SCOPE 13 - The Global Carbon Cycle

ABSTRACT

The mean global mixing ratios for carbon monoxide and methane are 0.11 ppm (Seiler, 1974) and 1.41 ppm (Ehhalt, 1974), respectively. These values are equivalent to a total amount of 530 x 10¹² g of CO (or 227 x 10¹² g C) and 4000 x 10¹² g of CH₄ (or 3000 x 10¹² g C) present in the atmosphere. Ehhalt (1974) has estimated that the CH₄ cycle contributes approximately 12% to the atmospheric carbon cycle, which applies even to the CO cycle because of photochemical coupling of both the CH₄ and the CO cycle. Recently, the cycles of these gases have been described in detail by Seiler (1974) and Jaffe (1975) for CO, and by Ehhalt (1974) for CH₄. Other trace gases containing carbon, such as non-CH₄ hydrocarbons, are usually present in the atmosphere by less than three orders of magnitude compared to CH₄ (Robinson et al., 1973). Recently, Duce (1978) has estimated the global tropospheric content of non-CH₄ organic carbon at about 50 x 10¹² g C, which is probably only accurate within a factor of 3.

4.1 CARBON MONOXIDE

Each year, more CO is released into the atmosphere than any other pollutant (excluding CO₂), and each year the quantity released increases. One would, therefore, expect a gradual increase in ambient CO levels, but the background levels of CO, based upon analyses of ice samples in the Arctic and Antarctic, do not appear to have fluctuated over a period of many centuries (Robbins et al., 1973). To maintain the global balance of CO, one or more major active sinks for CO must be operating in the troposphere.

4.1.1 Global CO Background Levels and Distribution

Background measurements of CO in marine air masses over the Atlantic and Pacific Oceans range from 0.04 to 0.30 ppm (Robinson and Robbins, 1968a, 1970; Swinnerton et al., 1969, 1970a, b; Seiler and Junge, 1970; Lamontagne et al., 1971; Seiler, 1974; Swinnerton and Lamontagne, 1974b). In continental air masses under clean-air conditions, values within the same range have been found (Cavanagh et al., 1969; Seiler and Junge, 1970). The CO concentrations in the northern hemisphere are considerably higher than in the southern hemisphere. Maximum background levels of 0.20 ppm at 40^° to 50^° N latitude are related to the larger northernhemisphere anthropogenic CO pollution sources. From this data and additional aircraft measurements in the upper troposphere and lower stratosphere, considering only marine air masses, Seiler (1974) has calculated the latitudinal distribution of the mean tropospheric CO mixing ratio. He has averaged the CO mixing ratios at 0.15 ppm for the northern and 0.06 ppm for the southern hemisphere with an overall global mixing ratio of 0.11 ppm. Identical values have been calculated by Robinson and Robbins (1970) from marine air measurements in the North and South Pacific.

The CO mixing ratio in the upper troposphere (Seiler, 1974) has not been found to be as uniformly distributed as assumed previously (Junge et al., 1971a). In the northern hemisphere, at middle latitudes, fluctuations ranging from 0.12 to 0.22 ppm with an average of 0.15 ppm have been observed. In the lower stratosphere, the CO mixing ratio decreases very rapidly to a value of about 0.05 ppm at 1 km above the tropopause. From the calculations of Hesstvedt (1970), a mixing ratio of 0.1 ppm, is obtained in the troposphere 2 km below the tropopause, while 2 km above the tropopause, the mixing ratio has dropped to 0.02 ppm. More recently, the CO mixing ratio in the upper stratosphere between 40 and 50 km altitude has been determined at 0.05 ± 0.01 ppm using a cryogenic sampling technique (Ehhalt et al., 1975). From these measurements, it is obvious that the stratosphere acts as a sink for CO.

4.1.2 Anthropogenic Sources of CO

Carbon monoxide is produced globally in large quantities, due to incomplete combustion of fossil fuels. By far the largest single anthropogenic source of CO is motor-vehicle exhaust. Jaffe (1973, 1975) has estimated the total global anthropogenic CO emissions to be approximately 360 x 10¹² g for the calendar year 1970. This value has to be increased by at least about 25% to 450 x 10¹² g for the present time, due to the increased consumption of fossil fuels (United Nations Statistical Yearbook, 1975), neglecting emission control programmes which have reduced CO emissions. Considering additional anthropogenic CO sources and higher average CO emission factors, Seiler (1974) has estimated the total anthropogenic CO production to be 640 x 10¹² g per year. Using the annual consumption of petrol in both hemispheres as an indicator for the anthropogenic CO production, 85% of this is produced in the northern hemisphere, mainly between 40^° and 60^° N.

A. CO Concentration in Polluted Areas

The concentration of CO in metropolitan areas has been measured by several authors. The CO levels correlate remarkably well with traffic volume and traffic speed; the highest levels are found in places where vehicular traffic is heaviest and vehicular speed is low (Brief et al., 1960; Georgii and Weber, 1962, Lawther et al., 1962; Brice and Roesler, 1966; Chovin, 1967; Johnson et al., 1968, Colucci and Begeman, 1969; Bové and Siebenberg, 1970; Wolf, 1971). Two daily peaks occur, corresponding to the morning and evening traffic rush hours. Peak concentrations are higher on weekdays than on Saturdays and Sundays, corresponding to the relative traffic volumes. Meteorological factors, such as atmospheric stability and wind speed or mechanical turbulence, play a role in the rate of dispersion of ambient CO (McCormick and Xintaras, 1962; Jaffe, 1968; Wright et al., 1975; Chock and Levitt, 1976). Lynn et al. (1967), in a study of CO exposures in a number of U.S. cities, have found that people in moving vehicles, particularly those in heavy traffic, are at times exposed to levels of 50 ppm CO or more. In the Los Angeles basin, some evidence of a general downtrend of ambient CO levels has been found since 1966, due to emission control programmes (Tiao et al., 1975).

