Optical sensors are passive systems that measure reflected solar radiation in the visible, near-infrared and short-wave infrared regions. They are called optical because they sense light in a similar way to the human eye or a camera lens, and the images often correspond to what we see. The spectrum of reflected radiation depends on the radiation absorbed by the surface (and is therefore not reflected). Different surfaces produce individual spectral signatures, or signals, because they absorb and reflect sunlight in different ways.

In active systems, a specific electromagnetic radiation signal is transmitted from the instrument and the sensor detects the component of this signal that is reflected or back-scattered by the surface or atmosphere. Active systems include synthetic aperture radar (SAR) and LiDAR (light detection and ranging).

Passive systems detect the short-wave electromagnetic radiation that is reflected or long-wave radiation that is emitted back to space from the Earth’s surface and atmosphere. That is, natural radiation is the measurement source. The MODIS (Moderate Resolution Imaging Spectroradiometer) on NASA’s Terra and Aqua satellites and the MSI (MultiSpectral Instrument) on the European Space Agency’s (ESA) Sentinel-2 satellite (Figure 2.1.24(a)) are examples of passive instruments.

What are the possible fates of electromagnetic radiation as it passes through the atmosphere?

The classification of the electromagnetic spectrum into different regions is based on what property?

Earth observation data are usually in the form of digital imagery. This may be an image similar to a photo, depicting a view or scene familiar to what we see. A digital image, however, is any image composed of several picture elements, or pixels, that have numeric values assigned to them representing the intensity of some measured quantity.

In Study session 1.4.5 you learned how the introduction of satellite monitoring revolutionised Arctic sea ice science. Earth-observing satellites were first launched in the 1960s and today there over 1000 active EO satellites in orbit. These range from small CubeSats (as small as 10 cm cubed ) to large multi-sensor platforms, such as NASA’s Terra satellite, which has five instruments on board, all generating data on Earth system processes on land, sea and in the atmosphere. EO is used to monitor a range of phenomena, such as land cover and vegetation changes, land and sea surface temperature, wildfires, ice volumes, water resources, atmospheric constituents and processes for weather forecasting.

We are in a data-rich world, and environmental science is no exception. Data are being collected at an increasing rate from different types of sensors and platforms that measure and monitor Earth system processes across space and timescales. An area that is expanding rapidly and contributing significantly to our understanding of the environment is remote sensing. This is the observation and analysis of an object or area from a distance, without direct physical contact, in contrast to in situ measurements or on-site observations. When studying Earth remotely, the term Earth observation (EO) is commonly used, typically referring to data collected from Earth-observing satellites. EO can also include data acquisition from aircraft and, increasingly, drone-based platforms.

ome observations you might have made include the following:Oceans tend to have more positive net radiation than land masses in the summer in the mid latitudes, meaning they are absorbing more radiation.Some land masses during summer have low net radiation, such as North Africa, due to the high albedo of the light-coloured desert sands of the Sahara.The polar regions continue to have a negative net radiation in summer due to the ice-covered surfaces (this is seen on Greenland in the Arctic and on the Antarctic continent).

However, this is not the case for every point on the Earth’s surface. Large-scale patterns of variation exist, and one fundamental source of variation is with latitude. Video 2.1.4 in the following activity is an animation of monthly net radiation across the globe between 2000 and 2015. The data are from the Clouds and the Earth’s Radiant Energy System (CERES) sensors on NASA’s Aqua and Terra satellites. Net radiation is the difference between incoming solar radiation and outgoing terrestrial radiation. Positive values mean incoming solar radiation exceeds outgoing terrestrial radiation at that location, and negative values mean outgoing terrestrial radiation exceeds incoming solar radiation.

Evapotranspiration from the surface (land and oceans) accounts for 80 W m super negative two of the surface-absorbed radiation. This is transferred to the atmosphere. How does this compare to other processes transferring energy to the atmosphere?

Therefore, the global average solar radiation arriving at the top of the atmosphere is one-quarter of the solar constant.

The energy budget can be determined at different timescales but is usually averaged over one or more years to account for seasonal effects. This is known as the steady-state condition. It does not mean that the atmosphere is unchanging, but that these variations average out and it is not in a rapid transition to a different state.

The greenhouse effect acts in a way to ‘trap’ energy near the Earth’s surface. Solar energy that is absorbed by the Earth’s surface and emitted as infrared radiation energy is mostly reabsorbed by the atmosphere, which in turn radiates infrared radiation itself, both upwards and downwards. This radiation is reabsorbed and re-emitted by other levels in the atmosphere, perhaps many times, depending on how transparent the atmosphere is at that wavelength. Radiation may be absorbed and re-emitted several times by gases at different temperatures and through the atmospheric column before escaping to space. At low altitudes, the atmosphere is denser and warmer, which intercepts and re-emits more radiation. At higher altitudes, the upward-emitted radiation mostly escapes to space, as the atmosphere above is sufficiently transparent. At high altitudes, the atmosphere is also colder, so emits less radiation. The net result is that more radiation returns to the surface by being emitted downwards from the lower atmosphere than escapes from the atmosphere by being emitted upwards from the higher atmosphere. This recycling of radiation energy in the lower atmosphere means the surface receives energy from the Sun more than once, first directly as solar radiation and then as re-emitted infrared radiation from the atmosphere, raising its temperature (Figure 2.1.19).

