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