Chloromethane (CH3Cl) is an important provider of chlorine to the
stratosphere but detailed knowledge of its budget is missing. Stable isotope
analysis is a potentially powerful tool to constrain CH3Cl flux
estimates. The largest degree of isotope fractionation is expected to occur
for deuterium in CH3Cl in the hydrogen abstraction reactions with its
main sink reactant tropospheric OH and its minor sink reactant Cl atoms. We
determined the isotope fractionation by stable hydrogen isotope analysis of
the fraction of CH3Cl remaining after reaction with hydroxyl and
chlorine radicals in a 3.5 m3 Teflon smog chamber at 293 ± 1 K.
We measured the stable hydrogen isotope values of the unreacted CH3Cl
using compound-specific thermal conversion isotope ratio mass spectrometry.
The isotope fractionations of CH3Cl for the reactions with hydroxyl and
chlorine radicals were found to be -264±45 and -280±11 ‰, respectively. For comparison, we performed similar
experiments using methane (CH4) as the target compound with OH and
obtained a fractionation constant of -205±6 ‰ which is in good
agreement with values previously reported. The observed large kinetic isotope
effects are helpful when employing isotopic analyses of CH3Cl in the
atmosphere to improve our knowledge of its atmospheric budget.
Introduction
Chloromethane (often called methyl chloride) is the most abundant chlorine-containing trace gas in the Earth's atmosphere, currently with a global mean
mixing ratio of ∼ 540 ± 5 parts per trillion by volume (pptv) and
an atmospheric lifetime of 1.0–1.2 years (Carpenter et al., 2014). The
global emissions of CH3Cl have been estimated to be in the range of 4 to
5 Tg yr-1 (1 Tg = 1012 g) stemming from predominantly
natural but also anthropogenic sources (Montzka and Fraser, 2003; WMO, 2011;
Carpenter et al., 2014). However, current estimates of the CH3Cl global
budget and the apportionment between sources and sinks are still highly
uncertain. Known natural sources of CH3Cl include tropical plants
(Yokouchi et al., 2002, 2007; Umezawa et al., 2015),
wood-rotting fungi (Harper, 1985), oceans (Moore et al., 1996; Kolusu et al.,
2017), plants of salt marshes (Rhew et al., 2000, 2003), aerated
and flooded soil (Redeker et al., 2000; Keppler et al., 2000), senescent
leaves and leaf litter (Hamilton et al., 2003; Derendorp et al., 2011) and
wildfires. Anthropogenic CH3Cl release to the atmosphere comes from the
combustion of coal and biomass with minor emissions from cattle (Williams et
al., 1999) and humans (Keppler et al., 2017). In addition, it has been
reported that emissions from industrial sources, particularly in China, might
be much higher than previously assumed (Li et al., 2017).
The dominant sink for atmospheric CH3Cl results from the reaction with
photochemically produced hydroxyl radicals (OH), currently estimated at about
2.8 Tg yr-1 (Carpenter et al., 2014). Furthermore, in the marine
boundary layer the reaction of CH3Cl with chlorine radicals (Cl)
represents another sink estimated to account for up to 0.4 Tg yr-1
(Khalil et al., 1999; Montzka and Fraser, 2003). Microbial CH3Cl
degradation in soils may be a relevant additional global sink (McAnulla et
al., 2001; Harper et al., 2003; Miller et al., 2004; Jaeger et al., 2018a),
but its impact on the global CH3Cl budget is still highly uncertain. The
microbial CH3Cl soil sink strength has been estimated to range from 0.1
to 1.6 Tg yr-1 (Harper et al., 2003; Keppler et al., 2005; Carpenter
et al., 2014). Moreover, small proportions of tropospheric CH3Cl are
lost to the stratosphere (146 Gg yr-1, 1 Gg = 109 g) and to
cold polar oceans (370 Gg yr-1) though oceans in total are a net
source (Carpenter et al., 2014). Loss of tropospheric CH3Cl to the
stratosphere is a result of turbulent mixing and the transport process itself
is not thought to cause substantial isotope fractionation (Thompson et al.,
2002).
