The atmosphere of Earth is a layer of gases surrounding the planet and retained by Earth's gravity. In general, air pressure and density decrease with altitude. However, temperature has a more complicated profile with altitude and may remain relatively constant or even increase with altitude in some regions. Because the general pattern of the temperature-altitude profile is constant and recognizable, the temperature behavior provides a useful metric to distinguish between atmospheric layers. In this way, Earth's atmosphere can be divided into five main layers. From lowest to highest, these layers are: troposphere, stratosphere, mesosphere, thermosphere and exosphere.

Troposphere: The troposphere begins at Earth's surface and extends to the tropopause at an average height of about 12 km, although this altitude varies from about 9 km at the poles and 17 km at the equator, with some variation due to weather. The tropopause is mostly heated by transfer of energy from the surface, so on average the lowest part is warmest and temperature decreases with altitude. The troposphere contains roughly 80% of the mass of the atmosphere and basically all weather-associated cloud types. The tropopause is the boundary between the troposphere and stratosphere.

Stratosphere: The stratosphere extends from the tropopause at an altitude of about 12 km above sea level to the stratopause at an altitude of 50 to 55 km. The atmospheric pressure at the top of the stratosphere is roughly 1/1000 the pressure at sea level. Temperature increases with height due to increased absorption of ultraviolet radiation by the ozone layer, which restricts turbulence and mixing. Although the temperature may be -60 ^oC at the tropopause, the top of the stratosphere is much warmer and may be near freezing.

Mesosphere: The mesosphere extends from the stratopause at an altitude of about 50 km to the mesopause at 80-85 km above sea level. It is the layer where most meteors burn up entering the atmosphere. Temperature decreases with height in the mesosphere. The mesopause, the temperature minimum that marks the top of the mesosphere, is the coldest place on Earth with an average temperature of about -85 ^oC, and may drop to -100 ^oC.

Thermosphere: Temperature increases with height in the thermosphere from the mesopause at an altitude of about 80 km up to the thermopause. The altitude of the thermopause averages 700 km but varies with solar activity, ranging from about 500-1000 km. The temperature inversion is a result of absorption of highly energetic solar radiation. Temperatures are highly dependent on solar activity and can rise to 1500 ^oC, though the gas molecules are so far apart that temperature in the usual sense is not well defined.

Exosphere: The exosphere is the outermost layer of Earth's atmosphere, ranging from the thermopause, or exobase, to thousands of kilometers into space. It is mainly composed of hydrogen, helium and some heavier molecules such as nitrogen, oxygen and carbon dioxide closer to the exobase. The particle density is too low to behave as a gas. Temperatures are constant with height and are dependent on solar activity, ranging from about 500-2000 K.

The turbopause marks the altitude in Earth's atmosphere below which turbulent mixing dominates. The region below the turbopause is known as the homosphere, where the chemical constituents are well mixed and the composition of the atmosphere remains constant. The region above the turbopause is the heterosphere, where molecular diffusion dominates and the residual atmospheric gases sort into strata according to molecular mass. The turbopause lies near the mesopause at an altitude of roughly 110 km.

Ideal Gas Law

The gases that make up a planetary atmosphere can be treated like ideal gases, meaning that they obey the ideal gas law (also called perfect gas law). The ideal gas law is the equation of state of a hypothetical ideal gas. It is a good approximation of the behavior of many gases under many conditions. The common form of the ideal gas law is,

(1)	P × V = n × R* × T

where P is the absolute pressure of the gas (Pa), V is the volume of the gas (m³), n is the amount of gas (kmol), T is the absolute temperature of the gas (K), and R* is the universal gas constant (8314.4621 J/kmol-K).

Additional forms of the equation include,

(2)	P = ρ × R* × T / M

(3)	P = N × R* × T / NA

(4)	P = N × k × T

And from equations (2) and (3) we find,

(5)	ρ = N × M / NA

where ρ is the density of the gas (kg/m³), M is the mean molecular weight (kg/kmol), N is the total number density (m^-3), N_A is the Avagadro constant (6.0221413×10²⁶ kmol^-1), and k is the Boltzmann constant (1.3806488×10^-23 J/K).

