Effective temperature

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https://en.wikipedia.org/wiki/Effective_temperature 

The effective temperature (aka ET) of a body such as a star or planet is the temperature of a black body that would emit the same total energy as electromagnetic radiation. Effective temperature is often used as an estimate of a body's surface temperature when the body's emissivity curve (as a function of wavelength) is not known.

When the star's or planet's net emissivity in the relevant wavelength band is less than unity (less than that of a black body), the actual temperature of the body will be higher than the effective temperature. The net emissivity may be low due to surface or atmospheric properties, such as the greenhouse effect.

Star

The effective temperature of the Sun (5778 kelvins) is the temperature a black body of the same size must have to yield the same total emissive power.

The effective temperature of a star is the temperature of a black body with the same luminosity per surface area (F_Bol) as the star and is defined according to the Stefan–Boltzmann law F_Bol = σT_eff⁴. Notice that the total (bolometric) luminosity of a star is then L = 4πR²σT_eff⁴, where R is the stellar radius. The definition of the stellar radius is obviously not straightforward. More rigorously the effective temperature corresponds to the temperature at the radius that is defined by a certain value of the Rosseland optical depth (usually 1) within the stellar atmosphere. The effective temperature and the bolometric luminosity are the two fundamental physical parameters needed to place a star on the Hertzsprung–Russell diagram. Both effective temperature and bolometric luminosity depend on the chemical composition of a star.

The effective temperature of the Sun is around 5,778 K. The nominal value defined by the International Astronomical Union for use as a unit of measure of temperature is 5,772±0.8 K. Stars have a decreasing temperature gradient, going from their central core up to the atmosphere. The "core temperature" of the Sun—the temperature at the centre of the Sun where nuclear reactions take place—is estimated to be 15,000,000 K.

The color index of a star indicates its temperature from the very cool—by stellar standards—red M stars that radiate heavily in the infrared to the very hot blue O stars that radiate largely in the ultraviolet. Various colour-effective temperature relations exist in the literature. Their relations also have smaller dependencies on other stellar parameters, such as the stellar metallicity and surface gravity. The effective temperature of a star indicates the amount of heat that the star radiates per unit of surface area. From the hottest surfaces to the coolest is the sequence of stellar classifications known as O, B, A, F, G, K, M.

A red star could be a tiny red dwarf, a star of feeble energy production and a small surface or a bloated giant or even supergiant star such as Antares or Betelgeuse, either of which generates far greater energy but passes it through a surface so large that the star radiates little per unit of surface area. A star near the middle of the spectrum, such as the modest Sun or the giant Capella radiates more energy per unit of surface area than the feeble red dwarf stars or the bloated supergiants, but much less than such a white or blue star as Vega or Rigel.

Planet

Blackbody temperature

Main article: Planetary equilibrium temperature

To find the effective (blackbody) temperature of a planet, it can be calculated by equating the power received by the planet to the known power emitted by a blackbody of temperature T.

Take the case of a planet at a distance D from the star, of luminosity L.

Assuming the star radiates isotropically and that the planet is a long way from the star, the power absorbed by the planet is given by treating the planet as a disc of radius r, which intercepts some of the power which is spread over the surface of a sphere of radius D (the distance of the planet from the star). The calculation assumes the planet reflects some of the incoming radiation by incorporating a parameter called the albedo (a). An albedo of 1 means that all the radiation is reflected, an albedo of 0 means all of it is absorbed. The expression for absorbed power is then:

${\displaystyle P_{\mathrm{a}\mathrm{b}\mathrm{s}}=\frac{L{r}^2(1-a)}{4D^2}}$

The next assumption we can make is that the entire planet is at the same temperature T, and that the planet radiates as a blackbody. The Stefan–Boltzmann law gives an expression for the power radiated by the planet:

${\displaystyle P_{\mathrm{r}\mathrm{a}\mathrm{d}}=4\pi r^2\sigma T^4}$

Equating these two expressions and rearranging gives an expression for the effective temperature:

${\displaystyle T=\sqrt[4]{\frac{L(1-a)}{16\pi \sigma D^2}}}$

Where ${\displaystyle \sigma }$ is the Stefan–Boltzmann constant. Note that the planet's radius has cancelled out of the final expression.

The effective temperature for Jupiter from this calculation is 88 K and 51 Pegasi b (Bellerophon) is 1,258 K. A better estimate of effective temperature for some planets, such as Jupiter, would need to include the internal heating as a power input. The actual temperature depends on albedo and atmosphere effects. The actual temperature from spectroscopic analysis for HD 209458 b (Osiris) is 1,130 K, but the effective temperature is 1,359 K. The internal heating within Jupiter raises the effective temperature to about 152 K.

