In order to understand solar radiation and the complex interplay between solar photons and the Earth’s atmosphere, it should be recognised that none of these elements or occurrences are independent, but rather form part of a “bigger picture”. It is known from the First Law of Thermodynamics, a.k.a. the Law of Conservation of Energy, that energy cannot be created or destroyed but rather only be converted into or transferred to another form of energy. Radiation is also governed by this law. A balance is maintained between the radiation received by the Earth and the radiation re-emitted and reflected from the Earth. This balance is known as the Earth radiation budget [1].

Earth Radiation Budget

More specifically, the Earth radiation budget is the balance between incoming short-wave radiation and reflected short-wave radiation and re-emitted long-wave radiation. Approximately 30% of the incoming short-wave radiation is reflected, while 70% is absorbed by the atmosphere, Earth’s surface and clouds, which re-emits at longer wavelengths.

A simplified (and idealised) representation of the spatial-temporal average Earth radiation budget, adapted from Rees [1].

The Earth radiation budget is both spatial and temporal variant, which includes local and seasonal variation resulting from surface albedo, cloud cover and weather variations [2]. The Earth radiation budget is, as a result, not only a function of the top of the atmosphere (TOA) and ground surface radiation but also the properties of the Earth’s surface and atmosphere. It should also be noted for future reference that many of these parameters which influence the Earth radiation budget is wavelength-dependent, as implied by the relationship between the incoming and reflected/re-emitted radiation. This wavelength dependence is critical in understanding the scattering and absorption of radiation throughout the atmosphere.

Although the character of the incident radiation will be discussed in a future article, considering the basics of radiative interaction here will foreground some important concepts for these future discussions.

Radiative-Atmospheric Interactions

As radiation traverses the atmosphere it will interact with aerosols and gasses, in which case it will either be absorbed and converted into potential or molecular kinetic energy (again here we find the Law of Conservation of Energy), or scattered/reflected into a different direction without radiant energy loss [3]. The absorption and conversion of energy is the source of the re-emitted radiation at longer wavelengths discussed earlier, and whether radiant energy is absorbed or scattered depends on the wavelengths of the incident radiation (on the particle) versus the energy of the bonds of that particle. Without going into too much detail on the charge of the electrons of the particle, it should be noted that if the frequency (wavelength) of the incident radiation is similar to the resonant (natural) frequency of the particle (molecule), it will very likely be absorbed by the particle.

So, it can be said that incident radiation with wavelengths λ < 100 nm have enough energy to ionise molecules/atoms (charge to absorb and convert radiation energy) in the thermosphere, while radiation with longer wavelengths (and less energy) interact less with gasses and may, therefore, penetrate further into the atmosphere [3]. In some cases, the radiation (200 < λ < 300 nm) does not have enough energy to ionise but may have enough energy to dissolve some molecular bonds, causing secondary reactions in the meso- and stratosphere. One of the most important secondary reactions in the stratosphere is the formation of oxygen (O3).

A representation of a cross-section of the atmosphere and the wavelength dependence of solar radiation throughout the atmosphere, adapted from Lamb and Verlinde [3].

Incident radiation with wavelengths λ > 300 nm may traverse the atmosphere towards the Earth’s surface without significant energy loss, while other complex absorbing and scattering processes in the troposphere influence both the TOA incident radiation, as well as the re-emitted radiation from the Earth’s surface (terrestrial radiation). The emission of terrestrial radiation back to space is possible since the atmosphere is essentially “transparent” to radiation with wavelengths in the IR range (8 < λ < 12 μm), i.e. the interaction between the atmosphere and the radiation within this wavelength range is minimal [3].

To complicate radiative transfer even further, clouds form within the troposphere and in some rare cases, the stratosphere. Clouds introduce additional complexity in the radiative transfer process since it contains molecules/atoms that both absorb and scatter radiation at a much broader wavelength range, while clouds and the formation thereof is inherently complex and dependent on both local and larger weather scales.

[1] G. Rees, “Earth Radiation Budget,” in Remote Sensing Data Book, Cambridge University Press, 1999,

[2] G.E. Musgrave, A. M. Larsen, and T. Sgobba, “Combined Albedo and Planetary Infrared Effects” in Safety Design for Space Systems, Elsevier, 2009,

[3] D. Lamb and J. Verlinde, “The atmospheric setting” in Physics and Chemistry of Clouds, Cambridge University Press, 2011,