Note that in equations containing both R* and M, and where M is constant, we may sometimes substitute the specific gas constant R, where R = R*/M.

Hydrostatic Equation

The hydrostatic equation states that whenever there is no vertical motion, the difference in pressure dP between two levels dz is caused by the weight of the layer of air. Mathematically, this statement can be written as,

(6)	dP = –ρ × g × dz

where g is the height-dependent acceleration of gravity. The inverse-square law of gravitation provides an expression for g as a function of altitude,

(7)	g = go × [ ro / (ro + z) ]2

where g_o is the sea-level value of the acceleration of gravity (9.80665 m/s²) and r_o is the effective radius of Earth at a specific latitude.

Using equations (2) and (6), the hydrostatic equation can be written as,

(8)	dP = –(P × M / (R* × T)) × g × dz

Integrating from level z₁ to level z₂ results in the hypsometric equation,

(9)	P2 = P1 × e [ –g × M / (R* × T) × (z2 – z1) ]

If we define scale height H as follows,

(10)	H = R* × T / (M × g)

equation (9) then becomes,

(11)	P2 = P1 × e [ –(z2 – z1) / H ]

Geopotential Height

It has been customary in standard-atmosphere calculations to effectively eliminate the variable portion of the acceleration of gravity by the transformation of geometric height z to geopotential height h, thereby simplifying both integration of the hydrostatic equation and the resulting expression for computing pressure. The relationship between geometric and geopotential altitude depends upon the concept of gravity.

The unit of measurement of geopotential is the standard geopotential meter (m'), which represents the work done by lifting a unit mass 1 geometric meter through a region in which the acceleration of gravity is uniformly 9.80665 m/s². The geopotential of any point with respect to mean sea level (assumed zero potential), expressed in geopotential meters, is called geopotential altitude. Therefore, geopotential altitude h is given by,

(12)

and is expressed in geopotential meters (m') when the unit geopotential g_o' is set equal to 9.80665 m²/(s²m'). With geopotential altitude defined as above, the differential may be expressed as,

(13)	go' × dh = g × dz

This expression is used in the hydrostatic equation to reduce the number of variables prior to its integration, thereby leading to an expression for computing pressure as a function of geopotential height.

The relationship between geopotential height h and geometric height z are given by the following expressions,

(14)	h = ro × z / (ro + z)

(15)	z = ro × h / (ro – h)

Molecular-Scale Temperature

The molecular-scale temperature T_M at a point is defined as the product of the kinetic temperature T and the ratio M/M_o, where M is the molecular weight of air at that point, and M_o is the sea-level value of M. Analytically,

(16)	TM = T × Mo / M

The principle virtue of the parameter T_M is that it combines the variable portion of M with the variable T into a single new variable, in a manner somewhat similar to the combining of the variable portion of g with z to form the new variable h. When both of these transformations are introduced into the hydrostatic equation, and when T_M is expressed as a linear function of h, the resulting differential equation has an exact integral. Under these conditions, the computation of P versus h becomes a simple process not requiring numerical integration.

Hydrostatic equation in terms of h and T_M

With geopotential height h and molecular-scale temperature T_M defined, equations (10) and (11) are re-written as,

(17)	H = R* × TM / (Mo × go')

(18)	P2 = P1 × e [ –(h2 – h1) / H ]

Part I — Homosphere

Molecular Weight

In that part of the atmosphere below the turbopause, mixing is dominant and the effect of diffusion and photochemical processes upon mean molecular weight is negligible. In this region the fractional composition of each species is assumed to remain constant and mean molecular weight M remains constant at its sea-level value M_o.

