Troposphere

The term troposphere derives from the Greek words tropos (rotating) and sphaira (sphere) indicating that rotational turbulence mixes the layers of air and so determines the structure and the phenomena of the troposphere. The rotational friction of the troposphere against the planetary surface affects the flow of the air, and so forms the planetary boundary layer (PBL) that varies in height from hundreds of meters up to 2 km (1.2 mi; 6,600 ft). The measures of the PBL vary according to the latitude, the landform, and the time of day when the meteorological measurement is realized. Atop the troposphere is the tropopause, which is the functional atmospheric border that demarcates the troposphere from the stratosphere. As such, because the tropopause is an inversion layer in which air-temperature increases with altitude, the temperature of the tropopause remains constant. The layer has the largest concentration of nitrogen.

Structure

Composition

The Earth's troposphere is composed of a mixture of gases, which when dry is by mole fraction 78.08% nitrogen as N₂, 20.95% oxygen as O₂, 0.93% argon, 0.04% carbon dioxide as CO₂, and other trace gases. Additionally, the troposphere contains the vast majority of water vapor in the atmosphere. The flux of water vapor in the troposphere can slightly change these numbers, but even at the most humid, water vapor constitutes less than 5% of the local tropospheric composition. Local phenomena like combustion, volcanic eruptions, and industrial air pollution can also affect the local tropospheric composition.

Pressure

The maximum air pressure (weight of the atmosphere) is at sea level and decreases at high altitude because the atmosphere is in hydrostatic equilibrium, wherein the air pressure is equal to the weight of the air above a given point on the planetary surface. The relation between decreased air pressure and high altitude can be equated to the density of a fluid, by way of the following hydrostatic equation:

${\displaystyle \frac{dP}{dz}=-\rho g_n=-\frac{mP{g}_n}{RT}}$

where:

• g_n is the standard gravity
• ρ is the density
• z is the altitude
• P is the pressure
• R is the gas constant
• T is the thermodynamic (absolute) temperature
• m is the molar mass

Temperature

The planetary surface of the Earth heats the troposphere by means of latent heat, thermal radiation, and sensible heat. The gas layers of the troposphere are less dense at the geographic poles and denser at the equator, where the average height of the tropical troposphere is 13 km, approximately 7.0 km greater than the 6.0 km average height of the polar troposphere at the geographic poles; therefore, surplus heating and vertical expansion of the troposphere occur in the tropical latitudes. At the middle latitudes, tropospheric temperatures decrease from an average temperature of 15 °C (59 °F) at sea level to approximately −55 °C (−67 °F) at the tropopause. At the equator, the tropospheric temperatures decrease from an average temperature of 20 °C (68 °F) at sea level to approximately −70 to −75 °C (−94 to −103 °F) at the tropopause. At the geographical poles, the Arctic and the Antarctic regions, the tropospheric temperature decreases from an average temperature of 0 °C (32 °F) at sea level to approximately −45 °C (−49 °F) at the tropopause.

Altitude

The temperature of the troposphere decreases with increased altitude, and the rate of decrease in air temperature is measured with the environmental lapse rate (${\displaystyle -dT/dz}$), which is the numeric difference between the temperature of the planetary surface and the temperature of the tropopause divided by the altitude. Functionally, the ELR equation presumes that the planetary atmosphere is static and that there is no mixing of the layers of air by either vertical atmospheric convection or winds that could create turbulence.

The difference in temperature derives from the planetary surface absorbing most of the energy from the sun, which then radiates outwards and heats the troposphere (the first layer of the atmosphere of Earth) while the radiation of surface heat to the upper atmosphere results in the cooling of that layer of the atmosphere. The ELR equation also assumes that the atmosphere is static, but heated air becomes buoyant, expands, and rises. The dry adiabatic lapse rate (DALR) accounts for the effect of the expansion of dry air as it rises in the atmosphere, and the wet adiabatic lapse rate (WALR) includes the effect of the condensation-rate of water vapor upon the environmental lapse rate.

