Review of progress in isotope studies of atmospheric carbon monoxide

Chemosphere: Global Change Science 1 (1999) 33±52

Review of progress in isotope studies of atmospheric carbon monoxide C.A.M. Brenninkmeijer *, T. R ockmann, M. Br aunlich, P. J ockel, P. Bergamaschi Max Planck Institute for Chemistry, Air Chemistry Division, Mainz, Germany Received 20 April 1998; accepted 3 September 1998 This paper is dedicated to Charles M. Stevens for his pioneering research on isotope variations in atmospheric carbon monoxide and methane.

Importance of this Paper: Isotope analysis is important in geochemistry, but has had a low impact in atmospheric chemistry. This is changing. Isotope variations in atmospheric trace gases are studied more frequently. Carbon monoxide is an important trace gas, and gives excellent examples of the information stable isotope analysis can give. We call attention to 14 CO, which is a unique tracer in the atmosphere. We also mention the important reaction CO + OH, for which a new isotope e€ect has been found recently. This review tries, without too much ado, to present the highlights of this exciting research. Abstract Progress made since the start of isotopic analysis of the important atmospheric trace gas carbon monoxide is reviewed. A systematic discussion of the main processes leading to variations in stable isotope ratios for 13 C and 18 O is given. Also the recently discovered deviation in the 17 O content of atmospheric CO is discussed. Further the role of 14 CO in atmospheric chemistry as a gauge for atmospheric hydroxyl is elucidated. Ó 1999 Elsevier Science Ltd. All rights reserved. Keywords: Carbon monoxide; Isotopes; Carbon-13; Oxygen-18; Oxygen-17; Carbon-14; Atmospheric chemistry; Fractionation

1. Introduction In this paper we review the most important advances in the application of isotope analysis to atmospheric carbon monoxide. Whereas concen*

Corresponding author.

tration and ¯ux measurements generally provide the primary data for constructing trace gas budgets, isotope measurements give complementary, sometimes unique information. Nevertheless, considering the greater complexity of isotope analyses compared to concentration measurements, it is obvious that isotope analysis is only worthwhile if substantial new information and insights

1465-9972/99/$ ± see front matter Ó 1999 Elsevier Science Ltd. All rights reserved. PII: S 1 4 6 5 - 9 9 7 2 ( 9 9 ) 0 0 0 1 8 - 5

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are obtained. For another review on CO isotopes we refer to Conny (1998), and for a general review about isotope variations in atmospheric trace gases to Kaye (1987). In the review in front of you, we focus on the latest developments, place the isotope work in context with atmospheric chemistry, and highlight the most important and useful ®ndings. For CO we have on the one hand a trace gas with a poorly de®ned budget, and on the other hand several particularly promising isotope applications, as we will see later on. Together this is the rationale for going into the trouble of collecting air samples for subsequently performing the isotope analyses in the laboratory. In situ isotope analysis is hardly feasible in view of the low abundance of 13 C, 17 O, 18 O and particularly 14 C (about 10ÿ12 ), combined with the low mixing ratio of CO itself (about 10ÿ7 ). Generally remote measurements of isotope ratios in trace gases are beset with large errors. At the same time however, the isotope effects tend to be very small. A corollary to the low abundance of the isotopes of C&O is of course that the chemistry as such is not noticeably in¯uenced by isotopic variations. The isotopes are mere tracers for chemical and physical processes. Note however that recently very large isotope e€ects have been discovered, for instance in the formation of ozone where reactions for certain isotopic combinations are up to 50% faster. Thus in principle it is possible that isotope e€ects do a€ect the chemistry whenever isotopic abundances are high, as is for instance the case for chlorine. The various sources of CO do show di€erent isotopic signatures, and in an ideal case the CO present in an air mass can be apportioned to its sources provided (a) the sources have clearly distinct isotopic signatures, (b) the number of sources does not exceed the number of independent isotope ratios available. However, during the ongoing removal process of CO primarily by reaction with OH radicals, the isotope ratios for the fraction remaining in the atmosphere are modi®ed. This socalled isotope fractionation thus clearly introduces complications. Nonetheless at the same time the isotopic modi®cation during the exposure of CO to the oxidation by OH also introduces useful signals.

The aging of CO due to exposure to OH can be closely followed by the isotopic variations incurred. At the outset we emphasize that there is a fundamental di€erence between the stable isotope versions (isotopomers) of CO (13 CO, C18 O, and C17 O) and the radioactive 14 CO. Curiously, this has less to do with the radioactivity of 14 C than perhaps expected. Obviously, the half life of 14 C far exceeds the chemical lifetime of CO, which is only a few weeks in the tropics. Furthermore, since the advent of particle accelerator mass spectrometry even the detection of 14 C is not anymore based on its radioactive decay. However, 14 CO can be considered in its own right as an independent ultra trace species in the atmosphere because it has primarily a cosmogenic source. The 14 C share it derives from the biogenic CO sources is small. Thus 14 CO is only weakly coupled to CO via the common biogenic source. For instance, when 14 CO is abundant at 12 molecules per cm3 at standard temperature and pressure, and CO at 100 nmole/mole, of which we assume for the sake of argument that 20 nmole/ mole is of fossil origin, and the remaining 80 nmole/ mole of biogenic origin with a speci®c activity of 100 pMC (percent modern carbon), 3 out of the 12 molecules are actually of biogenic origin. These 14 CO molecules contain 14 C atoms which have been recycled through the biosphere. At the same time 14 CO shares of course with CO the soil and OH sink. As a result of all this, when dealing with 14 CO, not the speci®c activity of CO, but rather the abundance of 14 CO itself will be used. We state again that the abundance of 14 CO is given as molecules per cm3 air, at STP. The conventional units for atmospheric hydroxyl (OH) are cmÿ3 , i.e. space. We also note that the reaction rate constant for 14 CO with OH is about 1% smaller than that for 12 CO, based on the about 0.5% smaller reaction rate for 13 CO. Whereas isotope e€ects of this magnitude are the essence of the stable isotope applications, for 14 CO applications this does not play a signi®cant role. To de-convolute the complex signal of spatial and temporal variations in the isotope ratios which are due to transport e€ects, the wide range of sources, and the e€ect of the OH and soil sinks, clearly requires 3D modeling with adequate resolution in time and space. One intrinsic problem is

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of course that the source isotope signatures are not always well known and in some cases may remain ill de®ned. Are we introducing more unknowns to try to help to better constrain the CO cycle? Another problem is that the magnitude of isotopic fractionation in the reaction CO + OH depends on pressure, and thus varies with altitude. At the time of this review, very little modeling has been performed, rendering the interpretation of the measurement results often rather qualitative. Nevertheless, this situation is changing rapidly with more isotope data becoming available and 3D modeling for CO speci®cally and isotope applications in general becoming more common. We will give simpli®ed examples of isotopic changes of CO due to oxidation by OH as a heuristic model.

