Introduction
Compared to the troposphere's main oxidant OH (hydroxyl radical), the role of
Cl (atomic chlorine) for CH4 is small. A recently published detailed
model-based estimate attributes ∼ 2.6 % of methane's photochemical
tropospheric loss to Cl (Hossaini et al., 2016). Because this loss
constitutes only a small term in the methane budget, it might be deemed
irrelevant. Nevertheless, growing spatial and temporal coverage in
CH4 observational data allows for top-down estimates of changes in
the source–sink budget to the order of ∼ 1 %. Moreover,
considering that the photochemical sink is the dominant and best-known term
in the global methane budget, it makes sense to improve our calculations. The
grateful aspect of this endeavour clearly is that one does not need an
accurate estimate of Cl as a global tropospheric sink of CH4 as such.
It would already be helpful to have independent estimates of the upper limit
for this interesting sink of CH4, whose rise in the Anthropocene thus
far has contributed one-fifth to global warming.
Irrespective of the implications for the CH4 budget, it stands to
reason to fully understand tropospheric Cl and its chemistry in different air
masses, from marine boundary layer air to strongly polluted air masses, and
several studies address these complex processes. It is also clear that the
budget of a species as fickle as atomic chlorine is hard to determine in
general terms (which forms a less grateful aspect of “assessing chlorine”).
Nevertheless, a new effort – in assessing chlorine's role on a larger than
regional scale, on the basis of trace gas measurements, may be useful.
Even more so than for OH, estimates of the abundance of Cl atoms are chiefly
based on indirect evidence. Direct measurements of OH concentrations ([OH])
are difficult and rare, and for [Cl] this is even much more so. Therefore, the
method (by choice or opportunity) is indirect. Not only are indirect
measurements easier, the use of trace gases that react with OH and Cl also
has the advantage that space- and time-averaged estimates are obtainable. In
this case, one can select for instance two hydrocarbons, one of which has a
comparatively high reactivity to Cl. The change in ratio between the two
hydrocarbon concentrations gives information on [Cl] relative to [OH].
Using stable isotope ratio information offers another such indirect method.
The intrinsic advantage here is that one can use a single trace gas, a single
hydrocarbon, or even the much studied greenhouse gas CH4 itself.
Although the rate coefficient for the reaction of OH with 12CH4 is
only ∼ 4 ‰ faster than that with 13CH4
(Saueressig et al., 2001), for Cl + CH4 the difference is much
larger (Saueressig et al., 1995; Crowley et al., 1999), viz.
63–75 ‰ (at the range of tropospheric temperatures). Broadly
speaking, the presence of 13C-enriched CH4 points to
reaction with Cl. If this were not enough, one could measure the D / H
ratio of CH4 and obtain additional valuable information because of
the large isotope fractionation (KIE, kinetic isotope effect, formerly and
still expressed using the kinetic fractionation constant ε= α-1) and the differences between the KIEs for 13C and
D. A recent paper (Whitehill et al., 2017) reports changes in the clumped
isotopic composition of CH4 in reaction with Cl based on laboratory
experiments, raising hope that clumped isotope measurements (which are very
difficult) may in an additional way assist to further assess the role of Cl
in the oxidation of CH4 in the atmosphere.
An advantage is that the “stable isotope method” in principle removes the
uncertainty about the variability induced by having to use two different
trace gas species, each of which may have an independent, variable source.
Routinely overlooked is another (principle) advantage of stable isotope
analysis offered in the case of atmospheric CH4 → CO
conversion, namely measurement of the isotopic composition of the reaction
product CO. Even though variations in [CO] may not be resolvable due to the
large spatio-temporal variability of its sources and sink, its
13C / 12C ratio may well tell a clearer story. This
is the added advantage of the stable isotope method (we note that the
lifetime of 14C is sufficiently long to render much of what is
stated to also apply to this well-known radioisotope, but there are
complications on which we cannot dwell here).
In this way the presence of Cl during Antarctic ozone hole conditions could
be inferred in an independent fashion (Brenninkmeijer et al., 1996). Not only
did the CH4 inventory become slightly enriched in 13C due
to the large KIE in Cl + CH4, the CO ensuing from CH4
resulted in strong depletions in the background 13CO. There are at least three reasons for the strong isotope
depletion. First, CO concentrations are low in the stratosphere and the in
situ produced CO had a large impact. Second, the 13C content of
CH4 is characteristically low due to its chiefly bacterial origin.
Third, and this is an important point mentioned above, the 13C KIE
for Cl + CH4 happens to be very large. The combination of these
effects renders the stable isotope analysis of CO a sensitive indicator.
Dealing with tropospheric Cl, the same principle has been applied during
springtime tropospheric ozone depletion events in the Arctic. Short-term
bursts of free Cl could be inferred from concomitant decreases in
δ13C(CO) within a per mil range (Röckmann et al., 1999).
We record that there also is a removal of CO by reaction with Cl atoms, with
the rate constant typically being 6 times smaller than that of CO + OH.
