Question
describe insulation and radiation balance on earth
Answers
Answered by
GPT-5 mini
I’ll assume you mean “insolation” (incoming solar radiation). Below is a concise explanation of how incoming solar energy and thermal radiation determine Earth’s radiation (energy) balance.
1) Insolation — what Earth receives
- The Sun’s power at Earth’s distance (the solar constant) is ≈ 1361 W/m² at the top of the atmosphere on a surface perpendicular to the Sun’s rays.
- Averaged over the whole spherical Earth that becomes S0/4 ≈ 1361/4 ≈ 340 W/m² because only a disk faces the Sun while energy is spread over the sphere.
- Not all incoming shortwave (visible/near‑IR) radiation reaches and warms the surface: some is reflected to space by clouds, aerosols and the surface (planetary albedo ≈ 0.30 on average), and some is absorbed by the atmosphere.
2) Earth’s radiative balance — the basic idea
- At steady state, the global average incoming solar energy absorbed by the Earth system equals the outgoing longwave (thermal infrared) radiation emitted to space.
- Simple balance: S0/4 × (1 − A) = σ Te^4 where A is albedo, σ is the Stefan–Boltzmann constant and Te is the planet’s effective radiating temperature. Using A ≈ 0.3 gives Te ≈ 255 K (−18 °C). That is the temperature a blackbody would need to emit the observed outgoing IR.
3) Why the surface is warmer (greenhouse effect)
- The atmosphere is partly transparent to solar shortwave but absorbs and re‑emits thermal longwave radiation. Greenhouse gases (H2O, CO2, CH4, etc.) and clouds trap some outgoing IR, so the surface+lower atmosphere must be warmer than Te to produce the same outgoing flux to space.
- This greenhouse effect raises the global mean surface temperature to ≈ 288 K (≈ 15 °C), about 33 K warmer than the effective radiating temperature.
4) Typical global average energy budget numbers (rounded)
- Incoming solar at top of atmosphere: ~340 W/m² (average).
- Reflected to space (albedo): ~100 W/m² (~30%).
- Absorbed by Earth system: ~240 W/m².
- Outgoing longwave radiation to space: ~240 W/m² (in balance with absorbed sunlight).
- At the surface: net solar absorbed ≈ 160–170 W/m², downward longwave (back radiation) from atmosphere ≈ 300–330 W/m², upward surface emission ≈ 390–400 W/m². The imbalance is closed by sensible + latent heat fluxes (~100 W/m²) from surface to atmosphere.
5) Spatial and temporal variations
- Insolation varies strongly with latitude, season, time of day and cloud cover. Tropics receive more net solar energy than polar regions.
- The latitudinal imbalance is compensated by heat transport by the atmosphere and oceans (winds, currents) from low to high latitudes.
- Diurnal and seasonal cycles create large local changes (e.g., night/day, summer/winter), but the global annual average is approximately steady.
6) Feedbacks and radiative forcing
- Climate sensitivity depends on feedbacks: ice‑albedo feedback (less ice → lower albedo → more absorption), water vapor feedback (warming → more water vapor → more greenhouse effect), cloud feedbacks (complex: can cool or warm).
- Radiative forcing is the perturbation to the balance (e.g., increased CO2 reduces outgoing longwave at certain wavelengths). A positive radiative forcing leads to net warming until a new balance is reached.
7) Why this matters
- Small changes in the radiative balance (fractions of a watt per square meter averaged globally) can produce significant climate changes because the system integrates energy over long times and large areas.
- Human emissions of greenhouse gases are altering the outgoing/returning radiation balance, producing a persistent positive forcing and global warming.
If you want, I can:
- Show the basic Stefan–Boltzmann calculations step‑by‑step,
- Draw a simple numeric energy‑budget diagram (incoming, reflected, atmospheric absorption, surface fluxes), or
- Explain one component in more depth (clouds, ocean heat transport, or radiative forcing by CO2).
1) Insolation — what Earth receives
- The Sun’s power at Earth’s distance (the solar constant) is ≈ 1361 W/m² at the top of the atmosphere on a surface perpendicular to the Sun’s rays.
- Averaged over the whole spherical Earth that becomes S0/4 ≈ 1361/4 ≈ 340 W/m² because only a disk faces the Sun while energy is spread over the sphere.
- Not all incoming shortwave (visible/near‑IR) radiation reaches and warms the surface: some is reflected to space by clouds, aerosols and the surface (planetary albedo ≈ 0.30 on average), and some is absorbed by the atmosphere.
2) Earth’s radiative balance — the basic idea
- At steady state, the global average incoming solar energy absorbed by the Earth system equals the outgoing longwave (thermal infrared) radiation emitted to space.
- Simple balance: S0/4 × (1 − A) = σ Te^4 where A is albedo, σ is the Stefan–Boltzmann constant and Te is the planet’s effective radiating temperature. Using A ≈ 0.3 gives Te ≈ 255 K (−18 °C). That is the temperature a blackbody would need to emit the observed outgoing IR.
3) Why the surface is warmer (greenhouse effect)
- The atmosphere is partly transparent to solar shortwave but absorbs and re‑emits thermal longwave radiation. Greenhouse gases (H2O, CO2, CH4, etc.) and clouds trap some outgoing IR, so the surface+lower atmosphere must be warmer than Te to produce the same outgoing flux to space.
- This greenhouse effect raises the global mean surface temperature to ≈ 288 K (≈ 15 °C), about 33 K warmer than the effective radiating temperature.
4) Typical global average energy budget numbers (rounded)
- Incoming solar at top of atmosphere: ~340 W/m² (average).
- Reflected to space (albedo): ~100 W/m² (~30%).
- Absorbed by Earth system: ~240 W/m².
- Outgoing longwave radiation to space: ~240 W/m² (in balance with absorbed sunlight).
- At the surface: net solar absorbed ≈ 160–170 W/m², downward longwave (back radiation) from atmosphere ≈ 300–330 W/m², upward surface emission ≈ 390–400 W/m². The imbalance is closed by sensible + latent heat fluxes (~100 W/m²) from surface to atmosphere.
5) Spatial and temporal variations
- Insolation varies strongly with latitude, season, time of day and cloud cover. Tropics receive more net solar energy than polar regions.
- The latitudinal imbalance is compensated by heat transport by the atmosphere and oceans (winds, currents) from low to high latitudes.
- Diurnal and seasonal cycles create large local changes (e.g., night/day, summer/winter), but the global annual average is approximately steady.
6) Feedbacks and radiative forcing
- Climate sensitivity depends on feedbacks: ice‑albedo feedback (less ice → lower albedo → more absorption), water vapor feedback (warming → more water vapor → more greenhouse effect), cloud feedbacks (complex: can cool or warm).
- Radiative forcing is the perturbation to the balance (e.g., increased CO2 reduces outgoing longwave at certain wavelengths). A positive radiative forcing leads to net warming until a new balance is reached.
7) Why this matters
- Small changes in the radiative balance (fractions of a watt per square meter averaged globally) can produce significant climate changes because the system integrates energy over long times and large areas.
- Human emissions of greenhouse gases are altering the outgoing/returning radiation balance, producing a persistent positive forcing and global warming.
If you want, I can:
- Show the basic Stefan–Boltzmann calculations step‑by‑step,
- Draw a simple numeric energy‑budget diagram (incoming, reflected, atmospheric absorption, surface fluxes), or
- Explain one component in more depth (clouds, ocean heat transport, or radiative forcing by CO2).