VGT):
VGT = (/ „ - /B)/(ZB -
where /n - /v - temperature difference at the lower and upper levels, ° С; ZB - ZH- height difference, m. Usually, the VGT is calculated for 100 m of height.
In the surface layer of the atmosphere, the VGT can be 1000 times higher than the average for the troposphere
The value of the VGT in the surface layer depends on weather conditions(in clear weather it is more than in cloudy), time of year (more in summer than in winter) and time of day (more during the day than at night). The wind reduces the VGT, since when the air is mixed, its temperature is equalized at different heights. Above moist soil, WGT sharply decreases in the surface layer, and over bare soil (fallow field) WGT is greater than over dense crops or meadows. This is due to differences in the temperature regime of these surfaces (see Chap. 3).
As a result of a certain combination of these factors, the VGT near the surface in terms of 100 m of height can be more than 100 ° C / 100 m. In such cases, thermal convection occurs.
The change in air temperature with altitude determines the sign of the UGT: if the UGT > 0, then the temperature decreases with distance from the active surface, which usually happens during the day and in summer (Fig. 4.4); if VGT = 0, then the temperature does not change with height; if VGT< 0, то температура увеличивается с высотой и такое распределение температуры называют инверсией.
Depending on the conditions for the formation of inversions in the surface layer of the atmosphere, they are divided into radiative and advective.
1. Radiative inversions occur during radiative cooling of the earth's surface. Such inversions during the warm period of the year are formed at night, and in winter they are also observed during the day. Therefore, radiative inversions are divided into night (summer) and winter ones.
Night inversions are set in clear calm weather after the transition of the radiation balance through 0 for 1.0...1.5 hours before sunset. During the night, they intensify and reach their maximum power before sunrise. After sunrise, the active surface and the air warm up, which destroys the inversion. The height of the inversion layer is most often several tens of meters, but under certain conditions (for example, in closed valleys surrounded by significant elevations) it can reach 200 m or more. This is facilitated by the flow of cooled air from the slopes into the valley. Cloudiness weakens the inversion, and the wind speed of more than 2.5...3.0 m/s destroys it. Under the canopy of dense herbage, crops, as well as forests in summer, inversions are also observed during the day.
Night radiation inversions in spring and autumn, and in some places in summer, can cause a decrease in soil and air surface temperatures to negative values(frost), which causes damage to many cultivated plants.
Winter inversions occur in clear, calm weather under conditions short day when the cooling of the active surface continuously increases every day; they can persist for several weeks, weakening a little during the day and increasing again at night.
The radiative inversions are especially intensified with a sharply inhomogeneous terrain. Cooling air flows down into depressions and basins, where weakened turbulent mixing contributes to its further cooling. Radiative inversions associated with the features of the terrain are usually called orographic.
2. Advective inversions are formed during the advection (movement) of warm air onto a cold underlying surface, which cools the layers of advancing air adjacent to it. These inversions also include snow inversions. They arise during the advection of air having a temperature above 0 "C onto a surface covered with snow. A decrease in temperature in the lowest layer in this case is associated with heat costs for melting snow.
INDICATORS OF THE TEMPERATURE REGIME IN THIS AREA AND THE NEEDS OF PLANTS FOR HEAT
When evaluating temperature regime large area or a separate point, temperature characteristics are used for a year or for separate periods (vegetation period, season, month, decade and day). The main of these indicators are as follows.
The average daily temperature is the arithmetic mean of the temperatures measured during all periods of observation. At meteorological stations Russian Federation air temperature is measured eight times a day. Summing up the results of these measurements and dividing the sum by 8, the average daily air temperature is obtained.
The average monthly temperature is the arithmetic average of the average daily temperatures for the entire day of the month.
The mean annual temperature is the arithmetic mean of the mean daily (or mean monthly) temperatures for the entire year.
The average code air temperature gives only a general idea of the amount of heat, it does not characterize annual course temperature. So, the average annual temperature in the south of Ireland and in the steppes of Kalmykia, located at the same latitude, is close (9 ° C). But in Ireland, the average January temperature is 5 ... 8 "C, and the meadows are green all winter, and in the steppes of Kalmykia, the average January temperature is -5 ... -8 ° C. In summer, it is cool in Ireland: 14 ° C, and the average July temperature in Kalmykia is 23...26 °C.
