Stratospheric ozone depletion: CFC chemistry, the Antarctic ozone hole, and the Montreal Protocol

Intuition Beginner

High above the ground, between about 15 and 30 kilometers, a thin veil of ozone gas shields the surface of the Earth. Ozone is a molecule built from three oxygen atoms. In this part of the atmosphere, called the stratosphere, ozone absorbs the most damaging part of sunlight: ultraviolet radiation. Without this shield, ultraviolet rays would reach the ground in doses that would burn skin, damage eyes, harm crops, and kill the microscopic plankton that anchor ocean food webs.

In the 1970s, scientists began to suspect that human-made chemicals could eat away at this shield. The chemicals were chlorofluorocarbons, known as CFCs, used in refrigerators, air conditioners, foam packaging, and spray cans. CFCs were prized because they were stable and non-toxic. That very stability became the problem. CFCs survive for years, drift up to the stratosphere, and there, struck by intense sunlight, they break apart and release chlorine atoms.

Each chlorine atom acts as a catalyst. It destroys an ozone molecule, is regenerated, and destroys another, over and over. A single chlorine atom can wreck about 100,000 ozone molecules before it is finally locked away into a stable compound. In 1985 a team measuring the atmosphere over Antarctica found that nearly half the ozone had vanished each spring. The discovery shocked the world and led to a treaty, the Montreal Protocol, that phased out CFCs.

The ozone layer is now recovering. The Montreal Protocol is often called the most successful environmental agreement in history. It worked because the underlying chemistry was well understood, the chemicals came from a small set of industries, and affordable replacements were within reach. This unit tells the full story: the chemistry, the Antarctic hole, the diplomatic response, and why the same physics still matters for climate today.

Visual Beginner

Stage	What happens	Where	Key species
Formation	Ultraviolet light splits O2; fragments join O2 to make O3	Mid stratosphere	O, O3
Shielding	O3 absorbs ultraviolet-B and ultraviolet-C	15-30 km altitude	O3
CFC release	Stable CFCs drift upward for years	Troposphere to stratosphere	CF2Cl2
Catalytic loss	One Cl atom destroys ~100,000 O3 molecules	Global stratosphere	Cl, ClO
Polar hole	Cold clouds activate chlorine; spring sun triggers loss	Antarctic vortex	Cl2, ClO
Recovery	Montreal Protocol phases out CFCs	Worldwide policy	HCFCs, HFCs

Worked example Beginner

How much ozone can a single chlorine atom destroy? The catalytic cycle turns chlorine into a reusable razor: each pass consumes one ozone molecule but regenerates the chlorine atom, ready for the next pass. The chlorine is not used up; it is recycled.

Suppose a chlorine atom completes 100,000 cycles before a terminating reaction locks it into hydrogen chloride. Then that one atom has destroyed about 100,000 ozone molecules. The catalytic cycle is the reason a trace amount of chlorine, measured in parts per trillion, can dismantle a planetary ultraviolet shield.

A molecule of CFC-12 ($C{F}_{2}C{l}_2$) holds two chlorine atoms. If both are freed by sunlight in the stratosphere, the single CFC molecule can destroy roughly $2\times 100,000=200,000$ ozone molecules over its lifetime. This multiplier is why a gas present in vanishingly small concentrations can punch a hole in a layer that shields all life from ultraviolet radiation.

Check your understanding Beginner

Formal definition Intermediate+

Stratospheric ozone is the triatomic oxygen ($O_3$) residing between roughly 15 and 30 km altitude, maintained in photochemical steady state by the competition between production by solar ultraviolet photolysis of $O_2$ and destruction by catalytic cycles. The column abundance is measured in Dobson units (DU), where 1 DU corresponds to a layer of pure ozone $10\mu \mathrm{m}$ thick at standard temperature and pressure; a healthy global column is about 300 DU.