A potentially powerful tool in the investigation of the budgets of
atmospheric volatile organic compounds is the use of stable isotope ratios
(Brenninkmeijer et al., 2003; Gensch et al., 2014). The general approach is
that the atmospheric isotope ratio of a compound (e.g., CH3Cl) is
considered to be equal the sum of isotopic fluxes from all sources corrected
for kinetic isotopic fractionations that happen in sink processes:
δ2Hatm=∑i=lnΦisource×δ2Hisource+∑j=lnΦjsink×εjsink,
where δ2Hatm and
δ2Hisource are the hydrogen isotope values of
CH3Cl in the atmosphere and of the different sources i in per mil.
Φi and Φj are the CH3Cl flux fraction for each source
and sink. εj is the isotope fractionation of each sink j in
per mil.
The isotopic composition of atmospheric compounds might be altered by the
kinetic isotope effects of physical, chemical or biological loss processes.
The kinetic isotope effect (KIE) is usually defined as
KIE=k1k2,
where k1 and k2 are the reaction rate constants for loss of the
lighter and the heavier isotopologues, respectively. The KIE is typically
expressed as isotope fractionation ε (also termed isotope
enrichment constant) or isotope fractionation constant α.
First approaches of an isotope mass balance regarding stable carbon isotopes
of CH3Cl have been provided by Keppler et al. (2005) and Saito and
Yokouchi (2008). Several studies have investigated the stable carbon isotope
source signature of CH3Cl produced via biotic and abiotic processes;
however, for a more detailed overview we refer readers to the studies of
Keppler et al. (2005) and Saito and Yokouchi (2008). Moreover, researchers
have measured the KIE of stable carbon isotopes of CH3Cl during
oxidation and during biodegradation by bacterial isolates (Miller et al., 2001;
Nadalig et al., 2013; Nadalig et al., 2014), and in soils under laboratory
conditions (Miller et al., 2004; Jaeger et al., 2018a). The first, and so
far, only available analysis of the KIE for reaction of CH3Cl with OH
has been reported by Gola et al. (2005) and this revealed unexpectedly
large stable carbon isotope fractionation. The experiments were carried out
in a smog chamber using long path Fourier transform infrared spectroscopy
(FTIR) detection. However, we consider it important to confirm this result
using another measurement technique such as stable isotope ratio mass
spectrometry (IRMS).
So far most isotopic investigations of CH3Cl have focused on stable
carbon isotope measurements, but stable hydrogen isotope measurements
including both sources and sinks of CH3Cl have also recently become
available (Greule et al., 2012; Nadalig et al., 2013, 2014;
Jaeger et al., 2018a, b). Moreover, relative rate
experiments have been carried out for three isotopologues of CH3Cl and
their reactions with Cl and OH. The OH and Cl reaction rates of CH2DCl
were measured by long-path FTIR spectroscopy relative to CH3Cl at
298 ± 2 K and 1 atm (Sellevåg et al., 2006; Table 1).
Reported hydrogen isotope enrichment constants for the reaction of
CH3Cl with OH radicals and with Cl atoms and the reaction of CH4
with OH radicals.
Reactionε / ‰Method and remarksReferenceCH3Cl + OH-264± 45experimental: 3.5 m3 smog chamber at 293 ± 1 K; IRMSExp. 1 to 3, this studyCH3Cl + OH-410± 50experimental: smog chamber, long-path FTIR spectroscopy relative to CH3Cl at 298 ± 2 KSellevåg et al. (2006)CH3Cl + OH-330 to -430theoretical calculationsSellevåg et al. (2006)CH3Cl + Cl-280± 11experimental: 3.5 m3 smog chamber at 293 ± 1 K; IRMSExp. 4, this studyCH3Cl + Cl-420± 40experimental: smog chamber, long-path FTIR spectroscopy relative to CH3Cl at 298 ± 2 KSellevåg et al. (2006)CH3Cl + Cl-540 to -590theoretical calculationsSellevåg et al. (2006)CH4+ OH-205± 6experimental: 3.5 m3 smog chamber at 293 ± 1 K; IRMSExp. 5, this studyCH4+ OH-227± 11experimental: at 296 K, IRMS and tunable diode laser absorption spectroscopySaueressig et al. (2001)CH4+ OH-231± 45experimental: at 277 KGierczak et al. (1997)CH4+ OH-251± 10ab initio at 298 KXiao et al. (1993)CH4+ OH-145± 30experimental: at 298 KDeMore et al. (1993)CH4+ OH-294± 18experimental: smog chamber, long-path FTIR spectroscopy relative to CH3Cl at 298 ± 2 KSellevåg et al. (2006)CH4+ OH-60 to -270theoretical at 298 KSellevåg et al. (2006)
In this manuscript, using a 3.5 m3 Teflon smog chamber and IRMS
measurements, we present results from kinetic studies of the hydrogen isotope
fractionation in the atmospheric OH and Cl loss processes of CH3Cl.
