COSMOGENIC ¹⁴CO FOR ASSESSING THE OH-BASED SELF-CLEANING CAPACITY OF THE TROPOSPHERE

Abstract

An application of radiocarbon (¹⁴C) in atmospheric chemistry is reviewed. ¹⁴C produced by cosmic neutrons immediately forms ¹⁴CO, which reacts with hydroxyl radicals (OH) to ¹⁴CO₂. By this the distribution and seasonality (the lifetime of ¹⁴CO is ∼1 month) of the pivotal atmospheric oxidant OH can be established. ¹⁴CO measurement is a complex but unique application which benefitted enormously from the realization of AMS, bearing in mind that ¹⁴CO abundance is of the order of merely 10 molecules per cm³ not only provides ¹⁴CO an independent measure for the OH based self-cleansing capacity of the troposphere, but also enabled detection of ¹⁴C production due to high energy solar protons in 1989. Although its production takes place throughout the atmosphere and does not have the character of a point source, transport processes in the atmosphere affect the distribution of ¹⁴CO. Vertical mixing in the troposphere renders gradients in its production rate less critical, but considerable meridional gradients exist. One question has remained open, namely confirmation of calculated ¹⁴C production by direct measurement. A new sampling method is proposed. The conclusions are a guide to future work on ¹⁴CO in relation to OH and atmospheric transport.

INTRODUCTION

Before a fresh cosmogenic ¹⁴C atom (produced by ¹⁴N(n,p)¹⁴C) enters the main carbon cycle, it almost instantaneously forms ¹⁴CO and a such partakes in atmospheric chemistry for a period of a month or so. This brief intermezzo of not yet being fully oxidized in an otherwise fairly long life is of interest to atmospheric chemists. Bernard Weinstock had been a student of Harold Urey and worked on the Manhattan project. Subsequently he was for 20 years research scientist at Ford Motor Company, Dearborn, Michigan, a company with a brilliant scientist researching air pollution and the impact of combustion engine exhaust gases. In 1969, Weinstock published a brief paper in Science in which he estimated the atmospheric residence time of carbon monoxide. Pandow et al. (1960) already had focused on the fate of ¹⁴C recoil atoms (∼95% forms ¹⁴CO). Subsequently, MacKay et al. (1963) achieved the amazing first ¹⁴CO measurements (by Minze Stuiver) using CO samples obtained from an air liquification plant (huge air volumes needed to be processed) and Mackay suggested ¹⁴CO to be a useful atmospheric tracer. But Weinstock (1969) managed to use these ¹⁴CO data to estimate the lifetime of ¹⁴CO and thus of CO itself (¹²CO and ¹⁴CO for this purpose behave identically). When one of us, as a Dutch student (Groningen Radiocarbon Lab under W. G. Mook and P. M. Grootes) in New Zealand (Institute of Nuclear Science under A. Rafter) in 1976, had written Bernard Weinstock about his approach, proposing to do more ¹⁴CO measurements, an encouraging answer arrived within two weeks by airmail. This was an exposure to the generous attitude of American scientists.

Why was this estimate of the lifetime of CO so important? Concern about air pollution (e.g., Los Angeles) was certainly a factor, but moreover; when the number of automobiles increases strongly and each produces about 100 g CO (or more) per gallon gas the question is, “Will this toxic gas accumulate in the global atmosphere?” It was early days for atmospheric chemistry. Christian Junge (1959) had listed a mere 9 atmospheric trace gases in his 1959 “Air Chemistry” paper (he noted: “It is very likely that only a few of the trace gases are known”). Furthermore, consensus was that CO had only an anthropogenic source and that soils and the stratosphere were its only sinks. Having a guess of the input (it was thought that CO was mostly of fossil fuel origin) and of its atmospheric inventory did give a lifetime estimate. A totally independent lifetime calculation however, was worth a lot. Weinstock estimated using a rate of production of ¹⁴C at 1.3 × 10¹⁹ atomsec^–1, a lower limit of 0.1 year. An excellent number. The big problem in those days was, however, that it was becoming clear that neither the stratosphere nor soils (the known sinks) could take up tropospheric CO (and thus ¹⁴CO) at the rate necessary to obtain such a short lifetime.

The fast developments in early atmospheric chemistry led to the landmark paper by Levy (1971) in which he showed that high concentrations of the hydroxyl radical OH do occur in the troposphere. High levels of formaldehyde were predicted too, which would be a source of CO, a partly natural source. Prior to this paper by Levy, it was assumed that OH abundance was negligible. But suddenly, the troposphere turned out to be a true chemical reactor, alike a slowly burning cold flame. OH is the troposphere’s pivotal oxidant. For instance, nearly the entire annual input of CH₄ is removed by this radical, in fact >500 Tg/year. Likewise, CO and most Volatile Organic Carbons are removed by OH. It is hard to imagine how a biological-life-friendly atmosphere could function without OH.

