Introduction
Volcanic activity is one of the major natural forcers of the Earth's
climate, as volcanic emissions alter the chemical composition and radiative
properties of the atmosphere, at local, regional and even global scales
. Beyond their environmental impacts,
sulfuric acid aerosols have adverse effects on human health since they are
linked to cardiovascular and respiratory diseases
. Moreover, sulfate aerosols
can lead to acid rain causing damage to vegetation and to urban
infrastructures. Over the last decades, our understanding of volcanic
emissions in the atmosphere has greatly improved, thanks to satellite and
field measurements, and to more sophisticated physical-chemical models
.
The main gases emitted to the atmosphere by volcanic activity are
H2O, CO2, SO2, H2S and halogen
species, such as HCl, HBr and HF
. In addition, measurements at
crater rims of volcanoes also suggest direct emissions of small amounts of
sulfate aerosols .
Among all the compounds emitted, volcanic sulfur gases and in particular
SO2, are considered to be the most effective in affecting climate.
Climatic perturbations from volcanic emissions are principally caused by
conversion of sulfur gases into sulfate aerosols, which can then interact
with solar and terrestrial radiation via scattering and absorption
. Once injected into the troposphere, volcanic SO2
is converted in few days typically to H2SO4 by a range of gas-phase
and liquid-phase reactions taking place in volcanic plumes and clouds
. In the atmosphere, depending on the
oxidation pathway, H2SO4 is produced either in the gas phase or
liquid phase. When generated in the gas-phase, volcanic H2SO4
condenses very rapidly onto pre-existing particles, or it may even form very
small sulfate particles by nucleation. In the boundary layer, sulfate
aerosols have a residence time much shorter than a week because of the fast
wet and dry depositions. However, at higher altitudes, such as in the free
troposphere, removal is much slower; consequently, volcanic sulfate aerosols
can have a much longer residence time of up to a few weeks
. In addition, the
residence time of volcanic aerosols in the stratosphere can reach lifetimes
of about a year .
Nowadays, anthropogenic SO2 emissions outweigh those from natural
sources . Volcanic quiescent degassing and eruptions is an
important natural source of SO2, notably to the free troposphere
. Volcanic emissions release about
10–13 Tg y-1 of SO2 to the atmosphere and
contribute to up to 10 % of total sulfur emissions to the atmosphere
. Remarkably, volcanic emissions also have a bigger
impact on the tropospheric aerosol burden than other sulfur sources
because volcanoes tend to emit SO2 at higher
altitudes than most other surface sulfur emissions, where the lifetime is
longer.
Most of the tropospheric sulfate is generated in the liquid phase
via oxidation of aqueous SO2 by dissolved
oxidants of the atmosphere, such as H2O2, O3, O2
catalysed by transition metal ions (Fe(III) and Mn(II)) and,
possibly, HOBr and HOCl
.
Note that the importance of the halogen oxidation pathway remains unclear. A
significant amount of tropospheric H2SO4 is formed in the gas phase
via the termolecular reaction between SO2 and hydroxyl radicals
(OH) . In presence of liquid water and for typical
pH values of atmospheric water droplets (3.0 < pH < 5.6), SO2
is quickly oxidized by dissolved H2O2, and the two species rarely
coexist in liquid phases
. At acidic pH
values, synergism among transition metal ions (TMI) enhances the rate of
SO2 oxidation by dissolved O2 ,
which can thus compete with the other SO2 oxidation channels.
Particular attention has been paid recently to this heterogeneous oxidation
pathway, since its contribution could have been underestimated in previous
budget assessments of sulfate production in the troposphere
. During eruptive events, volcanoes
emit large quantities of ash and coarse material rich in iron-minerals
(mainly glass, and in lesser extents magnetite and hematite), which can
easily dissolve in water because of the high acidity reached in volcanic
cloud droplets and aerosols .
As a consequence, the O2/TMI heterogeneous oxidation reaction may be
more significant than previously thought.
Quantifying the importance of the different SO2 oxidation pathways is
challenging. It requires the quantification of, among other things, the rates
of the different oxidation processes. Conventional methods rely mostly on
models that are evaluated and constrained with atmospheric concentration
measurements of oxidants, because there is no direct means of measuring
chemical fluxes associated with individual reactions .
Simultaneous measurements of SO2 oxidants in both the gas- and liquid
phases in the atmosphere, let alone specifically in a volcanic plume, would
be experimentally challenging. Alternative approaches need to be considered
to reduce the uncertainty in the relative contributions from the different
oxidation pathways. Isotopic approaches can provide such constraints
. Isotopic ratios, indeed, provide
direct insights into the nature and importance of individual oxidation fluxes
.
Thanks to peculiar distribution of isotopes among its three oxygen atoms,
ozone and its chemistry provides a useful tool of investigation for
atmospheric processes using isotopic signatures. Ozone bears a very
significant non-mass dependent (also called mass-independent) isotopic
fractionation, which is due to its formation mechanism
. Since oxygen atoms in
tropospheric oxygen-bearing species sometimes originate directly or
indirectly from ozone via multiple photochemical reactions, a variety of
atmospheric species carry anomalous isotopic mass-independent fractionations
(MIFs ). For oxygen-bearing species, the anomalous oxygen
MIF (Δ17O, O-MIF) is calculated with respect to a reference
standard:
Δ17O=δ17O-0.52×δ18O.
Where δ17 O and δ18 O represent deviations
to the reference standard isotopic ratios (Rstd):
Rx=xO16Ox=17;18,
and
δxO=RxRstd-1.