Compression and expansion

A parcel of air rises and expands because of the lower atmospheric pressure at high altitudes. The expansion of the air parcel pushes outwards against the surrounding air, and transfers energy (as work) from the parcel of air to the atmosphere. Transferring energy to a parcel of air by way of heat is a slow and inefficient exchange of energy with the environment, which is an adiabatic process (no energy transfer by way of heat). As the rising parcel of air loses energy while it acts upon the surrounding atmosphere, no heat energy is transferred from the atmosphere to the air parcel to compensate for the heat loss. The parcel of air loses energy as it reaches greater altitude, which is manifested as a decrease in the temperature of the air mass. Analogously, the reverse process occurs within a cold parcel of air that is being compressed and is sinking to the planetary surface.

The compression and the expansion of an air parcel are reversible phenomena in which energy is not transferred into or out of the air parcel; atmospheric compression and expansion are measured as an isentropic process (${\displaystyle dS=0}$) wherein there occurs no change in entropy as the air parcel rises or falls within the atmosphere. Because the heat exchanged (${\displaystyle dQ=0}$) is related to the change in entropy (${\displaystyle dS}$ by ${\displaystyle dQ=TdS}$) the equation governing the air temperature as a function of altitude for a mixed atmosphere is: ${\displaystyle \frac{dS}{dz}=0}$ where S is the entropy. The isentropic equation states that atmospheric entropy does not change with altitude; the adiabatic lapse rate measures the rate at which temperature decreases with altitude under such conditions.

Humidity

If the air contains water vapor, then cooling of the air can cause the water to condense, and the air no longer functions as an ideal gas. If the air is at the saturation vapor pressure, then the rate at which temperature decreases with altitude is called the saturated adiabatic lapse rate. The actual rate at which the temperature decreases with altitude is the environmental lapse rate. In the troposphere, the average environmental lapse rate is a decrease of about 6.5 °C for every 1.0 km (1,000m) of increased altitude. For dry air, an approximately ideal gas, the adiabatic equation is: ${\displaystyle p(z)[T(z){]}^{-\frac{\gamma }{\gamma -1}}=\text{constant}}$ wherein ${\displaystyle \gamma }$ is the heat capacity ratio (${\displaystyle \gamma \approx }$7⁄5) for air. The combination of the equation for the air pressure yields the dry adiabatic lapse rate:${\displaystyle \frac{dT}{dz}=-\frac{mg}{R}\frac{\gamma -1}{\gamma }=-{9.8}^{\circ }\mathrm{C}/\mathrm{k}\mathrm{m}}$.

Environment

The environmental lapse rate (${\displaystyle dT/dz}$), at which temperature decreases with altitude, usually is unequal to the adiabatic lapse rate (${\displaystyle dS/dz\ne 0}$). If the upper air is warmer than predicted by the adiabatic lapse rate (${\displaystyle dS/dz>0}$), then a rising and expanding parcel of air will arrive at the new altitude at a lower temperature than the surrounding air. In which case, the air parcel is denser than the surrounding air, and so falls back to its original altitude as an air mass that is stable against being lifted. If the upper air is cooler than predicted by the adiabatic lapse rate, then, when the air parcel rises to a new altitude, the air mass will have a higher temperature and a lower density than the surrounding air and will continue to accelerate and rise.

Tropopause

The tropopause is the atmospheric boundary layer between the troposphere and the stratosphere, and is located by measuring the changes in temperature relative to increased altitude in the troposphere and in the stratosphere. In the troposphere, the temperature of the air decreases at high altitude, however, in the stratosphere the air temperature initially is constant, and then increases with altitude. The increase of air temperature at stratospheric altitudes results from the ozone layer's absorption and retention of the ultraviolet (UV) radiation that Earth receives from the Sun. The coldest layer of the atmosphere, where the temperature lapse rate changes from a positive rate (in the troposphere) to a negative rate (in the stratosphere) locates and identifies the tropopause as an inversion layer in which limited mixing of air layers occurs between the troposphere and the stratosphere.