The most compelling case for understanding the cycle and tropospheric distribution of OH is the budget of the important greenhouse gas methane. CH₄ has been monitored on a large scale for decades and changes in its natural and anthropogenic sources cause variations that are difficult to trace back. Time and again growth rate changes of the global atmospheric burden CH₄ have surprised us. The main factor in the budget of CH₄ is of course its sink, namely OH. Numerous detailed studies address the role of changes in OH (e.g., Zhao et al. 2019), which—to complicate matters—are partly attributable to CH₄ emission changes and their global pattern.

The UV photolysis of tropospheric O₃ leads to O₂ molecules and extremely short-lived excited O(¹D) radicals. These can react with H₂O to form 2 OH radicals. These stable but highly reactive OH radicals live for about a second before reacting with either CH₄, CO, organics and so on. One can by measurement of ¹⁴CO calculate the abundance of OH radicals, provided one knows the reaction rate constant for ¹⁴CO + OH (The rate is pressure dependent, at 200 hPa it is ∼30% lower than at sea level). This is the “¹⁴CO method” for determining OH, the subject of this article.

What makes ¹⁴CO relevant for OH determination? First of all, OH has a lifetime of seconds only. Its formation by daylight produces a strong diurnal cycle. Its concentration is much higher in summer than winter leading to a seasonal cycle. It is much more abundant in the tropics than at high latitudes. Its abundance varies with altitude due to mainly water vapour and UV flux gradients. In short, OH is extremely variable in space and time at different scales. Peak values are below 10⁷ radicals per cm³, sub parts per trillion levels. This rarity combined with its high reactivity made it initially impossible to detect OH, but successful detailed experiments to use tracer ¹⁴CO in a flow reactor with ambient air were done (Felton et al. 1992). This radiochemical method proved to be problematic. To date, in-situ measurement of OH still pose an experimental challenge in atmospheric chemistry (Cho et al. 2021). And even though presently one can obtain in-situ OH data, no large-scale distribution of OH can be reconstructed this way as too many detection systems globally distributed would be needed. Indirect methods for gauging OH are therefore called upon. Instead of trying to measure OH, one observes the removal of certain trace gases that react specifically with OH. The indirect methods have the advantage of mapping the effect of OH on reactive gases, all of which have different reaction times, sources and distributions. The in-situ OH detection is on the other hand essential in achieving a detailed understanding of radical chemistry. Atmospheric chemistry models allow us to construct OH fields, but for verification, measurements are needed. The indirect methods use a trace gas with known input into the atmosphere that predominantly reacts with OH. Initially (Singh 1977; Prinn et al. 1983) this was methylchloroform (CH₃CCl₃). However, methylchloroform has been almost eliminated from the atmosphere thanks to the Montreal Protocol measures. A recent paper by Patra et al. (2021) notes: “Estimations of tropospheric OH trends and variability based on budget analysis of methyl chloroform (CH₃CCl₃) and process-based chemistry transport models often produce conflicting results.” In their paper they show that no major change in the chemical removal is likely.

Apart from this problem of phasing out, an inherent problem remains that uncertainties about the rate of release exist, which also applies to other industrial trace gases that potentially may be applied in constraining OH estimates. The conclusions about OH from the use of tracers like methylchloroform, often in painstaking work and finest analytical work and modeling support over decades, are not always in agreement.

Alternatively, ¹⁴CO is a geophysical tracer that will be around for a long time. Work is even in progress to use ¹⁴CO measurements in air extracted from firn and ice to reconstruct OH concentrations in the past. We review the application of ¹⁴CO in atmospheric science with some explanation for a broader readership. We also mention a few less successful studies to help to avoid pitfalls in future research. We do not deal with in-situ production of ¹⁴C in solids (e.g., ice).

APPLICATION

Distribution of Tropospheric OH Concentration

Figure 1 shows the zonally averaged distribution of OH over one year derived with an atmospheric chemistry model. These ever more sophisticated models incorporate hundreds of chemical reactions that take place in the atmosphere. They are in essence atmospheric transport/climate models with embedded chemistry modules dealing with trace gases, particles, their sources, clouds, precipitation chemistry and so on. The tantalizing aspect is that these models show us a distribution of OH that we cannot measure, in contrast for instance to the model derived distribution of O₃, which is the most widely measured reactive trace gas. Different models have different OH distributions. Nicely et al. (2020) show that there are still large differences between OH distributions generated in current models. The main feature of model results is the high concentration of OH in the tropospheric tropics (the tropics are the atmosphere’s cooking pot). Basically, the amount of UV light, ozone and water vapor determine OH. With most anthropogenic emissions (which are considerable compared to natural emissions in this period of the Anthropocene; Crutzen 2021), including the key catalyst NO that sustains the oxidating cycle, being in the northern hemisphere, most models produced a higher OH abundance in the Northern Hemisphere. The publication by Spivakovsky et al. (1991) on OH has been frequently cited, underscoring the key role of OH. Work on using trace gases for assessing OH distribution is of course ongoing. For instance, Wolfe et al. (2019) mapped the variability of OH throughout the global remote troposphere via the synthesis of airborne and satellite formaldehyde observations.