Surface temperature of a planet

The surface temperature of a planet can be estimated by modifying the effective-temperature calculation to account for emissivity and temperature variation.

The area of the planet that absorbs the power from the star is A_abs which is some fraction of the total surface area A_total = 4πr², where r is the radius of the planet. This area intercepts some of the power which is spread over the surface of a sphere of radius D. We also allow the planet to reflect some of the incoming radiation by incorporating a parameter a called the albedo. An albedo of 1 means that all the radiation is reflected, an albedo of 0 means all of it is absorbed. The expression for absorbed power is then:

${\displaystyle P_{\mathrm{a}\mathrm{b}\mathrm{s}}=\frac{L{A}_{\mathrm{a}\mathrm{b}\mathrm{s}}(1-a)}{4\pi D^2}}$

The next assumption we can make is that although the entire planet is not at the same temperature, it will radiate as if it had a temperature T over an area A_rad which is again some fraction of the total area of the planet. There is also a factor ε, which is the emissivity and represents atmospheric effects. ε ranges from 1 to 0 with 1 meaning the planet is a perfect blackbody and emits all the incident power. The Stefan–Boltzmann law gives an expression for the power radiated by the planet:

${\displaystyle P_{\mathrm{r}\mathrm{a}\mathrm{d}}=A_{\mathrm{r}\mathrm{a}\mathrm{d}}\epsilon \sigma T^4}$

Equating these two expressions and rearranging gives an expression for the surface temperature:

${\displaystyle T=\sqrt[4]{\frac{A_{\mathrm{a}\mathrm{b}\mathrm{s}}}{A_{\mathrm{r}\mathrm{a}\mathrm{d}}}\frac{L(1-a)}{4\pi \sigma \epsilon D^2}}}$

Note the ratio of the two areas. Common assumptions for this ratio are ⁠1/4⁠ for a rapidly rotating body and ⁠1/2⁠ for a slowly rotating body, or a tidally locked body on the sunlit side. This ratio would be 1 for the subsolar point, the point on the planet directly below the sun and gives the maximum temperature of the planet — a factor of √2 (1.414) greater than the effective temperature of a rapidly rotating planet.

Also note here that this equation does not take into account any effects from internal heating of the planet, which can arise directly from sources such as radioactive decay and also be produced from frictions resulting from tidal forces.

Earth effective temperature

See also: Stefan–Boltzmann law § Effective temperature of the Earth

Earth has an albedo of about 0.306 and a solar irradiance (L / 4 π D²) of 1361 W m⁻² at its mean orbital radius of 1.5×10⁸ km. The calculation with ε=1 and remaining physical constants then gives an Earth effective temperature of 254 K (−19 °C).

The actual temperature of Earth's surface is an average 288 K (15 °C) as of 2020. The difference between the two values is called the greenhouse effect. The greenhouse effect results from materials in the atmosphere (greenhouse gases and clouds) absorbing thermal radiation and reducing emissions to space, i.e., reducing the planet's emissivity of thermal radiation from its surface into space. Substituting the surface temperature into the equation and solving for ε gives an effective emissivity of about 0.61 for a 288 K Earth. Furthermore, these values calculate an outgoing thermal radiation flux of 238 W m⁻² (with ε=0.61 as viewed from space) versus a surface thermal radiation flux of 390 W m⁻² (with ε≈1 at the surface). Both fluxes are near the confidence ranges reported by the IPCC.

Planetary equilibrium temperature

From Wikipedia, the free encyclopedia

https://en.wikipedia.org/wiki/Planetary_equilibrium_temperature 

The planetary equilibrium temperature is a theoretical temperature that a planet would be if it were in radiative equilibrium, typically under the assumption that it radiates as a black body being heated only by its parent star. In this model, the presence or absence of an atmosphere (and therefore any greenhouse effect) is irrelevant, as the equilibrium temperature is calculated purely from a balance with incident stellar energy.

Other authors use different names for this concept, such as equivalent blackbody temperature of a planet. The effective radiation emission temperature is a related concept, but focuses on the actual power radiated rather than on the power being received, and so may have a different value if the planet has an internal energy source or when the planet is not in radiative equilibrium.

Planetary equilibrium temperature differs from the global mean temperature and surface air temperature, which are measured observationally by satellites or surface-based instruments, and may be warmer than the equilibrium temperature due to the greenhouse effect.