Given this very low rate coefficient and the low Cl / OH ratio, only an
extremely large KIE in the CO + Cl reaction could impact significantly on
δ13C of the CO inventory. In contrast, the rate constant for
CH4 + Cl is typically 20 times larger than that for
CH4 + OH. Cl is not expected to play a significant role in
atmospheric CO removal, except possibly at polar sunrise (Hewitt et al.,
1996) and in some stratospheric chemistry analyses (see, for example, Müller et
al., 1996; Sander et al., 2011b). None of a few of papers on tropospheric CO
thus mentions Cl as a sink for CO because of its negligible share;
fortunately, because the reaction product is not so nice.
In this brief account we cannot do justice to all tropospheric Cl-related
papers in the literature and we refer to the recent model-based paper by
Hossaini et al. (2016) and references therein. In comparison with OH, which
is recycled in about two of three reactions in the troposphere (Lelieveld et
al., 2016), the role of recycling of Cl is lower and not well known. The
presence of Cl in the marine boundary layer has been inferred using
hydrocarbon measurements (early study by Parrish et
al., 1993) and likewise during polar sunrise (Jobson et al., 1994),
Cl2 has been measured in situ in coastal air (Spicer et al., 1998)
and in the Arctic (Liao et al., 2014). ClNO2, which is an important
precursor, has been measured (Osthoff et al., 2008 and Thornton et al.,
2010), also by Young et al. (2012), although they found no Cl fingerprint in
hydrocarbon ratios.
Recently, Baker et al. (2016) inferred the presence of Cl in pollution
outflow from continental Asia using hydrocarbon measurements on air samples
collected at cruise altitude by the CARIBIC Lufthansa Airbus aircraft
observatory. Before that, Baker et al. (2011) had likewise inferred that Cl
is formed in an emission plume of the Eyjafjallajökull volcano probed by the
same CARIBIC A340 aircraft. All these and other publications discuss the
presence of Cl in a variety of tropospheric environments wrestling with the
complexity of its chemistry and paucity of experimental data.
The additional importance of revisiting the role of Cl radicals in the present
atmosphere actually surfaces in the reconstruction and understanding of the
budget of CH4 in the past. Changes in the tropospheric burden of
CH4 that occurred in the past (last glacial maximum to present) are
due to changes in CH4 sources and to a minor degree to changes in OH
chemistry (Levine et al., 2011b). One would a priori expect
δ13C(CH4) to provide additional information on
source changes, as it did for immediate past changes (Schaefer et al., 2016),
were it not the case that large changes in Cl abundance may well have affected the
δ13C(CH4) record (Levine et al., 2011a). If this is indeed
the case, changes in Cl abundance in the past may have not affected
the CH4 budget itself significantly, but may have invalidated to a
certain degree the δ13C(CH4) isotope method for
determining changes in sources (biogenic vs. biomass burning).
We turn our attention to a paradox concerning today's tropospheric Cl,
namely: if the presence of tropospheric Cl could be inferred from
13C isotope enrichment in CH4, why is this effect not
visible as concurrent isotope depletion in CO? Or, more explicitly stated, if
the δ13C(CO) isotope method for Cl detection works well for the
austral polar stratosphere in spring (Brenninkmeijer et al., 1996) and for
the polar sunrise in the Arctic (Röckmann et al., 1999), why not so for
the troposphere, or does it? Is a clear negative signal in δ13C(CO) indeed absent, and if so, does this absence allow us to cap
estimates of tropospheric Cl levels?
Data analysis
Chlorine in the Southern Hemisphere
Because the budgets of CH4 and CO in the Southern Hemisphere (SH) are
less complicated than in the Northern Hemisphere, as is shown by their
compact regular seasonal cycles at remote observatories, and because long records
of CO and CH4 including isotopic data are available, we focus on the
Southern Hemisphere. In the SH evidently the emphasis is on Cl generated in
the marine boundary layer (MBL).
We first revisit the information on Cl based on δ13C
measurements of CH4. Initially, mixing ratio and δ13C(CH4) values for shipboard collected air samples in the
Pacific pointed to a large apparent sink isotope fractionation (“apparent”
KIE) of 12–15 ‰ – well in excess of the aforementioned
4 ‰ from OH + CH4 – which led to the conjecture that a
fraction of CH4 is removed in the MBL by Cl atoms which discriminate
strongly against 13CH4 (Lowe et al., 1999; Allan et al., 2001).
Following several publications exploring this effect, Allan et al. (2007)
(hereinafter referred to as A07) using global modelling and observational
data from the extratropical Southern Hemisphere (ETSH), confirmed a large
apparent KIE and could estimate a global marine boundary layer based Cl sink
for CH4 averaging at 25 Tg(CH4) yr-1.
Given this number, a first-order estimate of the accompanying response of
δ13C of CO to the production of CO from Cl + CH4
can be made. Assuming a 100 % yield of CO from OH + CH4 (and
likewise Cl + CH4), the 25 Tg yr-1 CH4 sink
corresponds to a Cl-based annual CO production of 44 Tg yr-1, which is
∼ 1.8 % of the total CO budget. By using a δ13C
value of CO of -28 ‰ (annual tropospheric average), that of
CH4 of -48 ‰ and a KIE of 70 ‰,
(Cl + CH4) causes a negative shift in δ13C(CO)
of about 1.6 ‰. Considering that the lifetime of CO is much shorter
than that of CH4 and that Cl is concentrated in the MBL, the local
and/or seasonal effect on
δ13C(CO) would be even larger.