The Chapman cycle

The pure-oxygen photochemistry proposed by Sydney Chapman in 1930 ^{[Chapman 1930]} consists of four reactions:

$$O_2+h\nu \stackrel{J_2}{\to }2O(R1)$$ $$O+O_2+M\stackrel{k_2}{\to }{O}_3+M(R2)$$ $$O_3+h\nu \stackrel{J_3}{\to }O+O_2(R3)$$ $$O+O_3\stackrel{k_3}{\to }2O_2(R4)$$

Here $h\nu$ denotes a photon, $J_2$ and $J_3$ are photolysis rate constants, $k_2$ and $k_3$ are bimolecular (or termolecular, for $R2$ with third body $M$) rate constants. Reactions $R1$ and $R2$ produce ozone; $R3$ photolyzes it but regenerates an oxygen atom that can remake ozone; $R4$ is the only true loss of odd oxygen.

Catalytic destruction cycles

The dominant loss pathways are catalytic chains of the form

$$X+O_3\to XO+O_2,XO+O\to X+O_2,$$

with net reaction $O_3+O\to 2O_2$, identical to $R4$. The catalyst $X$ is regenerated, so a single $X$ atom mediates the conversion many times before being sequestered. The four canonical catalyst families are:

• Hydrogen oxide ($H{O}_x$): $X=OH,H{O}_2$, from water and methane photolysis.
• Nitrogen oxide ($N{O}_x$): $X=NO$, from nitrous oxide photolysis, identified by Crutzen in 1970 ^{[Crutzen 1970]}.
• Chlorine ($Cl{O}_x$): $X=Cl$, from CFC photolysis, predicted by Molina and Rowland in 1974 ^{[Molina and Rowland 1974]}.
• Bromine ($Br{O}_x$): $X=Br$, from halons and methyl bromide; bromine is even more efficient per atom than chlorine.

CFC photolysis

Chlorofluorocarbons, synthesized by Thomas Midgley in 1928, are fully halogenated alkanes (for example CFC-12 is $C{F}_{2}C{l}_2$). They have no tropospheric sink: they are insoluble, non-flammable, and unreactive with hydroxyl radical. In the stratosphere, ultraviolet photons with wavelengths below about 230 nm cleave a carbon-chlorine bond:

$$C{F}_{2}C{l}_2+h\nu \to C{F}_{2}Cl+Cl.$$

The liberated chlorine atom enters the catalytic cycle. Because each chlorine atom survives roughly $10^5$ catalytic turns before termination as hydrogen chloride ($HCl$) or chlorine nitrate ($ClON{O}_2$), the effective yield of ozone destruction per CFC molecule is enormous.

Polar stratospheric clouds and heterogeneous activation

Polar stratospheric clouds (PSCs) form when stratospheric temperatures fall below about 195 K, conditions met only inside the winter polar vortex. Two main types matter. Type I PSCs are nitric acid trihydrate ($NAT$, $HN{O}_3\cdot 3H_{2}O$) particles near 195 K. Type II PSCs are water-ice particles near 188 K. On the surfaces of these particles, heterogeneous reactions convert the inert chlorine reservoirs into active, photolabile chlorine:

$$HCl+ClON{O}_2\stackrel{\text{PSC}}{\to }C{l}_2+HN{O}_3,$$ $$ClON{O}_2+H_{2}O\stackrel{\text{PSC}}{\to }HOCl+HN{O}_3.$$

The nitric acid remains in the condensed phase, denitrifying the gas and preventing re-formation of $ClON{O}_2$. When sunlight returns in spring, $C{l}_2$ and $HOCl$ photolyze within hours to release massive amounts of atomic chlorine, driving the catastrophic ozone loss observed by Farman in 1985 ^{[Farman 1985]} and mechanistically explained by Solomon and colleagues in 1986 ^{[Solomon 1986]}.

Key result: the catalytic amplifier and the Chapman discrepancy Intermediate+

Two quantitative facts dominate stratospheric ozone chemistry. First, the pure-oxygen Chapman cycle over-predicts observed ozone by roughly a factor of two, and the missing sink is precisely the catalytic destruction introduced above. Second, a single chlorine atom destroys on the order of $10^5$ ozone molecules before being sequestered.