In fact, −20 °C is a good estimate of the average tropospheric temperature from the surface to the tropopause (recall Figure 2.1.8), where long-wave radiation is also emitted to space. But as we have seen, gases in the atmosphere are absorbing and re-emitting radiation. Carbon dioxide and water vapour are particularly effective at absorbing the long-wave infrared radiation that is emitted by the Earth, which is then re-emitted both to space and back to Earth, warming the surface. Therefore, the surface is actually about 35 °C warmer than the black-body temperature of the Earth because of infrared absorption and re-emission within the atmosphere. This is the natural greenhouse effect, which is one of the most important impacts of the atmosphere on the environment. Thus, the atmosphere is far more than a thin, almost transparent layer of gas; without it, the Earth’s surface would have an average temperature of −18 °C.

The reason is that an important assumption was made in this calculation that is not correct: that the Earth can be treated as a single system, taking the surface and atmosphere together. This is an oversimplification of the actual situation.

We can also turn this equation around and use this relationship to estimate the temperature of the Earth, on the understanding that the Earth is at steady state. At steady state, the outgoing power of infrared radiation equals the power of absorbed solar radiation, so we can set cap r to the value of absorbed radiation.

Energy is always conserved. Considering the Earth’s environment as a whole, this means that the energy that goes in must either come back out or be stored in it. The energy going in is almost all from absorbed solar (short-wave) radiation, and energy going out is almost all infrared (long-wave) terrestrial radiation. As you saw in the previous study session, just under a third of solar radiation is reflected or scattered back to space, and about a third of what is left is absorbed by the atmosphere. This leaves the planet’s surface to absorb roughly half of the total radiation incident at the top of the atmosphere, predominantly in the visible region.

In addition to back-scattering by clouds, the surface of the Earth can also scatter light by different amounts. For example, snow and ice, which have a high albedo, back-scatter much more light than forests or grassland. All of these factors contribute to how much of the incoming solar radiation is absorbed by the Earth. Clouds also scatter the outgoing terrestrial radiation, which can ‘close’ the atmospheric infrared window and keep the surface warmer. Their net effect is complex, and the role of clouds remains one of the largest uncertainties in climate model projections.

Light is also scattered by particles that are larger than its wavelength. A good example is scattering by clouds, which appear white or various shades of grey. Clouds scatter solar radiation in all directions, including back into space (back-scattering). For this reason, clouds have a significant effect on the Earth’s atmosphere.

We have seen that sky appears blue on a clear day due to the more extensive scattering of short-wavelength light. On the other hand, when looking at the Sun when it is close to the horizon at sunrise or sunset, the light has to pass through more of the atmosphere than when directly overhead, which increases the amount of Rayleigh scattering. More blue and violet light is scattered out from the direct beams travelling towards you, and proportionally more (longer-wavelength) red and orange light reaches your eyes (Figure 2.1.18).

When solar radiation encounters agents that are much smaller than the wavelength (about one-tenth the size), the radiation is dispersed in all directions. This phenomenon is called Rayleigh scattering after British physicist Lord Rayleigh (1842–1919). In the atmosphere it is caused by individual gas molecules. It primarily affects shorter wavelengths and is particularly effective for visible light.

A particularly important window in the atmospheric absorption spectrum is found at infrared wavelengths from roughly 8  mu m to 15  mu m , where there is relatively low absorption by air, except just below 10  mu m due to ozone. This is also the region where the radiation emitted by Earth is high. Transparency in this region of the spectrum is what allows the land to cool rapidly on a cloudless night, as most of the energy being emitted from Earth’s surface is being lost to space. This window can be ‘closed’ if clouds are present, since they absorb and scatter radiation over a wider range of wavelengths than clear air.

The atmosphere is, however, a relatively good absorber of long-wave (infrared) radiation, due principally to carbon dioxide and water vapour, and these gases absorb much of the long-wave radiation emitted by the Earth. Because the atmosphere is largely transparent to short-wave (solar) radiation but absorbs more long-wave radiation, the atmosphere is heated from the ground up. Water vapour, which is more concentrated near the Earth’s surface, absorbs about 60% of the radiation emitted by the Earth and is the gas mainly responsible for warm temperatures in the lower troposphere. As you move further away from the surface, the temperature drops, as we saw in Study session 2.1.2. The fact that the atmosphere receives most of its energy from the Earth’s surface, rather than directly from the Sun, is critical for driving weather processes.

There are two important regions of the spectrum where the atmosphere is relatively transparent: the visible region and part of the radio region. The fact that humans have evolved to see in the visible region and have developed technology that uses radio wavelengths to communicate long distances is of course no coincidence. These regions are known as ‘windows’ because electromagnetic radiation of these wavelengths can pass through the air without much absorption (the regions in Figure 2.1.17(b) where total absorption and scattering is near zero). Because the atmosphere is largely transparent to visible radiation, most of this energy reaches the Earth’s surface, and it does not have a role in heating the atmosphere.