Temperature

Traditionally, standard atmospheres have defined temperature as a linear function of height to eliminate the need for numerical integration in the computation of pressure versus height. The function T_M versus h is expressed as a series of successive linear equations, the general form of which is,

(19)	TM = TM,b + LM,b × (h – hb)

where L_M is the molecular-scale temperature gradient, also called lapse rate. The value of subscript b refers to each of the successive layers.

Let's consider an example:

where we have three layers — h₀ = 0 to h₁ = 12 km, h₁ = 12 km to h₂ = 24 km, and h₂ = 24 km to h₃ = 50 km.

The value of T_M,b for the first layer (b = 0) is equal to T_o, the sea-level value of T, since at this level M = M_o. With this value of T_M,b defined, and the set of values of h_b and the corresponding values of L_M,b defined in Table 2, the function T_M of h is completely defined from the surface to 50 km'.

Suppose we define T_o = 290 K, then for b =0 we have,

therefore,

We now calculate the value of T_M,b for the second layer (b = 1),

For the second layer, we have

And for the third layer (b = 2), we have

Pressure

Within the region of homogeneity, two different equations are used for computing pressure, in which pressure is a function of geopotential altitude. One equation is for the case when L_M,b for a particular layer is not equal to zero, and the other when the value L_M,b is zero. The first of these two expressions is,

(20)	P = Pb × [ TM,b / (TM,b + LM,b × (h – hb)) ] (go' × Mo/(R* × LM,b))

and the latter is,

(21)	P = Pb × e [ –go' × Mo × (h – hb) / (R* × TM,b) ]

In these equations g_o', M_o, and R* are each defined single-valued constants, while L_M,b and h_b are each defined multi-valued constants in accordance with the value of b. The quantity T_M,b is a multi-valued constant dependent on L_M,b and h_b. The reference-level value of P_b for b = 0 is the defined sea-level value P_o. Values of P_b are obtained from the application of the appropriate equation for the case when h = h_b+1.

These two equations yield the pressure for any desired geopotential altitude from sea level to the maximum h_b. Pressures for negative values of h may also be computed from the first equation when subscript b is zero.

Using the values given in Table 2, let's now consider an example.

First we define our single-valued constants:

Note that the unusual units of R* and g_o' is because we're expressing h in units of km' rather than m'.

For the first layer (b = 0) we have,

Since L_M,0 is not equal to zero, we use equation (20),

For the second layer (b = 1) we have,

Since L_M,1 is equal to zero, we use equation (21),

And for the third layer (b = 2) we have,

Mass Density

Having computed temperature and pressure, mass density can be found using the ideal gas law,

(22)	ρ = P × Mo / (R* × TM)

Speed of Sound

The speed of sound C is calculated using the equation,

(23)	C = (γ × R* × TM / Mo)1/2

where γ is the specific heat ratio. Specific heat ratio is the ratio of constant-pressure to constant-volume specific heat, C_p/C_v. Specific heat is a function of temperature; however, the specific heats of elemental gases (such as N₂, O₂ and Ar) vary little with temperature. Since these gases dominate Earth's atmosphere, γ for air can be assumed constant, where γ = 1.400 (dimensionless).

Part II — Heterosphere

Molecular Weight

In that part of the atmosphere above the turbopause, mean molecular weight decreases with increasing altitude. Two processes are primarily responsible for this: the first is the dissociation of molecular oxygen, and the second is diffusive separation, which becomes increasingly important relative to mixing in this height region.

The mean molecular weight M of a mixture of gases is,

(24)	M = Σ(ni × Mi) / Σni = Σ(Fi × Mi)

where

Fi = ni / Σni

where n_i and M_i are the number density and defined molecular weight, respectively, of the individual gas species. Alternatively, F_i is the fractional volume of the individual gas species.

Temperature

In the view of the necessity for computing individual density-altitude profiles for each atmospheric species in the heterosphere, the use of molecular-scale temperature T_M becomes impractical, and in this region kinetic temperature T is used as a governing parameter. In addition, the use of a linearly segmented temperature-height function, with discontinuous first derivatives, is often terminated in favor of one in which the first derivative is continuous. Furthermore, geometric altitude replaces geopotential altitude as the argument of the temperature-height function.