For the first fact, set the Chapman reactions to steady state (derived in full in the proof set). The steady-state ozone density in the mid-stratosphere is

$$[O_3{]}_{ss}={\left(\frac{J_2{k}_2[O_2{]}^2[M]}{k_3{J}_3}\right)}^{1/2}.$$

Plugging in representative values ($J_2\sim 10^{-11}{\mathrm{s}}^{-1}$, $k_2\sim 10^{-33}\mathrm{c}{\mathrm{m}}^6\mathrm{molecul}{\mathrm{e}}^{-2}{\mathrm{s}}^{-1}$, $k_3\sim 10^{-15}\mathrm{c}{\mathrm{m}}^3\mathrm{molecul}{\mathrm{e}}^{-1}{\mathrm{s}}^{-1}$, $J_3\sim 10^{-3}{\mathrm{s}}^{-1}$, with ambient $[O_2]$ and $[M]$ near 25 km) yields a column that exceeds observations by about a factor of two. The discrepancy is resolved by adding catalytic loss terms $2{\sum }_X{k}_X[O_3][X]$ to the odd-oxygen budget; when the measured radical concentrations ($OH$, $NO$, $Cl$, $Br$) are inserted, the steady state matches the observed column. This is the central quantitative argument that catalytic chemistry, not transport alone, controls the ozone abundance.

For the second fact, the catalytic cycle turns over on a timescale of minutes to hours in the sunlit stratosphere, whereas reservoir formation (the terminating reactions $Cl+C{H}_4\to HCl+C{H}_3$ and $ClO+N{O}_2\to ClON{O}_2$) operates on a timescale of days to weeks. The ratio of these timescales gives a chain length of roughly $10^4$ to $10^5$ ozone molecules destroyed per active chlorine atom. This amplification is what converts the part-per-trillion mixing ratios of stratospheric chlorine into a globally significant loss of ozone column.

Bridge. The Chapman steady state builds toward the full catalytic-chemistry framework: because pure-oxygen photochemistry over-predicts column ozone by roughly a factor of two, the foundational reason we need the catalytic destruction cycles is that no purely gas-phase oxygen reaction can close the budget. This chain-mediated loss appears again in the heterogeneous PSC chemistry of the Antarctic vortex, where reservoir activation pushes the same $Cl/ClO$ cycle to catastrophic rates. The central insight is that a regenerative catalyst multiplies a single atom's impact by its chain length; putting these together, the global ozone budget and the polar hole are two regimes of one mechanism, and the same reasoning generalises to every stratosphere in the solar system.

Exercises Intermediate+

Advanced results Master

Heterogeneous activation kinetics on PSC surfaces

The heterogeneous reactions on PSCs are surface-catalyzed and follow a reaction probability $\gamma$ formalism: the rate is $\gamma \cdot (\text{collision flux})\cdot (\text{surface area density})$. Measured reaction probabilities for $ClON{O}_2+HCl$ on NAT are of order $10^{-3}$ to $10^{-1}$ depending on temperature and surface composition, rising sharply as Type II ice clouds form below 188 K. The temperature sensitivity is so steep that a few degrees of cooling can multiply the activation rate by an order of magnitude, which is why a colder-than-average Antarctic winter produces a markedly deeper hole. The product $C{l}_2$ is volatile and escapes the particle into the gas phase, while $HN{O}_3$ remains condensed, irreversibly denitrifying the gas parcel as long as the particles sediment downward.

Denitrification and the sequestration of nitrogen

Denitrification, the removal of $HN{O}_3$ from the gas phase into sedimenting PSC particles, is the second key ingredient of the polar hole. As long as gas-phase $HN{O}_3$ is present, photolysis releases $N{O}_2$, which reacts with $ClO$ to reform $ClON{O}_2$ and shut down the catalytic loss. By sequestering nitrogen, denitrification removes this brake: even after the PSCs evaporate, the gas-phase $N{O}_y$ is depleted, and $ClO$ remains active for weeks. Large NAT particles sediment efficiently and denitrify deeply; smaller particles remain suspended and release $HN{O}_3$ back upon evaporation, renitrifying the parcel. The balance between these regimes controls the interannual depth of the hole.