Therefore, for more complete characteristics annual temperature variation in this place use data on the average temperature of the coldest (January) and warmest (July) months.
However, all the averaged characteristics do not give an accurate idea of the daily and annual course of temperature, i.e., just about the conditions that are especially important for agricultural production. In addition to the average temperatures are the maximum and minimum temperatures, amplitude. For example, knowing minimum temperature in the winter months, one can judge the conditions for overwintering of winter crops and fruit and berry plantations. The maximum temperature data shows the frequency and intensity of thaws in winter, and the number of hot days in summer when grain damage is possible during the filling period, etc.
In extreme temperatures, there are: absolute maximum (minimum) - the highest (lowest) temperature for the entire observation period; average of absolute maximums (minimums) - arithmetic average of absolute extremes; average maximum (minimum) - the arithmetic average of all extreme temperatures, for example, for a month, season, year. However, they can be calculated as multi-year period observations, as well as for the actual month, year, etc.
The amplitude of the daily and annual temperature variation characterizes the degree of continental climate: the greater the amplitude, the more continental the climate.
A characteristic of the temperature regime in a given area for a certain period is also the sum of average daily temperatures above or below a certain limit. For example, in climate reference books and atlases, the sums of temperatures are given above 0, 5, 10 and 15 ° C, as well as below -5 and -10 "C.
A visual representation of the geographical distribution of temperature regime indicators is provided by maps on which isotherms are drawn - lines of equal temperature values or sums of temperatures (Fig. 4.7). Maps, for example, of the sums of temperatures are used to justify the placement of crops (plantings) of cultivated plants with different requirements for heat.
To clarify the thermal conditions necessary for plants, the sums of day and night temperatures are also used, since the average daily temperature and its sums level out thermal differences in the daily course of air temperature.
The study of the thermal regime separately for day and night has a deep physiological significance. It is known that all processes occurring in the plant and animal world are subject to natural rhythms determined by external conditions, that is, they are subject to the law of the so-called "biological" clock. For example, according to (1964), for optimal growth conditions tropical plants the difference between day and night temperatures should be 3 ... 5 ° C, for plants temperate zone-5...7, and for desert plants - 8 °С and more. The study of day and night temperatures acquires a special meaning for increasing the productivity of agricultural plants, which is determined by the ratio of two processes - assimilation and respiration, occurring in qualitatively different light and dark hours of the day for plants.
The average daytime and nighttime temperatures and their sums indirectly take into account the latitudinal variability in the length of the day and night, as well as changes in the continentality of the climate and the influence of various landforms on the temperature regime.
The sums of average daily air temperatures, which are close for a pair of meteorological stations located approximately at the same latitude, but differ significantly in longitude, i.e., located in different conditions of continental climate, are shown in Table 4.1.
In more continental eastern regions sums of daytime temperatures are 200...500°C greater, and sums of nighttime temperatures are 300°C less than in western and especially sea areas, which explains for a long time known fact- accelerating the development of agricultural crops in a sharply continental climate.
The need of plants for heat is expressed by the sums of active and effective temperatures. In agricultural meteorology, active temperature is the average daily air (or soil) temperature above the biological minimum of crop development. The effective temperature is the average daily air (or soil) temperature, reduced by the value of the biological minimum.
Plants develop only if the average daily temperature exceeds their biological minimum, which is, for example, 5 ° C for spring wheat, 10 ° C for corn, and 13 ° C for cotton (15 ° C for southern varieties of cotton). The sums of active and effective temperatures have been established both for individual interphase periods and for the entire growing season of many varieties and hybrids of major crops (Table 11.1).
Through the sums of active and effective temperatures, the need for heat of poikilothermic (cold-blooded) organisms is also expressed both for the ontogenetic period and for centuries. the biological cycle.
When calculating the sums of average daily temperatures characterizing the heat demand of plants and poikilothermic organisms, it is necessary to introduce a correction for ballast temperatures that do not accelerate growth and development, i.e., take into account the upper temperature level for crops and organisms. For most plants and pests temperate zone this will be the average daily temperature exceeding 20 ... 25 "C.
13. Change in air temperature with height.
The vertical distribution of temperature in the atmosphere is the basis for dividing the atmosphere into five main layers. For agricultural meteorology, the regularities of temperature changes in the troposphere, especially in its surface layer, are of the greatest interest.
