I. Origin of Atmospheres

    A. Primitive atmospheres. Most planets originally had a primitive atmosphere, having solar abundances of volatiles, enriched in hydrogen. Most of the hydrogen escaped into the solar nebula, leaving heavier gases such as argon, neon, and ammonia concentrated toward the surface. Neon (Ne), which is cannot be produced by radioactive decay,  is a good tracer for determining the amount primitive atmosphere originally surrounding a planet.  The current Ne on Earth suggests that the primitive atmosphere was only .9% of the mass of the current atmosphere.

    B. Secondary atmospheres. Secondary atmospheres are produced by outgassing from the interior of planets.Volcanic gasses account for most of the Earth's current atmosphere.  The primitive atmosphere of Earth was probably strongly reducing, enriched in H2, H2O, CO and H2S.  The time duration of this reducing atmosphere is not well known, but it was probably short.

    C. Processes important in the formation of secondary atmospheres include photodissociation (H2O to O and O2) and photosynthesis on Earth after green plants evolved. Higher temperatures on Venus contributed to the photodissociation of H2O, removing it from the atmosphere. The enriched N2 content of Earth is difficult to explain without biological processes.  Mars volatiles are difficult to inventory until the amount of water included in permafrost is known.

II. Structure of Atmospheres.

    A. Scale height (measures pressure variation with height). If the ideal gas law is substituted for density in the formula for the change inhydrostatic pressure with height (dP = r g dz), and the atmosphere is assumed to isothermal, then a scale height can be defined by integrating the equation for pressure over height z. The scale height is given by H = kT/m MH g, where g is the gravitational acceleration of the planet, k is Boltzmann's constant, MH is the atomic weight of hydrogen, m is the mean atomic weight, and T is the temperature (assumed constant with height for an isothermal atmosphere). For Earth, the scale height is about 6 km, meaning that the pressure decreases by 1/e, where e = 2.71, at 6 km altitude.  For example, at the height of Mt. Everest or K-2 one would be breathing an atmosphere that is about e-1 less in pressure than that at sea level.

B. Temperature structure by dT/dz.  The commonly used names for layers of an atmosphere are derived from the behavior of temperature gradients.  If the temperature gradients are steeper than the adiabatic gradient and temperature decreases with height near the surface, then this region is called the troposphere.  Since this gradient is steeper than the adiabatic gradient, this region strongly convects (has strong winds/weather). Regions typically lying above the troposphere having temperature increasing with height is called a mesosphere.  Transition regions, separating other layers having diferent types of temperature gradient behavior are called "pauses".  For example, the tropopause lies between the troposphere and stratosphere.  The stratosphere on Earth is a region of near zero gradient above the troposphere.  This region has little convection, and hence is the smoothest region for air travel on Earth (20,0000 to 40,000 feet). The term stratosphere is generally only used for Earth and is not favored in general planetary useage, which typically only classifies a troposphere, mesosphere, and thermosphere or ionosphere on plantets, separated by "pauses."

    1. The ozone layer on Earth is concentrated in the mesosphere, at 30 to 60 km altitude.Ultraviolet light is screened out by this layer. The ultraviolet energy incident on O2 in this layer photodissociates O2, allowing O to recombine in O3 (ozone).  The screening of ultraviolet radiation is important to life on Earth.

    2. The ionosphere is a region of the thermosphere (the outermost atmosphere in which temperature increases with altitude) in which atoms and molecures have had electrons stripped. Ions and free electrons have very high mean velocities in this layer, corresponding to kinetic temperatures of 700 to 1800 deg K.  An object would not get very hot, however, in this layer because heating depends on the number of collisions between atmosphere atoms and the objects. These collisions would be more infrequent in the low density thermosphere.  The concentration of ions and their velocities in the ionosphere are stongly affected by the solar wind and changes with the sunspot cycle. The state of the ionosphere affects the reflectivity of Earth radio waves.

    3. Why atmospheres are not isothermal?  Most atmospheres have some opacity, meaning portions of the visible spectrum are absorbed. The absorption heats the atmosphere. For example, on Earth, the ozone layer absorbs ultraviolet wavelengths, heating the mesosphere; the greenhouse gasses, which includes water vapor on Earth, of the troposhere absorbs infrared resulting in a strong positive temperature gradient with decreasing height.

C. Scattering properties

    1. Blue sky. Scattering of light by micron sized particles (aerosols) is in the domain of Rayleigh scattering, where scattered energy decreases as the inverse fourth power of wavelength, or increases as the fourth power of frequency.  Thus, the sky is blue because the blue (higher frequency) portion of the visible spectrum is more highly scattered than the red portion. On Mars, the reddish-tan appearance of its sky is the result of micron sized dust particles suspended in its atmosphere.

    2. At sunset, the sky appears reddish because the blue portion has been  subtracted by scattering at all angles away from the observer by the long path through the atmospheric aerosols.

    3. Fog. The water droplets in clouds and fog are larger  than visible wavelengths, not in the wavelength/particle size domain in which Rayleigh scattering occurs and we see all frequencies (white spectrum) of incident sunlight.

    3. The outer edge of the thermosphere is called the exosphere.  Thermostat molecures (good radiators of heat) in the exosphere are important in controling a planet's temperature. CO in the Martian exosphere helps keep Mars' surface cool. H can escape from the exosphere if accelerated by collisons and the solar wind heating to escape velocities. Photodissociation of H2O and H escape in  Venusian exosphere may account for the relative absence of H2O on Venus.

III. Climate Evolution. Atmosphere density and composition is affected by variations in volcanic outgassing and tectonic activity. Insolation (amount of solar radiation) is affected by variations in tilt of rotational axes variations in orbital parameters.  On Mars, evidence of water flow and erosion and erosion and depostion on polar ice caps suggests strong variations in outgassing in its early history, possibly extending to as recently as 0.4 billion years ago.  Strong seasonal varations are evident on Mars polar ice caps.  Dramatic climate changes are observed in Earth's climate, many of which are associated with the Milankovitz cycles of insolation.

IV. Dynamics of Atmospheres. Superadiabatic temperature gradients in the lowermost layer (troposphere primarily) results in convection (winds).  Strong insolation differences due to Mars' high orbital eccentrcity excite seasonal dust storms by convection in its atmosphere.  Simple planetary circulation models start with warm equatorial air rising upward and poleward and cold polar air sinking downward and toward the equator. Coriolis forces on a rotating planet create spiral motions of the atmosphere with opposite spiral direcitons in the northern and southern hemispheres.