The ionosphere is produced above about 90 km altitude by two principal sources of atmospheric ionization: photoionization by solar radiation and energetic particle “impact” ionization. Photoionization is primarily confined to the day-side hemisphere, although there is some sunlight scattered over the terminator (the day-night boundary), and on unmagnetized planets pressure gradient forces can cause night-ward flow of ionospheric ions produced on the dayside.
To ionize an atmospheric atom or molecule, solar photons must have energies at least equal to the binding energy of the electron. For most atmospheric gases, the necessary energies fall in the extreme ultraviolet (EUV) wavelength range rather than in the visible range. This wavelength range exhibits average flux variations of about a factor of two to three during the solar cycle, causing solar cycle variations in the dayside ionosphere.
To calculate the ionization produced by a flux of photons, one must first determine the solution to a radiative transport equation. This solution describes the attenuation with depth of the incident photon flux given knowledge of the cross sections for photon absorption. Years ago, Chapman studied the simple case of monochromatic radiation incident on a single constituent atmosphere having constant scale height h (the scale length for exponential fall-off of the atmospheric density with increasing height). The solutions for the resulting profiles of photon energy deposition and hence ion production Q (assuming one ion pair per 35 eV deposited) are the basis of the Chapman layer model of an ionosphere. Electron density profiles are derived from the production rates by assuming that recombination at a rate proportional to n2 is the only loss process, whereby n is proportional to Q½.
The Chapman layer model, which is described in most basic texts on ionospheric physics (e.g. Rishbeth and Garriott; Ratcliffe) can apply to radiation incident at any angle from the zenith, and hence can be used to compare ionosphere properties at different times of day and different latitudes. The constant neutral atmosphere scale height can also be varied to emulate the effects of changing solar activity conditions or the different gravitational fields, temperatures, and compositions at other planets. Here we have programmed the standard Chapman layer model in a way that allows you to experiment with both these variables.
On Earth the atmospheric scale height varies from about 5 km at 80 km altitude to 80 km at 500 km altitude. On Venus the carbon dioxide atmosphere and cold temperatures of the nighttime upper atmosphere can result in scale heights as low as 2 km.
The equations of Chapman layer theory will not be repeated here, given their accessibility in the literature.