An estimation of the 18 O / 16 O ratio of UT / LMS ozone based on artefact CO in air sampled during CARIBIC flights

1 2 3 4 5 6 7 8 9 10 11 12 13 14

1 Introduction [1] Successful determination of the atmospheric carbon monoxide (CO) content based on the collection of air samples depends on the preservation of the mixing ratio of CO inside the receptacle, from the point of sampling to the moment of physiochemical analysis in a laboratory.A well known example in our field of research is the filling of pairs of glass flasks at South Pole 1 Station for analysis at NOAA in Boulder, Colorado, USA (Novelli et al., 1998).There, the duplicate air sampling allowed for a degree of quality control which in view of the long transit times, especially during polar winter, was a perhaps not perfect, but certainly a practical measure.Here we deal with a different case: Using aircraft-based collection of very large air samples rendered duplicate sampling unpractical, yet analyses could be performed soon after the sampling had taken place because of the proximity of the aircraft's landing location to the laboratory involved.A presumption of the analytical integrity of the process was that the growth of CO in receptacles is gradual and takes its time.Reminding Thomas Henry Huxley's statement, "The great tragedy of Science -the slaying of a beautiful hypothesis by an ugly fact", it turned out, however, that for air we collected in stainless steel tanks in the upper troposphere/lowermost stratosphere (UT/LMS) higher CO values were measured in the laboratory than measured in situ during the collection of these air samples.Moreover, measurement of the stable oxygen isotopic composition of CO from these tanks revealed additional isotopic enrichments in 18 O of 10‰ or more.It was soon realised that this phenomenon was due to the formation of CO in these tanks and/or possibly in the sampling system and inlet tubing used, by reactions involving ozone ( Brenninkmeijer et al., 1999).
[2] Unexpectedly high 18 O/ 16 O ratios in stratospheric ozone (O 3 ) were discovered by Konrad Mauersberger using a balloon-borne mass spectrometer (Mauersberger, 1981), which has triggered a series of theoretical and experimental studies on atmospheric O 3 heavy isotope enrichments (see, e.g., Schinke et al. (2006) for a review).In view of the advances in theoretical and laboratory studies on the isotopic composition of O 3 atmospheric measurements are welcome, they do however form a challenge.In the stratosphere O 3 is abundant, but the remoteness of the sampling domain is a problem.In the troposphere, low O 3 concentrations are the main obstacle, as indicated by few experiments performed to date (Krankowsky et al., 1995;Johnston and Thiemens, 1997;Vicars and Savarino, 2014).Nevertheless, recent analytical improvements, namely the use of an indirect method of reacting atmospheric O 3 with a substrate that can be analysed for the isotopic composition of the O 3 -derived oxygen (Vicars et al., 2012), has greatly improved our ability to obtain information on the O 3 isotopic composition.
[3] Although the increase of CO concentrations in air stored in vessels is a well recognised problem, to our knowledge a specific O 3 -related process has not been reported yet.Here we discuss this phenomenon and turn its disadvantage into an advantage, namely that of obtaining a valid estimate of the oxygen isotopic composition of O 3 in the UT/LMS, an atmospheric do- main not yet covered by specific measurements.The air samples we examine in this study were collected onboard a passenger aircraft carrying an airfreight container with analytical and air/aerosol sampling equipment on long distance flights from Germany to South India and the Caribbean within the framework of the CARIBIC project (Civil Aircraft for the Regular Investigation of the atmosphere Based on an Instrument Container, http://www.caribicatmospheric.com).