4.1.3 Natural Sources of CO

Recent investigations for CO sinks, to explain why the ambient CO concentrations have not been increasing, have resulted in the discovery of new natural sources of CO whose total quantity far exceeds the total mass of anthropogenically produced CO. Stevens et al. (1972) have shown by ¹⁸O and¹³C isotopic analyses of atmospheric CO, that there are at least five major varieties, two with light oxygen and three with heavy oxygen. The two species (which are less ¹⁸O-enriched) are produced year-round and are assumed to originate from the atmosphere itself. Measurements at a number of locations indicate that a constant concentration (0.10 to 0.15 ppm) of these species occurs throughout the world. The three remaining minor species are heavy-oxygen varieties (more ^{1 8} O enrichment), which are produced within a specific season of the year.

A. Atmospheric Oxidation of Methane and Formaldehyde

Atmospheric methane and formaldehyde have lately been suggested as natural sources of CO. It is supposed that methane is converted to CH₃ by reaction with OH or with O(¹D) (with photolysis playing a role at high altitudes). By a threebody reaction, CH₃ is converted to CH₃O₂, which ultimately forms formaldehyde, HCHO. Formaldehyde is photodissociated to form HCO, which reacts with oxygen primarily to form CO. The CO production by this process has been calculated by several authors (McConnell et al., 1971; Weinstock and Niki, 1972; Wofsy et al., 1972; Levy, 1973) at 15005000 x 10¹² g/year, at least more than three times as much as from anthropogenic sources. The average concentration of OH radicals in the troposphere required to achieve this production data is approximately 2.5 x 10⁶ molecules/cm³ (Levy, 1971, 1972, 1973). Warneck (1974) has pointed out that the OH concentration in the troposphere seems to be less than this value by a factor of five. This would result in a lower CO production by the same factor, assuming a 1:1 conversion of methane to CO. Warneck (1976) has calculated the CO production at < 400 x 10¹² g/year. This value seems to be too low because other groups (Crutzen, 1974; Wang et al., 1975; Davis et al., 1976; Perrier et al., 1976) have generally found OH concentrations to be larger than 10⁶ molecules/ cm ³ in direct observations. Using isotopic data, Stevens et al. (1972) have estimated the CO production by methane oxidation to be more than 3000 x 10¹² g/year in the northern hemisphere.

B. CO Production in the Oceans

The oceans have been found to be a source of CO. A number of studies by Swinnerton and co-workers (Swinnerton et al., 1969, 1970a, b; Wilson et al., 1970; Lamontagne et al., 1971; Linnenbom et al., 1973; Swinnerton and Lamontagne, 1974b), and by Seiler and Junge (1970) and Seiler and Schmidt (1974) have shown that the CO content of the surface waters in the Atlantic and Pacific Oceans ranged up to 500 times the calculated atmospheric equilibrium CO. Seiler and Schmidt (1974) have averaged the supersaturation factor by continuous measurements to be 6.8 for the Northern and 21 for the Southern Atlantic Ocean. The highest supersaturations have been found in nutrient-rich water. Undoubtedly, the high supersaturations are due to a CO production in sea-water by microbiological processes. Various organisms, including bacteria (Junge et al., 1971b, 1972), marine algae (Loewus and Delwiche, 1963; Chapman and Tocher, 1966; Troxler and Dokos, 1973), and jellyfish (Wittenberg, 1960; Pickwell et al., 1964; Pickwell, 1970), which produce carbon monoxide have been found in sea-water.

From the supersaturation measurements in the Atlantic and Pacific Oceans, Seiler (1974) has estimated the CO concentration in sea-water to be an average of about 5 x 10^-5 ml CO/1 water, corresponding to an average equilibrium value of 2.5 ppm CO. The average CO mixing ratio in air above the sea-water, by contrast, varies between 0.04 to 0.20 ppm, so that the sea-water is supersaturated on the average by a factor of about 30. The net flux of CO into the atmosphere, resulting from this high supersaturation, has been calculated by Seiler (1974) and Seiler and Schmidt (1974), using the molecular diffusion model through the laminar boundary layer at the water surface (Broecker and Peng, 1971, 1974), to be about 100 x 10¹² g/year, which is divided into 60 x 10¹² g for the southern and 40 x 10¹² g for the northern hemisphere. Using the mean diffusivity D/Z = 1000 m/year (= 0.32 x 10^-2 cm/sec) for CO given by Broecker and Peng (1974), the net flux of CO would be reduced to about 35 x 10¹² g/year. Nearly the same value of 43 x 10¹² g/year has been calculated by Liss and Slater (1974) with another model, assuming that the surface sea-water is supersaturated by a factor of 23. In a more recent study, Seiler and Schmidt (1976) have estimated a net CO flux of between 20 and 120 x 10¹² g/year. Linnenbom et al. (1973), using the same average supersaturation factor of 28.2 as Seiler (1974), have estimated by other calculations a flux value of about 220 x 10¹² g/year with 130 x 10¹² g for the southern and 90 x 10¹² g for the northern hemisphere.

C. CO Production in Rainwater

High supersaturations of dissolved CO in rain-water, relative to the CO mixing ratio in air, have been reported by Swinnerton et al. (1971). The CO concentrations have varied during daytime between 6.0 x 10^-5 ml CO/1 water in unpolluted air (Hawaii) and 140 x 10^-5 ml/l with an average of about 30 x 10^-5 ml/1 in polluted air (Washington, D.C.). Swinnerton et al. assume a CO production by oxidation of organic material in rain-water, as shown by Wilson et al. (1970), or by electric discharge within clouds leading to dissociation of CO₂. Galbally (1972) postulates that photolysis of formaldehyde and higher aldehydes by ultraviolet light produces this excess of carbon monoxide in rain droplets. Seiler (1974), on the contrary, has detected only small amounts of dissolved CO (6 x 10^-5 ml/1) in rain-water collected at Mainz, West Germany, and analysed up to 10 minutes after rainfall. The most surprising finding in Seiler's analyses has been the rapid increase of the dissolved CO in the samples, by a factor of more than 6 within one hour after sampling. Thus, Seiler expects a CO concentration in rain droplets during rainfall lower than the measured one of 6 x 10^-5 ml/1. At present, global estimations on the CO production in rain-water seem to be doubtful.