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2. Atmospheric carbon monoxide The annual emission of CO into the atmosphere by land based sources (Table 1), that is mainly incomplete combustion, added to the in situ production, notably the oxidation of methane and other volatile organic carbons, amounts to about 2600 Tg. This is equivalent to roughly 500 kg CO for each person of earth, showing that certainly man can have an impact on this trace gas. For instance an average fuel use of several kg per day per person, and a CO emission factor of 10%, may have an impact on the source strength. Because nearly all CO is removed through oxidation by OH, according to CO + OH ® CO2 + H, approximately 60% of the atmosphere's OH based

Table 1 The tropospheric budget of CO, with the source strengths and respective isotopic composition and with the sink isotopic fractionation constants

a

Sources

Source/sink Tg/yr

d13 C (&) V-PDB

d18 O (&) V-SMOW

14

Fossil fuel combustion Biomass burning

300±550 300±700

0.0l 0.0l

400±1000

~125

0?

NMHC oxidation

200±600

ÿ32.2e

~110

0?

Ozonolysis Biogenic Oceans

60±160 20±200

23.5a;b ~16.3b ~18e ~0b;g ~15e ~14.9e ~0b;g 80±100h

0.0 ~115

CH4 oxidation

ÿ27.5a ÿ21.3c ÿ24.5d ÿ52.6f

~110 ~110

25±40l 0.0 0.0

Total Sources

1800±2700

Sinks Reaction with OH Soil uptake Loss to stratosphere

1400±2600 250±640 ~100

10k

4h

Total sinks

2100±3000

ÿ13.5i

5j ;h

ÿ10j;h

CO (pMC)

Stevens et al., 1972. Brenninkmeijer, 1993. c Conny et al., 1997. d Conny, 1998. e Stevens and Wagner, 1989. f Value based on the CH4 d13 C of ÿ47.2& (Quay et al., 1991), and the fractionation in CH4 + OH of 5.4&. g Brenninkmeijer and R ockmann, 1997. h R ockmann et al., 1998. i Manning et al., 1997. j At atmospheric pressure. k Inferred from the value for 13 C, assuming mass dependent behavior. l R ockmann et al., 1998b. b

D17 O (&)

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oxidative self-cleansing capacity is occupied by CO (Crutzen and Zimmermann, 1991). CO competes with CH4 , and many other species for available OH. Generally speaking, increased CO emissions cause lower OH levels, which in turn could lead to increased CH4 , which is a greenhouse gas. The coupling between CH4 , OH and CO is complex, and has been treated in several studies (Crutzen and Zimmermann, 1991; Isaksen, 1987; Thompson and Cicerone, 1986). We note that although CO turns over a major fraction of all available OH, and as such is a gross sink for OH, the net sink

e€ect is much smaller, basically because in the oxidation an HO2 radical is formed, leading to the regeneration of OH via H + O2 ® HO2 and NO + HO2 ® NO2 + OH (Crutzen 1994), depending on NO and O3 levels. NO is regenerated from NO2 by rapid photolysis, after which O3 can be formed again. The boundary layer background levels of CO range from approximately 35 nmole/mole in the remote summer southern hemisphere to 200 nmole/mole in the remote winter northern hemisphere. This is illustrated in Fig. 1 which shows a

Fig. 1. A comparison of the annual change in the isotopic composition of CO for the high southern and northern latitudes, Scott Base Antarctica (1992±1993) and Spitsbergen (1996±1997) (Brenninkmeijer and R ockmann, 1998b).

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comparison of the concentration and isotope data for Scott Base, Antarctica, and Spitsbergen. The large spatial and temporal variations in CO are due to (1) OH distributions and seasonality, (2) its relatively rapid reaction with OH, and (3) to a lesser extent to source variations as well. This combined with the various types of sources and their complex distribution pattern hinders the construction of a detailed global budget. Ongoing global measurements have shown that CO is decreasing over the last years (Khalil and Rasmussen, 1994; Novelli et al., 1994), yet the cause or causes are unclear. Isotope measurements of CO may well provide part of the answers in the future. 3. Isotope variations in carbon monoxide 13

C, 17 O, 18 O are abundant at levels of about 1%, 0.04% and 0.2%. The naturally occurring variations in these numbers are in the 1% range only, and the usual d notation as per mil deviation is used. For 13 C this is: d13 C ˆ Rs /Rr ÿ1, where s stands for sample and r for reference, and R denotes the respective 13 C/12 C ratio. For 13 C the reference material is V-PDB, Vienna PeeDee Belemnite. This is a marine limestone standard, relative to which all organic material appears to be depleted and has negative d values of around ÿ25&. Atmospheric CO2 has a value near ÿ8&, which is gradually decreasing due to CO2 input from fossil fuel combustion. For 18 O the reference material is V-SMOW (Vienna-Standard Mean Ocean Water), relative to which atmospheric and organic oxygen appears to be enriched. Most sources of CO, perhaps all, have positive d18 O values. Conventions for reporting stable isotope ratios, based on CraigÕs (1957) early work, are given by Gon®antini (1978). A highly interesting recent development is the ®nding that atmospheric CO exhibits small but clear anomalies in its 17 O content (Hodder et al., 1994; R ockmann et al., 1998b,c). Generally, 17 O analysis of oxygen bearing compounds has had little attention, because it was long believed that apart from certain meteoritic material the changes in 17 O always match the concomitant changes in 18 O according to d17 O  0.5 d18 O. The reasons for

37

this mass dependent behavior of isotope fractionation in virtually all processes that play a role in nature are well understood. Deviations from this, known as non-mass dependent fractionation or simply mass independent fractionation (MIF) have been found for atmospheric O3 (Mauersberger, 1987; Thiemens and Heidenreich, 1983) and stratospheric CO2 . But as pointed out, also CO exhibits a certain degree of MIF. This opens exciting new applications of isotope e€ects in the atmosphere. 14 C is abundant in contemporaneous organic matter and atmospheric CO2 at a level of about 10ÿ12 only. This is chie¯y the result of a steady state distribution in the atmosphere determined by the global cosmogenic production of about 2 atoms cmÿ2 sÿ1 and transfer into the various carbon reservoirs, ultimately mainly the deep ocean and carbonate deposits, accompanied by radioactive decay. The 14 C/12 C ratio in atmospheric CO2 has once nearly doubled within a matter of years due to the surface testing of big nuclear bombs mainly in the early sixties, but otherwise recovered in recent years to close to natural levels. At the same time 14 CO2 levels are continuously being forced down below natural levels due to the combustion of fossil, i.e. 14 C free, fuels. 14 C levels in atmospheric methane are arti®cially somewhat high due to the nuclear industry (Lowe et al., 1988). 14 C levels in atmospheric CO are inherently much higher than in CH4 and CO2 due to the interesting feature that 14 C, nearly all of which is produced via the capturing of slow and thermal neutrons generated in cosmic radiation according to 14 N(n,p)14 C, for about 95% leads to 14 CO and not 14 CO2 (McKay et al., 1963). In e€ect, virtually all neutrons in the atmosphere are captured to produce 14 C atoms. About 7.5 kg 14 C is produced in the atmosphere per annum, assuming a production rate of 2 14 C atoms cmÿ2 sÿ1 . Assuming a chemical lifetime of CO and 14 CO of about 1 month, the abundance can be estimated to be 10 molecules per cm3 air. Thus for comparison, 1 m3 air contains about 107 14 CO, 1010 14 CO2 and 5 ´ 107 14 CH4 molecules. This again shows that atmospheric CO is naturally rather radioactive. Whereas CO2 values are typically just above 100 pMC, values for CO of up to 900 pMC have been