Calculation of equilibrium temperature

Consider a planet orbiting its host star. The star emits radiation isotropically, and some fraction of this radiation reaches the planet. The amount of radiation arriving at the planet is referred to as the incident solar radiation, ${\displaystyle I_o}$. The planet has an albedo that depends on the characteristics of its surface and atmosphere, and therefore only absorbs a fraction of radiation. The planet absorbs the radiation that isn't reflected by the albedo, and heats up. One may assume that the planet radiates energy like a blackbody at some temperature according to the Stefan–Boltzmann law. Radiative equilibrium exists when the power supplied by the star is equal to the power emitted by the planet. The temperature at which this balance occurs is the planetary equilibrium temperature.

Derivation

The solar flux absorbed by the planet from the star is equal to the flux emitted by the planet:

${\displaystyle F_{\mathrm{a}\mathrm{b}\mathrm{s}}=F_{\mathrm{e}\mathrm{m}\mathrm{i}\mathrm{t}}}$

Assuming a fraction of the incident sunlight is reflected according to the planet's Bond albedo, ${\displaystyle A_B}$:

${\displaystyle (1-A_B)F_{\mathrm{s}\mathrm{o}\mathrm{l}\mathrm{a}\mathrm{r}}=F_{\mathrm{e}\mathrm{m}\mathrm{i}\mathrm{t}}}$

where ${\displaystyle F_{\mathrm{s}\mathrm{o}\mathrm{l}\mathrm{a}\mathrm{r}}}$ represents the area- and time-averaged incident solar flux, and may be expressed as:

${\displaystyle F_{\mathrm{s}\mathrm{o}\mathrm{l}\mathrm{a}\mathrm{r}}=I_o/4}$

The factor of 1/4 in the above formula comes from the fact that only a single hemisphere is lit at any moment in time (creates a factor of 1/2), and from integrating over angles of incident sunlight on the lit hemisphere (creating another factor of 1/2).

Assuming the planet radiates as a blackbody according to the Stefan–Boltzmann law at some equilibrium temperature ${\displaystyle T_{eq}}$, a balance of the absorbed and outgoing fluxes produces:

${\displaystyle (1-A_B)\left(\frac{I_o}{4}\right)=\sigma T_{\mathrm{e}\mathrm{q}}^4}$ where ${\displaystyle \sigma }$ is the Stefan-Boltzmann constant.

Rearranging the above equation to find the equilibrium temperature leads to:

${\displaystyle T_{\mathrm{e}\mathrm{q}}={\left(\frac{I_o\left(1-A_B\right)}{4\sigma }\right)}^{1/4}}$

${\displaystyle T_{\mathrm{e}\mathrm{q}}={\left(\frac{L_o\left(1-A_B\right)}{16\sigma \pi d^2}\right)}^{1/4}}$ where ${\displaystyle L_0}$ is the luminosity of the Sun (${\displaystyle 3,828\cdot {10}^{26}}$ W), and ${\displaystyle d}$ the distance between the planet and the Sun, then :

${\displaystyle T_{\mathrm{e}\mathrm{q}}=1.07652\cdot {10}^8{\left(\frac{\left(1-A_B\right)}{d^2}\right)}^{1/4}}$ (with ${\displaystyle d}$ in metres), or :

${\displaystyle T_{\mathrm{e}\mathrm{q}}=3,404.255\cdot {\left(\frac{\left(1-A_B\right)}{d^2}\right)}^{1/4}}$ (with ${\displaystyle d}$ in million kilometres).

Calculation for extrasolar planets

For a planet around another star, ${\displaystyle I_o}$(the incident stellar flux on the planet) is not a readily measurable quantity. To find the equilibrium temperature of such a planet, it may be useful to approximate the host star's radiation as a blackbody as well, such that:

${\displaystyle F_{\mathrm{s}\mathrm{t}\mathrm{a}\mathrm{r}}=\sigma T_{\mathrm{s}\mathrm{t}\mathrm{a}\mathrm{r}}^4}$

The luminosity (${\displaystyle L}$) of the star, which can be measured from observations of the star's apparent brightness, can then be written as:

${\displaystyle L=4\pi R_{\mathrm{s}\mathrm{t}\mathrm{a}\mathrm{r}}^2\sigma T_{\mathrm{s}\mathrm{t}\mathrm{a}\mathrm{r}}^4}$ where the flux has been multiplied by the surface area of the star.