D. Other Natural CO Sources

Carbon monoxide is also produced by several different natural CO sources, first of all by forest fires and by open burning of agricultural waste. Using data given by Robinson and Robbins (1968b) and other reported data for the U.S.A., Seiler (1974) has estimated the global CO production by these two processes to be 34 x 10¹² and 30 x 10¹² g/year, respectively. The CO production by agricultural burning has been calculated by Jaffe (1973, 1975) to be about 36 x 10¹² g/year, which is included in his global anthropogenic CO production data.

There are indications that CO is also produced by the biosphere. Small quantities of CO are formed by vegetation during seed germination and seedling growth of higher plants (Wilks, 1959; Siegel et al., 1962), and by certain brown marine algae or kelps (Loewus and Delwiche, 1963; Chapman and Tocher, 1966; Delwiche, 1970). Microorganisms have been shown to produce CO from plant flavonoids (Westlake et al., 1961). An important biospherc CO source could be the degradation of chlorophyll in dying plant material. Crespi et al. (1972) have estimated that the annual degradation of chlorophyll yields 60 x 10¹² g/year of CO globally. Together with the bilin biosynthesis (Troxler et al., 1970), plants may be the source of 100 x 10¹² g/year of CO globally. Stevens et al. (1972) have interpreted their isotopic data by an autumnal burst of CO from chlorophyll destroyed in trees and other plants. They have calculated the corresponding CO emission for the northern hemisphere to be 200500 x 10¹² g over a 1.5-month period.

Carbon monoxide is produced, furthermore, in living animals (Landaw, 1970) and men, probably by decomposition of haemoglobin. The production rates vary between 0.5 mg for normal man and 4.3 mg CO/man-hour for patients with haemolytic anaemia (Coburn et al., 1964; Coburn, 1970; White, 1970), which is negligible compared with other natural sources.

Another natural source of CO is the result of photochemical oxidation of hydrocarbons, e.g., terpenes, by O₃ or NO₂ . Went (1966) has estimated that about 1000 x 10¹² g of volatile organics of plant origin are dispersed globally each year. Robinson and Robbins (1968b) have estimated that these amounts result in a CO production of about 60 x 10¹² g/year.`

4.1.4 Sinks of Atmospheric CO

Until 1969, knowledge on the removal processes of CO in the troposphere and lower stratosphere was unsatisfactory. Jaffe (1968) and Robinson and Robbins (1968b) discussed possible sink mechanisms but could not find any of importance. In the meantime, several processes have been reported which can decompose CO in signficant amounts: by the oxidation of CO by OH radicals, the consumption of CO on soil surfaces by microorganisms, and the uptake by plants.

A. Oxidation of CO by OHRadicals

It has been suggested that hydroxyl radicals account for both the production of CO by methane oxidation and for the CO destruction by oxidation to CO₂. All calculations result in a higher CO destruction rate. Weinstock and Niki (1972) have calculated an average CO destruction of 5000 x 10¹² g/year. Similar values can be derived from OH concentrations given by McConnell et al. (1971) and Wofsy et al. (1972). Levy (1973) has published CO loss rates as a function of altitudes and CO mixing ratios in the troposphere, which result in an annual CO loss of 2700 x 10¹² g, assuming an average CO-mixing ratio of 0.11 ppm. With the lower OH concentrations assumed, Warneck (1976) has estimated the CO loss rate to be 680 x 10¹² g/year.

The chemistry of these reactions is quite complex and is based largely on theoretical models. Conflicting publications of many other authors have been reported (Westenburg and de Haaf, 1972; Davis et al., 1973; Goldman et al., 1973; Kummler and Baurer, 1973; Shimazaki and Cadle, 1973; Whitten et al., 1973; Warneck, 1974; Weinstock and Chang, 1974; Sze, 1977). The resolution of production and sink mechanisms for CO by OH radicals must await actual field measurements.

B. Stratospheric Sink for Tropospheric CO

With an average CO mixing ratio of 0.13 ppm at the tropopause, and a constant CO mixing rate of 0.05 ppm in the stratosphere, Warneck et al. (1973) have calculated, by an eddy diffusion model, the CO flux into the stratosphere in the northern hemisphere to be 2.4 x 10¹⁰ molecules/cm² sec, which agrees with estimations by Seiler and Warneck (1972). The corresponding size of the stratospheric sink of CO in the northern hemisphere is 90 x 10¹² g/year. Due to the lower CO mixing ratio of 0.07 ppm in the upper troposphere of the southern hemisphere, the stratospheric sink in the southern hemisphere is considerably lower and amounts to 20 x 10¹² g/year.

C. CO Uptake on the Soil Surface

The oxidation of carbon monoxide to CO₂ by soil bacteria has been found by several authors. Therefore, it has been suggested (Jaffe, 1968) that the soil may be an effective sink for CO. Liebl (1971) has determined an average uptake rate of 1.5 x 10^-11 g/cm² sec over several types of soil and at a soil temperature of 15 ^°C. In similar experiments, Inman et al. (1971) have found uptake rates of between 6 and 47 x 10^-11 g/cm² sec with an average of 23 x 10^-11 g/cm² sec over various soils in the U.S.A. Later studies of the same group (Ingersoll et al., 1974) have shown similarly high uptake rates of between 6 and 73 x 10^-11 (average 32) for soils under natural vegetation and of between 3 and 21 x 10^-11 g/ cm² sec (average 7) for soils under cultivation. The discrepancy between the lower value of Liebl (1971) and the other higher values is caused by the different CO concentrations used. Liebl (1971) has used test atmospheres with a CO mixing ratio of 0.20 ppm,' comparable with the normal CO mixing ratio in air over the continent, while the other group has used an atmosphere with 100 ppm CO. In recently performed detailed measurements on different soil types in the area of Mainz-Wiesbaden, West Germany, similar lower values have been found, from which a global CO consumption rate by soil of 500 x 10¹² g/year has been estimated (Liebl and Seiler, 1976; Seiler et al., 1977).