Whole air sampling
[4] CARIBIC−1 (Phase #1, abbreviated hereafter "C1") was operational from November 1998 until April 2002 using a Boeing 767-300 ER operated by LTU International Airlines (Brenninkmeijer et al., 1999).Using a whole air sample (WAS) collection system, twelve air samples were collected per flight (of ~10 hours duration at cruise altitudes of 10−12 km) in stainless steel tanks for subsequent laboratory analysis of the abundances of various trace gases, including 14 CO.Large air samples were required in view of the ultra-low abundance of this mainly cosmogenic tracer (10−100 molecules cm −3 STP, about 40−400 amol/mol).Each C1 WAS sample (holding ~350 litres of air STP) was collected within 15−20 min intervals representing the integral of the compositions encountered along flight segments of ~250 km.The overall uncertainty of the measured WAS [CO] is less than ±1% for the mixing ratio and ±0.1‰/±0.2‰for δ 13 C(CO)/δ 18 O(CO), respectively (Brenninkmeijer, 1993;Brenninkmeijer et al., 2001).Isotope compositions are reported throughout this manuscript using δ i = ( i R/ i R st −1) relating the ratio of rare over abundant isotopes i R of interest (i denotes 13 C,
[5] CARIBIC−2 (Phase #2, referred to as "C2") started operation in December 2004 with a Lufthansa Airbus A340-600 fitted with a new inlet system and air sampling lines, including PFA lined tubing for trace gas intake (Brenninkmeijer et al., 2007).No flask CO mixing/isotope ratio measurements are performed in C2. i.e. concentration or mixing ratio, of the respective species).In situ CO analysis in C1 is done using a gas chromatography (GC)-reducing gas analyser which provides measurements each 130 s with uncertainty of ±3 nmol/mol (Zahn et al., 2000).In C2, a vacuum ultraviolet fluorescence (VUV) instrument with lower measurement uncertainty and higher temporal resolution of ±2 nmol/mol in ~2 s (Scharffe et al., 2012) is employed, respectively.

On-line instrumentation
Furthermore, the detection frequency for O 3 mixing ratios has also increased, viz.from 0.06 Hz in C1 to 5 Hz in C2 (Zahn et al., 2002;Zahn et al., 2012).

Results
[ however, is expected from our knowledge of UT/LMS CO sources (plus their isotope signatures) and available in situ observations (ibid., shown with triangles), as elucidated by Brenninkmeijer et al. (1996) (hereafter denoted as "B96").That is, the more stratospheric CO is, the greater fraction of its local inventory is refilled with the photochemical component stemming from methane oxidation with a characteristic δ 18 O signature of ~0‰ or lower (Brenninkmeijer and Röckmann, 1997).This occurs because the CO sink at ruling UT/LMS temperatures proceeds more readily than its production, as the reaction of hydroxyl radical OH (fractionation about ~11‰ at pressures below 300 hPa, Stevens et al., 1980;Röckmann et al., 1998b).
[9] It is beyond doubt that the enhancements of C1 C 18 O originate from O 3 , whose large enrichment in heavy oxygen (above +60‰ in δ 18 O, Brenninkmeijer et al., 2003) is typical and found transferred to other atmospheric compounds (see Savarino and Morin (2012) for a review).In Fig. 2 it is also notable that not only the LMS compositions are affected but elevations of (3−10)‰ from the bulk δ 18 O(CO) values are present in more tropospheric samples with [CO] of up to ~100 nmol/mol.These result from the dilution of the least affected tropospheric air with high mixing ratios by CO-poor, however substantially contaminated, stratospheric air, sampled into the same WAS tank.Such sampling-induced mixing renders an unambiguous determination of the artefact source' isotope signature rather difficult, because neither mixing nor isotope ratios of the admixed air portions are known sufficiently well (see below).
[10] Differences between the WAS and in situ measured [CO] -a possible indication that the δ 18 O(CO) contamination pertains specifically to the WAS data -average at Δ ¯WAS− = 5.3±0.2(1 SD of the mean, n = 408) and happen to be random with respect to any operational parameter or measured characteristic in , irrespective of CO or O 3 abundances.
The quoted mixing ratio discrepancy remained after several calibrations between the two systems had been performed, and likely results from the differences in the detection methods, drifts of the calibration standards used (see details in ) and a short-term production of CO in the stainless steel tanks during sampling.[13] (ii) Mixing effects can also occur artificially, originating from sampling peculiarities or data processing.Since the CARIBIC platform is not stationary, about 5 s long sampling of an in situ air probe in C1 implies integration of the compositions encountered along some hundred metres, owing to the high aircraft speed.This distance may cover a transect between tropospheric and stratospheric filaments of much different compositions.The effect of such 'translational mixing' can be simulated by averaging the sampling data with higher temporal frequency over longer time intervals.In this respect, the substantially more frequent CO data in C2 (<1 s) were artificially averaged over a set of increasing intervals to reckon whether the long sampling period in C1 could be the culprit for skewing its CO−O 3 distribution.As a result, the original C2 [15] Importantly, since we can quantify the contamination strength using only the O 3 abundance, the continuous in situ C1 [O 3 ] data allows to estimate the integral artefact CO component in each WAS sample and, if the isotope ratio of contaminating O 3 is known, to derive the initial δ 18 O(CO).The latter, as it was mentioned above, is subject to strong sample-mixing effects, which is witnessed by δ 18 O(CO) outliers even at relatively high [CO] up to 100 nmol/mol.Accounting for such cases is, however, problematic since it is necessary to distinguish the proportions of the least modified (tropospheric) and significantly affected (stratospheric) components in the resultant WAS sample mix.In reality, however, this information is not available, therefore we applied an ad hoc correction approach (which is capable of determining the contamination source (i.e., O 3 ) isotope signature as well), as described in the following.

