Gases in Volcanic Lake Environments

Abstract

Volcanic lake systems derive their gases from four distinct sources. Of greatest importance from a hazard perspective, and those which set these limnic systems apart from non-volcanic lakes, are gases derived directly from magmatic sources feeding the volcano, including CO2, SO2, H2S, HCl, HF and a myriad of minor species. The major gases are acidic in nature, and when dissolved into ground water, lead to the development of aggressively acidic solutions. Hydrolysis reactions wiThenclosing rocks, systematically alter the magmatic gas compositions towards more benign hydrothermal signatures, and this process usually leads to precipitation of permeabilityreducing mineral assemblages. Ground and lake waters carry dissolved atmospheric constituents into these environments, whereas lakes are wellknown biotic environments, whose populations may also leave their mark on solute gas compositions through their normal metabolic processes. Apart from magmatic eruption events, at least two specific hazards are attributable to gases in volcanic lake environments, both of which have been responsible for loss of life near volcanic lakes. Physical and chemical processes extant in systems where magmas lie within 100s of metres of the surface have the propensity to form mineralogic seals beneath the lakes. Such sealing may foster over-pressuring and associated gas-driven phreatic eruptions of the type that has occurred recently at Ruapehu and Raoul Island, New Zealand, often with little or no precursory activity. On the other hand, where magmatic gas sources are deeply-seated, to such an extent where heat decouples from the rising gas stream (principally CO2), conditions are perfect for the formation of cold, gas-stratified lakes. Overturn of such lakes typically leads to violent release of the gas, as has occurred in the Cameroonian Lakes Nyos and Monoun, leading to the deaths of near-by inhabitants. Both situations are endmembers of a continuum of processes operating where volcanoes interact wiThearth’s hydrosphere.

Full text

Gases in Volcanic Lake Environments
B. Christenson and F. Tassi
Abstract
Volcanic lake systems derive their gases from four distinct sources. Of
greatest importance from a hazard perspective, and those which set these
limnic systems apart from non-volcanic lakes, are gases derived directly
from magmatic sources feeding the volcano, including CO
2
,SO
2
,H
2
S, HCl,
HF and a myriad of minor species. The major gases are acidic in nature, and
when dissolved into ground water, lead to the development of aggressively
acidic solutions. Hydrolysis reactions with enclosing rocks, systematically
alter the magmatic gas compositions towards more benign hydrothermal
signatures, and this process usually leads to precipitation of permeability-
reducing mineral assemblages. Ground and lake waters carry dissolved
atmospheric constituents into these environments, whereas lakes are well-
known biotic environments, whose populations may also leave their mark
on solute gas compositions through their normal metabolic processes. Apart
from magmatic eruption events, at least two specific hazards are attributable
to gases in volcanic lake environments, both of which have been responsible
for loss of life near volcanic lakes. Physical and chemical processes extant in
systems where magmas lie within 100s of metres of the surface have the
propensity to form mineralogic seals beneath the lakes. Such sealing may
foster over-pressuring and associated gas-driven phreatic eruptions of the
type that has occurred recently at Ruapehu and Raoul Island, New Zealand,
often with little or no precursory activity. On the other hand, where
B. Christenson (&)
GNS Science, National Isotope Centre, Lower Hutt,
New Zealand
e-mail: b.christenson@gns.cri.nz
F. Tassi
Department of Earth Sciences, University
of Florence, Via G. La Pira, 4, 50121 Florence, Italy
F. Tassi
CNR-IGG, Institute of Geosciences and Earth
Resources, Via G. La Pira, 4, 50121 Florence, Italy
D. Rouwet et al. (eds.), Volcanic Lakes, Advances in Volcanology,
DOI 10.1007/978-3-642-36833-2_5,
©Springer-Verlag Berlin Heidelberg 2015
125