D. CO Uptake by Plants

At the present time, there exist several contradictory publications concerning the influence of plants on the CO cycle. Krall and Tolbert (1957) demonstrate that CO is consumed by higher plants, for example, by leaves of barley. The uptake mechanism has been found to be light-dependent and results in the production of serine. In contradiction to this, tests with a wide variety of plants by Ingersoll (1971) have shown no change in the test atmospheres. On the other hand, Bidwell and Fraser (1972) have found an uptake of CO by bean leaves, using ¹⁴C-labelled CO. In light, CO is converted mainly to sucrose and protein, whereas, in darkness, CO is almost completely oxidized to CO₂. The uptake rates have varied with the species and CO mixing ratios. Using the value for bean leaves (0.06 µmole/ dm ² hour) at a CO mixing ratio of 2 ppm, Bidwell and Fraser (1972) have calculated a total uptake rate by plants of 650 to 6500 x 10¹² g/year.

Taking into consideration the balance of sources and sinks of CO over the continents, and ambient mixing ratios in continental and marine air masses, Seiler (1974) has concluded that these values are overestimated, which is in agreement with Ingersoll's (1971) experiments.

4.1.5 Balance of the Cycle and Residence Time of CO

The estimated CO sources and sinks are summarized in Table 4.1. From the total CO production and consumption data, the CO cycle appears to be balanced. However, the estimations of the CO sources and sinks are still unreliable due to incomplete data. Photochemical reactions contribute, to a great extent, to the global production and consumption rates. Changes in the estimated density of OH radicals in the troposphere will, therefore, influence the rates and the residence time, as Warneck (1974) has pointed out. Further corrections will become necessary if new data, especially on the influence of the biosphere and on the CO production in rain-water, become available.

Using the summarized global production rates and the amount of CO present in the atmosphere, the residence time of tropospheric CO is about 0.10.2 years. A similar residence time, 0.1 years, has been calculated by Weinstock and Chang (1974), using the ¹⁴C production and the specific ¹⁴ C activity of tropospheric CO. A higher residence time will result if lower production and consumption rates of CO by photochemical reactions can be confirmed by measurements of OH densities.

Table 4.1 Summary of estimated annual global CO production and consumption, and amount of CO present in the atmosphere

Sources (x 1015 g/year)
Anthropogenic	0.450.64
Oxidation of methane	1.505.00
Oceans	0.020.22
Rain-water	uncertain
Forest fires, burning	0.06
of agricultural waste
Degradation of chlorophyll	0.100.50
Oxidation of hydrocarbons	0.06
Total CO production	2.196.48
Sinks (x 1015 g/year)
Oxidation by OH	2.70-5.00
Stratospheric sink	0.11
Uptake by soil	0.50
Uptake by plants	uncertain
Total CO consumption	3.31-5.61
Amount of CO present in the	0.53
atmosphere based on 0.11 ppm
mixing ratio (x 1015 g)

4.2 METHANE

The methane content of the atmosphere is rather high, compared to other trace gases. The CH₄ is strongly coupled with the H₂ and CO cycles and contributes, as estimated by Ehhalt (1974), approximately 1% to the total atmospheric carbon cycle.

4.2.1 Global CH₄ Background Levels and Distribution

In more recent measurements, Stephens and Burleson (1969) have reported CH₄ background concentrations of 1.39 ppm for clean mountain air in southern California (35^° N). Ehhalt (1974) has averaged his measurements in the Rocky Mountains (40^° N) at 1.35 ppm. In marine air masses over the Atlantic, between 22 and 35^° N, Swinnerton et al. (1969) have measured an average concentration of 1.24 ppm. At higher latitudes, averages of 1.33 ppm over the Atlantic at 40^° N (Ehhalt, 1974) and of 1.40 ppm over the Norwegian and Greenland Seas (60^° N80^° N) (Larson et al., 1972) have been found. From these measurements, a slight gradient of CH₄ seems to exist with latitude. A decrease with lower latitude in the northern hemisphere is also shown by the northsouth profile of CH₄, presented by Lamontagne et al. (1974). Recently, Ehhalt (1978) has summarized all CH₄ measurements over the North Atlantic and Pacific, which indicate a difference of about 0.1 ppm between the northern and southern hemispheres.

Tropospheric CH₄ profiles by Ehhalt and Heidt (1973) have shown that the troposphere is rather well mixed vertically with a mean tropospheric average of 1.41 ppm. This value proves true only for the northern hemisphere, while in the southern hemisphere the CH₄ mixing ratio is 1.30 ppm (Ehhalt and Schmidt, 1978). In the stratosphere, Bainbridge and Heidt (1966) have found a strong decrease of the CH₄ mixing ratio with altitude. Ehhalt et al. (1972) have shown that the CH₄ concentration at a mean altitude of 50 km is 0.25 ± 0.02 ppm. More recently, Ehhalt et al. (1975) have reported a CH₄ mixing ratio in the upper stratosphere, between 40 and 50 km altitude, of 0.37 ± 0.01 ppm, using a cryogenic sampling technique. From these measurements, it is obvious that the stratosphere acts as a sink for CH₄.

4.2.2 Sources of Atmospheric CH₄

It is well known that most of the atmospheric CH₄ is produced by anaerobic bacterial decomposition of recent organic matter. Several ¹⁴C analyses of atmospheric CH₄ have shown that the average ¹⁴ C content is about 80% of that of recent wood (Libby, private communication; Bainbridge et al., 1961; Bishop et al., 1962), which indicates that 80% of the CH₄ is of recent biological origin and 20% is dead CH₄.