The observed temperature-height profiles usually show large gradients in the lower thermosphere. At greater altitudes, the gradients decrease with increasing height to the exobase, where the temperature approaches as asymptote (usually referred to as the exospheric temperature T_∞), which varies with solar activity, time of day, and several other parameters.

The temperature-altitude profile may be defined in terms of one or more successive functions. For example, the U.S. Standard Atmosphere, which we'll examine in detail later, uses three functions between 86 and 120 km — an isothermal layer, a layer in which T(z) has the form of an ellipse, and a constant positive gradient layer — and above 120 km, a layer in which T increases exponentially toward an asymptote.

For the first of these functions, no one-size-fits-all instructions can be given, as the function(s) are selected on a case-by-case basis as needed to best fit the observed temperature-altitude profile in this region. The latter exponential function, however, is the predominate form at thermospheric altitudes. The standard exponential form is,

(25)	T = T∞ – (T∞ – To) × EXP(–λξ)

where

λ = Lo / (T∞ – To)

ξ = (z – zo) × (ro + zo) / (ro + z)

the principal features of which are shown in Figure 1. In this case, z_o and T_o are the geometric altitude and kinetic temperature at the origin of the exponential function, and should not be confused with sea-level values, which typically carry the same denotation. L_o is the initial temperature gradient at the origin, and r_o is the planet radius.

Pressure

Total pressure P is equal to the sum of the partial pressures for the individual species,

(26)	P = ΣPi = Σni × k × T

where k is the Boltzmann constant (1.3806488×10^-23 J/K), T is T(z) as defined for the appropriate altitude regions, and Σn_i the sum of the number densities of the individual gas species comprising the atmosphere at altitude z. Neither n_i, the number densities of individual species, nor Σn_i, the sum of the individual number densities, is known directly. Consequently, pressures cannot be computed without first determining n_i for each of the significant species.

Mass Density

The total volumetric mass density is equal to,

(27)	ρ = Σ(ni × Mi) / NA

where n_i and M_i are the number density and defined molecular weight of the individual gas species, and N_A is the Avagadro constant (6.0221413×10²⁶ kmol^-1). As with pressure, densities cannot be computed without first determining n_i for each of the significant species.

Number Densities

On Earth, in the altitude region between approximately 85 and 120 km, the effect of height- and time-dependent, molecular oxygen dissociation, and the competition between eddy and molecular diffusion combine to complicate the study of the height distribution of the atmospheric species, such that the generation of numerical values for the altitude profiles of physical parameters necessitates a considerable amount of numerical computation. More specifically, atomic oxygen becomes appreciable above 85 km, and diffusive separation begins to be effective at an average height of about 100 km. Also, in the regime where molecular diffusion becomes significant (above about 85 km), the effect of vertical winds in the composition in important.

These conditions lead to a complex dynamically oriented expression, applicable to each individual gas species, which includes vertical transport and diffusive separation. Ideally, this set of equations should be solved simultaneously, since the number densities of all the species are coupled through the expressions for molecular diffusion. Such a solution would require an inordinate amount of computation. A simpler approach is desired, which is found using some simplifying approximations, and by calculating the number densities of the individual species one at a time. Even this simpler approach is lengthier and more involved than we want to delve into here. However, if you wish to read more, refer to the following excerpt from NASA document SP-398:

The World Meteorological Organization (WMO) has defined the Standard Atmosphere as follows:

"A hypothetical vertical distribution of atmospheric temperature, pressure and density which by international agreement and for historical reasons, is roughly representative of year-round, mid-latitude conditions. Typical usages are as a basis for pressure altimeter calibrations, aircraft performance calculations, aircraft and rocket design, ballistic tables and meteorological diagrams. Air is assumed to obey the perfect gas law and the hydrostatic equation which, taken together, relate temperature, pressure, and density with geopotential. Only one standard atmosphere should be specified at a particular time and this standard atmosphere must not be subjected to amendment except at intervals of many years."