Figure 1 Annual zonal mean OH concentration in 10⁶ cm^–3 (volume, not STP) as simulated with the EMAC chemistry climate model (Jöckel et al. 2016). The dashed line is the tropopause. Tropospheric OH is slightly higher the NH.

Over the years, much has been learned about OH, in particular its degree of recycling (Lelieveld 2004). For instance, when an OH radical reacts with CH₄, it is lost. However, some of the follow-on products are again radicals and with NO these can lead to the re-formation of OH. To “definitely” remove one OH radical takes up to 20 CO molecules (Gromov 2013).

Distribution of ¹⁴CO Production

Figure 2 shows the calculated production rate distribution of ¹⁴CO based on Lingenfelter (1963). The mismatch with the OH distribution is large. We see high ¹⁴CO production rates at high latitudes and higher altitudes determined by the interaction of cosmic rays with the atmosphere and the nudging by the Earth’s magnetic field. Lowest production is in the tropics. High production combined with the longer lifetime of ¹⁴CO in much of the stratosphere (by mass) implies that its rate of transport into the troposphere (overall variable with season and stronger in the NH) is a key factor. In the troposphere ¹⁴CO is transported as well, be it to a degree that depends on its regional lifetime, ranging from one (tropics) to several months (at high latitudes).

Figure 2 Zonal-annual-mean galactic cosmic ray ¹⁴CO production rate (Lingenfelter 1963). The unit is 10^–3 molec g⁻¹ s⁻¹ normalized to a global average production rate of 1 molec cm^–3 s⁻¹. The yield of ¹⁴CO from ¹⁴C oxidation is assumed to be 95%. The contour lines refer to ¹⁴C production by solar protons which is discussed later in the text. (Figure from Jöckel 2000.)

The result of the contrasting distributions of the rate of ¹⁴C production (Figure 2) and of OH (removal) is shown in Figure 3 as the resulting distribution of ¹⁴CO itself. A model takes transport and the pressure dependence of the reaction rate constant into account. ¹⁴CO mixing ratios are high in the maximum production region (Figure 2) and almost exactly there, OH bottoms out. Moreover, in the tropics where ¹⁴CO is rarer, annual variations are small. High concentrations occur at higher latitudes, accompanied by large annual variations reflecting OH seasonality. Figure 4 shows a high latitude annual cycle of ¹⁴CO measured at Spitsbergen and Alert (Ellesmere Island, Canada), as a result of the ongoing production of ¹⁴CO, its transport and mixing of air masses, the annual averaged level of OH and its strong local seasonal cycle. ¹⁴CO accumulates in the high latitude winter limited only by transport and mixing with air masses at mid latitudes (lower ¹⁴CO). In summer local photochemical removal by OH becomes important.

Figure 3 Zonal-annual-mean distribution of ¹⁴CO (color scale) as a result of its production distribution (Figure 2) and the OH field (shown black contour lines). The dotted line indicates the tropopause. ¹⁴CO is given as a mixing ratio. OH is given as in Figure 1, molec cm^–3.

Figure 4 Measured annual cycle of surface 14CO at Ny Ålesund (78.90°N, 11.88° E, circles) and Alert (82.45°N, 62.52°W, triangles). The annual cycle is large. Accumulation of CO (mainly pollution) and ¹⁴CO (subsidence of stratospheric air and local production) takes place during polar night (no photo-chemistry, low OH) attenuated depending on the degree of exchange and mixing with air masses from lower latitudes.

Not all atmospheric ¹⁴CO is of direct cosmogenic origin. When we consider an air mass its CO content is highly variable because of its relatively short lifetime and the occurrence of concentrated sources, for instance industry (fossil fuel) and forest fires. Forest fires (biomass burning in general) releases copious amounts of CO into the troposphere with a ¹⁴C content determined by that of the burned matter. Likewise, the ongoing global removal of CH₄ by mainly OH leads to ¹⁴CO production in the atmosphere itself. This “recycled ¹⁴C” in the form of biogenic ¹⁴CO forms a complication in the use of cosmogenic ¹⁴CO for reconstructing OH. Besides this biogenic source, there are poorly quantified sources of ¹⁴CO from the nuclear power industry and possibly from the use of ¹⁴C containing radio-chemicals/pharmaceuticals. There are no indications so far that these two types of sources play a significant role for the total atmospheric ¹⁴CO inventory, but no detailed research has been done so far. To correct for biogenic ¹⁴CO (∼20%) one has to subtract its contribution from measured ¹⁴CO values. Knowledge about the cycle of CO is necessary and measuring stable isotopic composition of CO namely ¹³C, ¹⁸O, and ¹⁷O (a mass independent effect exists (Röckmann et al. 1998)), is helpful. The contribution of recycled ¹⁴CO is higher in the NH and more use could be made of ¹³CO observations to improve the corrections. Detailed experimental work on CO isotopologues is available (Röckmann 2002) complemented by use of modeling (Gromov 2010, 2013). A rule of thumb is that 30 nmole/mole biogenic CO is equivalent to 1 molecule ¹⁴CO per cm³ (Brenninkmeijer 1993). A 10% uncertainty in the estimate of biogenic ¹⁴CO results in a ∼2.5% uncertainty of the inferred amount of cosmogenic ¹⁴CO.