Furthermore, we also measured the isotope fractionation for the reaction
between methane (CH4) and OH using a similar experimental design and
compared this value with those from previous studies.
Materials and methodsSmog chamber experiments with chloromethane
The isotope fractionation experiments were performed in a 3.5 m3 Teflon
smog chamber (fluorinated ethylene propylene, FEP 200A, DuPont, Wilmington,
DE, USA) with initial CH3Cl mixing ratio of 5 to 10 parts per million by
volume (ppmv). Atomic chlorine was generated via photolysis of molecular
chlorine (Cl2; Rießner Gase, 0.971 % Cl2 in N2) by a
solar simulator with an actinic flux comparable to the sun in mid-summer in
Germany (Bleicher et al., 2014). Hydroxyl radicals were generated via the
photolysis of ozone (O3) at 253.7 nm in the presence of water vapor
(RH = 70 %; produced by double-distilled water in a three-neck
bottle humidifier) and/or H2. To obtain efficient OH formation, a
Philips TUV lamp T8 (55 W) was coated with Teflon film (FEP 200) and mounted
inside the smog chamber. O3 was monitored by a chemiluminescence
analyzer (UPK 8001). The chamber was continuously flushed with purified,
hydrocarbon-free zero air (zero air generator, cmc Instruments,
< 1 ppbv of O3, < 500 pptv NOx,
< 100 ppbv of CH4) at a rate of 4 L min-1 to maintain a
slight overpressure of 0.5–1 Pa logged with a differential pressure sensor
(Kalinsky Elektronik DS1). The quality of the air inside the chamber in terms
of possible contamination was controlled by monitoring NO and NOx
(EcoPhysics CLD 88p, coupled with a photolytic converter, EcoPhysics PLC
860). Perfluorohexane (PFH) with an initial mixing ratio of ∼ 25 parts
per billion by volume (ppbv) was used as an internal standard to correct the
resulting concentrations for dilution. The temperature was set to
20 ± 1 ∘C and monitored, together with the relative humidity,
by a Teflon-cased sensor (Rotronic, HC2-IC102). To guarantee constant mixing
and small temperature gradients, a Teflon fan was mounted and operated inside
the chamber. More detailed specification of the smog chamber can be found
elsewhere (e.g., Wittmer et al., 2015). The mixing ratios of CH3Cl and
PFH were quantified by a Hewlett Packard HP 6890 gas chromatograph coupled to
a MSD 5973 mass spectrometer (GC-MS, Agilent Technologies, Palo Alto, CA)
with a time resolution of 15 min throughout the experiments. Two CH3Cl
reference gases from Linde (1006 ± 12 ppmv diluted in N2) and
Scott (1 ppmv) were used for calibration. The abundance of CH3Cl
relative to PFH was used to calculate the remaining fraction of CH3Cl
(Eq. 4). The relative standard deviation of this procedure was determined
prior to each experiment and also during the control experiment and ranged
between 1.3 and 1.9 %. Aliquots (5 mL) were withdrawn from the chamber
with a gas-tight syringe, injected into a stream of He (30 mL min-1)
and directed to a pre-concentration unit that was attached to the GC-MS. The
pre-concentration unit consisted of a simple eight-port valve (VICI Valco)
equipped with two cryotraps made of fused silica, which were immersed in
liquid nitrogen for trapping the analytes. Prior to each sample measurement,
a gaseous standard (5 mL of 100 ppmv CH3Cl in N2) was measured.
Figure 1 shows the design of the smog chamber used in our experiments.
Schematic of the experimental smog chamber.
In the CH3Cl and OH experiments (1 to 3) 2000 ppmv H2 was used to
scavenge chlorine atoms originating from the photolysis or oxidation of
formyl chloride (HCOCl), which forms as an intermediate in the reaction
cascade. Under the experimental conditions typically more than 70 % of
the CH3Cl was degraded within 6 to 10 h. From each experiment
(CH3Cl + OH and CH3Cl + Cl) 10 to 15 canister samples (2 L
stainless steel, evacuated < 10-4 mbar) were collected at
regular time intervals for subsequent stable hydrogen isotope measurements at
Heidelberg University. An overview of the experimental details (Table S1 in
the Supplement) and control measurements is provided in the Supplement.