The gases that are important for the absorption of incoming solar radiation are water vapour, oxygen and ozone. Although nitrogen is the main constituent of the atmosphere, it is a poor absorber of solar radiation. Water vapour is the dominant gas, absorbing and scattering radiation across many regions of the spectrum. Oxygen and ozone are very effective at absorbing short-wavelength, high-energy radiation, such that very little radiation less than 0.3 mm reaches the Earth’s surface. Recall that the temperature profile of the atmosphere (Figure 2.1.8) shows warming in the stratosphere, between about 10 and 50 km altitude, which is due to the absorption of ultraviolet radiation in this region. This will be covered in more detail in Part 4 of the block.

In Figure 2.1.17 the absorption features of gases are smoothed for clarity and are actually comprised of numerous extremely fine lines, which merge into the larger features seen on the curve. The peaks in a gas’s absorption spectrum correspond to specific vibrational and rotational transitions of its molecules. Each transition occurs at a characteristic energy and therefore at a specific wavelength (or frequency) of electromagnetic radiation.

The gases in the Earth’s atmosphere are selective absorbers, and emitters, of radiation. About 20% of the radiation that arrives at the top of the atmosphere is absorbed, with this absorption occurring in different regions and wavelengths of the electromagnetic spectrum due to the properties of the different gases present.

Visible light appears white but is composed of all colours. Surfaces that reflect all wavelengths therefore appear white, and surfaces that absorb all wavelengths appear dark. Surfaces or compounds that absorb less, and reflect or scatter more, of a particular wavelength appear the colour of that wavelength. Plants are green, for example, because the chlorophyll in leaves absorbs more blue and red light than green. In the atmosphere, gas molecules scatter the short-wavelength blue and violet light more effectively than the longer-wavelength red and orange light. That is why the sky appears blue on a clear day when looking in any direction other than directly at the Sun, as more of the shorter-wavelength radiation is being scattered by the atmosphere (Figure 2.1.16).

In total, about 30% of solar radiation that arrives at the top of the atmosphere is lost back to space by reflection and back-scattering. This energy does not have a role in heating the atmosphere or the Earth’s surface. Therefore, as a whole, the Earth has an albedo of 0.3, or 30%, which is largely determined by clouds in the atmosphere.

Scattering, on the other hand, produces a larger number of weaker rays travelling in different directions as the light is bent through a range of angles by particles or rough surfaces. Radiation is scattered in the atmosphere because of the presence of molecules, dust and aerosol particles. The appearance of smoke and mist are two everyday examples of such scattering by small particles or water droplets. Scattered and reflected light is also known as diffuse or indirect radiation. Because scattering results in light travelling in different directions, areas that are not in a direct line from the Sun can receive light, such as under the canopy of a tree. Scattering in which the path of radiation is changed by more than 90 degrees, meaning radiation which was moving downwards is now moving upwards, and vice versa, is known as back-scattering.

Reflection is a process in which all light striking the surface bounces back at the same angle and intensity at which it arrived. The term ‘albedo’ is used for the proportion of radiation that is reflected by a surface, and is an important component of the energy balance of an environment. You will be familiar with the concept of albedo from Study session 1.4.5, in the context of sea ice.

All these processes occur in the atmosphere, and what determines whether solar radiation is transmitted, absorbed, scattered or reflected by the gases or other particles in the atmosphere depends on the wavelength of the radiation and the size and nature of the material it encounters.

When radiation encounters matter, three things can happen to it. First, it may be absorbed, causing the molecules of the matter to vibrate faster and increasing its temperature. This is what is happening when you go outside and are warmed by the Sun – your skin is absorbing the radiation. Second, it may be redirected or bounce off the object, which includes being scattered or reflected. It’s because radiation is scattered that we can see objects, as the light can arrive from any direction and bounce off again in (almost) any direction. Finally, it may simply pass through, without being absorbed or redirected, which is the process of transmission. Air and water are transparent to certain wavelengths of radiation, meaning they transmit this energy.

Identify days that are clear (a nice, smooth pattern) and days with intermittent clouds that shade the sensor.

Observe the reduction in PAR in the sensor that is below the canopy compared to above the canopy, and the difference in the level of reflected solar radiation compared to the incoming solar radiation.

Observe the typical daily pattern and how maximum values change over the annual cycle.

The quantum sensor measures the total flux of photons across the visible range, which is expressed in mu mol photons m super negative two  s super negative one . A photon is the minimum quantity, or quantum, of radiation that is involved in physical interactions with some other entity, such as the absorption by photosynthetic pigments in leaves (e.g. chlorophyll) during photosynthesis.

The pyranometers are thermopile pyranometers, which have a sensor that converts thermal energy to electrical energy and are sensitive to most of the solar spectrum. The electrical signal is recorded by the data logger and converted to solar irradiance units of W m super negative two using a calibration factor.

There are three different radiation sensors measuring at the OpenLiving Lab: two pyranometers and a quantum sensor, which are shown in the image below. The pyranometers measure short-wave radiation, or solar radiation, in W m super negative two . One measures incoming short-wave radiation: that is, the solar radiation received at the surface. The other pyranometer is downward-facing and measures outgoing short-wave radiation, or the reflected solar radiation from the surface. The quantum sensor measures radiation in the visible range, from 400–700 nm. This is the photosynthetically active radiation described earlier in this session, so these sensors are commonly referred to as PAR sensors. In addition, in the urban woodland site, there is a second PAR sensor below the canopy.