38

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encountered in the troposphere. Stratospheric values for CO can be as high as 10,000 pMC.

14

C

4. Sources, sinks and fractionation The sources of atmospheric CO and their isotopic composition, together with the fractionation factors are shown in Table 1, in which many data are based on the various studies by Stevens and coworkers. In some cases, notably CO from CH4 + OH, important information is still missing. The numbers given for this reaction are based on estimates derived from atmospheric measurements (Brenninkmeijer and R ockmann, 1997; Stevens and Wagner, 1989), and not laboratory measurements. Also biomass burning derived CO is not that well isotopically de®ned yet, although laboratory experiments have been carried out. Likewise the value for high temperature combustion is to some degree based on atmospheric observations. Biomass 18 O values are lower because fractionation can play a role at lower combustion temperatures. Isotopic exchange between CO2 , CO and H2 O has to be considered, and the moisture content a€ects the d18 O values (unpublished results). The type of isotope fractionation we are dealing with is called kinetic isotope fractionation and is due to the di€erence in reaction speeds of the various isotopomers in the removal reaction CO + OH. The fractionation factor is given as the ratio between the two ®rst order reaction rate constants, i.e. a ˆ k12 /k13 . The fractionation constant is de®ned as e ˆ aÿ1. Measurements for 13 C have been made by Stevens et al. (1980), Smit et al. (1982) and R ockmann et al. (1998c). A modest fractionation exists, which declines with decreasing pressure. Already below 400 mbar the isotope effect for 13 C reverts into an inverse e€ect, where 13 CO reacts faster with OH than 12 CO. Thus the overall fractionation in the troposphere is moderate, although one has to bear in mind that the removal rates of CO in the upper troposphere also are lower due to the pressure dependence in the reaction with OH, which itself also tends to be less abundant at higher altitudes. Measurements for 18 O also have been made by Stevens et al. (1980) and R ockmann et al. (1998c) and show an inverse

slight pressure dependent e€ect. A value of e 10& is representative for most of the troposphere. Measurement for 17 O for CO + OH have been made by R ockmann et al. (1998c). Further below these e€ects are treated in more detail. Kinetic fractionation is most probably also associated with the soil sink. Two steps are involved here. One is the fractionation during the di€usion of CO into the soil pores. This gives an isotope fractionation equal to, or less than that is derived from the di€erences in the di€usion speeds. The fractionation for 18 O should be twice as large as that for 13 C. The fractionation factor equals the square root of the ratio of the reduced masses. The fractionation associated with the actual absorption by the active soil organisms is not known. Experiments on the soil sink are required, but altogether a small fractionation for this moderate sink is expected. The degree of fractionation will depend much on the role of di€usion. Because most CO is absorbed in the top few cm of soil, di€usion e€ects may be suppressed, thus reducing fractionation. For making the isotopic information somewhat more transparent we have used a box model to simulate the isotopic changes incurred when CO interacts with OH (Fig. 2). Air representative for the mid latitude northern hemisphere at 150 nmole/mole has been used in these examples. The change with and without the CH4 + OH source are shown. Also a source of 1.2 nmole/mole CO per day is included to simulated the continuous input of CO from combustion sources. One feature that can be gleaned immediately is that at OH ˆ 106 , a relatively high value for the 24 hr average, it still takes several months to approach equilibrium. This illustrates the diculty of interpreting CO isotope data without modeling. The typical meridional transport times in the troposphere and the seasonal changes in OH are comparable to its chemical lifetime and the time scale of isotopic changes. 5. Measurement A suitable technique for CO isotope analysis has been developed by Stevens and Krout (1972) and is still being used. A major problem is that the

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Fig. 2. The changes in concentration and isotopic composition for CO in a given air mass due to the e€ect of OH. The full lines are for a scenario without any sources. The dashed line results include the methane oxidation source, assuming 106 OH cmÿ3 . The scenario represented by the dotted lines includes both the methane oxidation source and a constant source of 1.2 nmole/mole per day of CO with d13 C ˆ ÿ27& and d18 O ˆ 15&. For 14 CO sink only and sink with a constant source of 3 ´ 10ÿ6 molecules sÿ1 cmÿ3 is included.

extraction of CO from air is problematic due to its low abundance coupled with a boiling point close to that of air itself. Molecular sieve based techniques for recovering of CO from air have been developed, but their application at low concentrations is doubtful. Anyway, because high precision mass spectrometry is based on CO2 , the principle Stevens developed is based on the thorough removal of CO2 from air, followed by the oxidation of the CO, and the ensuing trapping as CO2 . Also N2 O has to be removed, because of the interference even at low levels with the mass

spectrometry of CO2 . The use of solely molecular sieve to achieve the desired puri®cation is generally not recommended because of various reasons. Molecular sieve may trap some CO, and actually a large molecular sieve trap may act as a GC column. Second, the trapping of N2 O is inadequate. As a consequence one has to resort to cryogenic trapping at liquid nitrogen temperature, for which Stevens and Krout (1972) had developed very ef®cient capillary traps. Highly ecient, robust traps, known as Russian Doll cryogenic traps, containing a set of