Smog chamber degradation experiments with methane
The CH4 degradation experiments were carried out under the same conditions
as the CH3Cl degradation experiments but without PFH as an internal
standard. Instead we used the flushing flow rate of zero air to account for
the dilution during the experiment. The initial CH4 mixing ratio was
6 ppmv. Throughout these experiments CH4 and CO2 mixing ratios
were monitored with a Picarro G225i cavity ring-down spectrometer directly
connected to the chamber. For more details see information provided in the
Supplement.
Stable hydrogen isotope analysis using isotope ratio mass
spectrometryChloromethane
Stable hydrogen isotope ratios of CH3Cl were measured by an in-house
cryogenic pre-concentration unit coupled to a Hewlett Packard HP 6890 gas
chromatograph (Agilent Technologies, Palo Alto, CA) and an isotope ratio mass
spectrometer (IRMS; Isoprime, Manchester, UK) as described in detail by
Greule et al. (2012). Diverging from the method of Greule et al. (2012) a
ceramic tube reactor without chromium pellets at 1450 ∘C was instead
used for high-temperature conversion (HTC). A tank of high-purity H2
(Alphagaz 2, hydrogen 6.0, Air Liquide, Düsseldorf, Germany) with a
δ2H value of ∼-250 ‰ versus the Vienna Standard Mean
Ocean Water (VSMOV) was used as the working gas. The conventional delta
notation, expressing the isotopic composition of the sample relative to that
of VSMOW in per mil is used. All sample δ2H values were measured
relative to an in-house working standard of known δ2H value. The
CH3Cl working standard was calibrated against IAEA standards VSMOW and
SLAP using TC/EA-IRMS (elemental analyzer – isotopic ratio mass
spectrometer, IsoLab, Max Planck Institute for Biogeochemistry, Jena,
Germany) resulting in a δ2H value of -140.1±1.0 ‰
vs. VSMOW (n=10, 1σ). The H3+ factor, determined daily
during this investigation (two different measurement periods), was in the
range of 5.75–6.16 (first period) and 8.90–9.21 (second period). The mean
precision based on replicate measurements (n=6) of the CH3Cl working
standard was 2.1 and 3.8 ‰ for the first and second measurement
periods, respectively. Samples were analyzed 3 times (n=3), and the
standard deviations (SD) of the measurements were in the range of 1.2 to
103.8 ‰. Lowest SD were observed for samples with lowest
δ2H values (∼-140 ‰) and highest mixing ratios and
higher SD for samples with highest δ2H values (∼+800 ‰) and lowest mixing ratios.
Please note that the above-described one-point calibration of the
δ2H data might be affected by an additional error (“scale
compression”) and particularly might affect the uncertainties of the very
positive δ2H values. Unfortunately CH3Cl working standards
with distinct isotopic signatures spanning the full range of measured
δ2H values (-150 to ∼+800 ‰) are not currently
available to eliminate or minimize such an error.
Methane
Stable hydrogen isotope ratios of CH4 were analyzed using an in-house
cryogenic pre-concentration unit coupled to a Hewlett Packard HP 6890
gas chromatograph (Agilent Technologies, Palo Alto, CA) and an isotope ratio
mass spectrometer (DeltaPlus XL, ThermoQuest Finnigan, Bremen, Germany). The
working gas was the same as that used for δ2H analysis of
CH3Cl (cf Sect. 2.3.1).
All δ2H values obtained from analysis of CH4 were corrected
using two CH4 working standards (isometric instruments, Victoria,
Canada) calibrated against IAEA and NIST reference substances (not specified
by the company). The calibrated δ2H values of the working standard
in ‰ vs. V-SMOW were -144±4 and -138±4 ‰.
The H3+ factor determined daily during the 2-week measurement period
was in the range 2.38–2.43. The daily average precision based on replicate
measurements of the CH4 working standard was 4.9 ‰ (n=7).