These calculations confirm what was previously discussed regarding the solar radiation spectrum (Figure 2.1.12), but now we can compare this to terrestrial radiation. Most energy lost from the Earth-atmosphere system back to space is terrestrial radiation, primarily in the infrared region. A plot of the intensity of radiation emitted by a body against wavelength (or frequency) is known as the emission spectrum of that body. Figure 2.1.12 shows the emission spectrum of the Sun, and Figure 2.1.13 shows a simplified version of this plot together with the emission spectrum of the Earth. Note that the radiative flux, which is the rate of energy passing through a unit area in W m super negative two , arriving from the Sun is more than a million (10 super six ) times that leaving from the Earth and is scaled accordingly to be able to show both curves on the same plot. These plots depict what are known as the black-body emission spectra for bodies at these temperatures. Because terrestrial radiation has its maximum flux at a wavelength (10  mu m) about 20 times longer than solar radiation (0.5  mu m), it is often referred to as long-wave radiation, and solar radiation as short-wave radiation.

temperature (in K) and sigma is the Stefan–Boltzmann constant, which is equal to 5.67 multiplication 10 super negative eight cap w m super negative two cap k super negative four .

The relationship between the temperature and the emitted wavelength of the radiating body is Wien’s displacement law: lamda sub m times a times x equals cap c solidus cap t

This law states that as an object increases in temperature it loses energy at an increasing rate by emitting radiation. In fact, the rate of loss of energy, or power loss, is very sensitive to temperature, being proportional to its fourth power, cap t super four (i.e. cap t multiplication cap t multiplication cap t multiplication cap t ). So, if the temperature doubles, the power emitted as radiation increases by a factor of 2 super four  = 16. To determine the total power radiated from the object, cap r , the radiant exitance needs to be multiplied by the area, cap a , of the obje

jects that are good absorbers of radiation are also good emitters. The Earth’s surface and the Sun are nearly perfect radiators, meaning they behave similarly to a black body. A black body is an idealised physical body that absorbs all incident electromagnetic radiation, regardless of wavelength, and re-emits it in a predictable manner based on its temperature. The gases that make up our atmosphere, on the other hand, are selective absorbers and emitters of radiation.

As mentioned, all objects continually radiate energy over a range of wavelengths. Hotter objects emit more total energy per unit area and more energy as short-wave radiation than cooler objects. That is, both the intensity and the wavelength distribution of emitted radiation depends on the temperature of the object, and these relationships are expressed in the two laws presented below.

It is no coincidence that plants use the visible part of the spectrum for photosynthesis. As shown in Figure 2.1.12, this region is where the most energy is coming from the Sun. This is both energy arriving at the Earth’s surface and that arriving at the top of the atmosphere. There is a difference in the quantity and spectral distribution of the radiation arriving at the Earth’s surface compared to the top of the atmosphere due to the absorption of some wavelengths by gases in the atmosphere. We will return to this in the next study session. But first we need to consider some laws of radiation, and their relationship with temperature, that underlie what we have discussed so far. This will be important for understanding subsequent study sessions.

The Sun emits radiation across most of the electromagnetic spectrum, but in varying amounts across different regions. Over 95% of solar radiation is in the wavelengths between 100 and 2500 nm ( one multiplication 10 super negative seven and 2.5 multiplication 10 super negative six m ). Figure 2.1.12 shows the solar irradiance spectrum for solar radiation arriving both at the top of the atmosphere (ToA) and at the Earth’s surface.

Electromagnetic waves with wavelengths shorter than 400 nm are described as ultraviolet (UV) radiation, which continues down to about 10 nm ( one multiplication 10 super negative eight m ), at which point they start to become known as X-rays. Electromagnetic waves with wavelengths longer than 700 nm are described as infrared radiation, which extend up to about 1 mm ( one multiplication 10 super negative three m , or one multiplication 10 super six nm ), at which point they start to become known as microwaves. Ultraviolet, visible and infrared radiation are the most important regions to be aware of for understanding the atmospheric energy balance.

Wavelengths in Figure 2.1.11 are labelled using the base SI unit of metres, which is convenient at the radio end of the electromagnetic spectrum, but less so at infrared, visible and ultraviolet wavelengths. These are therefore often quoted in units of micrometres, mu m (10‍ super negative six  m), or nanometres, nm (10‍ super negative nine  m).

Frequency and wavelength are therefore inversely related; that is, short-wavelength electromagnetic radiation has a high frequency, and low-frequency radiation has a long wavelength. An important feature of radiant energy is that shorter wavelengths (high frequencies) are more energetic. Although we divide the radiation spectrum into these different regions, the waves are all the same phenomena, which differ only in their wavelength and frequency. It is the difference in wavelength and the properties of the object that absorbs radiation that determines the effects of the radiation.

Figure 2.1.11 shows the different bands of the electromagnetic spectrum according to wavelength, going from left to right by increasing wavelength. The symbol lamda (lambda) is used for wavelength. All electromagnetic radiation travels at the same speed in a vacuum, 3.0 multiplication 10 super eight m s super negative one , which is known as the speed of light, c . Electromagnetic radiation can therefore also be defined by its frequency, f , which is its speed divided by its wavelength. This gives the number of waves per second.