Because of the interest of aerospace engineers and atmospheric scientists in conditions at much higher altitudes than those currently being considered by the WMO, members of the U.S. Committee on Extension to the Standard Atmosphere (COSEA) agreed to add the following paragraph to the above definition in describing the U.S. Standard Atmosphere, 1976, which extends to 1000 km:

"The atmosphere shall also be considered to rotate with the earth, and be an average over the diurnal cycle, semi-annual variation, and the range of conditions from active to quiet geomagnetic, and active to quiet sunspot conditions. Above the turbopause (about 110 km) generalized forms of the hydrostatic equations apply."

The U.S. Standard Atmosphere is an idealized, steady-state representation of mean annual conditions of Earth's atmosphere from the surface to 1000 km at latitude 45^o N, as it is assumed to exist during a period with moderate solar activity. The defining meteorological elements are sea-level temperature and pressure, and a temperature-height profile to 1000 km. The air is assumed to be dry, and at heights sufficiently below 86 km, the atmosphere is assumed to be homogeneously mixed with a relative-volume composition leading to a constant mean molecular weight.

Adopted Constants

The following constants are adopted by this Standard for the purposes of computation. Note that the values of k, N_A and R* differ from currently accepted values.

0 km to 86 km

For altitudes below 86 km, the U.S. Standard Atmosphere uses the method and equations described in the section Atmospheric Modeling, Part I – Homosphere. Values of h_b and L_M,b are given in Table 3. The reference-level values of T_M,b and P_b for b = 0 are equal to the sea level values T_o = 288.15 K and P_o = 101325 Pa.

The complete set of equations used to compute atmospheric properties from sea level to a geometric altitude of 86 km are summarized in Table 4. Be advised that the given altitudes are geopotential. In flight analysis or simulation, it is probable that you'll be working in geometric altitude. If this is the case, the geometric altitude z must be converted to geopotential altitude h using equation (14). Then, using the calculated value of h, the applicable equations are selected from Table 4 and the atmospheric properties are computed.

86 km to 1000 km

For altitudes above 86 km, the COSEA Task Group decided to include as constituents only those species that are known to contribute significantly to the total number density in any portion of the atmosphere between 86 km and 1000 km. Those gases that appear never to contribute more than about 0.5% of the total composition at any point within this height region, or which for various reasons do not exhibit predictable behavior, were purposely omitted. Using these guidelines, the following gases are included: N₂, O, O₂, Ar, and He. Atomic hydrogen is included at heights 150 km and above, but not included in boundary-value considerations at 86 km.

Temperature

Below 80 km, where M is constant at M_o, T is equal to T_M in accordance with equation (16). However, from Table 5 it is seen that the defined number densities at 86 km leads to a value of M = 28.9522 kg/kmol, about 0.04% less than M_o. To produce a smooth transition from this value of M to M_o, the altitude profile of M has arbitrarily been defined to decrease in terms of the ratio M/M_o from 1.000000 at 80 km to 0.9995788 at 86 km, thus the values of T correspondingly decrease from those of T_M. At z₇ = 86 km (h₇ = 84.852 km') we find that T_M,7 = 186.946 K and T₇ = 186.8673 K.

At heights above 86 km, the temperature-altitude profile is defined in terms of four successive functions, each of which is specified in such a way that the first derivative of T with respect to z is continuous over the entire altitude region, 86 to 1000 km. These four functions begin successively at the first four base heights, z_b listed in Table 6, and are designed to represent the following conditions: (1) an isothermal layer from 86 to 91 km; (2) a layer in which T(z) has the form of an ellipse from 91 to 110 km; (3) a constant, positive gradient layer from 110 to 120 km; and (4) a layer in which T increases exponentially toward an asymptote, as z increases from 120 to 1000 km.