RESULTS

After Mackay’s pioneering data, Volz et al. (1981) carried out systematic ¹⁴CO measurements. They removed CO₂ from 200 m³ sample air at the sampling site, oxidized its CO content to CO₂ (20–40 mL) which was collected for proportional counting. They obtained the first seasonal cycles of ¹⁴CO and measured at different locations, including shipboard measurements. Using a global 2D model they derived a tropospheric OH concentration below 10⁶ per cm³.

Subsequently ¹⁴CO measurements were set up at the Institute of Nuclear Sciences in Lower Hutt, New Zealand. Roger Sparks and colleagues had turned a decommissioned Pelletron accelerator from Australia into a ¹⁴C AMS system (Sparks et al. 1984). Using AMS the required sample size for ¹⁴CO dropped from 200 to less than 1 m³. AMS was a breakthrough for ¹⁴CO applications, especially because air samples of that smaller size could be collected in pressurized cylinders anywhere for processing in the laboratory. The required air sample size can be estimated as follows. When a conservative estimate of the efficiency of ¹⁴C detection of 1% is used in relation to the initial amount of ¹⁴C in the CO₂ sample and a 1% precision is required, the presence of 10 ¹⁴CO molecules per cm³ air STP implies a minimum air sample size of 100 L. Besides ¹⁴C determination, the CO derived CO₂ samples are analysed by stable isotope mass spectrometry to determine the ¹³C/¹²C and ¹⁸O/¹⁶O ratio. These numbers are useful for reconstructing the local budget of atmospheric CO which helps to estimate the mentioned biogenic ¹⁴CO fraction.

The CO extraction method is described by Brenninkmeijer (1993). Sample air is cryogenically purified by means of “Russian doll” traps submerged in liquid nitrogen. These traps consisting of nested borosilicate glass fibre thimbles remove all condensable gases and achieve the reduction of the CO₂ content of the air to below 0.1 nmole/mole. Depending on the sample CO content, the specific activity of CO is up to 1000 pMC (lowermost stratosphere) at CO mixing ratios down to 50 nmole/mole or less. Residual traces of CO₂ would dilute the ¹⁴CO samples in terms of pMC. After this purification step, the CO content is specifically oxidized to CO₂ using an iodine pentoxide-based reactant. The resulting micromole CO₂ samples are measured manometrically which gives the value of the CO mixing ratio. After off-line stable isotope mass spectrometry, the small sample is diluted with “dead” CO₂ for obtaining a CO₂ sample of sufficient size for target preparation. The advantage of the oxidant used (I₂O₅) is that it does not scramble the original oxygen isotopic composition of the CO because it merely adds an oxygen atom of known isotopic composition to a CO molecule. Insofar air samples are collected in cylinders the CO content can increase in time. This spurious CO would generally affect the ¹⁴C result little because of a low or zero ¹⁴C content. Sampling and storage systems limit spurious CO “growth” to less than the percent level. Irrespective of this, delay between sampling and processing can lead to significant in situ produced additional ¹⁴CO, for which a correction must be applied (Lowe et al. 2002).

The accuracy and precision of a ¹⁴CO determination is composed of the AMS error, the manometric error in dilution of the CO sample (dilution up to a factor 10), and the error in determining the CO content of the air sample and can be kept at about 1%. Adding the larger uncertainty of the biogenic fraction estimation of 2%, gives a combined uncertainty that is below the variability of ¹⁴CO in most air masses (remote SH air showed the least variability) during a typical 1-hr air sample collection period using a high-pressure compressor. Recently, Petrenko et al. (2021) reported an air sample collection and processing method for much smaller air samples, namely less than 100 L. This avoids the use of high-pressure sample air cylinders, which are more expensive and time consuming to transport from the point of collection to the laboratory.

Using this method (with on-site CO extraction, however) Brenninkmeijer et al. (1992) had measured ¹⁴CO data at the clean air station Baring Head, New Zealand and compared the results for the seasonal cycle with those of Volz et al. (1981) for Schauinsland, Germany. Lowest summertime values at Baring Head were about 5 molec cm^–3, compared to 10 at Schauinsland. These are two totally different sites from an atmospheric chemistry point of view, namely clean mid-latitude maritime SH air against continental NH air in the Black Forest. After correcting for the effect of the solar cycle and the difference in latitude and accounting for the biogenic ¹⁴CO fraction, Brenninkmeijer et al. (1993) concluded that OH levels in the SH were higher.

Using a large set of ¹⁴CO data Jöckel and Brenninkmeijer (2002) developed a climatology by which Jöckel et al. (2002) could evaluate global atmospheric transport (including stratosphere-troposphere exchange) and the hydroxyl radical (OH)-based oxidation in three-dimensional (3-D) atmospheric chemistry transport models (CTMs) and general circulation models (GCMs). They showed that the correct representation of phase (annual cycle) and strength of the STE, including its asymmetry between NH and SH, is essential to reproduce the observation based seasonal cycle of ¹⁴CO. Moreover, their results did not support the previously inferred interhemispheric asymmetry in the OH abundance with an on average higher concentration in the SH.