Ozone is a key chemical reactive species of the troposphere. Its isotopic
anomaly is intrinsically generated (through photolysis and recombination
reactions) instead of being inherited by isotopes transfer like for most
atmospheric species . Other oxygen-bearing species in the
atmosphere can gain excess-17O by transfer of this ozone anomaly
via reactions with ozone itself, reactions with species that have already
inherited the ozone anomaly or via anomalous kinetic isotopic effect
. As a consequence, transfer of
oxygen MIF among atmospheric species is process-specific and can be used as a
signature to trace the chemistry of species as they react with specific
oxidants. Once the isotopic anomalies of the oxidants are characterized, the
resulting Δ17O of an end-oxidation product is simply a
linear combination of the isotopic signatures of all the oxidation channels
weighted by their respective contributions, to the total production of the
end-oxidation products. During the last decade, there has been an increasing
number of studies that have used O-MIF oxygen anomalies in oxidation products
to constrain oxidation channels, often coupling isotopic measurements and
photochemical isotopic modelling
.
The isotopic signature in sulfates generated in the troposphere, the
so-called secondary sulfate (by opposition to sulfate directly emitted in
the atmosphere, the so-called primary sulfate) reflects the competition
within different oxidation channels. In the liquid phase, sulfate oxygen MIF
is produced during sulfur oxidation by transfer of isotopic anomalies from
ozone and H2O2, whereas sulfate with O-MIF very close to
0 ‰ is produced in the liquid phase via O2/TMI oxidation
(i.e. -0.08 ‰). Mass-dependent (MIF anomaly = 0 ‰)
sulfate is generally produced via OH oxidation in the gas-phase
.
Oxygen isotopic composition of volcanic sulfates from different
tropospheric emissions of the present geological era.
Volcano & date of eruption
Sample distance (km)
Source
Δ17O (‰)
Reference
Popocatépetl (Mexico), 2008
25
ash
0.35
Spurr (Alaska, USA), 1992
265
ash
-0.14
Fuego (Guatemala), 1974
57
ash
-0.04
Negro Cerro (Nicaragua), 1947
12
ash
-0.06
Parícutin (Mexico), 1948
5
ash
0.13
Mt. St. Helens (USA), 1980
400
ash
0.02
Gjálp (Iceland), 1998
<30
ash
-0.07
Pinatubo (Philippines), 1991
<50
ash
-0.04
Pinatubo (Philippines), 1991
<50
ash
0.19
Spurr (USA), 1953
n.a.
ash
0.06
Vesuvius (Italy), 1872
n.a.
ash
-0.07
Popocatépetl (Mexico), 1997
n.a.
ash
-0.08
Spurr (USA), 1992
n.a.
ash
0.06
Fuego (Guatemala), 1974
55
ash
-0.03
Pinatubo (Philippines), 1991
n.a.
anhydrite from pumice
-0.01
Santorini (Greece), Minoan age
n.a.
pumice + ash
0.09
Masaya (Nicaragua), 2003
0
aerosols
0.1
Masaya (Nicaragua), 2003
0
aerosols
0.2
* Refer to for a more extensive description
regarding oxygen isotopic anomalies measured in tropospheric volcanic
sulfate of present and past geological eras.
Most present-day tropospheric sulfates have O-MIF anomalies
(Δ17O), typically of the order of 1 ‰
. However, there is some variability. For instance,
O-MIF of sulfate aerosols generated in marine environments are higher
compared to isotopic anomalies found in continental sulfates
. Very significant Δ17O have also been
found in volcanic sulfates collected from ash deposits dating back to the
Miocene and the Oligocene, whose values reach 3.5 ‰–5.8 ‰. These
peculiar isotopic anomalies have been linked to a different oxidative state
of the atmosphere at that time . Tropospheric volcanic
sulfates of the present era distinguish themselves from other tropospheric
sulfates by having a Δ17O often close to 0 (within the
measurement error of about 0.1 ‰). This feature is found all over
the world in sulfates collected from volcanic ashes of small and medium-size
tropospheric explosive eruptions, independent of location or geology of
ash deposits (see
Table ). Notably, this is often the case for
volcanic sulfate extracted from ash deposits which are found very far from
volcanoes, where secondary sulfate is expected to dominate. The only
exception is volcanic sulfates in ice cores originating from very large
volcanic eruptions. This sulfate had formed and transited through the
stratosphere .
The question is why tropospheric volcanic sulfate from volcanic ash deposits
does not appear to carry some isotopic O-MIF as is the case for other types of
tropospheric sulfates. One might expect part of sulfate produced by
tropospheric oxidation of volcanic SO2 to carry some MIF isotopic
anomaly because the dominant SO2 oxidants in the troposphere are
thought to be species carrying O-MIFs (O3 and H2O2) with
some contribution from O2/TMI . An important
difference between volcanic sulfur and most other sources of sulfur is that
it is often emitted within dense volcanic plumes whose chemical compositions
are radically different from the background air. The purpose of the present
box-modelling study is to explore in detail the oxidation and fate of
volcanic sulfur in dense volcanic clouds and the resulting isotopic MIF
signature in volcanic sulfate. The objective is to see to what extent the
chemical environment of dense volcanic plumes may affect sulfur dynamics and
pathways of oxidation and, hence, sulfate isotopic composition. The focus
here is on volcanic clouds that are rich in sulfur but poor in halogens,
such in the case of intra-plate and rifting plate volcanoes (e.g. Nyarogongo
in Congo, Erta'ale in Ethiopia, Kīlauea in Hawai'i)
.Volcanic eruptions with remarkable low
halogens to sulfur emissions are the Holuhraun (Bárðarbunga) eruption
of 2012–2014 in Iceland , and the
Kīlauea eruption of 2008 in Hawai'i . In particular,
HCl/SO2 ratios of the order of 10-2 have been observed for the
Kīlauea eruption of 2008 (i.e. HCl ≈ 10–50 ppbv).
The second section of this work describes the photochemical model, including
its sulfur heterogeneous chemistry scheme and the associated oxygen isotopic
scheme. The mass balance equations used to evaluate the transfer of MIF
oxygen anomaly from ozone to volcanic sulfate via different oxidation
pathways are also presented. The third section is devoted to the study of
individual and combined oxidation pathways and the resulting isotopic
signatures in numerical experiments for this work standard volcanic plume
conditions. The fourth section covers sensitivity model studies,
investigating how different parameters in volcanic plumes affect the final
isotopic anomaly in sulfate. Dominant oxidation pathways are identified and
the ability of the model to reproduce observed isotopic signatures of
volcanic sulfate is assessed.