A. Enteric Fermentation of Animals

The oldest estimate of CH₄ production is that of the enteric fermentation of animals. Hutchinson (1957) has estimated the amount of CH₄ produced daily by various large herbivores, and has derived from these values a global CH₄ production of 45 x 10¹² g/year. Due to the increase of the world cattle population, and on the assumption that 10% of the globally produced dry plant matter is eaten by herbivores, Ehhalt (1974) has estimated the present CH₄ production by enteric fermentation to be 100220 x 10¹² g/year.

B. CH₄ Production in Paddy Fields

Koyama (1963, 1964) has measured the CH₄ production rate, and its dependence on temperature, for a number of Japanese paddy fields. From this data, Koyama has extrapolated the average CH₄ production of Japanese rice paddies at 2.3 x 10¹² g/year, which corresponds to an average productivity of 80 g CH₄/ m² year. Koyama has deduced that the global CH₄ production by paddy fields is 190 x 10¹² g/year.

Due to the increase of the global area of rice paddies, Ehhalt (1974) has estimated the present annual CH₄ production of paddy fields to be 280 x 10¹² g/ year. This figure may represent an overestimation for the CH₄ actually released to the atmosphere, since Koyama has determined the anaerobic CH₄ production of soils rather than the release of CH₄ to the atmosphere. In nature, the oxidation of CH₄ by anaerobic bacteria will consume some of the CH₄ before it can reach the atmosphere.

C. CH₄ Production in Freshwater Lakes, Swamps, and Marshes

Several studies have shown that lake sediments, swamps, and marshes produce CH₄, but only a few quantitative studies on the release rate of CH₄ to the atmosphere are known. Conger (1943) has measured the amount of CH₄ reaching the surface of Great Fresh Creek, a lake of about 2 m depth in Maryland (U.S.A.), for August to be about 0.32 g/m² day, which, if maintained throughout the year, corresponds to a release of about 120 g/m² year to the atmosphere. Using additional CH₄ production data in Lake Erie by Howard et al. (1971), and taking into consideration oxidation rates of CH₄ in the water layer, Ehhalt (1974) has estimated that swamps, marshes and lake sediments below a water column not deeper than 10 m release 50100 g CH₄ /m² year to the atmosphere on a global and annual average. Recently, higher fluxes of up to 400 g CH₄/m² year for the major swamp regions have been calculated by Baker-Blocker et al. (1977) from direct measurements. Using this new data, Ehhalt and Schmidt (1978) have revised older global CH₄ production data of swamps and marshes to (190300) x 10¹² g/year. The additional CH₄ release from freshwater lakes has been estimated by Ehhalt (1974) to be (125) x 10¹² g/year.

D. CH₄ Production in Upland Fields, Forests, and Tundra

Koyama (1963, 1964) has also measured the CH₄ production of upland and forest soils and has averaged the annual global rates at 0.44 g/m² year and 0.9 x 10^-2 g/m² year, respectively. From this data, he has estimated an annual global CH₄ production of 10 x 10¹² g/year in grassland, brushland, and cultivated areas and 0.4 x 10¹² g/year in forest areas. Some measurements indicate that waterlogged tundra also releases CH₄ (Benoit, 1973). On the assumption that between 1 and 10% of the tundra is waterlogged in the warm season (4 months) and at a CH₄ production rate of 50 g/m² 4 months, Ehhalt (1974) has estimated the global CH₄ production of tundra to be (1.313) x 10¹² g/year. In later studies, Ehhalt (1976) and Ehhalt and Schmidt (1978) have revised these values to (0.33) x 10¹² g/year, using the measured CH₄ production rates for tundra of 10 g CH₄/m² year, published by Svensson (1973) and Svensson et al. (1975).

E. CH₄ Production in the Oceans

The CH₄ concentrations in surface waters of the open ocean (Swinnerton and Linnenbom, 1967; Swinnerton et al., 1969; Lamontagne et al., 1973) are in the range of (3.77.8) x 10^-5 ml CH₄ /l water, and have been averaged at 4.7 x 10^-5 ml CH₄ /l water. These values are in the same range as those observed for CO. Because of the higher atmospheric mixing ratio of CH₄, the supersaturation is smaller and amounts to about 1.3. Using the molecular diffusion model through the laminar boundary layer at the water surface (Broecker and Peng, 1971, 1974), Ehhalt (1974) has estimated the net flux of CH₄ into the atmosphere to be (46.7) x 10¹² g/year. Recently, Ehhalt and Schmidt (1978) have given estimates of between 1.3 and 16.6 x 10¹² g/year. Measurements by Lamontagne et al. (1974) have shown that the partially ice-covered areas of Antarctica are undersaturated in CH₄ by a factor of about 1.2. Lamontagne et al. (1974) have stated, therefore, that the role of the open ocean as a major source or sink of CH₄ is insignificant.

All values refer to surface waters of the open ocean. Shelf and bay concentrations can be 23 orders of magnitude higher than these open-ocean surface values (Lamontagne et al., 1973). Brooks and Sackett (1973) have observed CH₄ concentrations in polluted areas of the Gulf of Mexico of up to 250 x 10^-5 ml/1. Ehhalt (1974) has pointed out that the CH₄ production in continental shelf areas, where there exist sediments with organic matter and overlying water layers of not more than 10 m, may be similar to those in freshwater lakes. Using the same assumptions as for freshwater lakes, Ehhalt (1976) has estimated global CH₄ production in shelf areas to be (0.071.4) x 10¹² g/year. Estimations of Seiler and Schmidt (1974) for the total oceanic CH₄ production amount to 16 x 10¹² g/year.

Table 4.2 Estimated annual global production of ¹⁴C-free CH₄ from anthropogenic and other sources

Sources (x 1012 g/year)
Coal mining	6.3	22
Lignite mining	1.6	5.7
Industrial losses	7	21
Automobile exhaust	0.5
Volcanic emissions	0.2
Total 14C-free CH4	15.6	49.4
Maximal 14C-free CH4	210
from 14C content of CH4

In anoxic waters, extremely high CH₄ concentrations in greater water depths, due to the activity of anaerobic bacteria, have been reported by several authors. Lamontagne et al. (1973) have observed, in the Cariaco Trench below 300 m, maximum CH4 concentrations of about 0.3 ml CH4/1 water. Atkinson and Richards (1967) have reported similarly high supersaturations in the Black Sea at depths below 200 m. The concentrations, however, do not contribute, in any way, to the global CH₄ release into the atmosphere, because the CH4 is oxidized in the overlying water layers.