To find the incident stellar flux on the planet, ${\displaystyle I_x}$, at some orbital distance from the star, ${\displaystyle a}$, one can divide by the surface area of a sphere with radius ${\displaystyle a}$:

${\displaystyle I_x=\left(\frac{L}{4\pi a^2}\right)}$

Plugging this into the general equation for planetary equilibrium temperature gives:

${\displaystyle T_{\mathrm{e}\mathrm{q}}={\left(\frac{L\left(1-A_B\right)}{16\sigma \pi a^2}\right)}^{1/4}}$

If the luminosity of the star is known from photometric observations, the other remaining variables that must be determined are the Bond albedo and orbital distance of the planet. Bond albedos of exoplanets can be constrained by flux measurements of transiting exoplanets, and may in future be obtainable from direct imaging of exoplanets and a conversion from geometric albedo. Orbital properties of the planet such as the orbital distance can be measured through radial velocity and transit period measurements.

Alternatively, the planetary equilibrium may be written in terms of the temperature and radius of the star:

${\displaystyle T_{\mathrm{e}\mathrm{q}}=T_{\mathrm{s}\mathrm{t}\mathrm{a}\mathrm{r}}\sqrt{\frac{R}{2a}}{\left(1-A_B\right)}^{1/4}}$

Caveats

The equilibrium temperature is neither an upper nor lower bound on actual temperatures on a planet. There are several reasons why measured temperatures deviate from predicted equilibrium temperatures.

Greenhouse effect

In the greenhouse effect, long wave radiation emitted by a planet is absorbed by certain gases in the atmosphere, reducing longwave emissions to space. Planets with substantial greenhouse atmospheres emit more longwave radiation at the surface than what reaches space. Consequently, such planets have surface temperatures higher than their effective radiation emission temperature. For example, Venus has an effective temperature of approximately 226 K (−47 °C; −53 °F), but a surface temperature of 740 K (467 °C; 872 °F). Similarly, Earth has an effective temperature of 255 K (−18 °C; −1 °F),^[14] but a surface temperature of about 288 K (15 °C; 59 °F) due to the greenhouse effect in our lower atmosphere. The surface temperatures of such planets are more accurately estimated by modeling thermal radiation transport through the atmosphere.

Airless bodies

On airless bodies, the lack of any significant greenhouse effect allows equilibrium temperatures to approach mean surface temperatures, as on Mars, where the equilibrium temperature is 210 K (−63 °C; −82 °F) and the mean surface temperature of emission is 215 K (−58 °C; −73 °F). There are large variations in surface temperature over space and time on airless or near-airless bodies like Mars, which has daily surface temperature variations of 50–60 K. Because of a relative lack of air to transport or retain heat, significant variations in temperature develop. Assuming the planet radiates as a blackbody (i.e. according to the Stefan-Boltzmann law), temperature variations propagate into emission variations, this time to the power of 4. This is significant because our understanding of planetary temperatures comes not from direct measurement of the temperatures, but from measurements of the fluxes. Consequently, in order to derive a meaningful mean surface temperature on an airless body (to compare with an equilibrium temperature), a global average surface emission flux is considered, and then an 'effective temperature of emission' that would produce such a flux is calculated. The same process would be necessary when considering the surface temperature of the Moon, which has an equilibrium temperature of 271 K (−2 °C; 28 °F), but can have temperatures of 373 K (100 °C; 212 °F) in the daytime and 100 K (−173 °C; −280 °F) at night. Again, these temperature variations result from poor heat transport and retention in the absence of an atmosphere.

Internal energy fluxes

Orbiting bodies can also be heated by tidal heating, geothermal energy which is driven by radioactive decay in the core of the planet, or accretional heating. These internal processes will cause the effective temperature (a blackbody temperature that produces the observed radiation from a planet) to be warmer than the equilibrium temperature (the blackbody temperature that one would expect from solar heating alone).

For example, on Saturn, the effective temperature is approximately 95 K, compared to an equilibrium temperature of about 63 K. This corresponds to a ratio between power emitted and solar power received of ~2.4, indicating a significant internal energy source. Jupiter and Neptune have ratios of power emitted to solar power received of 2.5 and 2.7, respectively.

Close correlation between the effective temperature and equilibrium temperature of Uranus can be taken as evidence that processes producing an internal flux are negligible on Uranus compared to the other giant planets.

Earth has insufficient geothermal heating to significantly affect its global temperature, with geothermal heating supplying only 0.03% of Earth's total energy budget.

Supercontinent

From Wikipedia, the free encyclopedia

https://en.wikipedia.org/wiki/Supercontinent 

The supercontinent of Pangaea with the positions of the continents at the Permian-Triassic boundary, about 250 Ma. AR=Amuria; NC=North China; SC=South China; PA=Panthalassic Ocean; PT=Paleotethys Ocean; NT=Neotethys Ocean. Orogens shown in red. Subduction zones shown in black. Spreading centers shown in green.