Modelling approach
The photochemical box model used during simulation is the Cambridge
Tropospheric Trajectory model of Chemistry and Transport (CiTTyCAT), a
photochemical box model developed to simulate tropospheric chemistry
. It describes the standard
gas-phase photochemistry accounting for kinetics of tropospheric species
(bimolecular, termolecular and photodissociation reactions), and deposition
of gases and particles. Photolysis reaction rates are evaluated using the
Fast-J code . Kinetic data are taken from JPL's data sheets
. CiTTyCAT had already been used with success to constrain
seasonal pathways of reactive nitrogen species in the troposphere, through
the implementation of its chemical scheme with an isotopic transfer scheme
accounting for Δ17O production in nitrates
. We have extended the capabilities of the model by
including parameterizations of the transfer of soluble species between liquid
and gas phases, of SO2 heterogeneous chemistry, of pH in the liquid
phase and of MIF of oxygen atoms in sulfates.
General continuity equations
The model resolves differential coupled mass balance equations (continuity
equations) describing the time evolution of species concentrations in the
troposphere. For given initial (e.g. initial concentrations of species) and
environmental conditions (e.g. pressure, temperature), mass balance equations
are solved for each species, accounting for production and loss as follows:
dCidt=Pi-Li,
where Ci is the concentration of species i, Pi the sum of physical
and chemical production rates for i, and Li the sum of the physical and
chemical loss rates of i.
Production and loss terms are calculated using chemical reaction kinetics,
where time evolution of concentrations of chemical species depends on the
relevant rate constants (ki) and on concentrations of reactants. They also
include liquid–gas transfer and deposition. In addition, mixing of air
between the volcanic sulfur cloud and the outside background air is also
accounted for. It is parameterized by a simple linear relaxation scheme
resulting in an exponential decay of plume concentrations towards background
concentrations .
dCidtmixing=Kmix×(Ci-Ci(bck.)),
where Kmix is a first-order mixing rate coefficient representing
all the processes mixing volcanic air with the background atmosphere and
C(i,bck.) is the concentration of species i in the background
air. Kmix is set to 0.1 d-1, a value typical of low mixing in
the free troposphere and corresponding to a characteristic mixing timescale
of 10 days .
Liquid–gas mass transfer
Concentrations of relevant soluble species are calculated taking into account
its partition between the gas and liquid phases. The transfer in both
directions (evaporation, condensation) is dynamically computed. At each time
step, rates of transfer are defined as follows:
d[C(aq.)]dt=Ji×C(i)-C(i,s),
where C(i) is gaseous concentration of species i far from liquid
droplets, C(i,s) is the gaseous concentration of species i at the
surface of droplets (which is assumed to be the equilibrium saturation vapour
of i over the liquid) and Ji is the coefficient of condensation (from
gas phase to liquid droplet) for species i, which is calculated using the
Dahneke's expression to cover mass-transfer from the
continuum to the kinetic regime (see pg. 502 of Seinfeld and Pandis, 2016).
Throughout all the model simulations, droplets are assumed to be very large,
with a radius of 5.0 µm. The sensitivity of the results to the
assumed amount of liquid phase is explored varying the concentration of water
droplets (and hence the liquid water content) instead of varying the size of
droplets. It is also possible that emitted water condenses onto ash
particles. Our treatment does not discriminate between liquid droplets and
liquid phases at the surface of solid particles.
Diagram of the new sulfur scheme implemented in CiTTyCAT.
![]()
Sulfur aqueous equilibria.
Equilibrium
K (M-1),
k298(forward) (M-1s-1),
Ea/R(K),
k298(backward) (M-2s-1),
SO2(aq.) + H2O ⇌ HSO3- + H3O+
3.13×10-4
6.27×104
-1940
2×108a,c
HSO3- + H2O ⇌ SO32- + H3O+
6.22×10-8
3110
-1960
5×1010a,c
H2SO4 + H2O → HSO4- + H3O+
∞
HSO4- + H2O ⇌ SO42- + H3O+
1.02×10-2
1.02×109
-2700
1×1011b,c
a ; b ; c .
Gaseous and heterogeneous sulfur chemistry
The model already describes the SO2 gas-phase chemistry. Since
SO2 is a mildly soluble species undergoing acid-base equilibrium in
the liquid phase, we have added the gas-liquid transfer and the chemical
reactions and equilibrium associated with its presence in the liquid phase
(see Table ). The extent of SO2
dissolution into water droplets is controlled by the pH. The oxidation of
S(IV) species (HSO3-, SO32-, SO2(aq.))
by reactions with H2O2, O3 or O2 in the liquid phase
pushes the gas-liquid partition towards dissolution of gaseous SO2. A
diagram of the sulfur chemical model is presented in
Fig. . Since the model CiTTyCAT resolves continuity
equations for species with gas-phase concentration units, liquid phase
concentrations (e.g. M) and rate constants have to be expressed into
gas-phase units in the code in order to be treated by the CiTTyCAT chemistry
solver .
The species involved in the acid-base equilibriums of SO2 and
H2SO4 are often grouped together according to their oxidation
state:
S(IV)=SO2(g.)+SO2(aq.)+HSO3-+SO32-,S(VI)=H2SO4(g.)+HSO4-+SO42-.
Sulfur chemistry scheme.
Gaseous reaction
k
units
SO2 + OH + M → HOSO2 + M
4.62×10-31×(T/298.0)-3.90
cm6molecule-2s-1a
HOSO2 + O2 → HO2 + SO3
1.30×10-12×(-330/T)-3.90
cm3molecule-1s-1a
SO3 + H2O → H2SO4
9.10×10-13
cm3molecule-1s-1a
Aqueous reaction
k(aq.)
units; (T)
SO2(aq.) + O3 → S(VI) + O2
2.4×104
Ms-1b
HSO3- + O3 → S(VI) + O2
3.7×105
Ms-1b
SO32- + O3 → S(VI) + O2
1.5×109
Ms-1b
HSO3- + H2O2 → S(VI) + H2O
kH2O2×[H+]1+K(eq.)×[H+]
Ms-1b
with K(eq.) = 13
M-1b
and kH2O2 = 7.5×107
M-2s-1b
S(IV) + 12 O2 ⟶TMIS(VI)
750× [Mn(II)] + 2600× [Fe(III)] + 1.0×1010[Mn(II)][Fe(III)]
s-1c
a ; b ; c .