F. Anthropogenic and Other Sources of ¹⁴C^-Free CH₄

Hitchcock and Wechsler (1972) have estimated the production rates of other ¹⁴C-free CH₄ sources, which are summarized in Table 4.2. The production rates are between 15.6 and 49.4 x 10¹² g/year. An upper limit can be obtained from the ¹⁴C content in CH₄, since ¹⁴C measurements indicate that the emission of ¹⁴C-free CH₄ is about 25% of the biogenic production, or 210 x 10¹² g CH₄ /year. 

4.2.3 Sinks of Atmospheric CH₄

Measurements over the past 20 years have indicated that the tropospheric CH₄ concentration is practically constant with time. Since the residence time of tropospheric CH₄ is in the area of a few years, the CH₄ cycle is in a steady state, and the CH₄ production must be balanced by an equivalent destruction.

A. Atmospheric Oxidation of Methane by OH-Radicals

The atmospheric oxidation of methane by OH radicals, or excited O(¹ D) atoms, has been discussed in Section 4.1.3 for the production of CO. The estimated CO production rates of 15005000 x 10¹² g/year correspond to CH₄ destruction rates of 8602860 x 10¹² g/year, assuming a 1:1 conversion of methane to CO. It has been pointed out that these rates are uncertain, because of the lack of experimental measurements of OH densities on the global average.

B. Stratospheric Sink for Tropospheric CH₄

From stratospheric profile measurements of CH₄, Ehhalt (1974) has calculated, by an eddy diffusion model, the global CH₄ loss at (45150) x 10¹² g/year. Ehhalt (1974) has estimated an additional CH₄ loss to the stratosphere of 60 x 10¹² g/year, due to the Hadley cell circulation in the tropics, in which air penetrates from the troposphere into the lower stratosphere.

C. CH₄ Uptake on the Soil Surface

It has been speculated that atmospheric CH₄ can be oxidized by soil bacteria. However, in experiments by Ingersoll et al. (1974) and Ehhalt (1974), no uptake of CH₄ by soils has been found. Therefore, soil is probably not a significant sink for atmospheric CH₄.

4.2.4 Balance of the Cycle and Residence Time of CH₄

The estimated CH₄ sources and sinks are summarized in Table 4.3. From the total CH₄ production and consumption data, the photochemical destruction of CH₄ by OH radicals seems to be overestimated. It is noticeable that the release of biogenic CH₄ to the atmosphere equals or exceeds the annual production of CH₄ from natural gas wells which was about 520 x 10¹² g CH₄/year in 1965 (Ehhalt, 1974). Furthermore, the biogenically released CH₄ amounts to as much as 0.5% of the annual production of dry organic matter (Ehhalt, 1974), which has been estimated by Woodwell (1970) to be 165 x 10¹⁵ g/year. Since the energy content of CH₄ is about 3 times that of cellulose, about 12% of the solar energy fixed by photosynthesis is lost to the atmosphere as CH₄ and escapes the biological food chain. The CH₄ cycle makes a small, but significant contribution to the global carbon cycle. Between 1.5 and 2.8 x 10¹⁵ g CO₂ /year pass through it annually. This amounts to 12% of the annual CO₂ uptake by land plants, which totals about 130 x 10¹⁵ g/year (Ehhalt, 1974).

Table 4.3 Summary of estimated annual global CH₄ production and consumption, and amount of CH₄ present in the atmosphere 

Sources (x 1015 g/year)
Enteric fermentation of animals	0.10	0.22
Paddy fields	0.28
Swamps and marshes	0.19	0.30
Freshwater lakes	0.001	0.025
Upland fields, forests, and tundra	0.011	0.013
Oceans	0.001	0.018
14C-free CH4	0.016	0.21
Total CH4 production	0.60	1.07
Sinks (x 1015 g/year)
Oxidation by OH	0.86	2.86
Stratospheric sink	0.10	0.21
Uptake by soil	negligible
Total CH4 consumption	0.96	3.07
Amount of CH4 present in the atmosphere	4.0
based on 1.41 ppm mixing ratio (x 1015 g)

4.3 NON-CH₄ ORGANIC GASES

Very few reliable measurements of non-CH₄ organic gases have been undertaken. Besides air pollution effects, the content of these gases seems to be strongly dependent on the environment. As yet, no global atmospheric cycle of these gases can be given. Not only paraffinic, olefinic, and aromatic hydrocarbons, but also alcohols, acetaldehyde, and acetone have been detected in considerable amounts.

4.3.1 Background Levels of Non-CH₄ Hydrocarbons and Other Gases

Available data on the total amount of particulate and vapour-phase organic carbon in the troposphere have recently been summarized by Duce (1978). Marine background measurements by Rasmussen (1974) in Hawaii and the North Atlantic, and by Wade and Quinn (1974) in Bermuda indicate levels of the vapour-phase C₃C₁₂ and C₁₄C₃₂ hydrocarbons in the range of 416 and 0.050.5 µg of organic carbon (OC) /m³ STP, respectively. Levels in continental non-urban areas are usually higher and amount to 5-30 for C₂C7 (Robinson et al., 1973) and to 27210µg OC/m³ STP for total non-CH₄ hydrocarbons (Chatfield and Rasmussen, 1976; Singh et al., 1977; Rasmussen et al., 1977). Levels of atmospheric particulate organic carbon (POC) in these areas are lower by at least one order of magnitude than those of vapour-phase hydrocarbons. Marine background levels of atmospheric POC range from 0.13 to 1.8 µg POC/m³ STP for the northern hemisphere (Hidy et al., 1974; Hoffman and Duce, 1974, 1977; Barger and Garrett, 1976; Ketseridis et al., 1976) and from 0.07 to 0.53 µg POC/m³ STP for the southern hemisphere (Barger and Garrett, 1976; Hoffman and Duce, 1977). Atmospheric POC concentrations in continental non-urban areas are between 0.6 and 5.0 µg POC/m³ STP (Ketseridis et al., 1976; Duce, 1978).