Samples were analyzed 3 times (n=3), and the SD of the measurements were
in the range of 1.4 to 40.9 ‰. Lowest SD were observed for samples
with lowest δ2H values (∼-180 ‰) and highest mixing
ratios and higher SD for samples with highest δ2H values (∼+300 ‰) and lowest mixing ratios.
Kinetic isotope effect, fractionation constant α and the isotope enrichment constant ε
In this study the isotope fractionation constant α and the isotope
enrichment constant ε are derived from the slope of the Rayleigh
plot according to Clark and Fritz (1997) and Elsner et al. (2005) and Eq. (2):
lnRtR0=δ2Ht+1δ2H0+1=ln(δ2H0+Δδ2H+1)(δ2H0+1),≅α-1⋅lnf=ε⋅lnf
where Rt and R0 are the 2H / 1H ratios in CH3Cl
or CH4 at the different time points and time zero, respectively, and f
is the remaining CH3Cl or CH4 fraction at the different time
points. Negative values of ε indicates that the remaining
CH3Cl or CH4 is enriched in the heavier isotope and corresponds to
a α < 1, meaning that over the entire experiment, the
heavier CH2DCl or CH3D reacts by this factor more slowly than the
lighter CH3Cl or CH4.
The kinetic isotope effect is then calculated as
KIE=1α.
To correct for ongoing analyte dilution the remaining fraction f has been
calculated as follows
f=cxt⋅ci0/(cx0⋅cit),
where cx0 and cxt are the mixing ratios of CH3Cl at time zero
and time t and ci0 and cit are the respective concentrations of the
internal standard PFH.
Results
Three experiments of CH3Cl degradation with OH were performed between
25 February 2014 and 3 February 2015. Under the experimental conditions (see Sect. 2 and Supplement) more than 70 % of the CH3Cl was degraded
within 6 to 10 h. The results from these experiments are shown in Fig. 2.
Both the trends of changes in δ2H values of CH3Cl as well as
the remaining fraction of CH3Cl observed in the three independent
experiments are in good agreement (Fig. 2a). The calculated ε
values for experiments 1 to 3 are -264±7, -219±6 and -308±8 ‰, respectively (Fig. 2b), with
a correlation coefficient R2 of the slope of the regression line of 0.99
for all three experiments.
Reaction of CH3Cl and OH. Three independent
experiments (triangles, dots and squares) were carried out using an initial
mixing ratio of 5 to 10 ppmv CH3Cl. More than 70 % of the CH3Cl
was degraded within 6 to 10 h. (a) Measured δ2H values (filled
circles, triangles and squares) of CH3Cl versus residual fraction (open
circles, triangles and squares) of CH3Cl (calculated from changes of
CH3Cl and PFH). Error bars of δ2H value of CH3Cl indicate
the standard deviation (SD) of the mean of three replicate measurements. Some
error bars lie within the symbol. (b) Rayleigh plot (Eq. 3). Error bars
were calculated by error propagation including uncertainties in δ2H
values of CH3Cl and the remaining fraction. Dashed lines represent
95 % confidence intervals of the linear regressions (bold lines).
The CH3Cl degradation with Cl experiment was conducted on
18 February 2014. Here, over 90 % of CH3Cl was degraded during reaction with
Cl radicals within 7 to 8 h (Fig. 3a). The calculated ε of
experiment 3 is -280±11 ‰ (Fig. 3b) with a correlation
coefficient of the slope of the regression line of 0.99. Due to limited
analytical resources it was not possible to repeat this experiment.
Reaction of CH3Cl and Cl. Initial mixing ratio of
CH3Cl was ∼ 10 ppmv. More than 90 % of the CH3Cl was
degraded within 7 to 8 h. (a) Measured δ2H values (filled circles)
of CH3Cl versus residual fraction (open diamonds) CH3Cl. Error bars
of δ2H values of CH3Cl indicate the standard deviation (SD) of
the mean of three replicate measurements. Some error bars lie within the
symbol. (b) Rayleigh plot (Eq. 3). Data are expressed as the mean ±
standard error of the mean, n=3. Error bars were calculated by error
propagation including uncertainties in δ2H values of CH3Cl.
Dashed lines represent 95 % confidence intervals of the linear regressions
(bold line).
The experiment to determine the isotope enrichment constant of the
degradation of CH4 by hydroxyl radicals was conducted on 2 February 2015. Over 80 % of CH4 was degraded during reaction with OH
radicals within 7 h (Fig. 4a). The calculated ε of
experiment 4 is -205±6 ‰ (Fig. 4b) with a correlation
coefficient of the slope of the regression line of 0.99.