40

C.A.M. Brenninkmeijer et al. / Chemosphere: Global Change Science 1 (1999) 33±52

nested concentric borosilicate glass ®ber thimbles have been developed since (Brenninkmeijer, 1991). A further improved version Russian doll trap also appears to be suitable for trapping, and releasing, the microliter quantities of CO derived CO2 (Brenninkmeijer and R ockmann, 1996). In this paper large isotope e€ects for CO at liquid nitrogen temperatures are noted. The amount of air to be processed is generally large. If 14 C analysis is required, a quick calculation can be made. Assuming a precision of 1% is desired, and that Poisson statistics apply, AMS detection requires 10,000 counts. Because the overall eciency of AMS is about 1%, 106 14 CO molecules are required. Given an abundance of for instance 10 molecules cmÿ3 , at least 100 l of air must be processed. If no 14 C analysis is required, sample amounts are prescribed by the mass spectrometry. For the conventional dual inlet systems, using a micro volume, several microliter of CO2 are necessary. Note that for the 17 O analysis, CO2 has to be converted to O2 because in CO2 based mass spectrometry both 13 C and 17 O variations contribute to the detected mass 45 changes. Generally somewhat larger amounts of sample are required. A fast, elegant method for the required conversion of CO2 to O2 has recently been developed (Brenninkmeijer and R ockmann, 1998a). Enormous progress in stable isotope analysis has been made possible by the development of continuous ¯ow mass spectrometry (Merritt, 1993; Merritt et al., 1995), which is presently being developed for CO (Mak and Yang, in press). Here gas chromatography and isotope ratio mass spectrometry are combined, allowing nanomole samples of CO to be analyzed. The major advantages are that it is not required to have large air samples available, and that the procedure is far less time consuming than bulk extractions. The actual oxidation step of CO to CO2 involves a special oxidant, being based on I2 O5 , utzeÕs H2 SO4 and silicagel, often referred to as Sch reagent (Sch utze, 1949; Smiley, 1965). This reagent merely adds an oxygen atom to the CO molecule at room temperature, without changing the original isotopic composition. Also the formed CO2 does not exchange in contact with this material, which is due to its strong acidity. The reproducibility of

the oxidation step using Sch utze's reagent is impressive. The major drawback of this oxidant is that its active compound is hydrolyzed by water vapor, and only dry air should be processed. For details of the methodology we refer to the extensive descriptions given by Stevens et al. (1972), Brenninkmeijer (1993) and R ockmann (1998). Also a report from the AES Canada is available (Norman et al., 1997). The essential aspects are as follows: One is that the extraction procedure allows a volumetric, absolute determination of the CO concentration. Independent CO concentration measurements are useful of course. Otherwise the performance of the extraction system is checked using a calibration gas made by dynamic dilution. Another aspect is that for 14 C analysis the CO derived CO2 sample is diluted with 14 C free CO2 . The reason is that AMS laboratories require more CO2 for graphite target preparation than is provided by the sample CO2 . The dilution by a factor of 10 or often more of the sample CO2 , which has a relative high speci®c activity, is a critical step. Isotope dilution can be used to independently check the degree of dilution (Brenninkmeijer, 1993). Dilution factors should be kept as low as possible and the dilution gas should preferably be free of 14 C. Proper corrections for fractionation and blanks are essential, and not trivial. A critical issue is the calibration required for 18 O isotope analysis, and some discrepancies may have existed in the literature. One has to correct for the isotopic composition of the oxygen derived from the Sch utze reagent. Several options exist, like, disproportionation of CO to CO2 , in which the isotopic integrity is maintained, or the production of CO with known isotopic composition by either the reaction of CO2 with C (Brenninkmeijer and R ockmann, 1997) or the reaction of isotopically well-de®ned water with C to CO (Stevens and Krout, 1972). The isotopically wellde®ned CO thus obtained can be injected in a ¯ow of zero air, to closely approach conditions under which atmospheric samples are processed. Laboratory inter-comparisons are advisable for the 18 O correction. The range of d18 O values of atmospheric CO is from near ÿ10& to 25&, and such a magnitude requires special care to exclude artifacts like

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isotopic exchange of the CO2 formed. In this context we also point out that the Sch utze reagent not only oxidizes CO, but with di€ering eciencies also acetone, propane, ethylene, acetylene, toluene etc. Note that some of these may not be scavenged by the cryogenic cleaning traps and can a€ect d values for samples from polluted air. Because the collection and storage without contamination or isotopic change of large air samples is dicult, Brenninkmeijer initially used an in situ extraction system. In situ CO extraction systems have been described by Brenninkmeijer (1993), Norman et al. (1997), and Hu€ et al. (1997). By using treated aluminum cylinders (Scott Marrin), and special compressors (Mak and Brenninkmeijer, 1994), sampling could take place at any location, including aircraft (Brenninkmeijer and Roberts, 1994), (Brenninkmeijer et al., in press). 6. Carbon-13 variations Overall, the 13 C/12 C ratio of atmospheric CO closely re¯ects that of organic material, which after all is its main precursor. Not much isotopic fractionation appears to occur in the various processes leading to CO formation and the source value itself is the most important parameter. CO produced by internal combustion engines, or high temperature combustion in general has d13 C values between ÿ27& and ÿ25&. Although CO represents only a small fraction of the CO2 evolved in combustion, thus leaving ample room for isotope fractionation, high temperatures tend to reduce fractionation substantially. Rather di€erent values are expected for the combustion of natural gas d13 C  50& (Kato et al., 1999). With its increasing share in energy production, its isotopic signature may be visible in CO present in certain urban centers. A comprehensive survey of fuels and CO isotopic signatures was given by Stevens et al. (1972). The relevance of engine CO for the regional budgets has changed dramatically in many countries due to the implementation of ecient exhaust gas catalysts. Our own attempts to isotopically characterize car exhaust in Mainz were thwarted repeatedly by the near absence of CO in

41

the exhaust of most vehicles. It appears to us that it is still necessary to better characterize combustion sources, although for many important applications in particular natural sources of CO need to be isotopically better de®ned. For instance for the photochemical oxidation of isoprene and terpenes scanty information is available only. A distinctly 13 C depleted source of CO is the atmospheric oxidation of CH4 , although experimental evidence based on laboratory measurements is blatantly absent. CH4 is largely a biogenic product for which its atmospheric d value is about ÿ47&. The combined sources have an inferred d13 C value of about ÿ52&, because in its oxidative removal initiated by OH a fractionation of about 5& is incurred. In turn, CO from CH4 oxidation should again have a similar value of near ÿ50&, assuming that after the formation of CH3 radicals no further fractionation takes place. However, because 5±30% of the carbon is not converted to CO, there is some room for fractionation. For instance the photolysis of the formaldehyde (HCHO) intermediate may be isotope selective. The e€ect of the light methane oxidation source should be most clearly visible in the southern hemisphere, where the other sources are weak compared to the northern hemisphere, as witnessed by the north south gradient in CO. Indeed, average d13 C values for Scott Base, Antarctica and Baring Head, New Zealand, are lower (Brenninkmeijer, 1993) than the values obtained by Stevens et al. for Illinois, and even approach ÿ32& in late summer. Fig. 1 illustrates the pronounced di€erence between background air in the Arctic and Antarctic. During summer, d13 C values should increase due to the increase in OH, which accelerates removal. Indeed a small initial decrease is seen in Spitsbergen. In Antarctica the e€ect of the increased apportionment towards the CH4 source is overriding, and d13 C decreases by 4±5& in summer. The aircraft data from Mak and Brenninkmeijer (1998) show that with increasing latitude towards the south pole, upper tropospheric d13 C values actually increase to ÿ29&, with a concurrent decrease in CO, down to levels below 40 nmole/mole. The only explanation for this e€ect is the fractionation of the CO during transport. The