Although not a supercontinent, the current Afro-Eurasian landmass contains about 57% of Earth's land area.

In geology, a supercontinent is the assembly of most or all of Earth's continental blocks or cratons to form a single large landmass. However, some geologists use a different definition, "a grouping of formerly dispersed continents", which leaves room for interpretation and is easier to apply to Precambrian times. To separate supercontinents from other groupings, a limit has been proposed in which a continent must include at least about 75% of the continental crust then in existence in order to qualify as a supercontinent.

Moving under the forces of plate tectonics, supercontinents have assembled and dispersed multiple times in the geologic past. According to modern definitions, a supercontinent does not exist today; the closest is the current Afro-Eurasian landmass, which covers approximately 57% of Earth's total land area. The last period in which the continental landmasses were near to one another was 336 to 175 million years ago, forming the supercontinent Pangaea. The positions of continents have been accurately determined back to the early Jurassic, shortly before the breakup of Pangaea. Pangaea's predecessor Gondwana is not considered a supercontinent under the first definition since the landmasses of Baltica, Laurentia and Siberia were separate at the time.

A future supercontinent, termed Pangaea Proxima, is hypothesized to form within the next 250 million years.

Theories

The Phanerozoic supercontinent Pangaea began to break up 215 Ma and this distancing continues today. Because Pangaea is the most recent of Earth's supercontinents, it is the best known and understood. Contributing to Pangaea's popularity in the classroom, its reconstruction is almost as simple as fitting together the present continents bordering the Atlantic ocean like puzzle pieces.

For the period before Pangaea, there are two contrasting models for supercontinent evolution through geological time.

Series

The first model theorizes that at least two separate supercontinents existed comprising Vaalbara and Kenorland, with Kenorland comprising Superia and Sclavia. These parts of Neoarchean age broke off at ~2480 and 2312 Ma, and portions of them later collided to form Nuna (Northern Europe and North America). Nuna continued to develop during the Mesoproterozoic, primarily by lateral accretion of juvenile arcs, and in ~1000 Ma Nuna collided with other land masses, forming Rodinia. Between ~825 and 750 Ma Rodinia broke apart. However, before completely breaking up, some fragments of Rodinia had already come together to form Gondwana by ~608 Ma. Pangaea formed through the collision of Gondwana, Laurasia (Laurentia and Baltica), and Siberia.

Protopangea–Paleopangea

The second model (Kenorland-Arctica) is based on both palaeomagnetic and geological evidence and proposes that the continental crust comprised a single supercontinent from ~2.72 Ga until break-up during the Ediacaran period after ~0.573 Ga. The reconstruction is derived from the observation that palaeomagnetic poles converge to quasi-static positions for long intervals between ~2.72–2.115 Ga; 1.35–1.13 Ga; and 0.75–0.573 Ga with only small peripheral modifications to the reconstruction. During the intervening periods, the poles conform to a unified apparent polar wander path.

Although it contrasts the first model, the first phase (Protopangea) essentially incorporates Vaalbara and Kenorland of the first model. The explanation for the prolonged duration of the Protopangea–Paleopangea supercontinent appears to be that lid tectonics (comparable to the tectonics operating on Mars and Venus) prevailed during Precambrian times. According to this theory, plate tectonics as seen on the contemporary Earth became dominant only during the latter part of geological times. This approach was widely criticized by many researchers as it uses incorrect application of paleomagnetic data.

Cycles

A supercontinent cycle is the break-up of one supercontinent and the development of another, which takes place on a global scale. Supercontinent cycles are not the same as the Wilson cycle, which is the opening and closing of an individual oceanic basin. The Wilson cycle rarely synchronizes with the timing of a supercontinent cycle. However, supercontinent cycles and Wilson cycles were both involved in the creation of Pangaea and Rodinia.

Secular trends such as carbonatites, granulites, eclogites, and greenstone belt deformation events are all possible indicators of Precambrian supercontinent cyclicity, although the Protopangea–Paleopangea solution implies that Phanerozoic style of supercontinent cycles did not operate during these times. Also, there are instances where these secular trends have a weak, uneven, or absent imprint on the supercontinent cycle; secular methods for supercontinent reconstruction will produce results that have only one explanation, and each explanation for a trend must fit in with the rest.

The following table names reconstructed ancient supercontinents, using Bradley's 2011 looser definition, with an approximate timescale of millions of years ago (Ma).