Reaction of CH4 and OH. Initial mixing ratio of
CH4 was ∼ 6 ppmv. More than 80 % of the CH4 was degraded
within 7 h. (a) Measured δ2H values of CH4 versus residual
fraction of CH4. Error bars of δ2H values of CH4 indicate the standard deviation (SD) of the mean of three replicate
measurements. Some error bars lie within the symbol. (b) Rayleigh plot
(Eq. 3). Error bars were calculated by error propagation including
uncertainties in δ2H values of CH4 and the remaining
fraction. Dashed lines represent 95 % confidence intervals of the linear
regressions (bold line).
Discussion
Chloromethane reacts with both hydroxyl and chlorine radicals in the
atmosphere. The first degradation step of CH3Cl in both reactions is
the abstraction of a hydrogen atom to yield CH2Cl and H2O or HCl,
respectively (Spence et al., 1976; Khalil and Rasmussen, 1999). In both
reactions hydrogen is directly present in the reacting bond, and thus
influenced by the so-called primary isotope effect (Elsner et al., 2005).
Particularly for hydrogen these primary kinetic isotope effects are in
general large as they involve a large change in relative mass of the atoms
being abstracted. In the following we would like to discuss and compare our
results with (i) previous work conducted by Sellevåg et al. (2006),
(ii) with OH degradation experiments of CH4 and (iii) with the very
recent report of biochemical degradation of CH3Cl in soils and plants
(Jaeger et al., 2018a, b).
Although our experimental results show relatively large hydrogen isotope
fractionations with ε values of -264±45 (mean result from
three independent experiments ±SD) and -280±11 ‰ (mean result from three replicate analytical
measurements of the same sample ±SD) for reaction of CH3Cl with OH
and Cl radicals, respectively, they are smaller than the isotope
fractionations previously measured and theoretically calculated by
Sellevåg et al. (2006; Table 1). These researchers employed smog
chamber experiments at 298 K and used FTIR measurements to determine the
stable hydrogen isotope fractionation of CH3Cl and reported ε values of -410 and -420 ‰ for the reaction of CH3Cl with OH
and Cl radicals, respectively. They also performed theoretical calculations
of ε for the reactions of CH2DCl with OH and Cl radicals
and reported ε values in the range of -330 to -430 and -540 to
-590 ‰ , respectively (Table 1). Whilst we do not know the reasons
for the discrepancies in the experimental ε values observed here
and those reported by Sellevåg et al. (2006), we suggest that they may be
due to different measurement techniques employed in each of the studies. For
further discussion regarding differences of the experimental and analytical
design and protocols of the two studies we refer the reader to the
Supplement. However, we also conducted similar smog chamber experiments
for the degradation of CH4 with hydroxyl radicals (see Sect. 2
and Fig. 4) and calculated an ε value of -205±6 ‰ for the reaction of CH4 with OH radicals at a
temperature of 293 ± 1 K. In Table 1 we compare our results with those
from a number of previous studies (Saueressig et al., 2001; Sellevåg et
al., 2006; DeMore, 1993; Gierczak et al., 1997; Xiao et al., 1993), which
were conducted at temperatures ranging from 277 to 298 K (Table 1). The
ε values for the reaction of CH4 with OH radicals from all
studies ranged from -145 to -294 ‰ with a mean value of -229±44 ‰ and the most negative ε value of -294±18 ‰ reported by Sellevåg et al. (2006). The
ε value found in this study (-205±6 ‰) was in good agreement with previous experimentally reported values
conducted at similar temperatures. This finding gave us confidence that our
experimental design and the measurements made using GC-IRMS were reliable.
Compared to primary isotope effects, changes in bonding are much smaller in
the case of secondary isotope effects, where positions adjacent to the
reacting bond are only slightly affected by the proximity to the reaction
center (Elsner et al., 2005; Kirsch, 1977). It was suggested that for the
same element, secondary isotope effects are generally at least 1 order of
magnitude smaller than primary isotope effects (Kirsch, 1977; Westaway,
1987; Merrigan et al., 1999).