Unfortunately, a negative shift in δ13C(CO) is
unwelcome in attempts to close the SH CO budget using δ13C.
As Manning et al. (1997) have pointed out, budget closure is only possible
when the yield of CO from CH4 + OH (denoted hereinafter
as λ) is assumed to be merely about 0.7. In other words, even without
incorporating the formation of CO from Cl + CH4, the
CH4-derived 13C-depleted fraction of CO (which is high in
the ETSH at above 40 %) appeared to be too dominant and had to be reduced
by assuming lower yields of CO from CH4. Soon thereafter,
Bergamaschi et al. (2000) also encountered this problem in a 3-D inverse modelling
study using the isotopic composition of CO and could best reconcile data and
model by reducing λ to about 0.86. They do mention that
incorporating CO from Cl + CH4 would require λ values as
low as 0.71. Platt et al. (2004), who discuss mechanisms for the
production of Cl in the marine boundary layer, also allude to the necessity to have
to reduce the assumed CO yield of OH + CH4.
One difficult feature of the δ13C(CH4)-based Cl
estimate was a large inter-annual variability that could not be explained.
A07 identified two periods of different Cl abundance in the ETSH, namely
1994–1996, with MBL values of 28 × 103 atoms cm-3
(high-Cl period, “HC”) and 1998–2000 with much lower values, viz.
9 × 103 atoms cm-3 (low-Cl period, “LC”). The nearly
3-fold drop in the resulting Cl + CH4 sink rate (37 to
13 Tg(CH4) yr-1, or 6.4 to 2.2 % of the total,
respectively) inferred from δ13C(CH4) for the two periods
is not discernible in the simultaneous δ13C(CO) record (see
Sect. 2.2).
Later, Lassey et al. (2011) investigated the apparent KIE in detail and found
that it can differ markedly from both the seasonal and mass-balanced KIEs. In
other words, the apparent KIE derived from the seasonal changes in
[CH4] and δ13C(CH4) values appeared not to
properly represent the respective effects of the two KIEs. The implication is
that the inferred very large range of [Cl] may be in error, and the absence
of a corresponding signal in δ13C(CO) is in that respect an
experimental confirmation. Below we will go into detail.
Observations in the ETSH
We scrutinise the mixing and 13C / 12C ratios of
CH4 and CO in the MBL air at Baring Head, New Zealand
(41.41∘ S, 174.87∘ E, 85 m a.s.l., denoted
hereinafter “BHD”) and at Scott Base, Antarctica (77.80∘ S,
166.67∘ E, 184 m a.s.l., denoted “SCB”) provided by
the National Institute of Water and Atmospheric Research (NIWA, 2010).
Examined in the A07 study on CH4, these data are the result of
laboratory analyses of large air samples collected on a monthly to weekly
basis. The collection strategy (using wind direction, CO2 mixing
ratio temporal stability and back-trajectory analysis) allows air masses that
represent background ETSH air to be selected. Established over two decades,
these time series confer the longest continuous records of 13CH4
and 13CO observations to date. The reported overall uncertainties of the
CH4 mixing ratio and δ13C do not exceed
±0.3 % (about ±5 nmol mol-1) and
±0.05 ‰ (Lowe et al., 1991). For CO, the respective
uncertainties are ±4 % / ±0.2 ‰ (prior to 1994,
Brenninkmeijer, 1993) and ±7 % / ±0.8 ‰ (since 1994,
NIWA, 2010). The CO records from BHD and SCB exhibit small variations in
annual (minimum-to-maximum) span and no significant long-term trend in both
mixing and isotope ratios throughout 1990–2005 (see Gromov, 2013,
Sect. 4.1.1). In contrast to this, the concomitant [CH4] values have
increased on average by about 5 % within the same period, which is
consistent with other observational records (Lassey et al., 2010). It can be
concluded that such augmentation of atmospheric burden of the major (and
largely depleted in 13C) in situ sources of CO remains
statistically indiscernible in the ETSH δ13C(CO) record,
because of more perceptible variations caused by changes in sink and/or the
other (foremost biomass burning) sources of CO.
We subsequently regard the statistics of the two subsets of observational
data falling into the HC and LC periods, as shown in Fig. 1. For testing the
robustness of our comparison against the timing of the air sampling, we
“bootstrap” the data by selecting only the pairs of CH4 / CO
samples collected within 1-week windows (shown with solid boxes in Fig. 1).
This operation has virtually no effect on CO distributions, as its statistic
is smaller (total of 116 and 88 samples at BHD and SCB, respectively) and
controls the sub-sampling of the datasets. For CH4, no effect is
noted either, with an exception of significant (i.e. exceeding measurement
uncertainty) changes to the “bootstrapped” median CH4 mixing ratio
at BHD, which is some 6 nmol mol-1 lower during the HC. This is an
indication that the CO sampling times are likely to be more representative for
background air. Overall, we conclude that the CH4 and CO datasets
reflect variations in the composition of the same background air. Contrary to
CH4, there is no perceptible reduction in seasonal variations of
mixing and isotope ratios of CO at SCB throughout the HC period.