For the layer from z₇ = 86 km to z₈ = 91 km, the temperature-altitude function is defined to be isothermally linear with respect to geometric altitude, so that the gradient of T with respect to z is zero (see Table 6). Thus, the standard form of the linear function degenerates to

Since T is defined to be constant for the entire layer, z₇ to z₈, the temperature at z₈ is T₈ = T₇ = 186.8673 K, and the gradient at z₈ is L_K,8 = 0.0 K/km, the same as for L_K,7.

For the layer from z₈ = 91 km to z₉ = 110 km, the temperature-altitude function is defined to be a segment of an ellipse expressed by

where T_c = 263.1905 K, A = –76.3232 K, a = –19.9429 km, and z is limited to values from 91 to 110 km. The equation is derived from the basic equation for an ellipse, to meet the values of T₈ and L_K,8 derived above, as well as defined values T₉ = 240.0 K, and L_K,9 = 12.0 K/km, for z₉ = 110 km.

For the layer from z₉ = 110 km to z₁₀ = 120 km, T(z) has the standard form of the linear function, where subscript b is 9, such that T_b and L_K,b are, respectively, the defined quantities T₉ and L_K,9, while z is limited to the range 110 to 120 km. Thus,

The value of T₁₀ at z₁₀ is found from the above to be 360.0 K.

For the layer from z₁₀ = 120 km to z₁₂ = 1000 km, T(z) is defined to have the exponential form of equation (25),

where

In the above expressions, T_∞ is the exospheric temperature and equals the defined value 1000 K, which is assumed to represent mean solar conditions.

Pressure and Density

From equations (26) and (27) we see that pressure and mass density in this height region are computed from the number densities of the individual gas species. However, we also learned that computation of number density, particularly between the altitudes of 86 km and 120 km, is a lengthier topic than we want to cover in this web page. Instead, we can obtain values of pressure and density from tables in the following document:

For situations where no particular date or location is specified, the U.S. Standard Atmosphere is a good representation of Earth's atmosphere, and an excellent model for flight simulations, or for other purposes where atmospheric properties vs. altitude must be known. However, there may be circumstances when a more specific model is needed or desired. The following are some alternative models.

Below 90 km

The atmospheric models are defined by temperature-altitude profiles in which temperatures change linearly with respect to geopotential altitude. It is assumed that the air is dry, is in hydrostatic equilibrium, and behaves as a perfect gas. The molecular weight of air at sea level, 28.9644 kg/kmol, is assumed constant to 90 km. Actually, dissociation of molecular oxygen begins to take place near 80 km and molecular weight starts decreasing slowly with height. Consequently, the temperatures given in the tables for altitudes above 80 km are slightly but not significantly larger than the ambient kinetic temperature.

Numerical values for the various thermodynamic and physical constants used in computing the tables of atmospheric properties for these Reference Atmospheres are identical to those used in the preparation of the U.S. Standard Atmosphere, 1976, with two exceptions. Surface conditions for the atmospheres are based on mean monthly sea-level values of pressure and temperature for the appropriate latitude rather than on standard conditions. The accelerations of gravity at sea level for the latitudes were obtained from the following expression in which gravity g varies with latitude φ:

(32)	gφ = 9.780356 × (1 + 0.0052885 × sin2 φ – 0.0000059 sin2 2φ)

The relationship between geopotential altitude and geometric altitude is as expressed in equation (14).

Sets of mean monthly Reference Atmospheres for altitudes up to 90 km have been developed for 15-degree intervals of latitude, including the equator and north pole, to provide information on the seasonal and latitudinal variations in the thermodynamic properties of the atmosphere. Properties of the January, April, July, and October models are presented in the following document:

Overview

The atmosphere of Mars is similar to Earth's in that it is thin and relatively transparent to sunlight. Mars' spin rate and axial tilt are also Earthlike. Thus, the Martian atmosphere falls into the category of a rapidly rotating, differentially heated atmosphere with a solid lower boundary. However, there are also important differences. The Martian atmosphere is primarily carbon dioxide with a much lower surface pressure than Earth's (about 0.6% of Earth), and Mars does not have an Earthlike hydrological cycle.