Vertical temperature gradient
A change in air temperature per 100 m of altitude is called a vertical temperature gradient (VGT depends on a number of factors: the season (it is less in winter, more in summer), the time of day (less at night, more during the day), the location of air masses (if at any heights above a layer of warmer air is located in a cold layer of air, then the UGT reverses its sign.) The average value of the VGT in the troposphere is about 0.6 ° C / 100 m.
In the surface layer of the atmosphere, the VGT depends on the time of day, the weather, and the nature of the underlying surface. In the daytime, VGT is almost always positive, especially in summer over land, but in clear weather it is ten times greater than in cloudy weather. On a clear noon in summer, the air temperature near the soil surface can be 10 °C or more higher than the temperature at a height of 2 m. As a result, the WGT in this two-meter layer in terms of 100 m is more than 500°C/100 m. The wind reduces the WGT, since at When the air is mixed, its temperature at different heights is equalized. Reduce VGT cloudiness and precipitation. With moist soil, the WGT sharply decreases in the surface layer of the atmosphere. Above bare soil (fallow field), the VGT is greater than over a developed crop or meadow. In winter, above the snow cover, the VGT in the surface layer of the atmosphere is small and often negative.
With height, the influence of the underlying surface and the weather on the VGT weakens, and the VGT decreases compared to its values in the surface air layer. Above 500 m, the influence of the diurnal variations in air temperature is attenuated. At altitudes from 1.5 to 5-6 km, the UGT is in the range of 0.5-0.6 ° С / 100 m. At an altitude of 6-9 km, the VGT increases and amounts to 0.65-0.75 ° С / 100 m. In the upper troposphere, the VGT again decreases to 0.5–0.2°C/100 m.
Data on VGT in various layers of the atmosphere are used in weather forecasting, in meteorological services for jet aircraft and in launching satellites into orbit, as well as in determining the conditions for the release and distribution of industrial waste in the atmosphere. Negative VGT in the surface air layer at night in spring and autumn indicates the possibility of freezing.
17. Characteristics of air humidity. Daily and annual course of partial pressure of water vapor and relative humidity.
The elasticity of water vapor in the atmosphere - the partial pressure of water vapor in the air
The Earth's atmosphere contains about 14 thousand km 3 of water vapor. Water enters the atmosphere as a result of evaporation from the underlying surface. In the atmosphere, moisture condenses, moves by air currents and again falls in the form of various precipitations on the surface of the Earth, thus making a constant cycle of water. The water cycle is possible due to the ability of water to be in three states (liquid, solid, gaseous (vapor)) and easily move from one state to another. Moisture circulation is one of the most important cycles of climate formation.
To quantify the content of water vapor in the atmosphere, various characteristics of air humidity are used. The main characteristics of air humidity are water vapor pressure and relative humidity.
Elasticity (actual) of water vapor (e) - the pressure of water vapor in the atmosphere is expressed in mm Hg. or in millibars (mb). Numerically, it almost coincides with absolute humidity (the content of water vapor in the air in g / m 3), therefore elasticity is often called absolute humidity. Saturation elasticity (maximum elasticity) (E) - the limit of water vapor content in the air at a given temperature. The value of saturation elasticity depends on the air temperature, the higher the temperature, the more it can contain water vapor.
The daily course of humidity (absolute) can be simple and double. The first one coincides with the daily temperature variation, has one maximum and one minimum, and is typical for places with a sufficient amount of moisture. It is observed over the oceans, and over land in winter and autumn.
The double course has two maxima and two minima and is typical for the summer season on land: maxima at 09:00 and 20-21:00, and minimums at 06:00 and 16:00.
The morning minimum before sunrise is explained by weak evaporation during the night hours. With an increase in radiant energy, evaporation increases, the elasticity of water vapor reaches a maximum at about 9 hours.
As a result of surface heating, air convection develops, moisture transfer occurs faster than its inflow from the evaporating surface, therefore, a second minimum occurs at about 16 hours. By evening, convection stops, and evaporation from the heated surface is still quite intense and moisture accumulates in the lower layers, providing the second maximum at about 20-21 hours.