Why the Antarctic, not the Arctic

The Antarctic ozone hole is far more severe and reliable than its Arctic counterpart for three coupled reasons. First, the Antarctic continent is surrounded by a cold ocean with no topographic interruptions, producing a stronger, colder, more isolated polar vortex than the Arctic, whose vortex is distorted by mountain ranges and land-sea contrasts in the Northern Hemisphere. Second, Antarctic winter stratospheric temperatures regularly fall below the 195 K threshold for Type I PSCs and sometimes below 188 K for Type II ice, whereas Arctic temperatures hover near the threshold and frequently exceed it. Third, the more isolated Antarctic vortex resists injection of ozone-rich mid-latitude air, so the depleted polar air persists through spring. The Arctic shows episodic depletion in cold years (for example 2011 and 2020) but not the systematic annual hole of the south.

The Montreal Protocol and its amendments

The Montreal Protocol on Substances that Deplete the Ozone Layer was signed in 1987 and entered into force in 1989 ^{[Andersen and Sarma 2002]}. It mandated phased reductions in CFC production, with the London Amendment (1990) tightening the schedule and adding controls, the Copenhagen Amendment (1992) accelerating the phase-out and regulating methyl bromide and HCFCs, and subsequent amendments (Beijing 1999, Kigali 2016) extending controls to hydrochlorofluorocarbons and, eventually, hydrofluorocarbons. HFCs, the replacement refrigerants, do not deplete ozone but are potent greenhouse gases; the Kigali Amendment therefore links ozone protection to climate policy. Atmospheric chlorine loading peaked in the late 1990s at about 4 parts per billion and has declined since; the ozone layer is projected to recover to 1980 levels over Antarctica by roughly the 2060s.

Equivalent effective stratospheric chlorine and the ozone-depletion potential

The combined destructive impact of all halogens is summarized by equivalent effective stratospheric chlorine (EESC), which weights each species by its ozone-depletion potential (ODP), its atmospheric lifetime, and the number of halogen atoms released. CFC-11 has ODP 1.0 by definition; CFC-12 is about 1.0; halon-1301 is about 10; HCFCs are typically 0.01 to 0.1; HFCs and the chlorine-free replacements have ODP zero. EESC trajectories reconstructed from surface measurements track the protocol's success: the decline of EESC since 1996 is the direct atmospheric signature of the treaty's effect.

Coupling to climate: ozone-depleting substances as greenhouse gases

Ozone-depleting substances are also strong infrared absorbers. Their positive radiative forcing has been estimated at roughly $0.3$ to $0.4\mathrm{W}{\mathrm{m}}^{-2}$ at present, comparable in magnitude to the forcing from methane; conversely, stratospheric ozone depletion itself produced a small negative forcing. The net effect of the Montreal Protocol on climate has therefore been protective in two ways: by allowing ozone recovery and by phasing out gases that would otherwise have added substantial warming. Some replacements, notably HFC-23 and HFC-134a, carry large global warming potentials, motivating the Kigali phase-down. Ozone recovery and climate mitigation are thus coupled problems governed by overlapping gas inventories.

The ozone-climate feedback and the Brewer-Dobson circulation

A changing climate alters stratospheric transport. Greenhouse-gas cooling of the stratosphere (the surface warms while the stratosphere cools, because increased $C{O}_2$ enhances infrared emission to space aloft) paradoxically favors ozone recovery, because the catalytic loss reactions slow at lower temperatures outside the polar regions, and because the colder lower stratosphere shifts the $H{O}_x$ and $N{O}_x$ equilibria toward less destructive partitions. Meanwhile, a strengthened Brewer-Dobson circulation, driven by changing wave activity, accelerates the ascent of tropical ozone and its transport toward the poles. These feedbacks complicate the recovery trajectory and are the leading source of uncertainty in twenty-first-century ozone projections.

Synthesis. The ozone story builds toward a general theory of catalytic atmospheric chemistry: any long-lived radical that regenerates after each reaction will amplify its impact by the chain length. The foundational reason the hole is polar is that the Antarctic vortex provides the cold, isolated reactor in which heterogeneous activation releases the catalyst; this is exactly why a warmer, leakier Arctic shows only intermittent depletion. The central insight unifying gas-phase and heterogeneous chemistry is that reservoir species, not active radicals, set the budget. Putting these together, the same catalytic logic generalises to every planetary stratosphere and appears again in the climate-forcing role of the replacement HFCs; the bridge is the recognition that atmospheric self-cleansing and atmospheric self-destruction are the same machinery running in opposite directions.