In these equations, dissolved H2SO4 is assumed to be totally
dissociated. Ultimately, S(VI) in droplets ends up deposited at the
Earth's surface. In the model, the amount of sulfate deposited is evaluated
as a variable. The pH of volcanic water droplets is also a prognostic
variable because sulfur species reaction rates and partitioning are pH
dependent . It is dynamically calculated considering the
most significant species dissolved in droplets:
=[HSO3-]+2×[SO32-]+[HSO4-]+2×[SO42-].
The main aqueous equilibrium reactions and S(IV) oxidation reactions
added to the chemical scheme are summarized in Table .
The final continuity equation for single SO2 oxidation channels can
be expressed as follows:
-d[SO2]dt=kOH+SO2×[SO2][OH]+∑j(kj×[S(IV)]aq.[Cj]aq.),
where kOH+SO2 is the rate constant of the gas-phase reaction
between OH and SO2 , kj the rate constant of the
aqueous reaction between SO2 and species Cj, whose concentration
in the aqueous phase is expressed as [Cj]aq..
Similar continuity equations can easily be derived for all the sulfur
species. The continuity equation for atmospheric sulfate S(IV) can be
determined by summing all the individual continuity equations of S(IV)
species:
d[S(IV)]dt=-kOH+SO2×[SO2][OH]-∑j(kj×[S(IV)]aq.[Cj]aq.)-kd×[SO2(aq.)+HSO3-+SO32-]-Kmix×([SO2]-[SO2](bck.)),
where kj the rate constant of the aqueous reaction between oxidant Cj
and relevant [S(IV)] species (see the list of aqueous oxidation
reaction in Table ), and kd is the deposition
coefficient of dissolved sulfur species. Dry deposition as such is not
expected to be important in the plume itself compared to wet deposition for
our cloudy conditions. Since only wet deposition is considered, only species
dissolved in water phases such as aqueous S(IV)
(SO2(aq.) + HSO3- + SO32-) and
S(VI) (HSO4- + SO42-) species are
deposited in the model. The deposition is treated as a first order loss with
kd = 2×10-6 s-1, equivalent to a
characteristic time scale of 5.7 days .
The same approach can be used for S(VI) and deposited S(VI):
d[S(VI)]dt=kOH+SO2×[SO2][OH]+∑j(kj×[S(IV)]aq.[Cj]aq.)-kd×[HSO4-+SO42-]-Kmix×([S(VI)]-[S(VI)](bck.),)
d[S(VI)dep.]dt=kd×[HSO4-+SO42-],
where S(VI)dep. is the sulfate deposited at the surface.
Oxygen isotope signatures in sulfur oxidation
The mass-balance equation describing the production of S(VI) species
is coupled to an oxygen isotope transfer scheme in order to track the
evolution of Δ17O in sulfates in water droplets and in
sulfates deposited at the surface. Therefore, the specific isotopic anomaly
acquired by a S(VI) molecule (produced by the oxidation of a
S(IV) molecule by a specific oxidant) is first derived using isotopic
transfer equations. New S(VI) isotopes tracers are then created in
order to monitor the amount of isotopic anomaly carried out by sulfates in
water droplets and deposited at the surface. They are defined as anomaly
products (Δ17O × [S(VI)]), and introduced
in the model on the basis of the following continuity equation
:
ddt[S(VI)]×Δ17O(S(VI))=∑j[Pj×Δ17O(S(VI)prod)j]-kd×Δ17O(S(VI)),
where Δ17O(S(VI)) is the isotopic anomaly of
atmospheric sulfate, Δ17O(S(VI)prod.)j
is the O-MIF anomaly transferred to sulfate through the specific oxidation
channel j, and Pj is the oxidation rate of channel j. Δ17O(S(VI)prod.)j is fixed for ozone,
H2O2 and TMI oxidation pathways but it is a prognostic
variable for OH (see Table ).
O-MIF signatures of S(IV) oxidation pathways in the model.
Oxidant
O-MIF pathway (‰)
OH
calculated (0 to a maximum of 4.5)
H2O2
0.87
O3
9
O2/TMI
-0.09
As deposited sulfate is a variable in the model (S(VI)dep.),
the transfer of isotopic anomaly during deposition is also monitored
following a similar equation,
ddt[S(VI)dep.]×Δ17O(S(VI))=kd×[S(VI)]×Δ17O(S(VI)),
The value of oxygen isotopic anomaly (O-MIF) in sulfate depends on the
relative importance of individual SO2 oxidation pathways (Pj) and
their respective transfer of O-MIF (Δ17O(S(VI))j). Note that the continuity equations of
S(VI) and S(VI)dep. isotopes tracers are integrated with
a 4th order Runge–Kutta method algorithm instead of using the CiTTyCAT
chemistry solver with the oxidation rates (i.e. Pj in )
kept constant over a time step . This approach
allows keeping the chemistry module totally independent from the oxygen
isotopic module. The external integration tool does not affect significantly
the results. Throughout this study, it is assumed that both SO2 and
water vapour (H2O) are not carrying any initial O-MIF. The isotopic
composition of magmatic SO2, indeed follows mass-dependent
fractionations and no significant Δ17O, Δ34S and Δ36S have been measured so far
. Measurements show that tropospheric H2O does not
carry any O-MIF , and the same is found for atmospheric
SO2 . It is therefore assumed that the O-MIF found in
sulfates only originates from the transmission of isotopic anomaly during
the aforementioned reactions of sulfur oxidation.