To determine the significance of observed changes in CO using sufficient
statistics, we derive quasi-annual averages (QAAs) of CO mixing / isotope
ratio averages representing the HC, LC and long-term periods (all data and
from 1994 onwards). For the correct temporal weighting of the samples, we
first calculated quasi-monthly averages and their variances, which then
contributed equally to the QAA. Table 1 lists the results along with the
number of samples used in the calculation. Note that there are about twice as
many outliers in the entire BHD record (3.8 %)
compared to those in the SCB (2.2 %), which suggests that the estimated
difference between the HC and LC averages (HC - LC, denoted Δ) is
probably more influenced by regional sources at BHD. Except for
δ13C(CO) at SCB (with considerable significance of Δ
being negative, p value of 0.79), we conclude that all CO QAAs emerge as
statistically indistinguishable, also when compared to the long-term
averages. For CO mixing ratios, the Cl-driven difference should amount up to
1.2 nmol mol-1 (conservatively assuming up to 50 % of CO derived
from CH4 oxidation changed by 4.2 %), which is 2.5–3 times
smaller than the errors in Δ. At both stations, the Δ values
indicate changes to the atmospheric reservoir involving
13C-depleted CO, but in opposite directions (i.e. a removal at
BHD – which contradicts A07 – and an addition at SCB). It is important to
note that the CO + OH sink alters atmospheric CO in a similar fashion
(i.e. the remaining CO burden becomes enriched in 13C).
Statistics on quasi-annual average (QAA) mixing and isotope ratios of CO
observed and simulated at BHD and SCB.
Data
Period
BHD
SCB
n
CO (nmol mol-1)
δ13C(CO) (‰)
n
CO (nmol mol-1)
δ13C(CO) (‰)
HC
1994–1996
65
56.1 ± 2.0
-28.97 ± 0.25
51
50.5 ± 2.6
-29.31 ± 0.64
LCa
1998–2000
48
58.4 ± 2.1
-29.48 ± 0.36
35
49.7 ± 2.5
-28.57 ± 0.64
Δ
HC–LC
-2.2 ± 2.9
+0.51 ± 0.43
+0.8 ± 3.6
-0.74 ± 0.90
Significance (p value)b
0.12/0.002
0.79/0.28
All data
1989–2005
379(15/4)
59.2 ± 1.8
-29.52 ± 0.29
227(5/0)
51.7 ± 2.1
-29.21 ± 0.50
1994–2005
192(5/1)
57.8 ± 2.1
-29.38 ± 0.36
155(0/0)
50.8 ± 2.3
-29.13 ± 0.58
EMAC
1996–2005c
57.0 ± 3.5
51.3 ± 1.7
(incl. from CH4 oxidation)
24.8 ± 0.6
23.7 ± 0.3
Notes: Values in parentheses are the number of outliers
(mild/extreme; see the footnote 4); the latter were
excluded from the calculation of the long-term (up to 2005) averages. Quoted are
standard errors of quasi-annual averages (±1σ). a Time-interpolated value is used for February (no samples
are available at SCB during the LC period). b The p value is
estimated for the null hypothesis that Δ of δ13C(CO)
QAA is below 0/-2σ (left-tail test). c The aggregate of the emission inventories used in the simulation correspond
closest to 2000 (see details in Gromov et al., 2017).
Statistics on the CH4 and CO mixing and
13C / 12C ratios observed at Baring Head (BHD) and
Scott Base (SCB) throughout the high-Cl (HC, orange shaded) and low-Cl (LC,
grey shaded) periods hypothesised by Allan et al. (2007) (see text for
details). Panels (c, d) show statistics on the anomalies with
respect to the annual averages (denoted with “Δyr”). Panel
(g) displays the number of samples in each subset. The full time
series of the data are shown in the Supplement (Fig. S2). Boxes and whiskers
present the median and interquartile ranges and ±1σ (of the
population) of the data. Circles and minus symbols denote the averages and
samples falling outside ±1σ. Solid boxes denote the subset of
data when CH4 and CO samples were taken simultaneously (up to 7 days
apart); hatched boxes refer to all data.
EMAC model
For extending the interpretation of observed ETSH CO, we resort to the
results of simulations performed with the ECHAM5/MESSy Atmospheric Chemistry
(EMAC) general circulation model (Jöckel et al., 2010). EMAC includes all
relevant processes (atmospheric transport, calculation of chemistry kinetics,
photolysis rates, trace gas emissions, etc.) for simulating the current
global atmospheric state. The set-up we use resembles that of the EMAC
evaluation study (MESSy Development Cycle 2, Jöckel et al., 2010) and is
augmented with kinetic tagging tools (Gromov et al., 2010). These allow
direct quantification of the CO component stemming from CH4 oxidation
(and as corollary provide λ) by following the carbon (C) exchanges
through all intermediates (shown in Supplement Fig. S1) within a
comprehensive chemistry mechanism simulated by the MECCA sub-model (Module
Efficiently Calculating the Chemistry of the Atmosphere, Sander et al.,
2011a). The emission set-up contains only the standard emissions and
precursors of Cl and yields average MBL Cl concentrations in the order of
101–102 atoms cm-3 (see the detailed simulated budgets in
the Supplement, Table S1). These results are in line with MBL [Cl] of
(0.5–2) × 102 atoms cm-3 obtained by Hossaini et
al. (2016) in a similar model set-up (ORG2).