Contamination isotope signatures
[16] Practically we resort to the differential mixing model (MM, originally known as the "Keeling-plot"), because it requires only the estimate of the artefact component mixing ratio, but no assumptions on the (unknown) shares and isotope signatures of the air portions mixed in a given WAS tank.The MM parameterises the admixing of the portion of artefact CO to the WAS sample with the "true" initial composition, as formulated below: ( ) where indices a, c and t distinguish the abundances C and isotope compositions i δ (i may refer to 13 C or 18 O) pertaining to the analysed sample, estimated contamination and "true" composition sought ( i.e., C t and i δ t ), respectively.(Here the contamination strength C c is derived by integrating Eq. (1) using the in situ C1 [O 3 ] data for each WAS sample.)By rewriting the above equation w.r.t. the isotope signature of the admixed portion i δ c , one obtains: which signifies that linear regression of the measured i δ a as a function of the reciprocal of C c yields the estimated contamination signature i δ c at (C c ) −1 → 0. (The Keeling plot detailing the calculations with the MM is shown in Supplementary Material, Fig. S3.)The MM described by Eq. ( 2) provides adequate results only for the invariable initial compositions (C t , i δ t ), therefore we apply it to the subsets of samples picked according to the same reckoned C t (within a ±2 nmol/mol window, n > 7).Such selection, however, may be insufficient: Due to the strong sampling effects in the WAS samples (see previous Section), it is possible to encounter samples that integrate different air masses to the same C t but rather different average i δ t .The solution in this case is to refer to the goodness of the MM regression fit, because the R 2 intrinsically measures the linearity of the regressed data, i.e. closeness of the "true" values in a regarded subset of samples, irrespective of underlying reasons for that.
[17] Higher R 2 values thus imply higher consistency of the estimate, as demonstrated in Fig. 3 showing the calculated i δ c for C t below 80 nmol/mol as a function of the regression R 2 .The latter decreases with greater C t (i.e., larger sample subset size, since tropospheric air is more often encountered) and, conformably, larger variations in i δ t .Ultimately, at lower R 2 the inferred 18 O signatures converge to values slightly above zero expected for uncorrelated data, i.e.C1 δ 18 O(CO) tropospheric average.A similar relationship is seen for the 13 C signatures (they converge around −28‰), however, there are no consistent estimates found (R 2 is generally below 0.4).Since such is not the case for δ 18 O, the MM is not sufficiently sensitive to the changes caused by the contamination, which implies that the artefact CO δ 13 C should be within the range of the "true" δ 13 C(CO) values.Interestingly, the MM is rather responsive to the growing It is noteworthy that we have accounted for the biases in the analysed C1 WAS δ 13 C(CO) expected from the mass-independent isotope composition of O 3 (see details in Appendix B).
[18] We derive the "best-guess" estimate of the admixed CO 18 O signature at 18 O δ c = +(92.0±8.3)‰, which agrees with the other MM results obtained at R 2 above 0.75.Taking the same subsets of samples, the concomitant 13 C signature matches ing "ozone hole" conditions and carried extremely low 13 CO abundances, which was attributed to the reaction of methane with available free Cl radicals (Brenninkmeijer et al., 1996).