In order to constrain individual SO2 oxidation pathways from isotopic
information, it is first necessary to characterize the specific O-MIFs they
transfer to sulfate using isotopic transfer equations.
Oxidation by ozone
The few isotopic measurements of tropospheric ozone indicate values of
Δ17O (O3bulk), ranging from 20 ‰
to 40 ‰ with a mean value of about 25 ‰
. The location of
oxygen isotopes within the structure of ozone is not uniform and heavier
isotopes are mostly located at the extremities of the molecule
. Indeed, molecules that have
asymmetrical geometrical structures, and bearing heavier oxygen isotopes on
terminal sites, are more energetically stable than their symmetric
counterparts . This enrichment in heavy oxygen isotopes at
terminal locations of ozone is confirmed by laboratory measurements
. Ozone does not always react with other molecules
via terminal oxygen atoms, although this reaction mechanism is energetically
favourable since it requires the breaking of only one molecular bond. During
the oxidation of reactive nitrogen leading to production of atmospheric
nitrate, most of the oxygen atoms involved in the reaction are from terminal
sites . Multiple studies found a similar selective
reactivity indeed, as during photochemical reactions or for reactions of
ozone on solid substrates . Considering
the mean bulk O-MIF and terminal isotopic enrichments, a mean reactive ozone
O-MIF (Δ17O (O3*)) of 36 ‰ has been
derived . This value is used
throughout this study, since it is in accordance with parametrizations used
in previous successful model simulations
.
The value of O-MIF in sulfates generated during the aqueous oxidation by
ozone is determined by identifying the origins of each oxygen atom in
sulfate during the reaction of oxidation. Ozone transfers to sulfate only
one oxygen atom during aqueous sulfur oxidation, while another oxygen atom
derives from the water molecule forming aqueous S(IV). The equation
describing the transfer of O-MIF to sulfate during oxidation by ozone is as follows:
Δ17O(S(VI))O3+SO2=12×Δ17O(SO2)+14×Δ17O(H2O)+14×Δ17O(O3*).
This equation can be simplified because the O-MIF anomalies in SO2
and H2O are negligible:
Δ17O(S(VI))O3+SO2=14×Δ17O(O3*).
Therefore, the isotopic anomaly in atmospheric sulfates produced in the
model during the oxidation of dissolved SO2 through O3 is
Δ17O (S(VI))O3+SO2 = 9 ‰
.
Oxidation by hydroxyl radical
In the atmosphere, OH radicals are formed as a result of ozone
photolysis in presence of water vapour. In particular, ozone
photodissociation can produce an O1(D) radical, which react with a
water molecule to produce two OH radicals. Tropospheric OH
radicals are thought not to carry O-MIF anomaly because the exchange of
oxygen atoms with water vapour is so fast that it erases any inherited
isotopic anomaly in OH. Recall that tropospheric H2O does not
carry any O-MIF because the tropospheric H2O cycle is entirely
controlled by physical processes (condensation, evaporation) and not by
chemical processes involving ozone. As a result, the O-MIF signature in
OH radicals is expected be 0 (Δ17O
(OH) = 0.0 ‰ ). However, when the
humidity and hence H2O levels are very low (e.g. upper troposphere),
the rate of isotopic exchange between OH radicals and H2O
molecules decreases so much that freshly produced OH radicals may have
time to react with other molecules before losing their isotopic anomaly by
isotopic exchange with H2O . Under those conditions,
when the OH loss reactions and cycling compete with the isotopic
exchange with H2O, some of the initial O-MIF originating from ozone
is still present in reacting OH. It is also possible for OH
loss to compete with the H2O isotopic exchange when the rate of
OH loss is highly enhanced instead of having a reduced rate of
H2O isotopic exchange. This may be the case in volcanic plumes, when
SO2 levels are so high that the SO2+OH reaction become the
dominant chemical loss , accelerating the OH cycling.
In order to account for this possibility, instead of assuming a null O-MIF
for OH, the O-MIF in the steady-state OH (Δ17O
(OH)) is calculated explicitly using the approach developed by Morin
et al. . Δ17O (OH) is simply
derived from the competing balance between the O-MIF erasing isotopic
exchange and the total OH loss, typically the reactions with CO
and CH4 in the troposphere. Since we consider sulfur-rich volcanic
plumes and clouds, the reaction between OH and SO2 is also
taken into account.
Considering all the transfers of oxygen atoms, the isotopic mass-balance
equation for the OH pathway can be expressed as follows:
Δ17O(S(VI))OH+SO2=12Δ17O(SO2)×+14×Δ17O(OH)+14×Δ17O(H2O).
Since tropospheric H2O and volcanic SO2 are not thought to
carry any O-MIF, the equation can be simplified:
Δ17O(S(VI))OH+SO2=14×Δ17O(OH).
The O-MIF of OH can be derived using the following equation
Δ17O(OH)=x×Δ17O(OHprod.*),
with
Δ17O(OHprod.*)=12×Δ17O(O3*),
and
x=DD+kOH+H2O*×[H2O]
D=kOH+CO×[CO]+kOH+CH4×[CH4]+kOH+SO2×[SO2],
where kOH+H2O* is the rate constant for the oxygen atoms
exchange reaction between OH and H2O, and kOH+CH4 and kOH+CO are the reaction rate constants for the gas
phase reaction of OH with CH4 and CO, respectively.
In this approach x represents the competition between the O-MIF erasing
effect of isotopic exchange and the O-MIF retaining effect of OH
chemical loss; only important loss reactions for tropospheric OH are
considered here. Δ17O(OHprod.*) is the
O-MIF of the OH radical freshly produced, and it is assumed that
OH is mostly formed by the photolysis of ozone followed by the
reaction of O1(D) with H2O.
The O-MIF in OH (Δ17O(OH)) is determined by
this x factor. If OH chemical loss is much faster than the isotopic
exchange, Δ17O(OH) = 0.5 ×Δ17O(O3*) (i.e. x=1). If chemical loss is much slower
than the isotopic exchange, Δ17O(OH)≈0‰ (i.e. x≪1).