Full proof set Master

Proposition (Chapman steady-state ozone and its over-prediction)

Under the four Chapman reactions $R1$ to $R4$ with rate constants $J_2,k_2,J_3,k_3$, and assuming the oxygen atom is in fast photochemical steady state, the steady-state ozone number density is

$$[O_3{]}_{ss}={\left(\frac{J_2{k}_2[O_2{]}^2[M]}{k_3{J}_3}\right)}^{1/2},$$

which over-predicts the observed mid-stratospheric ozone column by approximately a factor of two.

Proof. Define odd oxygen $O_x=[O]+[O_3]$. The production of $O_x$ comes only from $R1$, which creates two oxygen atoms per photon absorbed, so

$$P(O_x)=2J_2[O_2].$$

The only loss of $O_x$ is $R4$, which consumes one $O$ and one $O_3$, removing two $O_x$ per reaction, so

$$L(O_x)=2k_3[O][O_3].$$

The oxygen atom is highly reactive and attains a fast internal steady state set by the balance of $R2$, $R3$, and the smaller contributions of $R1$ and $R4$. In the mid-stratosphere the dominant balance is between production of $O$ by $R3$ ($J_3[O_3]$) and loss by $R2$ ($k_2[O][O_2][M]$), because $R2$ and $R3$ are orders of magnitude faster than $R1$ and $R4$. Setting $d[O]/dt=0$ under this dominant balance,

$$J_3[O_3]=k_2[O][O_2][M]⟹[O]=\frac{J_3[O_3]}{k_2[O_2][M]}.$$

Now impose $O_x$ steady state, $P(O_x)=L(O_x)$:

$$2J_2[O_2]=2k_3[O][O_3]⟹J_2[O_2]=k_3[O][O_3].$$

Substitute the expression for $[O]$:

$$J_2[O_2]=k_3\cdot \frac{J_3[O_3]}{k_2[O_2][M]}\cdot [O_3]=\frac{k_3{J}_3[O_3{]}^2}{k_2[O_2][M]}.$$

Solving for $[O_3{]}^2$,

$$[O_3{]}^2=\frac{J_2[O_2]\cdot k_2[O_2][M]}{k_3{J}_3}=\frac{J_2{k}_2[O_2{]}^2[M]}{k_3{J}_3},$$

and taking the square root yields the stated expression. Inserting representative values for 25 km altitude, $J_2\approx 10^{-11}{\mathrm{s}}^{-1}$, $k_2\approx 10^{-33}\mathrm{c}{\mathrm{m}}^6\mathrm{molecul}{\mathrm{e}}^{-2}{\mathrm{s}}^{-1}$, $k_3\approx 10^{-15}\mathrm{c}{\mathrm{m}}^3\mathrm{molecul}{\mathrm{e}}^{-1}{\mathrm{s}}^{-1}$, $J_3\approx 10^{-3}{\mathrm{s}}^{-1}$, $[O_2]\approx 3\times 10^{17}\mathrm{c}{\mathrm{m}}^{-3}$, $[M]\approx 4\times 10^{17}\mathrm{c}{\mathrm{m}}^{-3}$, gives $[O_3{]}_{ss}\approx 5\times 10^{12}\mathrm{c}{\mathrm{m}}^{-3}$, roughly twice the observed $\sim 2$ to $3\times 10^{12}\mathrm{c}{\mathrm{m}}^{-3}$. The discrepancy, the famous factor-of-two Chapman over-prediction, is the principal empirical evidence that additional catalytic loss terms (the $H{O}_x$, $N{O}_x$, $Cl{O}_x$, $Br{O}_x$ cycles) are required to close the odd-oxygen budget. $\square$

Proposition (catalytic chain length of an active chlorine atom)

Let ${\tau }_c$ be the mean time for one turnover of the $Cl/ClO$ catalytic cycle and ${\tau }_r$ the mean time for the active chlorine to be sequestered into a reservoir. Then the mean number of ozone molecules destroyed before sequestration is

$$N\approx \frac{{\tau }_r}{{\tau }_c},$$

and over a full stratospheric residence, with reactivation, the cumulative yield reaches $\sim 10^5$.