The annual course of water vapor elasticity corresponds to the annual course of temperature. In summer, the elasticity of water vapor is greater, in winter - less.
The daily and annual course of relative humidity is almost everywhere opposite to the course of temperature, since the maximum moisture content increases with increasing temperature faster than the elasticity of water vapor. The daily maximum of relative humidity occurs before sunrise, the minimum - at 15-16 hours.
During the year, the maximum relative humidity, as a rule, falls on the most cold month, at least for the warmest month. The exceptions are regions in which moist winds blow from the sea in summer, and dry ones from the mainland in winter.
Absolute humidity = amount of water in a given volume of air, measured in (g/m³)
Relative Humidity = The percentage of the actual amount of water (water vapor pressure) to the vapor pressure of water at that temperature under saturation conditions. Expressed as a percentage. Those. 40% humidity means that at this temperature all the water can evaporate another 60%.
inversion
increase in air temperature with height instead of the usual decrease
Alternative descriptions
An excited state of matter in which the number of particles at a higher energy. level exceeds the number of particles at a lower level (physics)
Change of direction magnetic field Earth on the reverse, observed at time intervals from 500 thousand years to 50 million years
Changing the normal position of elements, placing them in reverse order
Linguistic term for changing the usual word order in a sentence
Reverse order, reverse order
Logical operation "not"
Chromosomal rearrangement associated with the rotation of individual sections of the chromosome by 180
Conformal transformation of the Euclidean plane or space
Permutation in mathematics
A dramatic device that demonstrates the outcome of the conflict at the beginning of the play
In metrology - abnormal change any parameter
The state of matter in which high levels the energies of its constituent particles are more "populated" by particles than the lower
In organic chemistry, the process of breaking down a saccharide
Changing the order of words in a sentence
Changing word order for emphasis
white trail behind the plane
Changing word order
Reverse order of elements
Changing the normal order of words in a sentence in order to enhance the expressiveness of speech
In the first sections, we got acquainted in general terms with the structure of the atmosphere along the vertical and with changes in temperature with height.
Here we consider some interesting features of the temperature regime in the troposphere and in the overlying spheres.
Temperature and humidity in the troposphere. The troposphere is the most interesting area, since rock-forming processes are formed here. In the troposphere, as already mentioned in Chapter I, the air temperature decreases with height by an average of 6° per kilometer of elevation, or by 0.6° per 100 m. This value of the vertical temperature gradient is observed most often and is defined as the average of many measurements. In fact, the vertical temperature gradient in the temperate latitudes of the Earth is variable. It depends on the seasons of the year, the time of day, the nature of atmospheric processes, and in the lower layers of the troposphere - mainly on the temperature of the underlying surface.
In the warm season, when the layer of air adjacent to the surface of the earth is sufficiently heated, a decrease in temperature with height is characteristic. With a strong heating of the surface layer of air, the value of the vertical temperature gradient exceeds even 1 ° for every 100 m uplift.
In winter, with a strong cooling of the surface of the earth and the surface layer of air, instead of lowering, an increase in temperature is observed with height, i.e., a temperature inversion occurs. The strongest and most powerful inversions are observed in Siberia, especially in Yakutia in winter, where clear and calm weather prevails, which contributes to the radiation and subsequent cooling of the surface air layer. Very often, the temperature inversion here extends to a height of 2-3 km, and the difference between the air temperature at the earth's surface and the upper boundary of the inversion is often 20-25°. Inversions are also characteristic of the central regions of Antarctica. In winter, they are in Europe, especially in its eastern part, Canada and other areas. The magnitude of the change in temperature with height (vertical temperature gradient) largely determines the weather conditions and types of air movement in the vertical direction.
Stable and unstable atmosphere. The air in the troposphere is heated by the underlying surface. Air temperature changes with altitude and with atmospheric pressure. When this happens without heat exchange with the environment, then such a process is called adiabatic. Rising air does work at the expense of internal energy, which is spent on overcoming external resistance. Therefore, when it rises, the air cools, and when it descends, it heats up.
Adiabatic temperature changes occur according to dry adiabatic And wet adiabatic laws.
Accordingly, vertical gradients of temperature change with height are also distinguished. Dry adiabatic gradient is the change in temperature of dry or moist unsaturated air for every 100 m raise and lower it by 1 °, A wet adiabatic gradient is the decrease in temperature of moist saturated air for every 100 m elevation less than 1°.