Oxidation by hydrogen peroxide
In the troposphere, H2O2 can quickly dissolve into liquid water
phases . In a volcanic plume, these phases can be either
water droplets or water condensed on solid particles, typically on ash
particles. Once in the aqueous phase, H2O2 oxidizes SO2 by
nucleophilic displacement, and its two oxygen atoms are transmitted to the
produced sulfate molecule .
The isotopic balance for the oxidation by H2O2 in the liquid phase
is as follows:
Δ17O(S(VI))H2O2+SO2=12×Δ17O(SO2)+12×Δ17O(H2O2).
Since volcanic SO2 is thought to carry no significant O-MIF, the
final O-MIF transfer can be simplified:
Δ17O(S(VI))H2O2+SO2=12×Δ17O(H2O2).
Isotopic measurements of Δ17O of tropospheric H2O2
range between 1.30 ‰ and 2.20 ‰ with a mean O-MIF of 1.70 ‰
. Using this mean value, sulfate produced
by the H2O2 oxidation is assumed to carry a Δ17O(S(VI))H2O2+SO2 = 0.87 ‰
.
Oxidation by <inline-formula><mml:math id="M475" display="inline"><mml:mrow class="chem"><mml:msub><mml:mi mathvariant="normal">O</mml:mi><mml:mn mathvariant="normal">2</mml:mn></mml:msub><mml:mo>/</mml:mo><mml:mi mathvariant="normal">TMI</mml:mi></mml:mrow></mml:math></inline-formula>
Isotopic measurements of atmospheric O2 indicate that its O-MIF
anomaly is rather small . Kinetic isotope
fractionation associated to the Dole effect and
stratospheric influx of O2 generates a slightly negative O-MIF in
tropospheric O2. As theoretical investigations suggest, a slight
depletion of 17O is indeed found in tropospheric O2,
which is accompanied by a slightly negative O-MIF .
Theoretical calculations predict Δ17O (O2) as low
as -0.344 ‰ or even, more recently,
-0.410 ‰ for tropospheric O2 . Other
theoretical calculations suggest a Δ17O (O2)
between 0.141 ‰ and -0.305 ‰ .
We assume a Δ17O (O2) of -0.340 ‰
. This value is chosen because it gives a reasonably good
agreement between isotopic measurements and models
. In addition, it has to be kept in mind
that there are large uncertainties associated with the exact reaction
mechanism of SO2 oxidation catalysed by TMI. We assume that
only one oxygen atom of O2 is transmitted to sulfate during the
SO2 oxidation .
With these assumptions, the isotopic mass-balance equation for SO2
oxidation by O2/TMI is given by
Δ17O(S(VI))O2+SO2=34×Δ17O(S(IV))+14×Δ17O(O2).
Since volcanic SO2 is thought to carry no significant O-MIF, we can
assume that initial S(IV) species do not carry any O-MIF.
Consequently, the isotopic signature associated to this oxidation pathway can
be simplified:
Δ17O(S(VI))O2+SO2=14×Δ17O(O2).
Δ17O (O2) being taken as -0.34 ‰ (see
above), sulfate produced through this pathway carries a O-MIF (Δ17O(S(VI))O2+SO2) almost null, of about
-0.09 ‰ . The O-MIF signatures of all the
S(IV) oxidation pathways used in the model are summarized in
Table .
Box model set-up
Standard case: initial conditions
All simulations are run for springtime conditions and start at 08:00 at
tropical latitudes (8.3∘ N). In order to reach stable chemical
compositions, notably for medium- and short-lived reactive species, the model
is run for 3 days before injecting SO2, then, the evolution of the
chemical composition is followed for 7 days. This timescale corresponds
approximately to the lifetime of a plume in the free troposphere, in
occurrence of low turbulence and low wind shear .
Since most of volcanoes are situated in remote areas with their peaks close
to the free troposphere, or, at least, with volcanic plumes often ending up
in the free troposphere, the environmental conditions are chosen to be
representative of the lower free troposphere with temperature set at
283.15 K, and pressure fixed at 640 mbar (about 3 km
altitude). Since we consider cloudy conditions, the relative humidity is set
to 100 %.
Furthermore, concentrations of reactive species are also set to typical
levels found in the tropical lower free troposphere:
O3 = 45 ppbv and H2O2 = 0.1 ppbv. Finally, initial
SO2 is set to a mean volcanic plume concentration of 1 ppmv, a value
typical of volcanic plumes during degassing
.
The initial pH of the aqueous phase is set to 4.5. It has no impact on the
overall model results because the pH is almost immediately driven by
SO2 uptake and sulfur oxidation. Preliminary simulations have shown
that the initial SO2 concentration is a critical input.
Due to the large amounts of water that can be injected during explosive
eruptions, in our simulations it is assumed that volcanic water vapour is
largely in excess compared to relative humidity of the free troposphere.
Moreover, due to low temperature and pressure of the lower free troposphere,
for our simulations it is assumed that volcanic water vapour would mostly
condense to produce cloud droplets or to coat ash particles. Therefore,
throughout this study relative humidity (RH) inside volcanic plumes is set at
100 %, the water saturation point corresponding to the pressure and
temperature of the background atmosphere. The Liquid Water Content (LWC)
parameterizes the amount of liquid water within plumes. High levels of LWC can be reached, indeed, within volcanic plumes from
explosive eruptions. Modelling simulations suggest that LWC as high as
1.6 gm-3 could be reached at the core of water-rich volcanic
clouds condensing in the troposphere . It is possible that,
during the first stages of medium-size eruptions, LWC within the plume could
be at least comparable to LWC values of growing cumulus clouds. For all the
simulations LWC is set to 1.0 gm-3, a value between experimental
measurements (e.g. meteorological clouds) and modelling studies of water-rich
volcanic plumes reaching the upper troposphere
. Like SO2, LWC is found to be
a critical model input.