Background levels for the individual non-CH₄ organic gases have only been reported in a few measurements by Cavanagh et al. (1969) and Robinson et al. (1973), which are given in Table 4.4. All values are less than a few ppb, except for equatorial air, which indicates the strong influence of biogenic processes for ambient air levels of non-CH₄ organic gases. The high values for n-butanol in Arctic air are probably also due to local biogenic processes. Cavanagh et al. (1969) have assumed a fermentation of starch in the tundra cover at summer time as the main source. Other products of fermentation are acetone and ethanol, identified in the samples. In marine air from Hawaii, no n-butanol has been detected.

Table 4.4 Background levels of some non-CH₄ organic gases by regions

	Level of gases in ppb
			Region
	A	B	C	D
Ethane + Ethylene	n.d.0.1	n.m.	n.m.	1944
Propane	n.m.	n.m.	n.m.	183247
Butane	n.d.0.2	0.13.6	n.d.1.5	5491
Isobutane	n.m.	n.d.1.7	n.m.	n.m.
2Methylbutane	n.m.	n.d.1.0	n.m.	n.m.
Pentane	n.d.0.3	n.d.0.5	n.d.0.3	135269
2,4Dimethylpentane	n.m.	0.35.1	n.m.	n.m.
Benzene	n.d.0.4	0.31.1	n.d.0.2	714
Toluene	n.m.	0.11.0	n.d.0.2	115
Methanol + Ethanol	n.d.1.2	n.m.	n.m.	n.m.
nButanol	51126	n.m.	n.m.	n.m.
Acetaldehyde	n.d.0.3	n.m.	n.m.	n.m.
Acetone	n.d.2.9	0.30.9	0.13	240
Total of non-CH4 carbon	> 10*	> (530)	628	12001600
in µg OC/m3 STP
*Excluding n-Butanol
n.d. - not detected
n.m. - not measured

 4.3.2 Concentrations of Non-CH₄ Hydrocarbons in Polluted Areas

The hydrocarbon composition of polluted atmospheres in metropolitan areas has been measured by several authors (Altshuller and Bellar, 1963; Altshuller et al., 1966, 1971, 1973; Stephens and Burleson, 1967, 1969; Gordon et al., 1968; Lonneman et al., 1974; Bos et al., 1977; Siddiqi and Worley, 1977). The content of methane (2.42.6 ppm) has been found, in Cincinnati and Los Angeles, to be approximately the same as that for non-CH₄ hydrocarbons (0.83.0 ppm). The non-CH₄ hydrocarbon concentrations show typical rush-hour traffic profiles; the methane source, on the other hand, is relatively constant (Altshuller et al., 1966). Reactive unsaturated hydrocarbons decrease faster between morning traffic peak hours and mid-afternoon than saturated hydrocarbons (Altshuller et al., 1971). A decrease of reactive hydrocarbons is even found in smog situations (Stephens and Burleson, 1969). A spectrum of hydrocarbons, identified in Los Angeles in 1967 by Altshuller et al. (1971), is given in Table 4.5. Due to emission control programmes, the ambient total hydrocarbon concentrations in Los Angeles air have been reduced, since 196365, to 51.8% in 1973 (Leonard et al., 1976).

Table 4.5 Overall average of values of aliphatic hydrocarbon and alkylbenzene concentrations in Los Angeles in 1967 (Altshuller et al., 1971) 

Level of gases in ppb
Methane	2400
Ethane + Ethylene	102
Acetylene + Propane	76
Propylene	10	.5
Butane	46
Isobutane	12
Butene(1) + Isobutene	5	.5
Butene(2)	2
Butadiene(1.3)	2
Pentane	21
Isopentane	35
2Methylbutene(1)	3
Nonane + Decane	5
Toluene	30
Ethylbenzene	5
Xylenes	23	.5
Propylbenzenes	4	.5
Ethyltoluenes	7	.5
Other C9 and C10 alkylbenzenes	12
Total of identified non-CH4 carbon	~950
in µg OC/m3 STP

Total atmospheric content of non-CH₄ hydrocarbons in polluted metropolitan areas has been measured up to 2400 µg OC/m³ STP (Singh et al., 1977). That of atmospheric POC in these areas ranges from 2 to 110 μg POC/m³ STP (Henry and Blosser, 1971; Cukor et al., 1972; Mueller et al., 1972; Gordon and Bryan, 1973; Patterson, 1973; Ciaccio et al., 1974; Lee and Hein, 1974; Appel et al., 1976; Ketseridis et al., 1976).

4.3.3 Concentrations of Non-CH₄ Hydrocarbons in Oceanic Surface Waters 

Oceanic surface concentrations of volatile non-CH₄ hydrocarbons have been measured by Brooks and Sackett (1973), Swinnerton and Lamontagne (1974a) and Lamontagne et al. (1974); these measurements are given in Table 4.6. During a north-south trip in the Pacific, Lamontagne et al. (1974) have found definite latitudinal variations of ethylene and propylene concentrations, with a broad peak between 10^° N and 15^° S and 40^° S. The highest values have been obtained in the ice-covered region at 76^° S. With the exception of an enrichment in the ice-covered region, methane, ethane, and propane values have remained fairly constant. Lamontagne et al. (1974) have attributed these peaks to biological activity, which is in accordance with experiments of Wilson et al. (1970), who have shown that unsaturated hydrocarbons increase with biological activity, while there is very little. production of saturated hydrocarbons. The enrichment in the ice samples is explained by large amounts of algae enclosed in the samples. Swinnerton and Lamontagne (1974a) have demonstrated that methane, ethane, and propane concentrations can be used for determining the degree of contamination of ocean water. The following ratios (concentration in contaminated to open ocean) have  been calculated from average values: methane, 25; ethane, 142; propane, 280; ethylene, 2.3; propylene, 2.0. It is shown that unsaturated hydrocarbons are only slightly affected by contamination; their concentrations can, however, be used as indicators for productivity.