The variation of temperature with height on Mars gives rise to a troposphere, a mesosphere, and a thermosphere. Mars does not have a stratosphere because it lacks an ozone layer. The tropopause on Mars is deep by comparison to Earth, extending to almost 60 km. In the Martian mesosphere, temperatures become nearly constant. In the thermosphere, temperatures increase because of heating due to the absorption of solar radiation. The base of the thermosphere is about 120 km. Mars, like Earth, has an isothermal exosphere, which begins at about 160 km to 250 km.

Mars is colder than Earth, experiences a much greater seasonal change in available insolation (40% compared to 6% for Earth), and has Earthlike diurnal and seasonal changes. The globally averaged surface temperature of Mars is approximately 215 K. However, because Mars lacks oceans, its surface temperatures undergo considerable seasonal, diurnal, and latitudinal variation. The lowest surface temperatures (~150 K) occur in polar regions during winter, the highest surface temperatures (~300 K) occur in the southern subtropics when Mars is closest to the Sun. In these same regions, diurnal variations can exceed 100 K.

Atmospheric Model - Introduction

Although complex models of the Martian atmosphere exist, for this web page we'll produce our own model.

To date, all spacecraft that have landed on Mars have done so in either the equatorial region or the northern hemisphere. The first three landers to accomplish the feat – Viking 1, Viking 2, and Mars Pathfinder – touched down at northern latitudes during the summer season. There are a couple reasons for this: (1) landing near the equator, or at higher latitudes during the summer, provides the solar illumination needed for the operation of solar cells and the production of electricity; and (2) most of the northern hemisphere, and some of the equatorial regions, are of generally lower elevation than the southern hemisphere, thus allowing a lander to descend into a deeper and denser atmosphere, greatly improving the effectiveness of parachutes.

Overview

Although Venus is quite similar to Earth in size and distance from the Sun, its atmosphere is strikingly different from Earth's. The atmosphere of Venus is much more massive, with a total mass 93 times that of Earth, and is composed primarily of carbon dioxide, along with some nitrogen and other trace gases. A hydrosphere is absent on Venus — there is a very low abundance of water in any form. These compositional differences point to a drastically different evolution of the Venusian and terrestrial atmospheres through geological history.

The current state of the Venusian atmosphere also shows great differences from that of Earth. There is a strong greenhouse effect that raises the surface temperature to about 735 K, despite the fact that cloud cover strongly reflects sunlight so that less heat flux is absorbed on Venus than on Earth. The Venusian clouds are featureless and opaque in the visible. The clouds are composed of droplets of sulfuric acid, produced by photochemistry in the upper atmosphere.

The solid body of Venus rotates very slowly, with a period of 243 days, in the opposite direction to the general rotation of the solar system. The general circulation of the Venus atmosphere is dominated by rotation of the atmosphere in the same direction as that of the solid planet but increasing in speed with height from the surface up to the cloud tops. It reaches maximum speed (~100 m/s) near the cloud tops, where the rotation period is about 4 days. Atmospheric dynamics are thus very different on Venus than on Earth.

The length of a solar day on Venus is 116.75 Earth days. In spite of this long time, there is very little day-night temperature contrast in the lower atmosphere. There is also very little latitudinal temperature contrast in the lower atmosphere of Venus. Within the clouds, however, there is a latitudinal temperature gradient, cooler toward the poles by about 20 K. At higher altitude the temperature gradient reverses.

The temperature-height profile of Venus gives rise to a troposphere, a mesosphere, and a thermosphere. The troposphere, which contains about 99% of the atmosphere by mass, begins at the surface and extends upwards to about 65 km. The mesosphere of Venus extends from about 65 km to 120 km in altitude, and the thermosphere begins at around 120 km, eventually reaching the upper limit of the atmosphere (exosphere) at about 220 km to 350 km.

Atmospheric Model - Introduction