TMI concentrations in the liquid phase are set to [Fe(III)] =
0.5 µM and to [Mn(II)] = 0,05 µM. These values
are at the lower end of typical tropospheric measurements with
[Fe(III)] concentrations ranging between 0.5 and 2 µM
. Because of uncertainties
associated with iron dissolution in volcanic plumes, our TMI concentrations
are lower than concentrations found in dust-rich polluted conditions where
[Fe(III)] can reach concentrations of around 5 µM
. TMI concentrations follow the same
relation throughout the whole study and for each simulation
[Mn(II)] = 0.1 × [Fe(III)] .
Model experiments
The objective of the first set of numerical experiments is to assess the
competition among oxidation pathways in SO2-rich plumes/clouds for
the standard case. Three simulations (S1–S3) are run with oxidation schemes
of increasing complexity. They simulate oxidation of SO2: (S1) by
OH in gas phase, (S2) by OH in gas phase, and H2O2 and
O3 in aqueous phase, and (S3) by OH in gas phase, and
H2O2, O3 and O2/TMI in aqueous phase.
Since initial SO2 levels, LWC and TMI concentrations in volcanic
plumes are relatively uncertain and are key model inputs, the sensitivity of
the results to varying conditions within plausible ranges is also explored in
additional simulations. Isotopic anomaly transfers are investigated for
atmospheric concentrations stretching from passive degassing and quiescent
conditions to sulfur-rich volcanic clouds with varying levels of TMI. The
intervals used for the different sensitivity studies are summarized in
Table .
The first set of sensitivity simulations is devoted to the sensitivity of
results to initial SO2 levels in the case of the S1 simulation. It is
designed to explore not only the impact of varying SO2 levels on
sulfate O-MIF produced by the OH oxidation pathway, but also on OH
isotopic signature itself. Recall that the OH isotopic signature
(Δ17O (OH)) is generally assumed to be 0 in the
literature (see Sect. 2.4.2).
Ranges of SO2, LWC and TMI explored in the sensitivity studies.
SO2
0.1–10.0 ppmv
LWC
0.1–2.5 g m-3
TMI
0.1–3.0 µM
It is widely recognized that SO2 is the compound emitted by volcanic
activity which causes the largest climatic impacts through its conversion
into sulfate aerosols .
Emissions of volcanic SO2 have been measured both in proximity of
volcanic vents and in aged plumes. It is possible to constrain a range of
concentrations, considering age of plumes and distance from points of
emissions. During the first stages of plume development concentrations of
SO2 in the range of 10–50 ppmv can be reached right in proximity of
volcanic vents , while
concentrations in the range of 0.1–1 ppmv can be found in aged plumes at
longer distances from points of emissions .
These results are confirmed by modelling simulations which can constrain
volcanic emissions by accounting for quick dilution after plume ejection from
the vent . Consequently, based on
atmospheric simulations and on in situ measurements, the SO2
concentration is set to 1.0 ppmv in the standard case, and is varied from
0.1 to 10 ppmv in the sensitivity simulations.
LWC plays a crucial role in aqueous oxidation of volcanic SO2. The
range of LWC considered has been chosen based on LWC observed for different
cloud typologies such as mean saturated clouds (0.1 gm-3),
water-rich cumulus clouds (0.5–1 gm-3), and cumulonimbus clouds
(1–2 gm-3)
. LWC is
set to 1 gm-3 in the standard case and is varied from 0.1 to
more extreme values of 2.5 gm-3 for sensitivity simulations
.
Aqueous concentrations of iron
([Fe(tot)] = [Fe(II)] + [Fe(III)]) can peak to
9–10 µM in the troposphere
with [Fe(III)] concentrations between 2.0 and 5.0 µM in
polluted conditions if photochemical cycling between
[Fe(II)]-[Fe(III)] is inhibited . Volcanic
eruptions inject large quantities of solid material into the atmosphere in
the form of ash. As a result, volcanic plumes/clouds are characterized by
high concentrations of ash and minerals . Ash particles
have sizes as large as few mm and they are mainly composed of silica and
crystalline minerals of magmatic origin. Glass, olivine, magnetite, hematite
and fayelite are among the most common minerals injected during eruptions
. These minerals are composed
in different proportions by Fe(II) and Fe(III), which are
entrapped in the crystalline structure of rocks in different morphologies and
compositions. Since large quantities of water are also injected during
eruptions, water can condense on mineral particles, especially as the
volcanic column reaches higher altitudes and lower temperatures in the
troposphere . Once mineral particles
are coated by water, dissolution of iron from the solid mineral surface to
the thin liquid water film may take place depending on the acidity of the
aqueous phase . Acidic conditions
(pH < 2.0) due to H2SO4 condensation or formation within the
liquid phase favour the solubility of minerals containing iron and
dissolution of [Fe(III)] . Up to a third
of total Fe at the ash surface can dissolve into the liquid phase coating
volcanic particles depending on rock composition and
gases in the volcanic clouds. Laboratory experiments on dissolution in acidic
water of iron from volcanic ashes suggest that [Fe(III)]
concentrations of up to 2 µM can be reached quickly in the liquid
phase when pH reaches ∼ 2 . Concentrations as high
as [Fe(III)] = 3 µM could be reached if pH reaches 1
. Mobilization of iron ions from ashes could be enhanced
for plumes reaching the upper troposphere and undergoing ice formation
. High concentrations of [Fe(III)] might
persist in the liquid phase depending on the lifetime of water droplets,
notably driven by evaporation and condensation cycles
.
Evolution of the gas-phase concentrations of atmospheric species
during the S1 simulation (see text). The simulation starts at 08:00 and
SO2 is injected after 3 days. During the S1 simulation the concentration
of injected SO2 drops from 1.5 ppmv to a final value of 1.27 ppmv.
![]()
Cloud properties are affected by evaporation and condensation cycles changing
the pH, the size and number of droplets, while formation of insoluble salts
at the surface of ash particles entrapped in cloud droplets can affect
mobilization of ions . Therefore, lower acidity
combined with the presence of insoluble salts may result in a reduced
availability of dissolved TMI in volcanic clouds as the volcanic cloud ages.