A question that had to be addressed was the response time of the ¹⁴CO concentration to changes in total ¹⁴C production caused by modulation through solar activity (Jöckel et al. 2000) which is related to, but not the same as the lifetime. Figure 5 shows model lifetimes of ¹⁴CO throughout the atmosphere. At the surface lifetimes are below 1 month for parts of the NH tropics.

Figure 5 Lifetime of ¹⁴CO in days (color scale) simulated with the EMAC chemistry climate model (year 2000, Jöckel et al. 2016). The troposphere shows a considerable asymmetry with higher OH abundance in the NH. (Please see electronic version for color figures.)

Quay et al. (2000) applied a 2D transport model to their measured and several other ¹⁴CO datasets. They found good agreement and emphasize that in their model an increase in advection and slower stratosphere-troposphere exchange in the SH would best explain the lower ¹⁴CO levels of Baring Head.

The properties of ¹⁴CO in relation to OH and transport have been characterized in detail using the TM5 global chemistry transport model and inverse modeling (Krol et al. 2008). It was established that in the tropics ¹⁴CO is sensitive to regional OH. Therefore, it is a useful indicator for the local “speed” of the rate of removal of reduced gases. At the hemispheric scale however, transport of ¹⁴CO plays a bigger role. Variability is introduced by synoptic scale processes. The model research also compares ¹⁴CO with CH₃CCl₃, which appears to be more sensitive to bulk tropical OH. The model results closely follow the annual cycle of ¹⁴CO at baring Head, NZ, with an offset probably due to the assumed production rate. Recommendation about optimizing locations for measurements can be gleaned from this work.

The success of the ¹⁴CO method for inferring changes in OH is best illustrated by Figure 6 from Manning et al. (2005). It shows the longest available record, namely that of Baring Head including data from Scott Base, Antarctica. Cylinders with sample air from Antarctica often were stored on site for months before shipment was possible. Lowe et al. (2002) corrected the ¹⁴CO content for in-situ production. One clearly sees the solar cycle induced change of about 20% for the period 1989 until 2003. The seasonal cycle due to OH is about ± 40%. When taking the ¹⁴CO production variations into account and applying an atmospheric model the residuals relative to a fitted curve are within 10%, except for two periods. The first one is in 1991, after the eruption of Mount Pinatubo. Reduced UV radiation in the troposphere did lead to a reduced production of OH (Dlugokencky et al. 1996) and thus an increase in ¹⁴CO. The second period is 1997/1998. During this period a strong El Niño led to intensive biomass burning (Southeast Asia, particularly Indonesia) thus burdening the troposphere with reduced gases apparently leading to a reduction in OH in this part of the SH. Also, radiative effects of a large atmospheric load with soot can have influenced OH levels. There are no other observations of this kind that document such changes in OH.

Figure 6 The calculated ¹⁴C production rate (uppermost curve) and the measured ¹⁴CO concentrations at Baring Head (blue circles) and Scott Base (red squares). The black curve is the best-fit simulated ¹⁴CO. The lower section shows the residuals with about 10% variability and two major events (cf. Manning et al. 2005).

Several experiments have addressed free tropospheric and lowermost stratospheric ¹⁴CO measurements using air samples collected using aircraft. Initially modified scuba diving compressors were used (Mak and Brenninkmeijer 1994) and later membrane compressors (Brenninkmeijer et al. 2007; Brenninkmeijer and Roberts 1994) with the low ambient pressure being a challenge considering the large amount of air required.

Detection of 1989 Solar Proton Events

Figure 6 shows the calculated ¹⁴C production rate during solar cycle 22. The effect of the modulation of the cosmic ray flux by solar activity is considerable. Apart from this 11-yr gradual modulation, short-term fluctuations in ¹⁴C production do occur. For calculating the actual variation in the production of ¹⁴C, neutron count rates were used (Lowe et al. 2002). Both these longer-term and short-term variations have to be taken into account when reconstructing the actual ¹⁴C production rate for application to OH.

Besides this, independent, sporadic bursts (lasting days only) of atmospheric ¹⁴C production are due to high energy solar protons from coronal mass ejections. Such events can occur throughout a solar cycle and lead to elevated neutron count rates at ground level (Ground Level Enhancements, GLE). Lingenfelter and Flamm (1964) estimated the average ¹⁴C production due to solar protons to be at most 5% of the cosmogenic production. For events to be of significance for ¹⁴C production the proton energy has to exceed ∼30MeV and the flux has to be sufficiently high. During solar cycle 22 several high energy proton events occurred. Three of these events between August and December 1989 carried sufficient high energy protons to produce—via additional neutrons—¹⁴C that was detectable above the background. The GLE on 29 September 1989 was the largest recorded since the giant event of 23 February 1956 (Hessel de Vries submitted a paper about its effects on background count rates in low level ¹⁴C proportional counters a mere 2 weeks later; de Vries 1956). Figure 2 (contour lines) shows that the solar protons in 1989, despite their high energy, produced ¹⁴C higher in the polar stratosphere than the typical maximum of the production by galactic cosmic rays. By analysing the Baring Head ¹⁴CO record, using modeling and satellite data for the proton flux, Jöckel et al. (2003) established that between 3 and 30% additional ¹⁴C production due to solar protons had occurred in 1989. Such an application of ¹⁴CO is a unique case. To extract the solar flare signal (Figure 7) the period June 1990–June 1991 was used a reference (no significant solar flares). The model calculated increase overlaps with the measured ¹⁴C enhancement.