Proof. The catalytic cycle consists of $Cl+O_3\to ClO+O_2$ (rate $k_a$) followed by $ClO+O\to Cl+O_2$ (rate $k_b$). The chlorine atom alternates between the $Cl$ and $ClO$ forms; one full cycle consumes one $O_3$ (and one $O$) and regenerates the $Cl$. The mean turnover time is set by the slower of the two steps in the relevant regime. In the sunlit lower stratosphere, the second step is typically rate-limiting because atomic oxygen is scarce, giving ${\tau }_c=1/(k_b[O])$ of order minutes to hours.

Separately, the active chlorine is lost to a reservoir by reactions such as $Cl+C{H}_4\to HCl+C{H}_3$ (rate $k_c[C{H}_4]$) and $ClO+N{O}_2+M\to ClON{O}_2+M$ (rate $k_d[N{O}_2][M]$), giving a sequestration time ${\tau }_r=1/(k_c[C{H}_4]+k_d[N{O}_2][M])$ of order days to weeks.

Assuming the catalytic and sequestration processes are Poisson and independent, the expected number of catalytic turns before a sequestration event is the ratio of the rates:

$$N=\frac{\text{catalytic rate}}{\text{sequestration rate}}=\frac{1/{\tau }_c}{1/{\tau }_r}=\frac{{\tau }_r}{{\tau }_c}.$$

With ${\tau }_r\approx 10^6\mathrm{s}$ (about two weeks) and ${\tau }_c\approx 10^3$ to $10^4\mathrm{s}$, one obtains $N\approx 10^2$ to $10^3$ molecules per active episode. Over a stratospheric residence of one to two years, the reservoir species $HCl$ and $ClON{O}_2$ are themselves photolyzed or react heterogeneously, reactivating the chlorine many times, so the cumulative yield per chlorine atom over its full atmospheric lifetime reaches $\sim 10^5$ molecules. This is the quantitative content of the statement that a single chlorine atom destroys roughly one hundred thousand ozone molecules. $\square$

Connections Master

The natural prerequisite is the climate-change foundation unit 27.07.01, which frames ozone depletion alongside the broader greenhouse problem and the distinction between ozone-layer protection (ultraviolet shielding) and climate forcing (infrared trapping). The two are often confused in public discussion but are mechanistically distinct; this unit sharpens that boundary while tracing where they couple (through the radiative effect of ozone-depleting substances and the HFC replacements).

The quantitative backbone of ozone chemistry is radiative transfer and the Stefan-Boltzmann energy balance developed in 27.07.02 pending. The Dobson unit, the photolysis rate constants $J_2$ and $J_3$, and the coupling of ozone-depleting substances to radiative forcing all rest on the radiative-transfer machinery of that unit; the catalytic-chemistry budget layered on top is what closes the odd-oxygen steady state that radiative transfer alone leaves over-predicted.

The atmospheric context, the structure of the troposphere and stratosphere, and the dynamics of the polar vortex are developed in 27.04.01. The vertical temperature profile that places the ozone layer at 15 to 30 km, the wind systems that form the isolated Antarctic vortex, and the general circulation that governs stratospheric residence times are all standard material of that atmosphere-and-weather unit.

The chemical kinetics of radicals, catalysis, and reaction chains generalize the gas-phase and surface chemistry studied under atmospheric chemistry 27.07.03 pending, and the recovery of the ozone layer is itself a measurable climate proxy whose trajectory is tracked with the paleoclimate and atmospheric-measurement techniques of that unit.

Historical & philosophical context Master

Chapman's pure-oxygen theory

Sydney Chapman proposed in 1930 that the ozone layer could be explained by four reactions among pure-oxygen species driven by solar ultraviolet ^{[Chapman 1930]}. The elegance of the theory was its prediction that ozone is a steady-state photochemical product, not a primordial residue. Its limitation, recognized only decades later, was the factor-of-two over-prediction: the Chapman scheme produced too much ozone. The missing ingredient, catalytic destruction, was supplied incrementally by the work of Crutzen on $N{O}_x$ ^{[Crutzen 1970]}, then by Bates and Nicolet on $H{O}_x$, and finally by Molina and Rowland on $Cl{O}_x$. Each discovery revealed that trace radicals, not bulk oxygen chemistry, controlled the layer.