When dry, or unsaturated, air rises or falls, its temperature changes according to the dry adiabatic law, i.e., respectively, falls or rises by 1 ° every 100 m. This value does not change until the air, when rising, reaches a state of saturation, i.e. condensation level water vapor. Above this level, due to condensation, the latent heat of vaporization begins to be released, which is used to heat the air. This additional heat reduces the amount of air cooling as it rises. A further rise in saturated air occurs already according to the humid adiabatic law, and its temperature does not decrease by 1 ° per 100 m, but less. Since the moisture content of air depends on its temperature, the higher the air temperature, the more heat is released during condensation, and the lower the temperature, the less heat. Therefore, the humid adiabatic gradient in warm air is smaller than in cold air. For example, at a temperature of rising saturated air near the earth's surface of +20°, the humid adiabatic gradient in the lower troposphere is 0.33-0.43° per 100 m, and at a temperature of minus 20° its values range from 0.78° to 0.87° per 100 m.
The wet adiabatic gradient also depends on the air pressure: the lower the air pressure, the smaller the wet adiabatic gradient at the same initial temperature. This is due to the fact that at low pressure, the air density is also less, therefore, the released heat of condensation is used to heat a smaller mass of air.
Table 15 shows the average values of the wet adiabatic gradient at various temperatures and values
pressure 1000, 750 and 500 mb, which approximately corresponds to the surface of the earth and heights of 2.5-5.5 km.
In the warm season, the vertical temperature gradient averages 0.6-0.7° per 100 m uplift.
Knowing the temperature at the surface of the earth, it is possible to calculate the approximate values of the temperature at various heights. If, for example, the air temperature at the earth's surface is 28°, then, assuming that the vertical temperature gradient is on average 0.7° per 100 m or 7° per kilometer, we get that at a height of 4 km the temperature is 0°. The temperature gradient in winter in the middle latitudes over land rarely exceeds 0.4-0.5 ° per 100 m: There are frequent cases when in separate layers of air the temperature almost does not change with height, i.e., isothermia takes place.
By the magnitude of the vertical air temperature gradient, one can judge the nature of the equilibrium of the atmosphere - stable or unstable.
At stable equilibrium atmospheric masses of air do not tend to move vertically. In this case, if a certain volume of air is shifted upwards, it will return to its original position.
Stable equilibrium occurs when the vertical temperature gradient of unsaturated air is less than the dry adiabatic gradient, and the vertical temperature gradient of saturated air is less than the wet adiabatic one. If, under this condition, a small volume of unsaturated air is raised by an external action to a certain height, then as soon as the action of the external force ceases, this volume of air will return to its previous position. This happens because the raised volume of air, having spent internal energy on its expansion, was cooled by 1 ° for every 100 m(according to the dry adiabatic law). But since the vertical temperature gradient of the ambient air was less than the dry adiabatic one, it turned out that the volume of air raised at a given height had a lower temperature than the ambient air. Having a greater density than the surrounding air, it must sink until it reaches its original state. Let's show this with an example.
Suppose that the air temperature near the earth's surface is 20°, and the vertical temperature gradient in the layer under consideration is 0.7° per 100 m. With this value of the gradient, the air temperature at a height of 2 km will be equal to 6° (Fig. 19, A). Under the influence of an external force, a volume of unsaturated or dry air raised from the surface of the earth to this height, cooling according to the dry adiabatic law, i.e., by 1 ° per 100 m, will cool by 20 ° and take a temperature equal to 0 °. This volume of air will be 6° colder than the surrounding air, and therefore heavier due to its greater density. So he starts
descend, trying to reach the initial level, i.e., the surface of the earth.
A similar result will be obtained in the case of rising saturated air, if the vertical gradient of the ambient temperature is less than the humid adiabatic one. Therefore, under a stable state of the atmosphere in a homogeneous mass of air, there is no rapid formation of cumulus and cumulonimbus clouds.
The most stable state of the atmosphere is observed at small values of the vertical temperature gradient, and especially during inversions, since in this case warmer and lighter air is located above the lower cold, and therefore heavy, air.