The electromagnetic fields are not visible, but their strength and characteristics can be measured with specific instruments. Human eyes are designed to detect the waves across a specific range of wavelengths – from 400 nm to 700 nm, what is known as visible light – and relay this information to the brain. Different wavelengths in the visible range are interpreted as colours, and white light is a mixture of all visible wavelengths. The visible region of the spectrum is also very important for plants, as these are the wavelengths used for photosynthesis. In this context, visible light is referred to as photosynthetically active radiation (PAR). A radio performs a similar function to the eye in a different part of the electromagnetic spectrum, detecting radio waves and relaying them to a speaker to make sound.

Electromagnetic radiation consists of a continuous spectrum of waves of different sizes, or wavelengths, where the wavelength is the distance between the peaks of these waves. These wavelengths range in length over many orders of magnitude, from picometre (less than a billionth of a centimetre) gamma rays to radio waves that are tens of kilometres long. We can identify different regions, or wavelength bands, that have different effects when interacting with an object, and therefore classify electromagnetic radiation on the basis of wavelength.

Radiation is the only form of energy that can be transferred without a transfer medium. That is, it can occur through the vacuum of space. Nuclear fusion reactions in the Sun are the ultimate source of energy that drives our weather and other atmospheric processes. Radiation energy transmission occurs in waves carrying energy in the form of oscillating electric and magnetic fields, hence the term electromagnetic radiation.

The values of irradiance discussed above are for solar radiation reaching the top of the atmosphere. Solar radiation then has to travel through the atmosphere and, even though this is a very small distance relative to the overall journey from the Sun, the atmosphere has an important effect on the nature of the radiation reaching the Earth’s surface and the overall energy balance of the Earth because of interactions solar radiation has with the constituents of the atmosphere. We’ll discuss these interactions and their implications in the next study session, but first we need to understand a bit more about the nature of solar radiation.

Another term when considering the solar radiation that reaches Earth is insolation, which is the energy arriving per unit area over a given time period (e.g. a day, month or year). That is, solar insolation is the cumulative solar irradiance over a given time period. You may also come across this term in the context of solar power installations and the potential energy capture of a given location.

The total energy output per second from the Sun, or radiative power, is an enormous 3.85 multiplication 10 super 26  W. This solar energy is emitted in all directions from the Sun, and the power of solar radiation that hits the top of the Earth’s atmosphere is a tiny fraction of it, about 1.735 multiplication 10 super 17  W (or 4.5 multiplication 10 super negative eight  %). This value represents an average over the year, as it varies by a few per cent with the distance of the Earth from the Sun, due to the elliptical orbit of the Earth around the Sun. The power per unit area of solar radiation (for a surface perpendicular to the Sun’s rays) is termed solar irradiance and has units of W m‍ super negative two . The solar irradiance when the Earth is at the average distance from the Sun (1 astronomical unit, AU) is termed the solar constant and is 1361 W m‍ super negative two . This is illustrated in Figure 2.1.10. By the time solar radiation reaches Earth, the rays are essentially parallel due to the large distance between the Sun and Earth, and the intensity of the radiation crossing an imaginary plane perpendicular to the Sun’s rays defines the solar constant.

Virtually all the energy available on Earth originates from the Sun in the form of solar radiation or from electromagnetic radiation. All objects, including stars, planets, living beings and inanimate objects at a temperature above 0 K also emit electromagnetic radiation and, on balance, most energy lost from the Earth-atmosphere system to space is terrestrial radiation (radiation emitted from the Earth). The intensity and wavelength distribution of emitted radiation depends on the temperature of the object.

adiation is the only mechanism of heat transfer that can occur through the vacuum of space and is therefore the way solar energy reaches the Earth.

Convection involves the actual movement or circulation of a fluid substance of a given temperature (such as air or water) and is the main process by which heat is moved around in the atmosphere. You may have observed convection when boiling a pot of water. The pot is heated from the bottom and heat transfers via conduction from the pot to the water at the bottom of the pan. This water expands, rises and bubbles to the top where it is replaced by cooler and denser water that sinks to the bottom, establishing convective circulation, which will continue as long as the pot is being heated. Similarly, the conductive warming of air adjacent to the Earth’s surface causes it to expand, become less dense and rise, convectively transferring heat to higher layers in the atmosphere. Convection also operates at much larger scales due to the uneven heating of the Earth’s surface, which drives global convective circulation in the atmosphere, to which we will return in Study session 2.2.2.

The transfer of energy can occur via three mechanisms: conduction, convection and radiation (Figure 2.1.9). Conduction is the transfer of heat via molecular collisions between adjacent molecules. Different materials have different abilities to conduct heat. If you leave a teaspoon in a hot cup of tea for a while, then the end gets warm because metal is good conductor – heat is being conducted along the spoon. Air is a poor conductor (or is an insulator). Therefore, the only situation where conduction is important in the atmosphere is for air that is in direct contact with the Earth’s surface.

Sensible heat, on the other hand, is heat transfer that we can feel, or can be ‘sensed’, and can be measured with a thermometer. That is, we can measure the temperature of the two bodies or components in a system exchanging heat and observe the temperature changes. Sensible heat transfer occurs via different mechanisms.

The energy that has been absorbed by the vapour molecules is known as latent heat, meaning ‘hidden’, because it does not result in a change in temperature of the evaporated water vapour molecules. This stored latent heat is eventually released during the process of condensation, when water vapour returns to liquid water. In the atmosphere, this happens during cloud formation, and considerable amounts of energy are transferred from the Earth’s surface to the atmosphere as latent heat through the processes of evaporation and condensation. We will return to water in the atmosphere in Study session 2.2.1.