The QAAs of [CO] simulated in EMAC for the period 1996–2005 in the grid
boxes enclosing the locations of BHD and SCB are also given in Table 1.
Despite the spatial and temporal averaging used (∼ 2.8∘
horizontal grid cell size at the T42L31ECMWF resolution, weekly averages),
model QAAs match observations well and have similar uncertainties (resulting
from monthly means variation; the observed and simulated seasonalities are
shown in the Supplement, Fig. S3). Due to longer lifetimes of CO and
CH4 in the well-mixed ETSH and, more importantly, their synchronous
sink and production via OH, we expect much lower (factor
∼ 1/5 compared to that of the total CO) variation in the
CH4-derived [CO] component. The fraction of the latter (denoted
γ, see Table 2) is proportional to the average tropospheric λ of 93 % (diagnosed simulated value). Depending on the zonal domain, Cl
atoms in EMAC initiate (0.15–0.25) % of CH4 sink in the
troposphere. The fraction of CH4 removed in the ETSH
(43 Tg(C) yr-1) is minor compared to that in the tropics
(271 Tg(C) yr-1). About 13 % of tropospheric sink occurs in the
boundary layer.
Additionally, we simulate the sink-effective 13C enrichment in CO
(denoted ηc) resulting from the 12C-preferential
CO + OH reaction and removal of the CH4 → CO chain
intermediates (dry or wet deposition, when γ < 1),
convoluted with atmospheric mixing and transport. The corresponding
ηc value at a given space–time point denotes how much higher
the δ13C of airborne CO is compared to the case when sink
KIEs were absent. Altogether, values of γ and ηc at
the stations and domain-wise integrals of CH4 sink (S) and λ (listed in Table 2) are used in the calculations that follow now.
Sensitivity of δ13C(CO) to the CH4 + Cl sink
Using the observational and model data, we attempt to estimate the
sensitivity of δ13C(CO) at a given station to supposed
inter-annual changes in the Cl-initiated CH4 sink. The QAA of δ13C(CO) (denoted δc) can be approximated as a
two-component mixture of CH4- and non-CH4-derived CO sources
augmented by the effective sink enrichment:
δc≅(1-γ)δn+γ(δm-εm)+ηc.
We refer the reader to Table 2 for the explanation of the parameters and
their values. In essence, we account for the fractionations induced in
atmospheric sinks (ηc in CO and εm in
CH4) and mix the sources in the proportion defined by γ.
Exemplifying the estimate from A07, SH Cl changes should cause
εm to drop from 15 ‰ to 7 ‰ between
the HC and LC, rendering δ13C of the carbon from CH4
arriving to CO of -62.2 ‰ and -54.2 ‰, respectively. By
rearranging Eq. (1) we derive the non-CH4 CO source
δ13C signature δn (see Table 2). Since
there are virtually no surface sources of CO south of 40∘ S in the
ETSH (see, for example, Gromov et al., 2017, Sect. 3.4), the difference in
δn at BHD and SCB could be driven only by poleward
13C enrichment of the non-CH4 in situ sources (e.g.
oxidation of higher hydrocarbons) and/or a stronger (than simulated in EMAC)
zonal gradient in ηc. Note that the station-wise
δn discrepancy scales with the εm
value, although not strongly: at εm of OH sink KIE
(3.9 ‰) it reduces from (2.2 ± 2.1) ‰ to
(1.5 ± 2.2) ‰. In a statistical sense, the derived
δn values reflect the same underlying source signature
(p value is 0.31).
Parameters used in calculus.
Species and parameter (unit)
Value
Station
CO
BHD
SCB
γa
CH4-derived component (%)
43 ± 3
46 ± 2
ηca
Eff. 13C sink fractionation (‰)
+4.2±0.2
+4.6±0.1
δnb
δ13C of non-CH4 sources (‰)
-15.0±1.7
-12.8±1.3
δc
Observed δ13C(CO) (‰)
-29.5±0.3
-29.2±0.5
Domain:
CH4
SH
ETSH
Sa,c
Total sink (Tg(C) yr-1)
187.8
52.5
δm
Observed δ13C(CH4) (‰)
-47.2
λa
Yield of CO from CH4
93 %
Periodd:
HC
LC
ΔS
Changes to S due to Cl variations
+18
0
(Tg(C) yr-1)
εm
Total CH4 sink KIE (‰)
15
7
Notes: Quoted QAAs and
standard errors (±1σ); the latter are omitted for the components
contributing to δc and δn errors
insignificantly. a Estimate based
on EMAC results.
b Derived at εm= 11 ‰
(average of the LC and HC periods).
c Includes the LC Cl sink term
from A07 (9.7 Tg(C) yr-1). For the SH, the sum of the ETSH and halved intra-tropical
integrals is taken.
d Estimates from A07.