Supercontinent name	Age (Ma)	Period/Era Range	Comment
Vaalbara	3,636–2,803	Eoarchean-Mesoarchean	Also described as a supercraton or just a continent
Ur	2,803–2,408	Mesoarchean-Siderian	Described as both a continent and a supercontinent
Kenorland	2,720–2,114	Neoarchean-Rhyacian	Alternatively the continents may have formed into two groupings Superia and Sclavia
Arctica	2,114–1,995	Rhyacian-Orosirian	Not generally regarded as a supercontinent, depending on definition
Atlantica	1,991–1,124	Orosirian-Stenian	Not generally regarded as a supercontinent, depending on definition
Columbia (Nuna)	1,820–1,350	Orosirian-Ectasian
Rodinia	1,130–750	Stenian-Tonian
Pannotia	633–573	Ediacaran
Gondwana	550–175	Ediacaran-Jurassic	From the Carboniferous, formed part of Pangaea, not always regarded as a supercontinent
Pangaea	336–175	Carboniferous-Jurassic

Volcanism

The causes of supercontinent assembly and dispersal are thought to be driven by convection processes in Earth's mantle. Approximately 660 km into the mantle, a discontinuity occurs, affecting the surface crust through processes involving plumes and superplumes (aka large low-shear-velocity provinces). When a slab of the subducted crust is denser than the surrounding mantle, it sinks to discontinuity. Once the slabs build up, they will sink through to the lower mantle in what is known as a "slab avalanche". This displacement at the discontinuity will cause the lower mantle to compensate and rise elsewhere. The rising mantle can form a plume or superplume.

Besides having compositional effects on the upper mantle by replenishing the large-ion lithophile elements, volcanism affects plate movement. The plates will be moved towards a geoidal low perhaps where the slab avalanche occurred and pushed away from the geoidal high that can be caused by the plumes or superplumes. This causes the continents to push together to form supercontinents and was evidently the process that operated to cause the early continental crust to aggregate into Protopangea.

Dispersal of supercontinents is caused by the accumulation of heat underneath the crust due to the rising of very large convection cells or plumes, and a massive heat release resulted in the final break-up of Paleopangea. Accretion occurs over geoidal lows that can be caused by avalanche slabs or the downgoing limbs of convection cells. Evidence of the accretion and dispersion of supercontinents is seen in the geological rock record.

The influence of known volcanic eruptions does not compare to that of flood basalts. The timing of flood basalts has corresponded with a large-scale continental break-up. However, due to a lack of data on the time required to produce flood basalts, the climatic impact is difficult to quantify. The timing of a single lava flow is also undetermined. These are important factors on how flood basalts influenced paleoclimate.

Plate tectonics

Global palaeogeography and plate interactions as far back as Pangaea are relatively well understood today. However, the evidence becomes more sparse further back in geologic history. Marine magnetic anomalies, passive margin match-ups, geologic interpretation of orogenic belts, paleomagnetism, paleobiogeography of fossils, and distribution of climatically sensitive strata are all methods to obtain evidence for continent locality and indicators of the environment throughout time.

Phanerozoic (541 Ma to present) and Precambrian (4.6 Ga to 541 Ma) had primarily passive margins and detrital zircons (and orogenic granites), whereas the tenure of Pangaea contained few. Matching edges of continents are where passive margins form. The edges of these continents may rift. At this point, seafloor spreading becomes the driving force. Passive margins are therefore born during the break-up of supercontinents and die during supercontinent assembly. Pangaea's supercontinent cycle is a good example of the efficiency of using the presence or lack of these entities to record the development, tenure, and break-up of supercontinents. There is a sharp decrease in passive margins between 500 and 350 Ma during the timing of Pangaea's assembly. The tenure of Pangaea is marked by a low number of passive margins during 336 to 275 Ma, and its break-up is indicated accurately by an increase in passive margins.

Orogenic belts can form during the assembly of continents and supercontinents. The orogenic belts present on continental blocks are classified into three different categories and have implications for interpreting geologic bodies. Intercratonic orogenic belts are characteristic of ocean basin closure. Clear indicators of intracratonic activity contain ophiolites and other oceanic materials that are present in the suture zone. Intracratonic orogenic belts occur as thrust belts and do not contain any oceanic material. However, the absence of ophiolites is not strong evidence for intracratonic belts, because the oceanic material can be squeezed out and eroded away in an intracratonic environment. The third kind of orogenic belt is a confined orogenic belt which is the closure of small basins. The assembly of a supercontinent would have to show intracratonic orogenic belts. However, interpretation of orogenic belts can be difficult.