42

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CH4 oxidation source apparently had little in¯uence, which may be attributed to the reduced reaction speed at the low temperatures. In the same work, also very low d13 C values of below ÿ32& are reported for the tropics, demonstrating the e€ect of temperature on increasing the CH4 derived CO contribution. The isotopic ``invisibility'' of local sources is to some degree coincidental and caused by the 13 C depleted CH4 oxidation source, in combination with fractionation. Assuming steady state, CO from fossil sources will be enriched by several & during its atmospheric life. The addition of 13 C depleted CO from CH4 oxidation, ameliorates the overall enrichment of the atmospheric mix, having the e€ect of making its d13 C value less distinct from that of CO emitted by combustion sources. Nevertheless Hopper et al. (in press) show that use of 13 C in regional studies can be made, in which local sources contribute CO that is about 5& enriched. Although the annual cycle at Baring Head (and Scott Base, Antarctica) shows inter-annual variability, the main features are suciently distinct to produce a characteristic annual pattern that has been used by Manning et al. (1997) to constrain the southern hemisphere CO budget. The model derived and measured annual cycle for CO and 13 C are shown in Fig. 3. One outcome of this study was that the d13 C values obtained are still somewhat too high to fully account for the CH4 oxidation input. Also, a substantial ocean source had to be incorporated to make the budget match. Conny et al. (1997) rely almost entirely on modeling to reconstruct the 13 C isotopic composition under the in¯uence of biomass burning in Brazil. Their useful exercise shows extreme CO variations due to widespread, regional biomass burning, with variations arising from 2 processes. Changes in 13 C are expected on the basis of the innate 13 C di€erences between forest and Savannah burning due to the di€erent distribution of C3 and C4 plants. Furthermore, the biomass burning CO from both types of plants di€ers strongly from background CO. The d13 C value predicted for background air CO is necessarily low at near ÿ34& , because due to the high temperatures in the tropics, the amount of CO derived from CH4

Fig. 3. Measurement and model results for the annual cycle of d13 C(CO) and CO in the southern hemisphere (Manning et al., 1997). The contribution to the atmospheric inventory from the following sources is represented: C, Methane oxidation. Note the relatively low constant contribution, which however has a pronounced e€ect on d13 C; O, Ocean emission; I, Industry, i.e. CO imported from the northern hemisphere; V, Isoprene oxidation (Vegetation); T, Terpene oxidation; F, Forest burning; S, Savanna burning.

oxidation is high. Moreover, the fact that OH is high in the tropics also leads to the low model background d13 C values because CO from competing sources is decreased. The contribution from CH4 oxidation however is high and una€ected, with more CO being oxidized, but simultaneously more being produced by the CH4 oxidation. It will be useful to obtain CO isotope measurements from the tropics, but as a marker of regional biomass burning, the actual CO concentration often will be a simpler indicator. Also 18 O variations will be a useful indicator. Extremely low d13 C values (Fig. 4) have been encountered during aircraft ¯ights in the lowermost stratosphere over Antarctica during ozone hole conditions (Brenninkmeijer et al., 1996).

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Fig. 4. d13 C of CO in the polar vortex during ozone hole conditions, October 1993. The extreme 13 C depletion for CO mixing ratios below 30 nmole/mole (stratosphere) can only be explained by the large fractionation in the reaction Cl + CH4 . The tropospheric values (60±100 nmole/mole) show the large variations in CO in remote air towards the end of the southern hemisphere biomass burning season. The concomitant variations in d13 C is small, and do not show any trend (for a discussion of the relative 13 C isotopic invisibility of CO from burning processes see text).

It was only after laboratory measurements by Saueressig et al. (1995) had revealed an enormous kinetic isotope e€ect of 60 to 70& in the reaction Cl + CH4 , that the measurements could be interpreted through the use of a box model describing the fate of an air mass circulating in the polar vortex (M uller et al., 1996). The interesting e€ect thus is the combination of CH4 as a depleted precursor for CO, with the availability of sucient free Cl and the strong fractionation in CH4 + Cl. The modeling showed that the low d13 C could be achieved without having to make extreme assumptions. Notwithstanding there is very little doubt that smaller fractionation factors for Cl + CH4 , as has been suggested in the literature based on ab initio calculations, would not suce to explain the hard facts of the measurements. Apart from the Cl related deviations, lower stratosphere d13 C values were found not to di€er much from tropospheric values (Brenninkmeijer et al., 1995). There is little fractionation in its removal at these pressure levels, whereas at higher altitudes further decreasing values are expected with CO from CH4 oxidation slowly replacing the CO imported from the troposphere.

43

The potential of 13 C in CO as a gauge of free Cl, which is extremely dicult to measure directly, has been demonstrated convincingly in the troposphere as well. R ockmann et al. (1999) have shown that during surface ozone loss events in the Arctic, sucient 13 C depleted CO is formed from Cl + CH4 to give a measurable signal. For three consecutive years they detected brief d13 C excursions of about ÿ0.5&, coinciding with ozone depletions. Their estimates of the amount of free Cl that had been available compares very well with the indirect estimates based on the selective rate of removal of certain hydrocarbons (Ramacher et al., 1997). The question that follows is of course whether free Cl in the marine boundary layer could be detected. Unfortunately, the not so low d13 C values encountered at the coastal site Baring Head in New Zealand almost seem to imply that the overall contribution from Cl + CH4 in the marine boundary layer must be negligibly small at 40° south. Episodic events of substantial Cl activation however could still occur and be detectable using 13 C. The main experience obtained with 13 C measurements can be summarized as follows: 1. Local combustion sources generally do not give a clear signal. 2. The annual cycle in the southern hemisphere shows the e€ect of enhanced methane oxidation in summer. 3. The summer minima in d13 C will be very sensitive towards the fraction of non-methane oxidation sources. 4. As a rule d13 C has a clear annual cycle with little scatter. 5. The kinetic isotope fractionation seems to generally have a small impact, and may be noticeable in the northern hemisphere where the e€ect of the CH4 oxidation plays a lesser role. 6. The reaction of CH4 + Cl can be detected in the stratosphere and troposphere thanks to the large kinetic isotope fractionation combined with the 13 C depletion of CH4 . 7. Generally d13 C values in the remote southern hemisphere are not as low as would be expected from the importance of the CH4 oxidation. There may be some fractionation in the CO formed relative to the parent CH4 . Before