Using Eq. (1) to define δc in the HC and LC periods, one obtains its sensitivity
(Δδc) to changes in the CH4 + Cl sink
(ΔS) and in the total sink KIE (Δεm):
Δδc=(λa/λ)LCγ((δm-HCεm-δn)μ-Δεm).
Here superscripts indicate the period the values are taken for, Δ
denotes the HC - LC difference (same as in Sect. 2.2 above) and μ=ΔS/LCS is the change in the total CH4 sink S
relative to the LC conditions. The value of S represents the tropospheric
column of a given domain, i.e. we assume that ΔS is distributed
homogeneously over the SH or ETSH. Formulated using γ, Eq. (2) allows
the projection of the results for the alternative CO yield value
λa (different from that obtained in EMAC), as our
simulations confirm that λ
directly proportionates γ and S in the tropospheric column
(but not in the MBL). Furthermore, Δδc is derived
under the assumption of constancy of ηc and
δn values. Whilst for ηc such is likely the
case (judging by the very similar observed CO mixing ratios, and hence
lifetimes, during HC and LC), for the latter an upper limit of
±1 ‰ can be put from the typical variation in the δ13C of the underlying sources (see Gromov et al., 2017, Table 5). This is
lower than the uncertainty associated with δn values derived
here (cf. Table 2); we
discuss the range of δn values required to concomitantly
mask the changes in δc below.
(a) Top: expected CH4 + Cl sink-driven changes
to δ13C(CO) between HC and LC periods at the ETSH stations
(Δδc) as a function of CH4-derived CO
fraction (γ, top axis) resulting from assumed yield values
(λa, bottom axis, approximate). Large symbols denote the
observed (ordinate) and simulated (abscissa, EMAC) values. Thick lines
present Δδc values calculated using Eq. (2) assuming
that hypothesised changes to the CH4 + Cl sink occur within the
entire SH (solid) and ETSH only (dashed). Thin dashed–dotted lines exemplify
the effect due to mere changes in CH4 sink KIE (Δεm). Bottom: average augmentation to the non-CH4
sources signature (Δδn) required to compensate
Δδc at the respective values and domains (note the
different axis shown in red). Errors bars and areas denote ±1σ of
the annual means and derived estimates. See Sects. 2.4 and 3 for details.
(b) Tropospheric yield of CO from CH4 oxidation reckoned in
the current and previous studies. Symbols (error bars) denote the best (range
of) estimates or the global (domain) averages. Abbreviations refer to the
following: L81 –
Logan et al. (1981), LC91 – Lelieveld and
Crutzen (1991), T92 – Tie et al. (1992), M97 –
Manning et al. (1997), B00 – Bergamaschi et al. (2000),
F06 – Folberth et al. (2006), D07 – Duncan et al. (2007), E10 – Emmons et
al. (2010), H11 – Hooghiemstra et al. (2011), G13 – Gromov (2013), GT14 –
Gromov and Taraborrelli, MPI-C (unpublished results using EMAC,
2014), F17 –
Franco et al. (2018), EMAC – current study.
Figure 2a shows the values of Δδc, calculated for
different stations and domains, as a function of γ (implicitly
scaling with arbitrarily chosen yield value λa). Very large
changes are expected for the ETSH, where μ is about 4 times that in the
SH. Importantly, the LCS value includes the Cl sink term from A07
(which is ∼ 29 times greater than the total tropospheric
CH4 + Cl sink simulated in EMAC); hence, we receive the “lowest
sensitivity” for the case when the Cl sink is added up to (instead of partly
replacing) the other CH4 sinks, e.g. that via OH. Alternatively,
Δδc will additionally intensify by -0.2 ‰
and -(1.8–2.1) ‰ in the SH and ETSH, respectively. By setting
μ= 0 in Eq. (2), we quantify the contribution of the CH4 sink
KIE (which increases by Δεm) only. Independent
from the assumptions on the Cl sink domain and magnitude, it demonstrates the
effect of lowering of δ13C of C arriving to CO from
CH4 and accounts for one-third to two-thirds of the total Δδc value (cf. Fig. 2, thin dashed–dotted line).
Finally, we estimate the equivalent increase in the δ13C
value of the non-CH4 sources (Δδn) that would
be required to mask the depleting effect of a hypothetical
CH4 + Cl sink increase. We subtract Eq. (1) written for the HC
and LC and solve it assuming Δδc= 0 (notation from
Eq. 2 is kept):
Δδn=(δn-(δm-LCεm))μ+(1+μ)Δεm((λa/λ)LCγ)-1-(1+μ).
Averages of Δδn at BHD and SCB are plotted in the lower panel of
Fig. 2a, respectively. Similar to Δδc, Δδn scales with the assumed domain and CH4 input to
CO, albeit stronger, because δn is closer to the
δ13C of the total CO source (δc-ηc) compared to that for CH4 (δm-εm). Thus, if we accept the EMAC-suggested
tropospheric CO yield in the SH of λ= 93 %, Cl-driven changes
to the δ13C(CO) at BHD and SCB are expected to be of at
least -(5.8–6.3) ‰ between the LC and HC, unless these are masked
by unrealistic concurrent increases in δ13C of the
non-CH4 sources of about +(11.6–13.5) ‰. If one assumes
the CH4 + Cl sink changes only within the ETSH, these estimates
scale to -(13.1–14.5) ‰ and +(46–61) ‰, respectively.