At unstable equilibrium of the atmosphere the volume of air raised from the earth's surface does not return to its original position, but retains its upward movement to a level at which the temperatures of the rising and surrounding air are equalized. The unstable state of the atmosphere is characterized by large vertical temperature gradients, which is caused by heating of the lower layers of air. At the same time, the air masses warmed up below, as lighter ones, rush upwards.
Suppose, for example, that unsaturated air in the lower layers up to a height of 2 km stratified unstable, i.e. its temperature
decreases with altitude by 1.2° for every 100 m, and above, the air, having become saturated, has a stable stratification, i.e., its temperature drops already by 0.6 ° for every 100 m uplifts (Fig. 19, b). Once in such an environment, the volume of dry unsaturated air will begin to rise according to the dry adiabatic law, i.e., it will cool by 1 ° per 100 m. Then, if its temperature near the earth's surface is 20°, then at a height of 1 km it will become 10°, while the ambient temperature is 8°. Being 2° warmer and therefore lighter, this volume will rush higher. At height 2 km it will be already 4° warmer than the environment, since its temperature will reach 0°, and the ambient temperature is -4°. Being lighter again, the considered volume of air will continue its rise to a height of 3 km, where its temperature becomes equal to the ambient temperature (-10 °). After that, the free rise of the allocated air volume will stop.
To determine the state of the atmosphere are used aerological charts. These are diagrams with rectangular coordinate axes, along which the characteristics of the state of the air are plotted.
Families are plotted on upper-air diagrams dry And wet adiabats, i.e., curves graphically representing the change in the state of air during dry adiabatic and wet adiabatic processes.
Figure 20 shows such a diagram. Here, isobars are shown vertically, isotherms (lines of equal air pressure) horizontally, inclined solid lines are dry adiabats, inclined dashed lines are wet adiabats, dashed lines are specific humidity.The above diagram shows curves of air temperature changes with a height of two points for the same observation period - 15:00 on May 3, 1965. On the left - the temperature curve according to the data of a radiosonde launched in Leningrad, on the right - in Tashkent. It follows from the shape of the left curve of temperature change with height that the air in Leningrad is stable. In this case, up to the isobaric surface of 500 mb the vertical temperature gradient averages 0.55° per 100 m. In two small layers (on surfaces 900 and 700 mb) isotherm was recorded. This indicates that over Leningrad at heights of 1.5-4.5 km there is an atmospheric front that separates the cold air masses in the lower one and a half kilometers from the thermal air located above. The height of the condensation level, determined by the position of the temperature curve with respect to the wet adiabat, is about 1 km(900 mb).
In Tashkent, the air had an unstable stratification. Up to height 4 km vertical temperature gradient was close to adiabatic, i.e., for every 100 m rise, the temperature decreased by 1 °, and higher, up to 12 km- more adiabatic. Due to the dryness of the air, cloud formation did not occur.
Over Leningrad, the transition to the stratosphere took place at an altitude of 9 km(300 mb), and over Tashkent it is much higher - about 12 km(200 mb).
With a stable state of the atmosphere and sufficient humidity, stratus clouds and fogs can form, and with an unstable state and a high moisture content of the atmosphere, thermal convection, leading to the formation of cumulus and cumulonimbus clouds. The state of instability is associated with the formation of showers, thunderstorms, hail, small whirlwinds, squalls, etc.
The so-called "chatter" of the aircraft, i.e., the throws of the aircraft during flight, is also caused by the unstable state of the atmosphere.
In summer, the instability of the atmosphere is common in the afternoon, when the layers of air close to the earth's surface are heated. Therefore, heavy rains, squalls and similar dangerous weather phenomena are more often observed in the afternoon, when strong vertical currents arise due to breaking instability - ascending And descending air movement. For this reason, aircraft flying during the day at an altitude of 2-5 km above the surface of the earth, they are more subject to "chatter" than during night flight, when, due to the cooling of the surface layer of air, its stability increases.
Humidity also decreases with altitude. Almost half of all humidity is concentrated in the first one and a half kilometers of the atmosphere, and the first five kilometers contain almost 9/10 of all water vapor.
