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Title
Biogeochemical aspects of atmospheric methane
Permalink
https://escholarship.org/uc/item/3xq3t703
Journal
Global Biogeochemical Cycles, 2(4)
ISSN
0886-6236
Authors
Cicerone, RJ
Oremland, RS
Publication Date
1988
DOI
10.1029/GB002i004p00299
Copyright Information
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GLOBAL BIOGEOCHEMICAL CYCLES, VOL. 2, NO. 4, PAGES 299-327, DECEMBER 1988
BIOGEOCHEMICAL ASPECTS OF ATMOSPHERIC
METHANE
R. J. Cicerone
National Center for Atmospheric Research,
Boulder, Colorado
R. S. Oreroland
U.S. Geological Survey, Menlo Park,
California
Abstract. Methane is the most abundant organic chemi-
cal in Earth's atmosphere, and its concentration is increas-
ing with time, as a variety of independent measurements
have shown. Photochemical reactions oxidize methane in
the atmosphere; through these reactions, methane exerts
strong influence over the chemistry of the troposphere and
the stratosphere and many species including ozone, hy-
droxyl radicals, and carbon monoxide. Also, through its
infrared absorption spectrum, methane is an important
greenhouse gas in the climate system. We describe and
enumerate key roles and reactions. Then we focus on two
kinds of methane production: microbial and thermogenic.
Microbial methanogenesis is described, and key organisms
and substrates are identified along with their properties
and habitats. Microbial methane oxidation limits the re-
lease of methane from certain methanogenic areas. Both
aerobic and anaerobic oxidation are described here along
with methods to measure rates of methane production
and oxidation experimentally. Indicators of the origin of
methane, including C and H isotopes, are reviewed. We
identify and evaluate several constraints on the budget
of atmospheric methane, its sources, sinks and residence
time. From these constraints and other data on sources
and sinks we construct a list of sources and sinks, identi-
ties, and sizes. The quasi-steady state (defined in the text)
annual source {or sink} totals about 310{+60} x 10 TM tool
{500{+95} x 10 TM g}, but there are many remaining uncer-
tainties in source and sink sizes and several types of data
that could lead to stronger constraints and revised esti-
Copyright 1988
by the American Geophysical Union.
Paper number 88GB03540.
0886-6236 / 88 / 88 C B-03540510.00
mates in the future. It is particularly difficult to identify
enough sources of radiocarbon-free methane.
1. INTRODUCTION
Methane (CH4) is the most abundant organic gas in
Earth's atmosphere. Much of the history of the detection
of atmospheric methane and of the earliest systematic
measurement data has been reviewed by Ehhalt [1974] and
Wofsy [1976] and will not be repeated here. Since the
early 1970s a variety of roles of methane in atmospheric
chemistry and climate have been identified. For example,
methane affects tropospheric ozone, hydroxyl radicals and
carbon monoxide concentrations, stratospheric chlorine
and ozone chemistry and, through its infrared properties,
Earth's energy balance (see section 2).
Recent evidence that atmospheric methane concentra-
tions are increasing globally has made it necessary and
more urgent to understand natural processes, both bio-
logical and physical, which control methane and to iden-
tify the human activities that are involved. This evi-
dence is now overwhelming. Methane increases have been
demonstrated at many different locations and with inde-
pendent measurement techniques, including flame ioniza-
tion gas chromatography [Rasmussen and Khalil, 1981;
Fraser et al., 1984; Blake and Rowland, 1988] and infrared
absorption analysis [Rinsland et al., 1985]. Figure 1, from
Blake and Rowland [1988] presents globally averaged data
gathered between 1978 and late 1987; an annual rate of
increase of 0.016 ppm/yr, or about 1% per year, is appar-
ent. Globally averaged methane mole fractions are almost
1.70 ppm. Rinsland et al. [1985], through reanalysis of
some solar absorption spectra taken at Jungfraujoch in
1951, deduced an atmospheric methane concentration of
1.14(5:0.08) ppm for April 1951. Similar instrumentation
and analysis techniques at Kitt Peak in February 1981
yielded 1.58(5:0.09) ppm.