Figure 7 Observed (black) and solar cycle corrected (red, scc) enhancement of ¹⁴CO for June 1989 to June 1990 relative to June 1990 to June 1991 (reference year) at Baring Head. The length of the three arrows corresponds to the estimated total solar flare ¹⁴C production. The blue lines show the model prediction for normal cut-off rigidity (lower line) and an 80% reduced cut-off rigidity (upper line). The purple line shows the observations smoothed with a time window of T=4 weeks and corrected for the solar cycle (scc). (Figure from Jöckel et al. 2003; Creative Commons Attribution-NonCommercial-ShareAlike 2.5 License.)

Calculated Production of ¹⁴CO in the Atmosphere

The application of ¹⁴CO requires accurate information about the ¹⁴C production distribution (on average about 7 kg ¹⁴C is produced annually). Its subsequent fate by oxidation to ¹⁴CO₂ is then determined by transport and OH chemistry, with a difficult aspect being its transport from the stratosphere (¹⁴C production but low OH) into the troposphere (also production but high OH). The calculation of ¹⁴C production by Lingenfelter (1963) (Figure 2) has been used most. Obviously, the position of the tropopause (the dynamic interface between the two spheres) is of no relevance to the ¹⁴C production field. All neutrons generated by nuclear reactions in the earth atmosphere lead to the production of ¹⁴C via ¹⁴N(n,p)¹⁴C. Lingenfelter used neutron count rates to derive 3D neutron production rates for solar maximum and minimum conditions. O’Brien et al (O’Brien et al. 1979) carried out calculations based on hadronic cascades in the atmosphere. Masarik and Beer (1999) developed a pure physical model. All three approaches give a few % more ¹⁴C production in the SH due to the Earth’s magnetic field distribution. In the given order of principal author 52 (Lingenfelter), 57 (O’Brien) and 64% (Masarik), respectively, is produced in the stratosphere (Jöckel 2000). Whereas SH-NH differences are small, the vertical distribution needs attention in view of the role of stratosphere-troposphere exchange. Jöckel et al. (1999) showed by means of a model that tropospheric ¹⁴CO concentrations are relatively insensitive to these source differences and therefore not critical for the ¹⁴CO method. More extensive and refined work on cosmogenic isotope production was carried out by Usoskin and coworkers (Poluianov et al. 2016). Not only did this work at long last close the gap between calculated global ¹⁴C production and observations in the carbon cycle but also revealed a larger solar modulation of ¹⁴C production than in previous works.

Figure 8 shows vertical profiles of production rates from the four approaches. There are significant differences of a factor 5 at the surface and in the free troposphere. The steep production rate gradient in the troposphere is a lesser concern because of the relatively rapid vertical exchange of air masses and thus mixing of ¹⁴CO in the vertical. For local ¹⁴CO production the gradient matters.

Figure 8 Vertical profile of the zonally averaged ¹⁴C production rate at 80°N for January 2000 (heliospheric potential of 752MV). The colour coding is for calculated results from Poluianov et al. (2016, black), Lingenfelter (1963, red), Masarik and Beer (1999, green), and O’Brien (1979, blue). The dashed horizontal lines delineate the tropopause. Please note that we use here a linear vertical scale and that the production rate is per volume unit, not per mass unit.

In-Situ ¹⁴CO Production Experiments

From an experimentalist’s point of view, it is comforting to know that (in principle) it is straightforward to measure the actual production of ¹⁴C (as ¹⁴CO) by exposing a quantity of air to cosmic radiation. Mak et al. (1999) reported on such experiments based on exposing cylinders with compressed air at different altitudes and latitudes (South Pole Station and Mount Cook, NZ) and at different heights on a tall tower (Boulder Atmospheric Observatory, USA). The fundamental problem is that production rates at the surface are low (production peaks at 100 hPa), which renders the result as a measure for the column integrated production intrinsically less informative. An additional complication is the enhancement of the neutron flux close to the surface. The purpose of the tall tower experiment was to deal with this. The results showed that measured production rates were consistently lower by ∼50% than values derived from Lingenfelter’s calculations.