To illustrate the daily observed nature of the change in temperature with height in the troposphere and lower stratosphere in different regions of the Earth, Figure 21 shows three stratification curves up to a height of 22-25 km. These curves were built based on radiosonde observations at 3 pm: two in January - Olekminsk (Yakutia) and Leningrad, and the third in July - Takhta-Bazar (Central Asia). The first curve (Olekminsk) is characterized by the presence of a surface inversion, characterized by an increase in temperature from -48° at the earth's surface to -25° at a height of about 1 km. During this period, the tropopause over Olekminsk was at a height of 9 km(temperature -62°). In the stratosphere, an increase in temperature with height was observed, the value of which is at the level of 22 km approached -50°. The second curve, representing the change in temperature with height in Leningrad, indicates the presence of a small surface inversion, then an isotherm in a large layer and a decrease in temperature in the stratosphere. At level 25 km the temperature is -75°. The third curve (Takhta-Bazar) is very different from the northern point - Olekminsk. The temperature at the earth's surface is above 30°. The tropopause is at 16 km, and above 18 km there is an increase in temperature with altitude, which is usual for a southern summer.
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The sun's rays falling on the surface of the earth heat it up. The air is heated from the bottom up, i.e. from the earth's surface.
The transfer of heat from the lower layers of air to the upper ones occurs mainly due to the rise of warm, heated air up and the lowering of cold air down. This process of heating air is called convection.
In other cases, the upward heat transfer occurs due to dynamic turbulence. This is the name of chaotic whirlwinds that arise in the air as a result of its friction against the earth's surface during horizontal movement or during the friction of different layers of air with each other.
Convection is sometimes called thermal turbulence. Convection and turbulence are sometimes combined common name - exchange.
The cooling of the lower layers of the atmosphere occurs differently than heating. earth surface continuously loses heat to the surrounding atmosphere by emitting heat rays that are not visible to the eye. Cooling becomes especially strong after sunset (at night). Due to thermal conductivity, the air masses adjacent to the ground also gradually cool, transferring this cooling to the overlying layers of air; at the same time, the lowest layers are most intensively cooled.
Depending on solar heating, the temperature of the lower layers of air changes during the year and day, reaching a maximum at about 13-14 hours. daily course air temperature in different days for one and the same place is inconsistent; its value depends mainly on the state of the weather. Thus, changes in the temperature of the lower layers of air are associated with changes in the temperature of the earth's (underlying) surface.
Changes in air temperature also occur from its vertical movements.
It is known that when air expands, it cools, and when compressed, it heats up. In the atmosphere during the upward movement of air, falling into areas of more low pressure, expands and cools, and, conversely, with a downward movement, the air, compressing, heats up. Changes in air temperature during its vertical movements largely determine the formation and destruction of clouds.
Air temperature usually decreases with altitude. Change average temperature with the height above Europe in summer and winter is given in the table "Average air temperatures over Europe". 
The decrease in temperature with height is characterized by a vertical temperature gradient. This is the change in temperature for every 100 m of altitude. For technical and aeronautical calculations, the vertical temperature gradient is assumed to be 0.6. It must be borne in mind that this value is not constant. It may happen that in any layer of air the temperature will not change with height.
Such layers are called layers of isotherm.
Quite often, a phenomenon is observed in the atmosphere when, in a certain layer, the temperature even increases with height. These layers of the atmosphere are called inversion layers. Inversions arise from various reasons. One of them is the cooling of the underlying surface by radiation at night or winter time at clear sky. Sometimes, in the case of calm or light winds, the surface layers of air also cool and become colder than the overlying layers. As a result, the air at altitude is warmer than at the bottom. Such inversions are called radiation. Strong radiative inversions are usually observed over the snow cover and especially in mountain basins, and also during calm. The inversion layers extend up to a height of several tens or hundreds of meters.
Inversions also arise due to the movement (advection) of warm air onto the cold underlying surface. These are the so-called advective inversions. The height of these inversions is several hundred meters.
In addition to these inversions, frontal inversions and compression inversions are observed. Frontal inversions occur when warm air masses flow onto colder air masses. Compression inversions occur when air descends from the upper atmosphere. At the same time, the descending air is sometimes heated so much that its underlying layers turn out to be colder.
Temperature inversions are observed at various heights of the troposphere, most often at altitudes of about 1 km. The thickness of the inversion layer can vary from several tens to several hundreds of meters. The temperature difference during inversion can reach 15-20°.