Estimate of δ 18 O(O 3 )
[19] The 18 O δ c signature inferred here ( was formed.Indeed, the molecular lifetime (the period through which the species' isotope reservoir becomes entirely renewed, as opposed to the "bulk" lifetime) of O 3 encountered along the C1 flight routes is estimated on the order of minutes to hours at daylight (H.Riede, MPI-C, 2010), thus the isotope composition of the photochemically regenerated O 3 resets quickly according to the local conditions.Virtual absence of sinks, in turn, leads to "freezing" of the δ 18 O(O 3 ) value during night in the UT/LMS.Verifying the current δ 18 O(O 3 ) estimate against the kinetic data, in contrast to the stratospheric cases, is problematic.The laboratory studies on O 3 formation to date have scrutinised the concomitant kinetic isotope effects (KIEs) as a function of temperature at only low pressures (50 Torr); the attenuation of the KIEs with increasing pressure was studied only at room temperatures (see Table 1, also Brenninkmeijer et al. (2003) for references).A rather crude attempt may be undertaken by conjecturing an inhibition of the formation KIEs proportional to that measured at ~320K, however applied to the nominal lowpressure values reckoned at (220−230)K.A decrease in δ 18 O(O 3 ) of about (5.9−7.6)‰ is expected from such calculation, yet accounting for a mere one-half of the (13.3−14.6)‰"missing" in 18 O δ c .
[20] Lower 18 O δ c values could result from possible isotope fractionation accompanying the production of the artefact CO.Although not quantifiable here, oxygen KIEs in the O 3 → CO conversion chain cannot be ruled out, recalling that the intermediate reaction steps are not identifiable and the artefact CO represents at most 4% of all O 3 molecules.Furthermore, the yield λ O3 of CO from O 3 may be lower than unity (see details in Appendix A).On the other hand, the inference that the contamination strength primarily depends on [O 3 ] indicates that the kinetic frac-tionation may have greater effect on the carbon isotope ratios of the artefact CO produced (the 13 C δ c values) in contrast to the oxygen ones.That is because all reactive oxygen available from O 3 becomes converted to CO, whilst the concomitant carbon atoms are drawn from a virtually unlimited pool whose apparent isotope composition is altered by the magnitude of the 13 C KIEs.
[21] Besides KIEs, selectivity in the transfer of O atoms from O 3 to CO affects the resulting Considering the alternatives of the O transfer in our case (listed additionally in Table 1), the equiprobable incorporation of the terminal and central O 3 atoms into CO should result in the δ 18 O(O 3 ) value in agreement with the "crude" estimate based on laboratory data given above.
[22] Furthermore, the conditions that supported the reaction of O 3 (or its derivatives) followed by the production of CO are vague.A few hypotheses ought to be scrutinised here.from Appendix A) must be on the order of 6•10 −15 /τ c [molec −1 cm 3 s −1 ], where τ c is the exposure time.Assuming the case of a gas-phase CO production from a recombining O 3 derivative and an unknown carbonaceous compound X, the reaction rate coefficient for the latter ( X k r in Eq. (A1) in Appendix A) must be rather high, at least ~6•10 −10 [molec −1 cm 3 s −1 ] over τ c = 1/100 s.This number decreases proportionally with growing τ c and [X], if we take less strict exposure conditions.Nonetheless, in order to provide the amounts of artefact CO we detect, a minimum abundance of 20 nmol/mol (or up to 4 µg of C per flight) of X is required, which is not available in the UT/LMS from the species readily undergoing ozonolysis, e.g.alkenes.
[23] Second, a more complex heterogeneous chemistry on the inner surface of the inlet or supplying tubing may be involved.Such can be the tracers' surface adsorption, (catalytic) decomposition of O 3 and its reaction with organics or with surface carbon that also may lead to the Oyama , 2000).Evidence exists for the dissociative adsorption of O 3 on the surfaces with subsequent production of the reactive atomic oxygen species (see, e.g., Li et al., 1998, also Oyama, 2000).It is probable that sufficient amounts of organics have remained on the walls of the sampling line exposed to highly polluted tropospheric air, to be later broken down by the products of the heterogeneous decomposition of the ample stratospheric O 3 .
Unfortunately, the scope for a detailed quantification of intricate surface effects in the C1 CO contamination problem is very limited.

Conclusions
[24] Recapitulating, the in situ measurements of CO and O 3 allowed us to unambiguously quantify the artefact CO production from O 3 likely in the sample line of the CARIBIC−1 instrumentation.Strong evidence to that is provided by the isotope CO measurements.We demonstrate the ability of the simple mixing model ("Keeling-plot" approach) to single out the contamination isotope signatures even in the case of a large sampling-induced mixing of the air with very different compositions.Obtained as a collateral result, the estimate of the δ 18 O(O 3 ) in the UT/ LMS appears adequate, calling, however, for additional laboratory data (e.g., the temperaturedriven variations of the O 3 formation KIE at pressures above 100 hPa) for a more unambiguous verification.