The collision of Gondwana and Laurasia occurred in the late Palaeozoic. By this collision, the Variscan mountain range was created, along the equator. This 6000-km-long mountain range is usually referred to in two parts: the Hercynian mountain range of the late Carboniferous makes up the eastern part, and the western part is the Appalachian Mountains, uplifted in the early Permian. (The existence of a flat elevated plateau like the Tibetan Plateau is under debate.) The locality of the Variscan range made it influential to both the northern and southern hemispheres. The elevation of the Appalachians would greatly influence global atmospheric circulation.

Climate

Continents affect the climate of the planet drastically, with supercontinents having a larger, more prevalent influence. Continents modify global wind patterns, control ocean current paths, and have a higher albedo than the oceans. Winds are redirected by mountains, and albedo differences cause shifts in onshore winds. Higher elevation in continental interiors produces a cooler, drier climate, the phenomenon of continentality. This is seen today in Eurasia, and rock record shows evidence of continentality in the middle of Pangaea.

Glacial

The term glacial-epoch refers to a long episode of glaciation on Earth over millions of years. Glaciers have major implications on the climate, particularly through sea level change. Changes in the position and elevation of the continents, the paleolatitude and ocean circulation affect the glacial epochs. There is an association between the rifting and breakup of continents and supercontinents and glacial epochs.^[19] According to the model for Precambrian supercontinent series, the breakup of Kenorland and Rodinia was associated with the Paleoproterozoic and Neoproterozoic glacial epochs, respectively.

In contrast, the Protopangea–Paleopangea theory shows that these glaciations correlated with periods of low continental velocity, and it is concluded that a fall in tectonic and corresponding volcanic activity was responsible for these intervals of global frigidity. During the accumulation of supercontinents with times of regional uplift, glacial epochs seem to be rare with little supporting evidence. However, the lack of evidence does not allow for the conclusion that glacial epochs are not associated with the collisional assembly of supercontinents. This could just represent a preservation bias.

During the late Ordovician (~458.4 Ma), the particular configuration of Gondwana may have allowed for glaciation and high CO₂ levels to occur at the same time. However, some geologists disagree and think that there was a temperature increase at this time. This increase may have been strongly influenced by the movement of Gondwana across the South Pole, which may have prevented lengthy snow accumulation. Although late Ordovician temperatures at the South Pole may have reached freezing, there were no ice sheets during the early Silurian (~443.8 Ma) through the late Mississippian (~330.9 Ma). Agreement can be met with the theory that continental snow can occur when the edge of a continent is near the pole. Therefore Gondwana, although located tangent to the South Pole, may have experienced glaciation along its coasts.

Precipitation

Though precipitation rates during monsoonal circulations are difficult to predict, there is evidence for a large orographic barrier within the interior of Pangaea during the late Paleozoic (~251.9 Ma). The possibility of the southwest–northeast trending Appalachian-Hercynian Mountains makes the region's monsoonal circulations potentially relatable to present-day monsoonal circulations surrounding the Tibetan Plateau, which is known to positively influence the magnitude of monsoonal periods within Eurasia. It is therefore somewhat expected that lower topography in other regions of the supercontinent during the Jurassic would negatively influence precipitation variations. The breakup of supercontinents may have affected local precipitation. When any supercontinent breaks up, there will be an increase in precipitation runoff over the surface of the continental landmasses, increasing silicate weathering and the consumption of CO₂.

Temperature

Even though during the Archaean solar radiation was reduced by 30 percent and the Cambrian-Precambrian boundary by 6 percent, the Earth has only experienced three ice ages throughout the Precambrian. Erroneous conclusions are more likely to be made when models are limited to one climatic configuration (which is usually present-day).

Cold winters in continental interiors are due to rate ratios of radiative cooling (greater) and heat transport from continental rims. To raise winter temperatures within continental interiors, the rate of heat transport must increase to become greater than the rate of radiative cooling. Through climate models, alterations in atmospheric CO₂ content and ocean heat transport are not comparatively effective.

CO₂ models suggest that values were low in the late Cenozoic and Carboniferous-Permian glaciations. Although early Paleozoic values are much larger (more than 10 percent higher than that of today). This may be due to high seafloor spreading rates after the breakup of Precambrian supercontinents and the lack of land plants as a carbon sink.

During the late Permian, it is expected that seasonal Pangaean temperatures varied drastically. Subtropic summer temperatures were warmer than that of today by as much as 6–10 degrees, and mid-latitudes in the winter were less than −30 degrees Celsius. These seasonal changes within the supercontinent were influenced by the large size of Pangaea. And, just like today, coastal regions experienced much less variation.