Latent heat is the energy released or absorbed when water changes from one state of matter (liquid water, ice or water vapour) to another. Energy is required to evaporate water (going from liquid to water vapour), to break the hydrogen bonds between molecules, and this energy is absorbed by the water molecules that escape the waterbody. As the more energetic molecules escape first, evaporation is a cooling process (i.e. the remaining body of water cools).

Heat is the term used to describe the energy transfer into or out of an object due to temperature differences between that object and its surroundings. For example, holding a hot cup of tea will warm your hand as the heat transfers from the cup to your hand. Heat flows from a region of higher temperature to a region of lower temperature, so if you hold an ice cube, the heat flows from your hand to the ice cube. Heat flow stops when temperatures are equal.

Earth, like the rest of the universe, is made up of matter and energy. Matter is anything that has mass and takes up space, from individual atoms to everything comprised of atoms. Energy is commonly defined as the ability to do work, including the movement of mass. Energy exists in various forms, such as radiant, thermal, chemical, electrical and nuclear, but all forms of energy fall into the categories of kinetic and potential energy. Kinetic energy is energy in use, or the energy of motion. Potential energy is energy stored or not yet used. The SI unit for energy is the joule (J), and the rate of energy release or transfer is known as power, which has the unit of watts (W), where 1 W = 1 J s

At the altitudes of the upper atmosphere, the air is so rarefied (thin) that collisions between molecules become less common and it no longer behaves like a familiar gas. In fact, it becomes difficult to assign a single temperature to the atmosphere, and different gas species separate out and must be treated independently. Although there is no distinct border, one definition chooses a boundary of 100 km altitude as the line between the atmosphere and ‘outer space’. This is known as the Kármán Line, proposed by the Hungarian-American engineer and physicist Theodore von Kármán in the 1950s. At around this altitude (the 100 km definition provides a round, practical value), aerodynamic flight becomes impractical because the air is so thin that an aircraft would need to travel faster than orbital velocity to generate enough aerodynamic lift from the atmosphere to support itself. The upper atmosphere is only mentioned briefly here for context. This module will focus principally on the troposphere and stratosphere.

The term ‘troposphere’ is derived from the Greek work tropos, meaning ‘turn’ or ‘churn’, and this is the layer in which the greatest amount of churning or mixing of air occurs. Vertical churning is a crucial driver for weather systems, and the troposphere is where the weather action mostly occurs. It contains 80% of the mass of the atmosphere and virtually all of the clouds and moisture in the atmosphere. The warming in the stratosphere is caused by internal heating of the stratosphere where a layer of ozone absorbs incoming solar ultraviolet radiation. The stratosphere is so named because it is a very stable, or stratified, region with little convection. We will go into more detail on the processes in the troposphere and stratosphere in later sections.

It is evident that there is more structure to the atmospheric temperature profile than the pressure profile, and from this structure we can distinguish regions, or layers, of the atmosphere, which are indicated in Figure 2.1.8. Note that these heights are approximate and based on this average global profile; they will differ for specific profiles at a given place or time. The troposphere is known as the lower atmosphere and the stratosphere and mesosphere can be referred to together as the middle atmosphere. The upper atmosphere expanse above the mesopause is called the thermosphere, where temperature increases again (not shown in this figure).

If the relationship between pressure and altitude were exactly exponential, this plot would be a straight line. It is not quite straight because the temperature of the atmosphere also comes into play, and temperature is also not constant with height. Pressure decreases with altitude less quickly where the atmosphere is warmer because the density is lower, and more quickly where the temperature is lower. However, these variations are not huge because the temperature range (in kelvin) is not large – about 213–288 K over the troposphere, for example – compared to pressure changes that span many orders of magnitude. We will look at temperature changes with altitude next.

The fundamental variables that can be measured in the atmosphere are often pressure and temperature, and height itself is derived from the pressure and temperature structure. In fact, meteorologists often plot the height of a given pressure surface as a dependent variable because it is related to the mean atmospheric temperature below that pressure level.

As discussed in the previous section, atmospheric pressure is high near the Earth’s surface and decreases with altitude. Figure 2.1.6 illustrates the way atmospheric pressure varies with altitude. The figure shows a global mean pressure profile, meaning it is an average over location and time of atmospheric pressure against height above sea level. It is based on climatological data for the lowest 50 km of the atmosphere.

The weather was likely to be calm and warm. High-pressure zones are due to descending air, which suppresses weather development and often leads to calm, clear and sunny conditions.

Defining the precise extent of the atmosphere is difficult, but we can also consider altitude in terms of pressure. If the pressure at the surface is 1000 hPa, then exactly half the mass of the atmosphere will be below the altitude where the pressure is 500 hPa, and half will be above that point. However, atmospheric pressure at the surface is not constant in space or time, so the height at which this half-mass point occurs is not constant. This close connection between mass and pressure is just one reason pressure as a physical variable is important, and why many atmospheric phenomena are discussed in relation to pressure rather than height.

Atmospheric pressure at a given altitude is the weight of all the air above that point, and is one of the most fundamental atmospheric properties. Pressure is defined as force per area, and the pressure of a gas is simply a force per unit area, exerted in all directions, resulting from collisions with the moving gas molecules.