Inversion layers play a big role in the weather. Because the air in the inversion layer is warmer than the underlying layer, the air from the lower layers cannot rise. Consequently, layers of inversions retard vertical movements in the underlying air layer. When flying under a layer of inversion, a rheme ("bumpiness") is usually observed. Above the inversion layer, the flight of the aircraft usually proceeds normally. So-called wavy clouds develop under the layers of inversions.
The air temperature affects the piloting technique and the operation of the materiel. At temperatures near the ground below -20 °, the oil freezes, so it has to be filled in in a heated state. In flight at low temperatures the water in the cooling system of the motor is intensively cooled. At elevated temperatures (above + 30 °), the motor may overheat. Air temperature also affects the performance of the aircraft crew. At low temperatures, reaching up to -56 ° in the stratosphere, special uniforms are required for the crew.
The air temperature is very great importance for weather forecast.
Measurement of air temperature during the flight on an aircraft is carried out using electric thermometers attached to the aircraft. When measuring air temperature, it must be borne in mind that due to high speeds modern aircraft thermometers give errors. The high speeds of the aircraft cause an increase in the temperature of the thermometer itself, due to the friction of its reservoir against the air and the effect of heating due to air compression. Friction heating increases with an increase in aircraft flight speed and is expressed by the following quantities:
Speed in km/h …………. 100 200 Z00 400 500 600
Friction heating ……. 0°.34 1°.37 3°.1 5°.5 8°.6 12°,b
Heating from compression is expressed by the following quantities:
Speed in km/h …………. 100 200 300 400 500 600
Heating by compression ……. 0°.39 1°.55 3°.5 5°.2 9°.7 14°.0
Distortions in the readings of a thermometer installed on an airplane, when flying in clouds, are 30% less than the above values, due to the fact that part of the heat that occurs during friction and compression is spent on the evaporation of water condensed in the air in the form of droplets.
Air temperature. Units of measure, change in temperature with height. Inversion, isothermy, Types of inversions, Adiabatic process.
Air temperature is a value that characterizes its thermal state. It is expressed either in degrees Celsius (ºС on a centigrade scale or in Kelvins (K) on absolute scale. The transition from temperature in Kelvin to temperature in degrees Celsius is performed by the formula
t=T-273º
The lower layer of the atmosphere (troposphere) is characterized by a decrease in temperature with height, amounting to 0.65ºС per 100m.
This change in temperature with height per 100m is called the vertical temperature gradient. Knowing the temperature near the earth's surface and using the value of the vertical gradient, it is possible to calculate the approximate temperature at any height (for example, at a temperature near the earth's surface of +20ºС at a height of 5000m, the temperature will be equal to:
20º- (0.65 * 50) \u003d - 12..5.
The vertical gradient γ is not constant and depends on the type air mass, time of day and season of the year, the nature of the underlying surface and other reasons. When the temperature decreases with height, γ is considered positive, if the temperature does not change with height, then γ = 0 the layers are called isothermal. Atmospheric layers where the temperature rises with height (γ< 0), называются inversion. Depending on the magnitude of the vertical temperature gradient, the state of the atmosphere can be stable, unstable, or indifferent to dry (not saturated) or saturated air.
The decrease in air temperature as it rises adiabatically, that is, without heat exchange of air particles with the environment. If an air particle rises, then its volume expands, while the internal energy of the particle decreases.
As the particle descends, it contracts and its internal energy increases. From this it follows that with an upward movement of the volume of air, its temperature decreases, and with a downward movement, it rises. These processes play important role in the formation and development of clouds.
The horizontal gradient is the temperature expressed in degrees at a distance of 100 km. During the transition from cold to warm VM and from warm to cold, it can exceed 10º per 100 km.
Types of inversions.
Inversions are delay layers, they dampen vertical air movements, under them there is an accumulation of water vapor or other solid particles that impair visibility, fog and various forms clouds. The layers of inversions are decelerating layers for horizontal air movements as well. In many cases, these layers are wind break surfaces. Inversions in the troposphere can be observed near the surface of the earth and on high altitudes. The tropopause is a powerful layer of inversion.
Depending on the causes of occurrence, there are the following types inversions:
1. Radiation - the result of cooling the surface layer of air, usually at night.
2. Advective - when warm air moves to a cold underlying surface.
3. Compression or subsidence - formed in the central parts of inactive anticyclones.
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