We therefore compared our results from chemical degradation experiments with
those from recently reported biochemical degradation experiments (Jaeger et
al., 2018a, b). So far, the only known pathway for
biochemical consumption of CH3Cl is corrinoid- and
tetrahydrofolate-dependent and is termed cmu (abbreviation for
chloro methane utilization). This
pathway was characterized in detail for the aerobic facultative
methylotrophic strain Methylobacterium extorquens CM4 (Vannelli et
al., 1999) and involves genes that were also detected in several other
chloromethane-degrading strains (Schafer et al., 2007; Nadalig et al., 2011, 2013). During degradation of CH3Cl the methyl group is
transferred to a corrinoid cofactor by the protein CmuA. In this case the
carbon–chlorine bond of CH3Cl is broken and thus since the hydrogen
atoms are adjacent to the reacting bond only a secondary isotope effect would
be expected. Indeed, the first ε values reported (Jaeger et al.,
2018a, b) for CH3Cl biodegradation by different soils
and plants (ferns) are in the range of -50±13 and -8±19 ‰, respectively, and thus show considerably
smaller kinetic isotope effects than for chemical degradation of CH3Cl
by OH and Cl radicals measured in either this study or reported by
Sellevåg et al. (2006).
Conclusions and future perspectives
We have performed experiments to measure the hydrogen isotope fractionation
of the remaining unreacted CH3Cl following its degradation by hydroxyl
and chlorine radicals in a 3.5 m3 Teflon smog chamber at 293 ± 1 K.
δ2H values of CH3Cl were measured using GC-IRMS. The
calculated isotope fractionations of CH3Cl for the reactions with
hydroxyl and with chlorine radicals were found to be smaller than either the
experimentally measured (by FTIR) or theoretical values reported by
Sellevåg et al. (2006). We also performed degradation experiments of
CH4 using the same smog chamber facilities yielding an isotope
enrichment constant for the reaction of CH4 with hydroxyl radicals of
-205±6 ‰ which is in good agreement with previously reported
results. Although stable hydrogen isotope measurements of CH3Cl sources
are still scarce, some recent studies have reported first data on
δ2H values of CH3Cl sources and ε values on sinks
(Greule et al., 2012; Jaeger et al., 2018, 2018b; Nadalig et
al., 2013, 2014).
We have summarized all available information regarding δ2H
values of environmental CH3Cl sources and their estimated fluxes in
Table 2. Furthermore, the strengths of known CH3Cl sinks and their
associated isotope enrichment constants are presented in Table 3. Eventually
Fig. 5 displays the global CH3Cl budget showing the known hydrogen
isotope signatures of sources and isotope enrichment constants associated
with sinks.
Schematic of major sources and sinks involved in the global
CH3Cl cycle (modified after Keppler et al., 2005) with known
(experimentally determined) corresponding δ2H values and isotope
enrichment constants, respectively. Red straight and dashed lines of arrows
indicate sources and sinks of CH3Cl, respectively. Size/thickness of
arrows indicates strength of fluxes in the environment. Question marks
indicate where currently no data exist. All values are given in per mil.
Known sources and strengths of tropospheric CH3Cl and
corresponding δ2H values.
SourcesSource (best estimate)aSource (full range)aMean δ2H valueUncertainty(Gg yr-1)(Gg yr-1)‰ vs. VSMOVδ2H value ± ‰Open field biomass burning355142 to 569-236b44Indoor biomass burning11356 to 169-236b44Tropical and subtropical plants20401430 to 2650-202c10Fungi145128 to 162?Salt marshes851.1 to 170?Coal combustion16229 to 295?Industrial chemical productiond363278 to 448-130e20Oceans700510 to 910?Othersf∼ 5827 to 86?Total sources3658 (4021)2601 to 5459
a Values for source (best estimate) and source (full range) were taken
from Carpenter and Reimann (2014), except for emissions associated with
chemical production by the industry which are from Li et al. (2017). The value
shown for total sources in parentheses includes chemical production by the
industry.
b Greule et al. 2012; please note that all values provided for
CH3Cl released from dried plants at elevated temperatures have been
corrected by -23 ‰ due to recalibration of the reference gas.
c Jaeger et al. (2018b).
d Li et al. (2017).
e taken from Greule et al. (2012), Nadalig et al. (2013) and Jaeger et
al. (2018a, 2018b); please note that values provided by Greule et
al. (2012) and Nadalig et al. (2013) for CH3Cl from sources of the
chemical industry have been corrected by -23 ‰ due to recalibration
of the reference gas.
f including mangroves, wetlands, rice paddies and shrublands.