•00 Cicerone and Oreroland: Atmospheric Methane
1.7 -
CH4
I I I I I I I I I I I
978 79 80 81 82 83 84 85 86 87 88
YEAR
Fig. 1. Globally averaged tropospheric methane mole
fractions (parts per million by volume, ppmv I measured
between January 1978 and September 1987 and a best
fit line through the data [from Blake and Rowland, 1988,
reproduced with permission of the American Association
for the Advancement of Science].
Whatever the causes of these recent methane increases,
they are very rapid in comparison to geological time scales.
Also, there is clear evidence from analysis of gases trapped
in dated ice cores that atmospheric methane has more
than doubled in concentration in the last 200 years [Craig
and Chou, 1982; Rasmussen and Khalil, 1984; Pearman
et al., 1986] and that its concentration stayed between 0.6
and 0.8 ppm over the last 3,000 years (Figure 2 I. The
data in Figure 2 are from near both poles, Greenland
and Antarctica. Similar results have been obtained by
Stauffer et al. [1985], who also tested for differences
due to gas extraction methods. Stauffer et al. also
succeeded in analyzing a very recent ice core (1955 •- 10
yearsl; in that sample the CH4 concentration was about
1.4 ppm, consistent with other evidence that methane has
been increasing in recent decades. While there have been
questions about the degree to which ice core air represents
the atmosphere [Craig and Chou, 1982], we believe that
these questions have been resolved, and that at least for
CH4, we may believe the results. On this point, the
•3C isotope data of Craig eta!. [1988] are particularly
compelling (see also section
The ice core methane record now extends back to nearly
160,000 years ago, as is illustrated in Figure 3. All
methane data points from over 100,000 years ago are from
the Raynaud et al. [1988] analysis of the ¾ostok ice core
(Antarctica). The solid line in Figure 3 is a smoothed ver-
sion of the temperature record that Jouzel et al. [1987]
constructed from deuterium amounts in the ¾ostok core.
One point (circled in Figure 3}, at 27,200 years (650 ppb)
is from Craig and Chou [1982]. All remaining points are
from Stauffer et al. [1988], who found that during the
last glaciation methane concentrations were about 350 ppb
as opposed to about 650 ppb afterwards. The Craig and
Chou point at 27,200 years is higher than the Stauffer et al.
values for the same glacial period. The ¾ostok core data
from Raynaud et al. show that methane increased from
about 320 ppb to 620 ppb between the end of the next
to last glaciation and the subsequent interglacial period,
about 160,000 to 120,000 years ago. These changes sug-
gest that exposing and warming ice-covered soils produced
more methane as glaciers retreated.
In any case, contemporary methane amounts and their
rate of increase are unprecedented, at least during the past
160,000 years. Human activities are clearly involved in
causes for the increase from 650 ppb to 1700 ppb. Accord-
ingly, there is added motivation to understand the factors
and practices that produce methane and release it to the
atmosphere, the atmospheric processes that methane af-
fects, and the atmospheric processes that limit the rise of
methane concentrations.
In this paper we discuss the roles of methane in atmo-
spheric chemistry and climate (section 21, the microbiol-
ogy of methane formation and microbial oxidation in soils
and waters (section 3), and information from carbon and
hydrogen isotopes on the origin of environmental methane
(section 4). In section 5 we examine data on atmospheric
methane amounts, turnover rates, and isotopes to derive
objective constraints on the annual budget (sources and
sinks} of atmospheric methane. Methane fluxes are stated
in units of 10 TM g CH4 or Tg CH4/yr (1 Tg-10 •2 g).
2. METHANE IN ATMOSPHERIC CHEMISTRY AND
CLIMATE
Atmospheric Chemistry
In this section we discuss the important chemical roles
of atmospheric methane and then how methane affects the
radiative energy balance and hence the climate of Earth.