44

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drawing any conclusions, laboratory measurements on CO from CH4 + OH are necessary. 7. Oxygen-18 variations The variations in d18 O are larger than in d13 C, and can play a useful role, even though the measurement error for 18 O is somewhat larger as well. d13 C being mainly in¯uenced by the methane source, d18 O has two major causes of variability, i.e. combustion sources, and kinetic fractionation. During combustion chie¯y air oxygen is incorporated in the formed CO, whilst high temperatures rule out fractionation. As was ®rmly established by Stevens et al. (1972), CO from combustion sources is indeed enriched in 18 O and nearly matches the isotopic composition of atmospheric oxygen which has a d value of 23.5&. Because combustion is the single large source of enriched CO, southern hemisphere d18 O values are the lowest. Fig. 1 shows clearly not only the high NH values, but also the accompanying large annual cycle of over 10&. Conversely, in the SH the e€ect of local combustion sources is most clearly discernible, and Brenninkmeijer (1993) could correlate changes in CO with corresponding changes in d18 O in Antarctica due to local sources like snow vehicles. Using a dilution model, Brenninkmeijer and R ockmann (1997) derive a clear inverse correlation for CO and d18 O, with an intercept of near 23& (Fig. 5). This agrees with the value for atmospheric oxygen, and con®rms the use of the proper correction for the isotope shift induced by the oxidation by the Sch utze reagent. More data on CO from combustion sources is desirable however. Kato et al. (1999) present a large number of measurements of city air with CO values up to 2 ppm, and d18 O values reaching 20&. As expected, the lowest d18 O values occur in the stratosphere in the southern hemisphere, due to the remoteness from combustion sources and the ongoing isotope depletion during transport due to the inverse kinetic oxygen isotope fractionation in CO + OH. C18 O reacts about 10& faster than C16 O, and this fractionation was measured by Stevens et al. (1980) and R ockmann et al. (1998c). A con®rmation of this laboratory value was given

Fig. 5. The e€ect on d18 O of fossil fuel derived CO mixing with background CO. Shown is a dilution behavior between background air in the remote southern hemisphere (low d18 O values) and local sources. Often the source signature for 18 O may not be as sharply de®ned as for the current case. At some distance from the polluting sources, the d18 O signature may have changed due to the kinetic isotope e€ect for CO + OH. (Brenninkmeijer, 1993; Brenninkmeijer and R ockmann, 1997).

by Brenninkmeijer and R ockmann (1997) who used the southern hemisphere aircraft data to calculate the fractionation. The value of about 8& derived by them is in reasonable agreement. The lowest d18 O value ever measured was ÿ9.5& for samples collected by aircraft in the lowermost stratosphere. However, because the oxidation of CH4 increasingly contributes to the vertically decreasing CO inventory in the lower stratosphere, the depletion should bottom out. The lower limit in d18 O would be reached by CO produced in equilibrium through the oxidation of CH4 by OH and O(1 D) at stratospheric temperatures. Biomass burning is a moderately enriched CO source, which most likely is related to the lower temperature of combustion. In particular during the smoldering combustion phase, which is rich in CO, temperatures may not be high enough to eliminate fractionation. Indeed, analysis of CO in cigarette smoke gives values of around only 15& (Brenninkmeijer, 1993). At high combustion temperatures, higher 18 O values are expected however. Interception of certain air masses during the southern hemisphere aircraft ¯ights with enhanced CO allowed an estimate of the d18 O value of CO from biomass burning after its arrival at the mid to

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higher latitudes. Values of about only 4& were found (Brenninkmeijer and R ockmann, 1997). This was interpreted as showing that the air masses had undergone considerable chemical changes, and that the CO had become fractionated. Also CO precursors emitted during the biomass burning may be a concomitant source of CO less enriched in 18 O, with photochemical oxidation of volatile organic carbons being a CO source with relatively low d18 O values. Isotopic analysis of CO from large biomass burning plumes in east Siberia in August 1996 (Bergamaschi et al., 1997) produced d18 O values of 8±10&. These air masses had been chemically little altered, as back trajectories indicate recent ®res as the source. Accordingly d18 O cannot have changed much, and we have here a relatively low d18 O value for biomass burning. It would be interesting to see if the d18 O value of groundwater can play any role in in¯uencing the d18 O value for biomass burning CO. It is clear that more measurements of CO from biomass burning are required. The d18 O values of CO from methane and hydrocarbon oxidation are not well known. These values are not easily measured in the laboratory, yet, both Stevens and Brenninkmeijer have attempted to estimate its value based on the atmospheric observations. Some uncertainty exists here as Brenninkmeijer and R ockmann (1997) estimate a value near 0 per mil. The earlier estimate by Stevens and Wagner (1989) was higher, mainly due to relatively high d18 O values measured for some samples from Australia. We cannot be a priori sure that the d18 O value for CO from OH + CH4 is constant. Stevens (1993) has suggested that formaldehyde may undergo oxygen isotopic exchange with precipitation. This could lead to lower values at higher latitudes, where 18 O values in precipitation are lower. Clearly some work is required here. At high northern latitudes a large d18 O annual cycle exists, with a summer minimum as low as ÿ2&, and a winter maximum of almost 10&, and such a clear signal is supposed to provide useful constraints when modeling is used. The reason for the large annual cycle is that during summer, the kinetic isotope e€ect has its largest impact as CO is broken down maximally during transport from source areas to the Arctic. In summer also isoto-

45

pically light natural sources i.e. CH4 and VOC oxidation are more dominant. In winter, CO building up in the boundary layer in absence of OH most closely re¯ects CO of the mid-latitude and higher latitude surface source areas, fossil fuel combustion with its high 18 O content. The main experience obtained with 18 O measurements can be summarized as follows: 1. Combustion sources possess a distinctly enriched signal which can be used with success for calculating their contribution accurately. 2. The large kinetic fractionation o€ers the possibility to assess the degree to which CO has been removed by OH. 3. The impact on d18 O from biomass burning derived CO in the remote southern hemisphere is relatively small. 4. The d18 O value of CO from CH4 and non-methane hydrocarbon oxidation needs to be measured under controlled conditions. There is a large degree of uncertainty of the value derived form atmospheric observations. 8. Oxygen-17 variations Precursory measurements (Hodder et al., 1994) established that CO has a small excess of 17 O compared to nearly all other substances on earth, although an explanation could not be given. This intriguing new e€ect of mass independent enrichment in CO was subsequently ®rmly established by measurements on the air samples from Spitsbergen and Alert (Fig. 6), and samples from the US (Hu€ and Thiemens, 1996). The mass independent enrichment could in a ®rst attempt be partly explained by R ockmann et al. (1998b) on the basis of the ozonolysis of unsaturated hydrocarbons like terpenes and isoprene. During ozonolysis CO is formed, the oxygen of which is derived from O3 . Because tropospheric O3 is strongly mass independently enriched (d17 O  80&, and d18 O  100&, thus D17 O  30&), a signi®cant source was identi®ed. However, it was also realized that this could not be the only source, as there is insucient CO produced via ozonolysis to satisfactorily explain the atmospheric 17 O values. Moreover, a source of sucient strength to explain the mass