It is important to note that we gauge the expected changes to the annual
averages of δ13C(CO), which do integrate seasonal variations. The
latter are observed at merely ±1.5 ‰ (cf. Figs. 1 and S2) and
should also increase strongly, if the Cl sink has a similar seasonal
variation to that of OH (although A07 used a seasonal cycle based on
dimethyl-sulfide-related species in the SH, which has a shorter summer
maximum).
Discussion
The photochemical yield of CO from CH4 constitutes a major factor of
uncertainty in the CO budget. Modelling studies to date agreed on values of
λ≥ 0.7 (see the overview in Fig. 2b). Several recent
studies (see D07, E10 and H11) suggest, however, that λ is close to
unity and by doing so contradict findings of 13CO-inclusive studies
(see M97, B00 and G13). Assuming that λ < 0.7 or that
λ ∼ 1 would be in conflict with basic principles, i.e.
photochemical kinetics and dry and wet removal processes affecting the
intermediates of the CH4 → CO chain, or their erroneous
implementation in the global atmospheric models.
Our estimates of Δδc bear the uncertainty of the
assumed λ value; nonetheless, they affirm that even if only 70 %
of reacted CH4 molecules yield CO, at least one-third of the changes
to the δ13C signature of this source (that is, (δm+εm) times 0.7) should be expressed in the
ETSH δ13C(CO). Since δm changed by about
+0.1 ‰ between the HC and LC periods (cf. Fig. 1b), we conclude
that εm could not change by more than +2 ‰
in the SH as well (with this estimate being lower for λ above
0.7). Furthermore, statistically significant
non-zero Δδc values (p value of 0.01) should appear
at very low λ, viz. above 0.05 (ETSH sink) and 0.12 (SH sink,
respectively). We regard these two atmospheric domains because observations
in the well-mixed ETSH may not single out the actual location of the
Cl + CH4 sink: The large part of sink-driven variations in the
mixing ratio and δ13C of CH4 and CO is merely
transported into the ETSH from the tropics, where almost three-quarters of
the total CH4 sink and accompanying CO production is expected (see
Table S1 for EMAC results, also Gromov, 2013, Sect. 6.2.3). Accordingly,
Hossaini et al. (2016) also assign a major fraction of the
CH4 + Cl sink to the lower latitudes. If such were not the case
(i.e. if varying Cl + CH4 sink were confined to the ETSH), the
estimated effect on δ13C(CO) would be roughly twice that
reckoned for the SH, i.e. extreme values.
There are a few remarks on the usability of the method used by A07, in
addition to the thorough theoretical enquiry by Lassey et al. (2011).
Evidence, or at least indications, for Cl in the ETSH is based on the
[CH4] vs. δ13C(CH4) Lissajous (a.k.a. phase)
diagrams being ellipses in the case of seasonal cycles. The slope of their
major axis gives the “apparent” KIE, from which the ratio Cl / OH can
be inferred when the individual KIEs are known. Clearly, Cl was not assessed
on the basis of the annual average value of δ13C(CH4) but
on the basis of its seasonal cycle, which is small. Using annual averages,
however, is still impeded by perceptible long-term trends in [CH4]
and δ13C(CH4), which neither A07 (who consider the
final 8 equilibrated years of the 40-year spin-up simulations) nor Lassey et
al. (2011) (who use a rather idealised model) have accounted for. For
example, presence and asynchronous evolution of [CH4] and
δ13C(CH4) long-term trends could result in different
mixing and transport of CH4 isotopologues compared to that resulting
from trend-free simulated seasonal variations. We note that while observed
[CH4] growth is similar throughout both HC and LC periods, such is
not the case for δ13C(CH4), which does not increase
in the LC (cf. Fig. S2a, c and, in particular, the seasonal time series
fits for CH4 at the NIWA
website). Furthermore, the latter is likely a
global signal of the 2000–2007 intermittent stop in tropospheric CH4
growth, which manifested itself in δ13C earlier than in
mixing ratios and terminated with the reversed
13C / 12C trend (see, for example, Nisbet et al.,
2016). Currently available observational data do not allow unambiguous
attribution of this global phenomenon to one or several causes proposed
(Turner et al., 2017), however.
Our incomplete information about the 13C isotopic composition of
CH4 sources presently prevents a Cl-induced input into the annual
average value of δ13C(CH4) being singled out, even
though it should be perceptible (about +1.5 ‰, assuming for
the sake of matter a 2.5 % Cl sink). The corresponding negative shift in
δ13C(CO) is about 1.6 ‰ (estimated in Sect. 2.1).