The chemical reactions that destructively oxidize atmo-
spheric methane affect the chemical state of the atmo-
sphere through the products of the reactions and through
consumption of the reactant species. The most important
reactant that destroys methane is the gas phase hydroxyl
radical, OH, a key radical in atmospheric photochemistry
(not to be confused with aqueous OH-). Methane oxi-
dation produces CO, CO2, H20, H2, and CH20, and it
consumes OH. These reaction pathways (see below) af-
fect tropospheric and stratospheric ozone amounts, and
they produce important quantities of H20 in the strato-
sphere. Also, stratospheric CH4 reacts with C1 atoms,
forming HC1, a reservoir species for C1 atoms. Finally, a
portion of the flux of hydrogen carried upward into the
stratosphere in CH4 escapes to space, mostly as H atoms.
Escape of H to space represents a source of atmospheric
oxygen just like burial of organic carbon in deep sea sedi-
ments [Broecker, 1970; Van Valen, 1971].
To be clear as to how these effects arise, consider the
key elementary reactions that occur in the atmosphere.
First is the production of OH, which stems from ozone,
ultraviolet (UV) light, and water vapor:
(R1) 03 + h•/ ' , O(•D) + 02 •<315nm
Most of the electronically excited oxygen atoms, O{•D},
that are produced in {R1} are quenched in collisions with
N2 and 02 as in (R2)
Cicerone and Oremland: Atmospheric Methane •01
$ 5 $.0
1750
Log (oge yr$ B.P)
2O
1500
-I- 1250
tOO0
750
•oo I 1
3000 2000 tOO0 700 500 300 200 tO0 0
Time (yr$ B.P.)
Fig. 2. Concentrations of methane in air extracted from dated ice cores representing 100 to 3000 years
before present (B.P.) from Rasmussen and Khalil [1984]. Solid circles are data from Greenland ice, and
triangles represent Antarctic ice cores. The solid line is a smoothed fit to the data and near-contemporary
air data are represented by diamonds.
(R2) O(•D) + N2
Ozone is reformed in (R3)'
(R3)
0 + 02 + M -• 03 + M
where M = N2, 02, or any third body with which collisions
stabilize the O3 product. But some of the O(1D), about
1%, reacts with water vapor to produce hydroxyl radicals:
(R4) O(•D) + H20 , 2OH
The biggest single sink of atmospheric methane is its
reaction with OH in the troposphere.
(RS) CH4 + OH ---• H20 + CH3
Perhaps 85% of the methane that is emitted into the
atmosphere is destroyed by (R5) in the troposphere. In
the stratosphere, almost all of the remaining methane is
destroyed by OH, by C1 atoms, and by O(•D) atoms. A
small fraction of methane goes through the stratosphere
to the mesosphere where an additional sink, very short
wavelength UV light (principally Lyman alpha radiation,
121.6 nm) destroys methane photolytically.
Complete oxidation of methane yields CO2 and H20.
Schematically, this can be represented by the combustion
of methane:
(R6) CH4 + 202 • CO2 + 2H20
While (R6) seems simple and clear enough, it does not
describe the mechanism through which the atmosphere
oxidizes methane. In the atmosphere the process is ini-
tiated by OH radicals, not by 02, and it requires light as
discussed below. A pioneering study of this process was
that of Levy [1971] (see also Levy [1973]). However, we
now know that the mechanism of methane oxidation and
the products that are formed are very different in the two
cases of high concentrations of nitrogen oxides (NOx) and
low NOx concentrations (defined below). For example,
the methane oxidation chain can either produce or con-
sume ozone. Figure 4 summarizes the principal reactions
4_d • [ [ I [ • [ [ I [ [ [ ] I 700
ß
2 ;%0 • o-Greenland
0 • ß ,,-Antarctica
• -4 • 500
• -6
-8 400
-IO - %•
-
-12 • • • • I • ' • • • * • • ' • • 3OO
O 50000 IOOOOO 15OOOO
AGE (years)
Fig. 3. Methane concentrations (ppb by volume) from
dated ice cores plotted as age of a• before present. Sold
circles are data from Greenland cores, and triangles rep-
resent Antarctic cores. Data •e from Raynaud et al.
[1988] and Stauffer et al. [1988]. The circled point from
27,200 BP is from Craig and Chou [1982]. The sold l•e
• a smoothed version of the temperature record deduced
by Jouzel et al. [1987]. Lowest concentrations are from
glaciated epochs, and highest values •e from •tergl•ial
times.