Further experiments were made using cylinders with compressed air stowed in passenger aircraft (Ramsey et al. 2007). Such experiments are ideal from the point of view that civil aircraft happen to cruise at altitudes where ¹⁴C production peaks. By simulating the aircraft flights over weeks between cylinder exchanges in a model of the ¹⁴C production field expected accumulated production could be calculated. Also, calculations were carried out to correct for local neutron enhancement due to aircraft structure and payload. But there also was a big obstacle. Namely, only cylinders that were certified to fly on passenger aircraft could be used. These aviation oxygen supply (mild steel) cylinders were not the treated inert aluminum cylinders normally used in ¹⁴CO air sampling and storage. During the experiment, the aviation cylinders had to undergo hydraulic pressure tests. Altogether 90 experiments could be carried out. Sets of cylinders were kept in the laboratory for background testing. Measured was not only ¹⁴CO but also ¹⁴CO₂. The results showed not a ¹⁴CO₂ to ¹⁴CO ratio of about 5%, as expected on the basis of the early radiochemistry experiment, but about 20%. No explanation could be given. Furthermore, production rates were well above those calculated using the Masarik and Beer distribution. Despite a considerable effort the results were less conclusive.

¹⁴CO as the Precursor of ¹⁴CO₂

Because soils and oceans take up negligible amounts of ¹⁴CO and the chemical fate of CO is its oxidation, nearly all primary cosmogenic ¹⁴CO molecules will form ¹⁴CO₂. In equilibrium both ¹⁴CO and ¹⁴CO₂ have exactly the same flux (some deviations occur because of ¹⁴C from the nuclear weapons tests and variations in the rate of uptake of ¹⁴CO₂ by the biosphere and oceans). As a result of their vastly different residence times however, the atmospheric inventory of ¹⁴C as ¹⁴CO is more than 1000 times smaller than ¹⁴C as ¹⁴CO₂. Also, their distributions in the atmosphere bear no resemblance. Kanu et al. (2016) calculated an independent cosmogenic production rate of ¹⁴C using data from air samples collected in the stratosphere up to above 30 km. Their ¹⁴CO₂ data show a compact correlation with decreasing N₂O mixing ratios, a gas that is destroyed in the stratosphere. ∆¹⁴CO₂ reached about 140 per mil. Brenninkmeijer et al. (1996) report ¹⁴CO values in the lowermost stratosphere over Antarctica of up to 150 molec cm³, 10 times tropospheric values. To our knowledge, no model analysis for ¹⁴CO₂ production from ¹⁴CO in the stratosphere have been published so far. Our preliminary model calculations show that ∼30% of ¹⁴CO produced in the stratosphere is exported to the troposphere. About 10% of the tropospheric production (which is less than 50% of the total) ends up in the stratosphere. Therefore, the net ¹⁴CO loss for the stratosphere is only 20%. Consequently, 80% of ¹⁴CO in the stratosphere is oxidized there to ¹⁴CO₂.

Firn Air Measurements of ¹⁴CO

The analysis of air extracted in-situ from firn allows reconstruction of past atmospheric composition of lesser reactive trace gases. In the case of ¹⁴CO hopes were high to be able to shed some light on changes in OH over past decades. Indeed, CO itself appeared to be stable in firn air. Although the in-situ production of ¹⁴C of the air retained in the firn layers is small, in situ production in the solid phase appeared to complicate matters. As can be seen in the paper by Assonov et al. (2005) ¹⁴CO levels in firn air (Dronning Maud Land, Dome Concordia and Berkner Island) were above atmospheric levels and increased considerably near pore close-off depth. This points to a very complicated situation when one wishes to reconstruct past OH concentrations. In-situ production of ¹⁴C (Jull et al. 1994) complicates the use of ¹⁴CO in firn air. A sufficiently accurate correction for in situ production in the ¹⁴CO firn air application, will be difficult, if possible, at all. Nevertheless, work on this is in progress (Petrenko et al. 2020).

Suggested Improvements for the ¹⁴CO Application

Given the lack of alternatives for OH and the improvements in modeling including the recent ¹⁴C production calculation by Usoskin’s group, ¹⁴CO work should continue with improved methods. The optimal number and location of sampling sites can be better determined now (Krol et al. 2008). From an air sampling method point of view, a degree of time integrated collection optimizes the results. Short term variations are average by this. Up to now it has been a major effort to collect, compress and ship a dry air sample, but why this fast sampling? There is a lot of air for sampling, in contrast to many other ¹⁴C AMS applications where smallest amounts need to be assayed. Volz et al. (1981) and Brenninkmeijer et al. (1992) used on-site extraction, which was cumbersome. Thereafter, Mak and Brenninkmeijer (1994) used the experience obtained in Boulder (NOAA) for using high-pressure air (diving air) compressors. This facilitated getting samples from remote locations enormously and many ¹⁴CO data were collected (e.g., Rom et al. 2000; Gros et al. 2001). An excellent example of a sampling technique is the pioneering work by Östlund (1970), who measured the Tritium content of atmospheric H₂ already in 1970 using automated collection systems (HT is rarer in the atmosphere than ¹⁴CO) at remote locations, including Baring Head.