Over the long term, these conditions can lead to concentration of
Fe(III) in water droplets of volcanic clouds which can be lower than
typical concentrations found in tropospheric clouds
. In this study, [Fe(III)] is set
to 0.5 gm-3 in the standard case and is varied from 0.1 to
3 µM in the sensitivity simulations to cover the wide range of
possible [Fe(III)] concentrations.
The resulting model Δ17O(S(VI)) (i.e. from standard
and sensitivity simulations) are compared to sulfate O-MIF found in
tropospheric volcanic sulfates extracted from ash deposits of small and
medium-size tropospheric explosive eruptions of the present geological era
, or in sulfate aerosols collected at
volcanic vents, almost certainly primary sulfate .
Time evolution of Δ17O(S(VI)) in produced and
deposited sulfates, and of the pH of the liquid phases in volcanic plumes during
simulation S1, following the injection of SO2 in the box. The change of
pH in water droplets is also reported as a function of time.
![]()
Time evolution of the O-MIF transfer from OH to
H2SO4(g) at two different initial concentrations of
SO2. The light green line represents initial concentration of
S(IV) = 1 ppbv (e.g. mean troposphere); the dark green line
represents an initial concentration of SO2 = 1 ppmv
(e.g. volcanic plumes and clouds). The upper figure shows concentration trends
for OH during the two different scenarios.
![]()
Conclusions
We use the tropospheric photochemical box model CiTTyCAT to
analyse why most oxygen isotopic measurements of tropospheric volcanic
sulfate indicate that volcanic sulfates are essentially mass-dependent
(i.e. O-MIF anomalies lying close to zero within measurement uncertainties
of ±0.1 ‰ typically). This is also observed for volcanic
sulfate collected very far from volcanoes where secondary sulfate (produced
by oxidation of volcanic sulfur precursors, mostly SO2) is expected
to vastly dominate. This lack of O-MIF in volcanic sulfate is rather
intriguing because secondary sulfates originating from other sources exhibit
significant O-MIF. A major difference between volcanic sulfur and other
sources is that it is often emitted within very dense volcanic plumes whose
chemical compositions are radically different from background air. The
chemical environment of the plumes may affect the oxidation pathways and
hence sulfate isotopic composition.
A new sulfur isotopic O-MIF scheme is implemented in the model in order to
monitor the transfer of O-MIF from oxidants to sulfate during the oxidation
of volcanic SO2. A range of simulations are performed in order to
explore in details the different pathways of SO2 oxidation (gas-phase
oxidation by OH and aqueous oxidation by O3, H2O2 and
O2/TMI) and, more importantly for O-MIF, their relative importance
for a range of possible volcanic conditions. The first salient finding is
that, according to the model calculations, OH should carry a very significant
O-MIF in sulfur-rich volcanic plumes. This implies that, when volcanic
sulfate is produced in the gas phase via SO2 oxidation by OH,
its O-MIF should have a very significant positive value. Since most isotopic
measurements of volcanic sulfate do not indicate the presence of O-MIF, the
OH oxidation pathway cannot be the dominant channel for volcanic
sulfur. Nonetheless, uncertainties in the rate constant of the isotopic
exchange between OH and H2O and, more
generally, on photochemical modelling are substantial . It
would be useful for this unexpected model predictions of O-MIF in OH,
and hence volcanic sulfate produced in gas phase, to be tested in a
controlled environment, ideally with laboratory experiments of SO2
oxidation with a well constrained OH chemical budget, especially in relation to the loss
processes. The second important finding from these simulations is that,
although H2O2 is a major oxidant of SO2 throughout the
troposphere, it is very rapidly consumed in sulfur-rich volcanic plumes.
Since H2O2 produced within the plume and the entrainment of
H2O2 from the atmospheric background also represent relatively weak
sources, H2O2 is found to be a minor oxidant for volcanic
SO2 whatever the liquid water content. According to the simulations,
oxidation of SO2 by O3 is negligible because volcanic aqueous
phases are too acidic. The model predictions of minor or negligible sulfur
oxidation by H2O2 and O3, two oxidants carrying large O-MIF,
are consistent with the lack of O-MIF seen in isotopic measurements of
volcanic tropospheric sulfate. The third finding is that oxidation by
O2/TMI in volcanic plumes could be very substantial and, in some
cases, dominant, notably because the rates of SO2 oxidation by
OH, H2O2, and O3 are vastly reduced in a volcanic
plume compared to the background air. Only cases where sulfur oxidation by
O2/TMI is very dominant can explain the isotopic measurements of
volcanic tropospheric sulfate. We stress that oxidation by O2/TMI is
poorly constrained in model simulations because of the lack of measurements
of TMI aqueous concentrations in volcanic plumes. It is worth pointing out
that our model results are only applicable to cloudy volcanic plumes.
Nonetheless, water clouds do not always form in volcanic plumes, notably
during passive degassing. It would be interesting to also consider cloud-free
plumes where the condensed phase is concentrated sulfuric acid within
sulfate aerosols. In particular, these particles have a chemical reactivity
radically different from water droplets.
A potentially significant limitation of the model simulations is the omission
of volcanic halogens. Indeed, volcanic halogens are known to undergo
multi-phase chemistry, resulting in ozone depletion and possibly impacting
the oxidation of volcanic SO2
.
Halogen species such as HOBr may also directly oxidize SO2 in
the aqueous phase , but this oxidation pathway has not been
quantified yet for volcanic plumes. Overall, the present simulations might
only be representative of degassing or eruptions with extremely low halogen
emissions, typically originating from intraplate and rift volcanic activity.
It is certainly worth exploring the potential impact of halogens in the case
of halogen-rich eruptions, notably for volcanic plumes where water does not
condense and hence only sulfate aerosols are present. Since the
heterogeneous conversion of halogen halides into radicals is known to be fast
on sulfate aerosols
, halogens
might significantly impact the plume chemistry and the isotopic composition
of secondary sulfate for halogens rich conditions.