Appendix A. Contamination kinetic framework
[25] We infer the O 3 -exclusive functional dependence of the contamination strength C c by discriminating the C1 outliers from respective C2 data in the following kinetic framework: where k c denotes the overall pseudo-first-order rate coefficient of the reaction chain leading to the artefact CO production with the respective yield λ O3 .The individual rate coefficients X k r and (Assonov and Brenninkmeijer, 2001).In effect for the C1 CO data, the artefact CO produced from O 3 had contributed with unexpectedly high C 17 O abundances that led to the overestimated δ 13 C(CO) analysed.Knowing the contamination magnitude C c and assuming the typical O 3 MIF composition being 17 O Δ c , the respective bias 13 C δ b is calculated using     δ t regression (above 0.9).See text for details.

[ 6 ]
In addition to the WAS collection systems, both C1 and C2 measurement setups include different instrumentation for on-line detection of [CO] and [O 3 ] (hereinafter the squared brackets [] denote the abundance,

(
OH) with CO, being primarily pressure-dependent, outcompetes the temperature-sensitive reaction of OH with CH 4 .Furthermore, as the lifetime of CO quickly decreases with altitude, transport-mixing effects take the lead in determining the vertical distributions of [CO] and δ 18 O(CO) above the tropopause, hence their mutual relationship.This is seen from the B96 data at [CO] below 50 nmol/mol that line-up in a near linear relationship towards the end-members with lowest 18 O/ 16 O ratios.These result from the largest share of the 18 O-depleted photochemical component and extra depletion caused by the preferential removal of C 18 O in reaction with Fig.1), whereas the C1 outliers were exclusively registered in some 12 flights during data and their averages (equivalent to the C1 CO sample injection time) differ negligibly, as do the respective ρ O3:CO values; the actual C2 CO−O 3 statistic in the region of interest ([O 3 ] of 540−620 nmol/mol) remains insensitive to integration of up to 300 s.Furthermore, a very strong artificial mixing with an averaging interval of at least 1200 s (comparable to C1 WAS sampling time) is required to yield the averages from the C2 data with ρ O3:CO characteristic for the C1 outliers.[14](iii) In view of the above, it is unlikely that any natural or artificial mixing processes are involved in the stratospheric [CO] discrepancies seen in C1.It therefore stands to reason to conclude that the sample contamination in C1 occurred prior the probed air reaching the analytical/sampling instrumentation in the container, since clearly elevated stratospheric CO mixing ratios are common to WAS and in situ data.Two more indications, viz.growing [CO] discrepancy with increasing O 3 abundance, and the strong concomitant signal in δ 18 O(CO), imply that O 3 -mediated photochemical production of CO took place.Further, by confronting the C1 and C2 [CO] measurements in a kinetic framework (detailed in Appendix A), we quantify the artefact CO component being chiefly a function of O 3 abundance as C c = b•[O 3 ] 2 , b = (5.19±0.12)•10−5 [mol/nmol], (1) which is equivalent to 8−18 nmol/mol throughout the respective [O 3 ] range of 400−620 i δ nmol/mol (see Fig. 1 (d)).Subtracting this artefact signal yields the corrected in situ C1 CO−O 3 distribution conform to that of C2 (cf.red symbols in Fig. 1 (a)).
fraction of the CH 4 -derived component in CO with increasing [O 3 ], as the 13 C δ c value of −(47.2±5.8)‰inferred at R 2 above 0.4 is characteristic for the δ 13 C of methane in the UT/LMS.
13 C δ c = −(23.3±8.6)‰,indeed at the upper end of the expected LMS δ 13 C(CO) variations of −(25−31)‰, which likely does not allow the MM to ascertain this result as pertaining to the contamination (the corresponding R 2 values are below 0.1).Upon the correction using the inferred 18 O δ c value, the C1 WAS δ 18 O(CO) data appear adequate (shown with red symbols in Fig. 2).That is, variations in the observed C 18 O are driven by (i) the seasonal/regional changes in the composition of tropospheric air and by (ii) the degree of mixing or replacement of the latter with the stratospheric component that is less variable in 18 O.This is seen as stretching of the scattered tropospheric values ([CO] above 60 nmol/mol) in a mixing fashion towards δ 18 O(CO) of around −10‰ at [CO] of ~25 nmol/mol, respectively.The corrected C1 δ 13 C(CO) data (shown in Supplementary Material , Fig. S4) are found to be in a ±1‰ agreement with the observations by B96, except for several deep stratospheric samples ([CO] below 40 nmol/mol).The latter were encountered dur- terminal O atoms in O 3 are enriched w.r.t. to the molecular (bulk) O 3 composition when the latter is above ~+70‰ in δ 18 O(Janssen, 2005;Bhattacharya et al., 2008), therefore an incorporation of only central O atoms into the artefact CO molecules should result in a reduced apparent 18 O δ c value.Such exclusive selection is, however, less likely from the kinetic standpoint and was not observed in available laboratory studies (see Savarino et al. (2008) for a review).For instance, Röckmann et al. (1998a) established the evidence of direct O transfer from O 3 to the CO produced in alkene ozonolysis.A reanalysis of their results (in light of findings of Bhattacharya et al. (2008)) suggests that usually the terminal atoms of the O 3 molecule become transferred (their ratio over the central ones changes from the bulk ~2:1 to ~1:0 for various species).
First, a fast O 3 → CO conversion must have occurred, owing to short (i.e., fraction of a second) exposure time of the probed air to the contamination.Accounting for the typical C1 air sampling conditions (these are: sampled air pressure of 240−270 hPa and temperature of 220−235K outboard to 275−300K inboard, sampling rate of ~12.85•10 −3 moles s −1 corresponding to 350L STP sampled in 1200 s, inlet/tubing volume gauged to yield exposure times of 0.01 to 0.1 s due to variable air intake rate, [O 3 ] of 600 nmol/mol), the overall reaction rate coefficient (k c in Eq. (A1)