The atmosphere thins rapidly with altitude, and we can use the resulting change in atmospheric pressure to determine the vertical extent of the atmosphere.

hat reduction involves donating electrons in chemical reactions, and oxidation involves accepting electrons. Reducing gases are hydrogen or hydrogen-containing gases, such as methane (CH4), ammonia (NH3) and hydrogen sulfide (H2S). Oxidising gases include oxygen, ozone (O3) and other oxygen-containing gases. Essentially, this has meant going from an atmosphere free of oxygen to the current level of 21%. Consequently, we can infer what past atmospheres were like, as these changes have been recorded in rocks and sediments through chemical reactions between the atmosphere and the Earth’s crust, and biological processes associated with life. These stratigraphic records (layers in the Earth’s crust) have led to the division of geological time into four eons, each lasting hundreds of millions to billions of years: the Hadean Eon (4.6‍–‍4.0 billion years ago, bya), the Archean Eon (4.0‍–‍2.5 bya), the Proterozoic Eon (2.5 bya‍–‍540 million years ago, mya) and the Phanerozoic Eon (540 mya–now).

Earth’s atmosphere has gone through some substantial changes since the formation of the planet 4.6 billion years ago, and the current atmosphere reflects the dynamic and complex evolution of the Earth system through geological time. These changes have both influenced and been shaped by the environmental conditions on Earth and the development of life. The change in the atmosphere over geological time has broadly involved a transition from a reducing atmosphere to an oxidising one.

The values in Table 2.1.1 represent average values for the lower atmosphere, but the exact proportion of each gas can vary with location, both horizontally (latitude and longitude) and vertically (altitude), and with time between seasons. As you can see, the amount of water vapour in the air is very variable, so scientists usually deal with this constituent separately and refer to the other constituents as dry air. In the lower atmosphere, the three ‘essentially constant’ gases listed in Table 2.1.1 are well mixed by the winds and the churning of the atmosphere, and composition does not vary much from place to place. Higher layers of the atmosphere have similar proportions of the two main gases, oxygen and nitrogen, but they can have quite different proportions of the trace gases. A well-known example of a gas found in variable concentration in different parts of the atmosphere is ozone. Its mixing ratio is greatest in the stratosphere, and the amount of stratospheric ozone varies strongly because of chemical reactions in the atmosphere. This will be discussed further in Part 4 of this block.

This mixing ratio definition is based on the number of molecules and not their proportion by mass, which is called the mass mixing ratio. The mass mixing ratio is different from the volumetric mixing ratio because the molecules each have a different mass.

The mixing ratio of a gas is the number of molecules (or atoms of monatomic species, such as argon) of that gas divided by the total number of molecules of all gases present in a given volume. For trace gases, these are given as either parts per million (ppm; 1 ppm is a mixing ratio of 10 super negative six ), parts per billion (ppb; 1 ppb is a mixing ratio of 10 super negative nine ) or parts per trillion (ppt; 1 ppt is a mixing ratio of 10 super negative 12 ) as this is a more convenient way of expressing small mixing ratios. Note that, as a volume mixing ratio, the units are expressed as ppmv, ppbv and pptv, but these are also frequently shortened to ppm, ppb and ppt respectively. Although they have small mixing ratios, many trace gases play a vital role in atmospheric processes, as you will see later in this block.

Table 2.1.1 lists the mixing ratios of some of the gases in air in the lower part of the atmosphere, including the most common gases and some trace gases you will encounter in this block. A trace gas is one that makes up only a small proportion of a sample of air, and tends to be more variable in its mixing ratio.

Many forms of life on Earth (all multicellular organisms and most single-celled ones) can use oxygen because they are able to break the oxygen-to-oxygen bond in the O2 molecule. In contrast, only a very few species can cleave the strong bond that binds the N atoms in the N2 molecule (e.g. certain specialised bacteria which can use atmospheric nitrogen in protein synthesis). Figure 2.1.2 is interactive and allows you to compare the basic molecular structure of O2 (a) and N2 (b).

The gas we call air is a mixture of many individual gases, but it is predominantly nitrogen with oxygen. Nitrogen makes up a little over 78% of the atmosphere (Table 2.1.1) and is in the form of nitrogen molecules – that is, a pair of nitrogen (N) atoms strongly bonded together. Atomic N has three unpaired electrons and is very reactive, hence the gas usually forms the triple-bonded molecular dinitrogen, or N2. Oxygen, which makes up 21% of the atmosphere, is composed of O2 molecules, in which two oxygen (O) atoms are bonded together, but with a double bond that is not as strong as the bond connecting the N atoms in N2 molecules.

The atmosphere is supplied with gases and small particles (known as aerosols when suspended in a gas) from the interior of the planet by volcanic eruptions, which modify the climate and the surface temperature of the Earth. This natural ‘greenhouse effect’ traps thermal radiation emitted by the Earth and keeps the planet’s surface about 35 °C warmer than it would otherwise be. Importantly, this allows much of the Earth’s water to remain as a liquid, rather than freeze, which is essential for sustaining life.

Without this thin, gaseous envelope, there would be no life on our planet. The lower atmosphere is where weather happens, and therefore determines the climatic conditions underlying the distribution and function of life.