46

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Fig. 6. The time series for CO, d18 O, and D17 O in Spitsbergen (closed symbols) and Alert (open symbols). During fall and winter, when removal by OH and consequently fractionation is small, and when fossil sources outweigh natural sources, CO and d18 O predictably increase. Concurrently, D17 O decreases however, hinting to either natural sources as its origin, or CO + OH (R ockmann, 1998; R ockmann et al., 1998b).

independent signal in 17 O, would simultaneously produce a strong 18 O signal, which however is not observed. The search for the main origin of the e€ect was successful (R ockmann et al., 1998c), and surprisingly it was identi®ed to be the sink process CO + OH itself, in which the remaining CO fraction slowly acquires excess 17 O. Although a theoretical explanation for the mass independent enrichment in CO is lacking at this moment, as it is for O3 , the fact that the origin is known and its fractionation constant now has been measured in the laboratory, opens exiting new possibilities for applying it to the global CO cycle. In e€ect, D17 O (the 17 O excess) is a direct measure for the exposure of CO to OH. As far as we know, all CO

emitted or produced in the atmosphere, except for the ozonolysis, has a D17 O value of zero. In other words, the 17 O excess is the only isotope e€ect for which the source value is exactly known and constant. The mass independent oxygen isotope enrichment in atmospheric CO has a contraindication as well. In all CO isotope work published up to date, the CO2 mass spectrometry used for d13 C and d18 O determination was based on the assumption of mass dependent behavior. Now it is known that CO is mass independently fractionated, corrections have to be applied, which range from about 0.1& to 0.3& (R ockmann and Brenninkmeijer, 1998). The main experience obtained with 17 O measurements can be summarized as follows: 1. CO in background air possess mass independent enrichment in the oxygen isotopes, which is witnessed by a small but clear excess in 17 O. 2. All previously published 13 C data need a correction. 3. The ozonolysis is a small source, producing comparatively strongly mass independently enriched CO. 4. The major origin for MIF is the reaction CO + OH, which makes this signal a direct measure for exposure to OH. 9. Atmospheric

14

CO

Weinstock (1969) from Ford motor company realized that the chemical lifetime of CO in the atmosphere could be calculated by considering the fate of 14 CO, which indeed has the same sinks as CO itself. Using few existing measurements of 14 CO, and knowing its production of cosmogenic origin, he derived a lifetime of only 0.1 yr. Junge et al. (1971) tried to understand the global CO budget using model calculations and 14 CO. Discrepancies surfaced mainly due to the in those days unrecognized important role of OH. Weinstock and Niki (1972) revisited the issue shortly thereafter and con®rmed their earlier estimate. This short lifetime perhaps did come as a relief for automobile manufacturers, but was depressing to those atmospheric chemists who had believed that

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chemistry only took place under the harsh radiative environment of the stratosphere. Volz et al. (1981) performed the ®rst systematic 14 CO study, processing samples as large as 20,000 l (using NaOH solutions to strip CO2 ) to extract enough 14 CO as 14 CO2 for applying proportional beta counting. In this way he constructed the annual 14 CO cycle for the mid northern latitudes, and showed that it agreed with existing OH distribution and seasonality. The development of accelerator mass spectrometry for 14 C, which revolutionized 14 C dating, enabled Brenninkmeijer et al. (1992) to make routine 14 C measurements in the southern hemisphere, using 1 m3 air samples. To their surprise mid latitude SH levels were much lower than the NH data from Volz et al. (1981), a fact which was con®rmed by further measurements (Brenninkmeijer and R ockmann, in press; R ockmann and Brenninkmeijer, 1998). Various mechanisms, like di€erences in production due to changes in solar activity, latitudinal gradients, troposphere stratosphere exchange, and biogenic 14 CO were not considered adequate to explain the large di€erence, leading to the suggestion that OH levels in the SH may be higher. This issue has not yet been resolved. The most fundamental aspect remains the stratosphere±troposphere exchange, which is the subject to much current research. Fig. 1 highlights the 14 CO di€erence between the high latitude northern and southern hemisphere, although the record shown is not complete. Solar activity was somewhat higher in the 1992± 1993 data for the SH, but this would not account for more than 1 molecule cmÿ3 . The higher NH data can partly be accounted for by a larger biogenic 14 CO contribution. The excess CO ranging from about 100 nmole/mole in winter to 50 nmole/ mole in summer could explain, depending on the assumed fossil fuel derived fraction, 2.5 to 2 molecules cmÿ3 . A substantial, yet unexplained NH SH di€erence in 14 CO from cosmogenic origin remains. The question to what extent 14 CO can help to constrain OH distribution and seasonality is important because OH is central in the chemistry of the troposphere, yet surrounded by some uncertainties. The current situation is that OH is

47

calculated using sophisticated 3D chemistrytransport models. The integrated OH values derived from the resulting ®elds are assumed to be correct to within 10±20%, and often distribution and integrated value are veri®ed using measurements of methylchloroform (Spivakovsky et al., 1990). However there are certainly uncertainties associated with this process, calling for independent assessments. Notably the seasonal variations in methylchloroform, and its interhemispheric gradient, are small due to its rather long lifetime. One obvious direct venue to verify the model calculations is to measure OH abundances in the atmosphere. However due to the diculty of these measurements and the high variability of OH with its lifetime of only seconds, this is still problematic on a large scale. The replacement HCFCÕs o€er yet another possibility to gauge OH indirectly, but also here some problems are foreseeable. 14 CO has its speci®c challenges as well. However, it is clear that because its source is largely beyond human interference, this fundamental geophysical parameter will keep its value. Moreover, the lifetime of 14 CO is ideal for assessing OH. A problem however is posed by the fact that its production is at 60±70% in the stratosphere, where chemical removal is slowly leading to a large gradient across the tropopause. Thus only models that properly describe the ¯ux from the stratosphere into the troposphere are adequate to predict tropospheric 14 CO levels. Also 14 CO production rates peak in the higher latitudes lower stratosphere, whereas its removal of course is dominated by the tropics where OH is abundant. Conversely, if the OH distribution is known with great con®dence and in detail, 14 CO may provide information about the rates of transport from the stratosphere. Another consequence of the vertical gradient in the 14 CO production rate, which leads to stratospheric levels as high as 150 molecules cmÿ3 STP (Brenninkmeijer et al., 1995), is that tropospheric levels do show scatter. For instance surface measurements in New Zealand showed changes for 12± 14 molecules cmÿ3 within days. Also inspection of the extensive data sets obtained by Mak et al. (1992) show considerable variability. Undoubtedly is part of this inherent to the relatively short