( 0 .
10−0.27) of the artefact CO in the sample.Accordingly, the reckoned δ 13 C(CO) biases span (0.5−0.9)‰.Although not large, these well exceed the δ 13 C(CO) measurement precision of ±0.1‰ and were corrected for, and therefore are taken into account in the calculations with the MM presented in Sect.3.1.

Fig. 2 .
Fig. 2. 18 O/ 16 O isotope composition of CO as a function of its reciprocal mixing ratio.Triangles present the data from the remote SH UT/LMS obtained by Brenninkmeijer et al. (1996) (B96).Colour refers to the concomitantly observed O 3 abundances; note the extremely low [O 3 ] encountered by B96 in the Antarctic "ozone hole" conditions.Filled and hollow circles denote the original and corrected (as exemplified by the dashed arrow) C1 WAS data, respectively, with the symbol size scaling proportional to the estimated contamination magnitude (see text).

Fig. 3 .
Fig. 3. Results of the regression calculation with the MM.Shown with symbols are the contamination source isotope signatures i δ c as a function of the respective coefficient of determination (R 2 ).Colour denotes the number of samples in each subset selected.Solid and dashed lines present the best guess ±1 SD for the δ 18 O(O 3 ) and δ 13 C(C c ) estimates.Dashed circles mark the values obtained at highest R 2 for 18 O )where17O Δ t denotes the natural, i.e. expected "true" value of Δ 17 O(CO).The remaining parameters pertain to the contamination kinetic framework (see Appendix A, Eq. (A1)).For the purpose of the current estimate it is sufficient to take 17 Δ n of +5‰ representing equilibrium enrichments expected in the remote free troposphere and UT/LMS.For the O 3 MIF signature 17 Δ c , the value of +30‰ (the average Δ 17 O(O 3 ) expected from the kinetic laboratory data at conditions met along the C1 flight routes, see Sect.3.2 and Table1) is adopted.The coefficient that (Assonov and Brenninkmeijer, 2001)ed for the CO with initially unaccounted MIF (e.g., the sample is assumed to be mass-dependently fractionated) and quantifies some extra +0.73‰ in the analysed δ 13 C(CO) per every +10‰ of Δ 17 O(CO) excess(Assonov and Brenninkmeijer, 2001).The most contaminated C1 WAS CO samples at [O 3 ] above 300 nmol/mol are estimated to bear Δ 17 O(CO) of (6−12)‰ corresponding to fractions of