For the ¹⁴CO application one should ironically not think about small air samples as we did so far. We propose here that sample air, say 5000 L (flowrate of 1 L/min) is dried using Nafion® and subsequently cryogenically purified using Russian Doll traps. The next step is oxidation of the CO content at ambient temperature using a CO selective catalyst (e.g., Sofnocat®). The resulting CO₂ is collected on a molecular sieve cartridge for shipment to the AMS lab. Alternatively, molecular sieve cartridges are used for the thorough pre-purification after sample air drying. A large volume cartridge with molecular sieve would lead to a CO gas chromatography separation effect. Smaller cartridges in parallel, that are sequentially heated and backflushed for re-generation, avoid such a problem. After this set of traps one final molecular sieve cartridge acts as a safety barrier before the air flow reaches the catalyst. Another advantage of using molecular sieve (instead of the Russian doll traps) is that cooling systems are not required. Moreover, the air processing occurs at above atmospheric pressure. Cryogenic purification requires sub-ambient pressures.

The advantages compared to the method developed by Brenninkmeijer et al. (1999) are (1) large, “robust” CO-derived CO₂ samples, (2) time-integrated air sampling, (3) samples can be shipped without restrictions, (4) storage will most likely not lead to in situ ¹⁴C formation, and (5) the difficult I₂O₅ oxidant need not to be used (this “Schütze reagent” is very sensitive to moisture and difficult to prepare). The disadvantage is that no δ¹⁸O(CO) information is obtained in this way. CO mixing ratios are already obtained at may sampling sites in the framework of Global Atmosphere Watch. What remains difficult is the collection of air samples using civil aircraft. The CARIBIC Boeing 737 used for ¹⁴CO work (Brenninkmeijer et al. 2007) had a typical cruise speed of well over 800 km per hour, which meant that with the given air sample collection time of about ½ hour, sampling in the tropopause region where civil aircraft most frequently fly at mid-latitudes an averaging between ¹⁴CO in tropospheric and stratospheric air masses was inevitable.

CONCLUSIONS

The distribution and possible changes of OH in the troposphere have been a bone of contention for years (one can find several lively discussions in the literature, e.g., Spivakovky 1991). It was (and still is) not known if OH does change over large parts of the troposphere or even globally on a scale of several years and if OH remained stable at all over the current decades of the Anthropocene. The clear detection of reductions in OH due to the 1991 Mount Pinatubo eruption and the 1998 El Nino by Manning et al. (2005) is an important demonstration of the use of ¹⁴CO. Over many years emphasis has been on using methyl chloroform records, but this substance disappears from the atmosphere after having been banned by the Montreal Protocol. Its mixing ratio dropped from over 100 to a few pmole/mole (lifetime about 5 years). Clarissa Spivakovsky wrote about CH₃CCl₃ based estimates of tropospheric OH “Last Offerings of a Departing Friend” (Spivakovsky et al. 2000). A main concern is uncertainty about its source strength, the role of the oceans with a dwindling atmospheric abundance. Nevertheless, a recent paper seems to have given methyl chloroform back a brief lease of life (Patra et al. 2021). Some other industrial trace gases also help to “constrain” OH, but there remain inherent uncertainties. It is difficult.

The initial interest in ¹⁴CO from an atmospheric application point of view was definitely the determination of OH. In the application of atmospheric models to support this use, ¹⁴CO also proved to be a tough test for the simulated transport processes (Jöckel et al. 2002). A difficult aspect in this is the cross-tropopause transport. At civil aircraft cruise-altitude we frequently measured changes from 10 to almost 100 ¹⁴CO molecules per cm³ STP when crossing tropopause folds (CARIBIC unpublished results). ¹⁴CO correlates closely with O₃ in the lowermost stratosphere. This strong tropopause gradient means that a cut-off low releases ¹⁴CO enriched air into the troposphere. Such information—detailed as it is—is difficult to model. Today’s models can have a higher vertical degree of resolution in the tropopause region, which is helpful.

Given the improved ¹⁴C production calculations and continued ¹⁴CO applications it is worthwhile to carry out improved in-situ production experiments by exposing air samples. It seems to be a trivial undertaking, but so far, no really useful conclusions have been reached.

¹⁴CO has the great advantage that it tracks closely the seasonal cycle of OH. Corresponding variations in CH₃CCl₃ are below 3% only. ¹⁴CO can help to derive OH levels on a regional scale. ¹⁴CO has a source that is probably not affected noticeably by industrial emissions, it is a geophysical tracer. From this point of view is stands to reason to continue ¹⁴CO observations, and indeed recently further improvements were just reported (Petrenko et al. 2021). Current atmospheric chemistry models are of great complexity and the reconstructed rate of removal of trace gases by OH remains an essential piece of output. The discussion about OH is ongoing. In this, measurement of ¹⁴CO provides a definite test option, in absence of other suitable trace gases. Support for long-term times series, namely Baring Head (the ¹⁴CO “Mauna Loa”), will be most helpful.

In summary, given a need to gauge tropospheric OH in a changing atmosphere and the lack of suitable trace gases, and given the improvements in cosmogenic ¹⁴C production knowledge, and given improvements in atmospheric modeling, and knowing an optimized sampling strategy and using an improved ¹⁴CO measurement methodology as suggested, continued progress is possible.