48

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lifetime of CO. Nevertheless the question arises whether sampling should be restricted to a certain point in time, or be integrated over longer periods. In the former case, high resolution modeling is necessary to interpret the results. In the latter case details may get lost. The relatively high cost of 14 C analyses should be considered as well. The problems to model 14 CO were borne out by the studies published so far. Using 2 di€erent 2D models Mak et al. (1992, 1994) have shown that it is dicult to reconcile the measurements and models. For both models transport had to be changed drastically. Moreover, predicted values were anyway far too high. The size of the discrepancies is a reminder that possibly even after more sophisticated modeling, OH values may need reconsideration. Because global OH is dominated by its high levels in the tropics, measurements there are very important. Mak and Southon (1998) evaluate 14 CO measurements from Barbados, and ®nd a larger than expected seasonal change. Part of the variations in 14 CO is due to transport from the mid latitudes, but even after correcting for this, some discrepancy in seasonality remains. Measurements in the tropics are particularly useful for 14 CO as the impact of OH is here at its greatest. One application of 14 CO that has not yet been mentioned thus far is its use as a tracer for certain sources. For instance, fossil fuel sources produce CO devoid of 14 C, and in this way, measurements in polluted areas can shed some light on the fossil source strength. Obviously, one tends to ®nd a predictable correlation between 14 C as pMC or speci®c activity, and CO concentration. Fig. 7, gives a clear example where local sources contaminated background air masses. However, when biomass burning is involved, the picture is di€erent due to its 14 C content, and CO derived from wood burning stoves, can be distinguished from that of fossil fuels (Klouda and Connolly, 1995; Klouda et al., 1988; Sakugawa and Kaplan, 1997). Clear isotopic evidence for extensive biomass burning was obtained by Bergamaschi et al. (1997) for polluted air masses in east Siberia during a TROICA (Trans-Siberian Investigation of the Chemistry of the Atmosphere) expedition with the Trans-Siberian railroad. Over a distance of near

Fig. 7. The e€ect of fossil fuel derived CO from local sources mixing with background CO, given as pMC, i.e. the speci®c activity relative to contemporaneous organic material. Shown is a simple dilution behavior between background air in the remote southern hemisphere and local fossil sources.

2000 km elevated CO levels peaking at 1500 nmole/mole were encountered in August 1996. North of the river Amur 14 CO abundance rose from the background value of 12 molecules cmÿ3 to 40, concurrent with CO increases from 100 to 900 nmole/mole. The slope of the 14 CO increase amounted to 38 molecules cmÿ3 per 1000 nmole/ mole of CO increase. This corresponds to a speci®c activity of 120 pMC, and identi®es biomass burning as the source. Tyler et al. (1999) report 14 CO, CO and d13 C data for Niwot Ridge, Colorado, for 1991±1992. 14 CO values range from about 10 molecules cmÿ3 in summer to near 20 in winter. Assuming that models will correctly describe transport and OH, the question of the exact 14 CO production rate and its distribution will become important again. We presently have at our disposal two di€erent production distributions. The main di€erence between the distribution calculated by OÕBrien (1979) as compared to the older calculations from Lingenfelter (1963) is that the former has considerably lower production near the surface due to a stronger attenuation of cosmic radiation with atmospheric depth. The sensitivity of the resulting 14 CO distribution in a 3D model to the di€erent distribution has been calculated and will be published soon (J ockel et al., 1999).

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One venue to ®nd an answer is to actually measure the in situ 14 C production. Experiments have been carried out by exposing cylinders with ultra pure air doped with 14 C free CO to cosmic radiation. Basically because inside the cylinders OH is absent, 14 CO levels continuously increase. Suitable sites for exposure obviously are higher latitudes and altitudes, and locations selected are South Pole Station, Mount Cook, New Zealand, and others. Results of these experiments are in process of being published. Also measurements of 14 CO produced inside aircraft tires have been carried out, showing elevated 14 CO. However the use for calibrating production rates is problematic (unpublished results). The main experience obtained with 14 CO measurements can be summarized as follows: 1. 14 CO measurements can be performed routinely on samples as small as 100 l. The lower limit is given by counting statistics. 2. 14 CO provides a very detailed picture of e€ects due to transport combined with its removal by OH. 3. Di€erent air masses have distinctly di€erent 14 CO values depending on OH and origin. 4. More modeling is required, in particular using high resolution models. 5. It is probable that the biogenic 14 CO contribution can be adequately assessed based on concentration and stable isotope measurements. 6. There is an as yet unexplained di€erence between the northern and southern hemisphere, but more measurements and comparisons are required. 7. Up to date no indications of signi®cant quantities of 14 CO from the nuclear industry have been obtained. 8. Experimental determination of the cosmogenic 14 C production rate would be useful. 9. A simple collection method for CO for 14 CO analysis would be very useful. 10. Concluding remarks

many useful ways complements the concentration data. Little is known about the fractionation in the uptake of CO by soils, and clearly some work has to be done here. Likewise we have insucient knowledge of the isotopic composition of CO from the oceans. Stevens (1996) has carefully re-analyzed his 1971 data in the light of new developments, and pointed out that 14 C can be used to help to estimate the fossil fuel derived CO fraction. This highlights the strength of using the combined isotope analysis of CO. The recent identi®cation of CO + OH as a process which causes mass independent enrichment in the surviving CO adds a new dimension to the existing isotope work, apart from raising the rather fundamental question as to its origin. It seems to us that in particular in view of the rather poorly de®ned cycle for this important atmospheric trace gas, isotopic analyses are warranted and desirable. In recent years an increasing number of researchers have embarked on this type of work, with the great bene®t of many more data becoming available. The continuous ¯ow mass spectrometry is bound to have a considerable impact. Necessarily not all work has been available for inclusion into this review, and we apologize for possible omissions. Once a certain momentum of experimental research has been reached, modeling e€orts will become worthwhile and without doubt lead to an improved understanding of the global CO cycle. Acknowledgements Part of the work carried out to complete this review has been supported by the European Commission, DG XII, as part of the contract COOH-EUROPE. We thank J.M. Conny for useful comments on a draft. References

13

17

18

The isotopic composition for C, O, O and C of CO contains considerable unique information about sources, sinks and transport, and in 14

49

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