In this respect, δ13C(CH4) and δ13C(CO)
are equally sensitive to Cl. Because oxidation of CH4 is a main
source of CO in the ETSH, and the isotopic composition of atmospheric
CH4 is better known than that of its sources, it may well be that
variation in the annual average value of δ13C(CO) is a more
useful variable for estimating [Cl]. The relatively long lifetime and small
seasonality in sources result in weak seasonal cycles of mixing ratio and
δ13C in CH4. In contrast, the seasonal cycle of
δ13C(CO) is dominated by the large difference in isotopic
composition of its sources, with the main driver being the switch between CO
from CH4 oxidation and that of the other sources. Since the presence
of Cl makes CH4 oxidation an even more 13C-depleted source,
the impact of CH4 oxidation on CO in the ETSH peaks and may render
the seasonal amplitude (in particular summer minima) of δ13C(CO) a
sensitive indicator for Cl. Unfortunately, deficit of observational data
(large uncertainties due to insufficient statistics) currently hinder such
application.
A fundamental problem remains that the ETSH δ13C(CO) budget
cannot be closed even when a Cl sink is excluded, unless a CO yield from
CH4 of 0.7–0.86 is assumed (Manning et al., 1997; Bergamaschi et
al., 2000). Yields below unity leave, however, the possibility that a
positive fractionation in the removal of the CH4 → CO
intermediates may be at play. Using λ= (0.7–0.86) and γ= 0.3 for the troposphere, one calculates that an average KIE of
(11–33) ‰ should escort the removal of intermediates in order to
offset the Cl input to δ13C(CO). This estimate is 3–8 times
higher than current parameterisations suggest (about 4 ‰, see
Gromov, 2013, Sect. 6.2.4) and is even higher in the SH, where γ is
above 0.4. Another complication is potentially present because one cannot
exclude that the room temperature laboratory data for the 13C KIE
for CO+OH reaction are not applicable to the bulk of the troposphere, even
though the reaction itself is little temperature- but mostly
pressure-dependent (see Gromov, 2013, Sect. 6.1.4). The unbalanced
13C(CO) budget may then be the consequence of underestimating the
CO sink KIE in the models, despite adequate estimates of the sources'
13C / 12C ratios.
Conclusions
We emphasise the value of long-term observations of CO isotopic composition,
especially at locations like Scott Base (Antarctica), where influence of
local sources is smallest and the fraction of photochemically produced CO is
largest. In combination with modelling (e.g. EMAC), δ13C(CO)
allows monitoring for intra-annual changes in the carbon isotopic composition
of CH4-derived CO, namely the δ13C value of reacted
CH4 modified by the total sink KIE (εm). Within
the range of probable λ values (0.7–0.93), we are able to cap the
potential changes in εm by +(2.0–1.5) ‰
between 1994–1996 and 1998–2000 in the ETSH, which contrasts the
+8 ‰ derived by Allan et al. (2007). Conversely,
δ13C(CO) may also be employed for “top-down” estimates of
δ13C values of CH4 sources, provided the
εm is equilibrated on a scale of a tropospheric
CH4 lifetime. This could be achieved in a differential mixing model
(also known as the “Keeling” plot) contrasting small variance in
CH4-derived [CO] and δ13C and largely varying input
from other CO sources (e.g. biomass burning).
We conclude that δ13C(CO) is particularly sensitive to the
CH4 + Cl sink. Its temporal variations, if they exist, may allow
an independent “bottom-up” [Cl] proxy to be calibrated, e.g. emissions of
Cl simulated in process-based models. For example, changes in observed
δ13C(CO) at SCB (see Table 1) allow variations of the Cl-driven
sink of CH4 not larger than (1.5 λa-1) % of
its total (assuming the yield λa of CO from CH4).
Projecting this figure onto EMAC results (Table S1, zonal tropospheric
integrals) implies that variations in mean ETSH chlorine abundance should
have not exceeded Δ[Cl] = (0.9 λa-1)×103 atoms cm-3 between 1994–1996 and 1998–2000. Regarding the
fact that Manning et al. (1997) and Bergamaschi et al. (2000) could only
close the SH 13C(CO) budget assuming λ values of 0.7 and
0.86, which are within the generally accepted range, it is unlikely that
tropospheric Cl is as high as assumed in the literature.
Although invoking isotopic information is often like opening a can of worms
(scientists' favourite diet), relevant conclusions emerge.
Lassey et al. (2011) exposed shortcomings of the phase
diagram method; we show here, using
a low- and high-Cl scenario, that unrealistic yield values of CO from
CH4 oxidation (λ below 0.12 in the SH) and/or implausible
increases in the δ13C of non-CH4 sources of CO
(exceeding +7 ‰ at realistic λ≥ 0.7) would have to be
assumed to explain the absence of concurrent inter-annual variations in
δ13C(CO) in the ETSH. This constitutes an independent,
observation-based evaluation of [Cl] variations envisaged by Allan et
al. (2007), from which we conclude that such variations are extremely
unlikely. Concerning estimates of background levels of Cl, even attributing
1 % of the total tropospheric sink of CH4 to Cl aggravates the
non-trivial problem of balancing the global 13C(CO) budget. It
follows that the role of tropospheric Cl as a sink of CH4 oxidation
(see, for example, Saunois et al., 2016, and references therein) is seriously
overestimated.