During the Jurassic, summer temperatures did not rise above zero degrees Celsius along the northern rim of Laurasia, which was the northernmost part of Pangaea (the southernmost portion of Pangaea was Gondwana). Ice-rafted dropstones sourced from Russia are indicators of this northern boundary. The Jurassic is thought to have been approximately 10 degrees Celsius warmer along 90 degrees East paleolongitude compared to the present temperature of today's central Eurasia.

Milankovitch cycles

Many studies of the Milankovitch cycles during supercontinent time periods have focused on the mid-Cretaceous. Present amplitudes of Milankovitch cycles over present-day Eurasia may be mirrored in both the southern and northern hemispheres of the supercontinent Pangaea. Climate modeling shows that summer fluctuations varied 14–16 degrees Celsius on Pangaea, which is similar or slightly higher than summer temperatures of Eurasia during the Pleistocene. The largest-amplitude Milankovitch cycles are expected to have been at mid-to high-latitudes during the Triassic and Jurassic.

Atmospheric gases

Plate tectonics and the chemical composition of the atmosphere (specifically greenhouse gases) are the two most prevailing factors present within the geologic time scale. Continental drift influences both cold and warm climatic episodes. Atmospheric circulation and climate are strongly influenced by the location and formation of continents and supercontinents. Therefore, continental drift influences mean global temperature.

Oxygen levels of the Archaean were negligible, and today they are roughly 21 percent. It is thought that the Earth's oxygen content has risen in stages: six or seven steps that are timed very closely to the development of Earth's supercontinents.

1. Continents collide
2. Super-mountains form
3. Erosion of super-mountains
4. Large quantities of minerals and nutrients wash out to open ocean
5. Explosion of marine algae life (partly sourced from noted nutrients)
6. Mass amounts of oxygen produced during photosynthesis

The process of Earth's increase in atmospheric oxygen content is theorized to have started with the continent-continent collision of huge landmasses forming supercontinents, and therefore possibly supercontinent mountain ranges (super-mountains). These super-mountains would have eroded, and the mass amounts of nutrients, including iron and phosphorus, would have washed into oceans, just as is seen happening today. The oceans would then be rich in nutrients essential to photosynthetic organisms, which would then be able to respire mass amounts of oxygen. There is an apparent direct relationship between orogeny and the atmospheric oxygen content. There is also evidence for increased sedimentation concurrent with the timing of these mass oxygenation events, meaning that the organic carbon and pyrite at these times were more likely to be buried beneath sediment and therefore unable to react with the free oxygen. This sustained the atmospheric oxygen increases.

At 2.65 Ga there was an increase in molybdenum isotope fractionation. It was temporary but supports the increase in atmospheric oxygen because molybdenum isotopes require free oxygen to fractionate. Between 2.45 and 2.32 Ga, the second period of oxygenation occurred, which has been called the 'great oxygenation event.' Evidence supporting this event includes red beds appearance 2.3 Ga (meaning that Fe³⁺ was being produced and became an important component in soils).

The third oxygenation stage approximately 1.8 Ga is indicated by the disappearance of iron formations. Neodymium isotopic studies suggest that iron formations are usually from continental sources, meaning that dissolved Fe and Fe²⁺ had to be transported during continental erosion. A rise in atmospheric oxygen prevents Fe transport, so the lack of iron formations may have been the result of an increase in oxygen. The fourth oxygenation event, roughly 0.6 Ga, is based on modeled rates of sulfur isotopes from marine carbonate-associated sulfates. An increase (near doubled concentration) of sulfur isotopes, which is suggested by these models, would require an increase in the oxygen content of the deep oceans.

Between 650 and 550 Ma there were three increases in ocean oxygen levels, this period is the fifth oxygenation stage. One of the reasons indicating this period to be an oxygenation event is the increase in redox-sensitive molybdenum in black shales. The sixth event occurred between 360 and 260 Ma and was identified by models suggesting shifts in the balance of ³⁴S in sulfates and ¹³C in carbonates, which were strongly influenced by an increase in atmospheric oxygen.

Proxies

Granites and detrital zircons have notably similar and episodic appearances in the rock record. Their fluctuations correlate with Precambrian supercontinent cycles. The U–Pb zircon dates from orogenic granites are among the most reliable aging determinants.

Some issues exist with relying on granite sourced zircons, such as a lack of evenly globally sourced data and the loss of granite zircons by sedimentary coverage or plutonic consumption. Where granite zircons are less adequate, detrital zircons from sandstones appear and make up for the gaps. These detrital zircons are taken from the sands of major modern rivers and their drainage basins. Oceanic magnetic anomalies and paleomagnetic data are the primary resources used for reconstructing continent and supercontinent locations back to roughly 150 Ma.