? denotes that no value has been provided.
Known sinks of tropospheric CH3Cl and the mean
isotope enrichment constant ε reported for each.
SinksSink (best estimate)aSink (full range)aIsotope enrichmentUncertainty(Gg yr-1)(Gg yr-1)constant ε / ‰ε± ‰Reaction with OH in troposphere28322470 to 3420-264b45b-410c50cLoss to stratosphere146?0d?Reaction with Cl in marine boundary layer370d180 to 550d-280b11b-420c40cMicrobial degradation in soil1058664 to 1482-50e13eLoss in ocean370296 to 4450f10fMicrobial degradation in plantsg??-8g19gTotal sinks4406 (4776)
a Values for sink strength (best estimate and full range) were taken
from Carpenter and Reimann (2014), except for the value of the reaction with
Cl radicals in marine boundary layer and for total sinks shown in parentheses
which includes the potential sink strength by Cl radicals in marine boundary
layer (Montzka and Fraser, 2003).
b this study.
c Sellevåg et al. (2006).
d Thompson et al. (2002) and discussion in this manuscript.
e Jaeger et al. (2018a).
f Nadalig et al. (2014).
g Jaeger et al. (2018b).
? denotes that no value has been provided.
Our results suggest that stable hydrogen isotope measurements of both
sources and sinks of CH3Cl and particularly the observed large kinetic
isotope effect of the atmospheric CH3Cl sinks might strongly assist
with the refinement of current models of the global atmospheric CH3Cl
budget. In contrast to the large hydrogen fractionation of CH3Cl by
chemical degradation of OH and Cl radicals, the isotope fractionation of
CH3Cl biodegradation is in the range of an order of magnitude lower.
This therefore has the possibility of improving our understanding of the
global CH3Cl budget once the δ2H value of atmospheric
CH3Cl has been measured. The stable hydrogen isotopic composition of
tropospheric CH3Cl depends on the isotopic source signatures and the
kinetic isotope effects of the sinks, primarily the reaction with OH and
consumption by soils and potentially plants.
Several attempts at modeling the global CH3Cl budget using stable
carbon isotope ratios have already been made (Harper et al., 2001, 2003; Thompson et al., 2002; Keppler et al., 2005; Saito and Yokouchi,
2008) but there are still major uncertainties regarding source and sink
strengths as well as the respective stable isotope signatures. Therefore, we
now suggest combining our knowledge of stable carbon and hydrogen isotopes
of CH3Cl in the environment. Such a two-dimensional (2-D) stable isotope
approach of hydrogen and carbon can be used to better understand the
processes of CH3Cl biodegradation and formation. Furthermore, when this
approach is combined with CH3Cl flux estimates it could help to better
constrain the strength of CH3Cl sinks and sources within the global
CH3Cl budget (Nadalig et al., 2014; Jaeger et al., 2018b).
We note that currently no data are available for the δ2H value of atmospheric CH3Cl. Although it will be a massive
analytical challenge to obtain this value, we strongly suspect that it
would likely lead to a better-refined isotopic mass balance for atmospheric
CH3Cl and thus to a better understanding of the global CH3Cl
budget.
The data used in this publication are available to the
community and can be accessed by request to the corresponding author.
The supplement related to this article is available online at: https://doi.org/10.5194/acp-18-6625-2018-supplement.
The authors declare that they have no conflict of
interest.
Acknowledgements
This study was supported by DFG (KE 884/8-1; KE 884/8-2, KE 884/10-1) and by
the DFG research unit 763 “Natural Halogenation Processes in the Environment
– Atmosphere and Soil” (KE 884/7-1, SCHO 286/7-2, ZE 792/5-2). We further
acknowledge the German Federal Ministry of Education and Research (BMBF) for
funding within SOPRAN “Surface Ocean Processes in the Anthropocene” (grants
03F0611E and 03F0662E). We thank John Hamilton and Carl Brenninkmeijer for
comments on an earlier version of the manuscript and Daniela Polag for
statistical evaluation of the data.
The article processing charges for this open-access publication were covered by the Max Planck Society.
Edited by:
Sergey A. Nizkorodov
Reviewed by: Matthew Johnson and two anonymous referees
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