The Molina-Rowland prediction and the DuPont response

In June 1974, Mario Molina and F. Sherwood Rowland published in Nature the prediction that chlorofluoromethanes would photolyze in the stratosphere, releasing chlorine that would catalytically destroy ozone ^{[Molina and Rowland 1974]}. Their argument was theoretical, resting on the known inertness of CFCs and the known catalytic chemistry of chlorine. The prediction was contested by the CFC industry, which argued that the chain length was overstated and that natural sinks would remove chlorine. A decade of atmospheric measurement vindicated Molina and Rowland decisively.

Farman's Antarctic shock and the Solomon mechanism

In 1985, Joe Farman, Brian Gardiner, and Jonathan Shanklin reported in Nature that the springtime total ozone over Halley Bay, Antarctica, had fallen by about 40 percent since the late 1970s, a loss no model had predicted ^{[Farman 1985]}. The Tomsat satellite had recorded the hole but its data had been rejected by quality-control algorithms as implausible. Susan Solomon and colleagues explained the phenomenon in 1986 by the heterogeneous activation of chlorine on polar stratospheric clouds ^{[Solomon 1986]}, a mechanism that united gas-phase catalytic theory with the thermodynamics of the polar vortex. In 1995, Rowland, Molina, and Crutzen shared the Nobel Prize in Chemistry for this body of work.

The Protocol as a model of precautionary governance

The Montreal Protocol, signed only two years after Farman's discovery, is studied in international relations as the paradigm of precautionary environmental governance ^{[Andersen and Sarma 2002]}. Several features distinguished it: a small number of producer firms, the availability of technical substitutes, a clear scientific consensus, and a financing mechanism (the Multilateral Fund) that supported developing-country compliance. It contrasts instructively with the climate regime, where the emitting industries are pervasive, the substitutes are contested, and the science, while robust, has been politically contested for decades. The ozone layer's recovery is the empirical demonstration that coordinated phase-out of a pollutant can reverse a global atmospheric change.

Bibliography Master

1. Chapman, S. (1930). "A theory of upper-atmospheric ozone." Memoirs of the Royal Meteorological Society, 3, 103-125.

2. Crutzen, P. J. (1970). "The influence of nitrogen oxides on the atmospheric ozone content." Quarterly Journal of the Royal Meteorological Society, 96, 320-325.

3. Molina, M. J. and Rowland, F. S. (1974). "Stratospheric sink for chlorofluoromethanes: chlorine atom-catalysed destruction of ozone." Nature, 249, 810-812.

4. Farman, J. C., Gardiner, B. G. and Shanklin, J. D. (1985). "Large losses of total ozone in Antarctica reveal seasonal ClOx/NOx interaction." Nature, 315, 207-210.

5. Solomon, S., Garcia, R. R., Rowland, F. S. and Wuebbles, D. J. (1986). "On the depletion of Antarctic ozone." Nature, 321, 755-758.

6. Anderson, J. G., Toohey, D. W. and Brune, W. H. (1991). "Free radicals within the Antarctic vortex: the role of CFCs in Antarctic ozone loss." Science, 251, 39-46.

7. Jacob, D. J. (1999). Introduction to Atmospheric Chemistry. Princeton University Press.

8. World Meteorological Organization (2018). Scientific Assessment of Ozone Depletion: 2018. WMO Global Ozone Research and Monitoring Project, Report No. 58.

9. Andersen, S. O. and Sarma, K. M. (2002). Protecting the Ozone Layer: The United Nations History. Earthscan.

10. Solomon, S. (1999). "Stratospheric ozone depletion: a review of concepts and history." Reviews of Geophysics, 37, 275-316.

11. Newman, P. A., Nash, E. R. and Rosenfield, J. E. (2001). "What controls the temperature of the Arctic stratosphere during the spring?" Journal of Geophysical Research, 106, 19999-20010.

12. Velders, G. J. M., Anderson, S. O., Daniel, J. S., Fahey, D. W. and McFarland, M. (2007). "The importance of the Montreal Protocol in protecting climate." Proceedings of the National Academy of Sciences, 104, 4814-4819.

13. Tarbuck, E. J. and Lutgens, F. K. (2018). Earth Science (15th ed.). Pearson.