Table 4.6 Concentrations of light hydrocarbons in surface waters

	Concentrations in 10-6 ml gas/l water
	A		B	C	D
Ethane		3.0	0.11.7(0.50)*	1.4650(71)	11.9
Ethylene			0.712.1(4.8)*	1.935(11)	11.2
Propane		1.0	0.051.4(0.34)	1.01110(95)	4.3
Propylene		n.m.	0.15.8(1.4)	0.116(2.8)	5.3
Butane + Isobutane	~	0.1	~0.05	198	5.5
Butene		n.m.	n.m.	n.m.	1.7
*The average is given in parentheses
n.m. = not measured

A comparison of the atmospheric background levels of non-CH₄ hydrocarbons (with the exception of those in the equatorial region) with the open-ocean surface water concentrations shows, by calculation, that ethane, ethylene, and probably propane and propylene, are also supersaturated in ocean water by a factor of at least 100. Due to the poor data available so far, no calculations for the oceanic source strength for non-CH₄ hydrocarbons have been undertaken. Duce et al. (1974) have even suggested an oceanic sink for non-CH₄ hydrocarbons.

4.3.4 Sources of Non-CH₄ Hydrocarbons

Hydrocarbons are released to the atmosphere as a result of both anthropogenic activity and natural processes. The most important anthropogenic source of hydrocarbons, resulting from the incomplete combustion of fuel, is motor-vehicle exhaust. The anthropogenic source contributions in Southern California air have recently been calculated by Mayrsohn et al. (1976, 1977) to be 53% for motorvehicle exhaust, 22% for gasoline and gasoline vapour, 5% for commercial natural gas, 19% for geogenic natural gas, and 1% for liquefied petroleum gas. Robinson and Moser (1971) have estimated the anthropogenic emissions of total hydrocarbons to be 88 x 10¹² g/year. About one-third, or 27 x 10¹² g/year, are classed as reactive olefins and aromatics (Robinson and Robbins, 1968b). Recently, Duce (1978) has assumed anthropogenic emissions of 80 x 10¹² g of non-CH₄ hydrocarbons/year or, with 80% OC, about 65 x 10¹² g OC/year, in 197374.

Vegetation is another important source of vapour-phase organic substances. The major reactive hydrocarbons emitted by trees are ethylene, monoterpene (C₁₀), and isoprene (C₅). Rasmussen (1972) has estimated that forests represent a global natural source of 175 x 10¹² g of reactive hydrocarbons/year, or 140 x 10¹² g OC/year. This emission rate is 6 times greater than that estimated for reactive anthropogenic hydrocarbons. Other estimates are in the range of (10350) x 10¹² g OC/year (Rasmussen and Went, 1965). A very high estimate of global vapour-phase organic emissions, with an average of 8000 x 10¹² g/year, or 6500 x 10¹² g OC/year, has been given by the Research Triangle Institute (1974). Westberg (1977) has pointed out the uncertainties in this estimate, but he has indicated that appreciable amounts of organic compounds, other than terpenes, are emitted from vegetation.

4.3.5 Sinks of Non-CH₄ Hydrocarbons

Various studies have shown that reactive hydrocarbons undergo rapid photochemical reactions (in the region of hours) in the presence of O, O₃, nitric-oxides, and SO₂. The kinetic mechanisms of these reactions are summarized by Altshuller and Bufalini (1965) and Hecht and Seinfeld (1972).

Quantitative data on the rates of these atmospheric reactions are rare. Hidy (1973) has suggested that 110% by weight of the emitted reactive hydrocarbons are converted to POC in the form of aerosols which are rapidly removed by scavenging or deposition. The remaining hydrocarbons are eventually oxidized to CO₂ and H₂O. Robinson and Robbins (1971) and Peterson and Junge (1971) have estimated that about one-third of the anthropogenic non-CH₄ hydrocarbons undergoes rapid conversion to POC. Recently, Gay and Arnts (1977) and Schuetzle and Rasmussen (1977) have studied the conversion of terpenoid hydrocarbons. From their data, Duce (1978) has assumed that 80% of the terpenoid hydrocarbons are rapidly converted.

The role of soils as sink or source for light hydrocarbons is not well established. Abeles et al. (1971) and Smith et al. (1973) have shown that soils absorb ethylene and acetylene due to microbial activity. On the other hand, soils can be the source of non-CH₄ hydrocarbons. A summary on the microbial formation of hydrocarbons is given by Bird and Lynch (1974).

4.3.6 Atmospheric Amount and Residence Time of Non-CH₄ Hydrocarbons

As yet, all data on global sources, atmospheric distribution and cycles of non-CH₄ hydrocarbons are highly uncertain. Nevertheless, Duce (1977) has made an attempt to estimate the atmospheric level of these gases. The data are summarized in Table 4.7. With the lower production values of terpenoid hydrocarbons given by Rasmussen (1972) and the estimated gas-to-particle conversion ratios, a residence time of about 0.7 years for the long-lived vapour-phase hydrocarbons follows. Duce et al. (1974) had earlier estimated a residence time in the range of 0.5 to 2.0 years. The residence time of POC is between 2 and 30 days, depending on the particle size (Duce, 1978).

Table 4.7 Summary of estimated annual global non-CH₄ production and amount of OC and POC present in the atmosphere

Sources (x 1015 g OC/year)
Anthropogenic	0.065
(30% conversion to POC)
Vegetation	0.010.35
(80% conversion to POC)	6.5 (?)
Atmospheric amount (x 1015 g)
OC over land	0.037
over sea	0.014
Total OC (estimated range)	0.051 (0.020.15)
POC over land	0.001
over sea	0.001
Total POC (estimated range)	0.002 (0.0020.006)

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