30• Cicerone and Oremland: Atmospheric Methane
co 2
O•
OH, O 2
OH
OH I0( I D)]
C.••i L I
ci j
CH302
H2+CO NO
h•
OH'
Fig. 4. Principal reactions and species in atmospheric methane oxidation (adapted from Ravishankara
[1988]) for the case of adequate to high NO and NO2 concentrations, defined in text. Dashed box indicates
species whose reactions were not included in Levy's early work; otherwise, the scheme is essentially that of
Levy [1971, 1973].
and species of methane oxidation for the case of high NOx
concentrations. With low NOx concentrations there are
substantial deviations from the pathways of Figure 4 (see
below and Crutzen [1987]). Also, the actual reaction chain
can be cut short in the atmosphere when products such as
CH20 are removed, for example, by rainfall before further
oxidation.
In high-NO• air columns, as in polluted or moderately
dirty tropospheric air and all of the stratosphere, methane
oxidation produces ozone and hydrogen oxides (HO -[-
HO2). The process begins with (RS) and proceeds first
to production of formaldehyde.
(RS) CH4 + OH , H20 + CH3
(R7) CH3 + O2 + M , CH302 + M
(R8) CH302 + NO , CH30 + NO2
(R9) CH30 + O2 , CH20 + HO2
(RI0) HO2 + NO , NO2 + OH
(Rll) 2[NO 2 + h• , NO + O] A < 400 nm
Net reaction
CH4 -•- 402 + 2hu • CH20 + H20 + 203
Let us designate this reaction sequence as RS1. In RS1,
NO and NO2 are catalysts. Formaldehyde is oxidized to
CO through three reaction pathways.
(R12a)
(R12b)
(R13)
(R14)
($1o
(R11)
CH20 + hu • CO + H2 A < 360 nm
CH20 + hy , H + HCO A < 360 nm
H+ O2 +M • HO2+ M
HCO + 02 • CO + HO2
2[HO 2 + NO ---, OH + NO2]
2[NO 2 + h• ---, NO + O]
2[0 + 02 + M -• 03 + M]
Net reaction
CH20 + 402 + 2hu • CO + 203 + 2OH
RS2 is defined as the reaction sequence (R12b), (R13),
(R•4), (R•0'), (R•), •na (R3'). The sequence RS2 pro-
duces two molecules of ozone and two OH radicals per
CH20 molecule that is consumed. The third pathway from
CH20 to CO is
(R15) CH20 + OH • HCO + H20
(R14) HCO + O2 , CO + HO2
(R10) HO2 + NO , OH + NO2
(Rl1') NO2 + hv , O + NO
O+O2+M ,O3+M
Net reaction
CH20 + 202 + hv -• CO + H20 + 03
The final step, still assumed to occur in the presence of
high NO• concentrations, is the oxidation of CO to CO2.
(R16)
(R13)
(R10)
(Rl1')
Net reaction
CO + OH ----, CO2 + H
H + O2 + M • HO2 + M
HO2 + NO ----, OH + NO2
NO2 + hv --• 0 + NO
O + O2 + M -• O3 + M
C0+202+hv ,C02+03
Thus the complete oxidation of CH4 in the presence of ad-
equate NO• produces 03 and depending on the relative
fractions of CH20 oxidized by (R12a), RS2, and RS3, can
produce OH radicals. Crutzen [1987] has calculated that
the averaged relative fractions of (R12a), RS2, and RS3 are
about 50-60%, 20-25%, and 20-30%, respectively, in the
troposphere. In this case, methane oxidation to CO2 and
H20 produces 3.7 03 molecules and 0.5 OH radicals per
methane molecule destroyed. Note that these numerical
relationships are stoichiometric between methane destruc-
tion and production of 03 and OH but are not ratios of
number density changes because other processes partially
control 03 and HOz concentrations, for example, surface
deposition and inflows of 03 from the stratosphere, and
there are also photochemical feedbacks in the system. For
this pathway (RS1 then (R12a) or RS2 or RS3, then RS4)
to proceed there must be enough NO present for HO2 to