When looking skyward, you might think that the atmosphere extends for considerable distance but, in fact, relative to the size of the Earth, the atmosphere is very thin (Figure 2.1.1). The diameter of the Earth is about 12 700 km, but 99% of the Earth’s atmosphere is within 30 km of the surface. The mass of the whole atmosphere is much less than one-millionth of the total mass of the Earth. The atmosphere becomes exponentially less dense with distance from the surface, and half of the total mass of the atmosphere lies within about 5.6 km of the Earth’s surface. Although the atmosphere does extend to over 100 km from the Earth’s surface, it becomes extremely thin by this point.

The three principles of tidy data are:each variable forms a columneach observation forms a rowdifferent types of observations are stored in separate tables.

every dataset is made up of values: the numbers or text that are recorded when we collect dataeach value is part of an observation: the thing that we collect data abouteach observation has one or more associated variables: the attributes of the observation.

Comparison or control: this describes the way the two variables are considered. If they are both continuous (e.g. age and body length), a relationship between the two variables could be appropriate. A difference is used if at least one variable is categorical (e.g. gender), if an experimental treatment is compared to a control (e.g. difference in fungal growth on apples treated with a fungicide versus the control where no fungicide was added) or if two sites or interventions are being compared (e.g. difference in soil pH under horse chestnut versus Scots pine trees).

It is clear from the projections depicted in Video 1.4.5 that there will be dramatic changes in the chemistry and biology of the oceans in coming decades, even if conditions do not change to the extent that coral reefs and the shells of other organisms in the surface oceans actually dissolve. It is for this reason that the planetary boundary for cap omega sub arag is set at ≥80% of the preindustrial average of 3.44. At the time of writing (2025), the best estimate of this measure is around 2.8, approximately 81% of the preindustrial value and just a fraction above the boundary of 2.75. This is one planetary boundary that is on the verge of being breached, and it is only a matter of time before that happens.

Values of cap omega sub arag greater than 1 favour precipitation of aragonite, while values less than 1 favour dissolution. The following video shows historic and projected global trends in surface water aragonite saturation state. Some parts of the oceans (primarily the poles) have historically low cap omega sub arag , but most of the temperate and tropical oceans have values greater than 3 (colour blue in the video). This changes over time, with much of the oceans forecast to be below 3 by the end of the century. Although cap omega sub arag <1 favours inorganic dissolution of aragonite, values <3 make the production of aragonite by marine organisms energetically much more expensive.

So, the ocean acidification planetary boundary relates to the saturation state of aragonite in the surface waters. The aragonite saturation state refers to the concentration of dissolved carbonate ions in relation to the solubility of aragonite. It is referred to by the symbol normal cap omega sub a times r times a times g (where normal cap omega is the Greek letter capital omega). It is calculated using the formula: cap omega sub arag equals left square bracket cap c a super two postfix plus right square bracket times left square bracket cap c cap o sub three super two postfix minus right square bracket divided by cap k sub sp super prime

where left square bracket cap c a super two postfix plus right square bracket and left square bracket cap c cap o sub three super two postfix minus right square bracket are the concentration of their respective ions and cap k sub sp super prime is the ‘apparent solubility product’ – the equilibrium constant for the dissolution of the compound, in this case aragonite. The important thing to take away from this is that, other things being equal, the saturation state is dependent on the concentration of calcium and carbonate ions, which, as you learned in Study session 1.3.1, vary with changing CO2 concentration and pH. Furthermore, cap k sub sp super prime increases with temperature, so in warmer seas (as expected with climate change), if calcium and carbonate ion concentrations stay the same, cap omega sub arag would decrease. Overall, however, changes in ion concentrations are expected to be the main influence on cap omega sub arag as our climate changes.

The crystal structure of the two minerals differs. Calcite forms blocky crystals while aragonite forms needle-like crystals. Calcite is the more stable form of CaCO3 in most conditions and is by far the most abundant form in rocks. It is the major component of most limestone. However, the presence of magnesium ions in solution in seawater alongside calcium ions favours the formation of aragonite. Although many organisms can form both calcite and aragonite in their shells and exoskeletons, going against the energetically favoured form in any environment requires greater energy expenditure by the organism. Marine conditions in Earth’s oceans have favoured organisms that use aragonite predominantly over calcite in their hard structures. This is important in understanding the effects of ocean acidification because aragonite is less stable and more prone to dissolution than calcite. Over geological time and under certain conditions, aragonite can convert to (or dissolve and re-precipitate as) calcite, which is one reason why limestone rocks, made from the bodies of marine organisms, predominantly contain calcite.

The global picture of ocean hypoxia matches the patterns evident from the two examples above. Coastal zones which drain large areas of croplands and those in shallow seas are where most hypoxia are found (Figure 1.4.7).

The second part of the interview (from 7 minutes 5 seconds) describes the processes involved in causing hypoxia at this location. The questions that follow focus on this part of the interview. You may wish to make notes on this part to help you answer them.

The first part of this interview reviews what ocean hypoxia is and describes some of the effects on marine organisms. Listen to the first part (up to 7 minutes 5 seconds) to set the scene for the activity.

series of thought experiments and mathematics. It is an example of deductive reasoning

many observations. It sought to explain the variety observed in nature at small and large scales. It is an example of inductive reasoning