magmatic gas sources are deeply-seated, to such an extent where heat
decouples from the rising gas stream (principally CO
2
), conditions are
perfect for the formation of cold, gas-stratified lakes. Overturn of such lakes
typically leads to violent release of the gas, as has occurred in the
Cameroonian Lakes Nyos and Monoun, leading to the deaths of near-by
inhabitants. Both situations are endmembers of a continuum of processes
operating where volcanoes interact with Earth’s hydrosphere.
Keywords
Dissolved gas composition Hyperacidic lake Lake monitoring Nyos-
type lake Gas/water reaction
1 Introduction
Gases are the lifeblood of volcanic systems.
From the onset of volatile exsolution from
vapour-saturated host magmas, and from depths
extending to deep crustal environments (e.g.,
Anderson 1975; Giggenbach 1996; Wallace
2005) through to their interaction with volcanic
lakes at the Earth’s surface, magmatic volatiles
play a key role in the physical and chemical
evolution of their enclosing magmatic-hydro-
thermal environment. Whereas exsolved mag-
matic gases not only augment density differences
between the melt phase and enclosing crustal
materials, thereby driving buoyant ascent of
magma, in the words of Giggenbach, they also
“rule, or at least witness”, many of the chemical
processes operating on both sides of the magma-
hydrothermal interface during ascent.
Interaction between magmatic volatiles and the
Earth’s hydrosphere leads to complex physico-
chemical processes, most of which are not obser-
vable to us on the surface, and many of which lead
to potential hazards for those living near such
volcanic systems. These hazards include, but are
not limited to, the development of hydrothermal
seals which typically lead to phreatic and/or phre-
atomagmatic eruptions, and the accumulation of
gases in lakes (e.g., Nyos-type hazards).
In this chapter, we briefly explore the nature
of magmatic volatiles which are released from
active volcanoes, and some of the key physical
and chemical consequences of their interaction
with meteoric surface waters, including lakes.
We also briefly discuss how these signatures are
opportunistically altered by living organisms in
lake environments. We will then look at two
representative case histories where gases have
contributed to hazardous natural events in vol-
canic lake settings, including a recent gas-driven
phreatomagmatic eruption from the active
andesitic massif of Mt Ruapehu (New Zealand),
and an overview of the key processes associated
with the accumulation and cataclysmic release of
gases from Cameroonian lakes (e.g. Nyos),
detailed discussions for which are found else-
where in this volume.
2 Origins of Gases in Volcanic
Lake Systems
Gases in volcanic lake systems can be described
in terms of four endmember components,
including two principle sources (magmatic and
meteoric) and two derived sources (hydrother-
mal and biogenic), as shown schematically in
Fig. 1. Of greatest interest to us from a hazard
stand point, and those which will therefore
receive most of our attention here, are gases
derived from the magma, and its enveloping
magmatic-hydrothermal environment. The
hydrosphere accommodates the other two sour-
ces, including atmospheric gases dissolved in
ground and lake water, and those derived from
metabolic processes of living organisms (the
biogenic fraction).
126 B. Christenson and F. Tassi

The degree to which a lake in a volcanic terrain
will be chemically or physically influenced by a
deeper magmatic system is dependent upon its
proximity to the underlying magma, the rate of
degassing, the permeability of the intervening
rock mass, and the degree to which convective
circulation (if any) is established between the two
environments. Maximum magmatic gas emission
rates will occur where gases transfer across an
open, freely degassing lava surface within the
magma conduit into a highly permeable vent
system, as portrayed in Fig. 1. At relatively shal-
low depths (say 1–3 km), pressure-temperature
and compositional (P-T-X) conditions are con-
ducive to the development of a single-phase
vapour envelope surrounding the magma, the size
of which will be proportional to the magma body,
and its heat/gas emission rate (Todesco et al. 2004,
this issue; Christenson et al. 2010). Enclosing this
zone of single-phase vapour will be a 2-phase
vapour-liquid region, which in turn is surrounded
by single-phase liquid (i.e., water). The vertical
extent of this topology will again be controlled
largely by heat and mass flow of volatiles from the
magma, and permeability within the conduit. Heat
and mass transport modelling of these systems
indicates that convective circulation develops
around the conduits, with hot, low density fluids
rising within and adjacent the conduit, and draw-
ing cooler marginal fluids into the flow regime.
The most extreme examples of this type of
lake system are found where the volcanoes are
active, with magma residing close to the surface,
and where the main eruption vents are covered
by the lake. Examples of such lakes are Aso
(Shinohara et al. this issue), Poás (Rowe et al.
1992), Ruapehu (Christenson and Wood 1993),
and Copahue (Varekamp et al. 2009).
A similar, but less expansive topology arises
in systems with less vigorously degassing
magma bodies at their core, or perhaps where
both crystallisation and magmatic degassing are
well advanced, and the remaining melt fraction is
enclosed by a hot, ductile carapace. A model for
gas release across quenched margins of magma
bodies has been elegantly described by Fournier
(1999), and a possible example of this type of
lake system is located at Raoul Island (e.g.,
Weissberg and Sarbutt 1966; Christenson et al.
2007). Similarly, and for reasons which will be
discussed below, systems where the magma body
resides at great depth (such as Laacher See and
Lake Nyos), have rather simpler gas composi-
tions, consisting predominantly of CO
2
.
Water-rock interaction, and the transition
from the magmatic to the hydrothermal envi-
ronment, initiates from the moment that con-
densation begins in the 2-phase liquid-vapour
environment enveloping the conduit and single-
phase vapour region, and proceeds across this
zone and to the thermal margins of the single
crater
lake
biospheric
component
sources
hydrothermal
component
sources
magmatic
source
components
acid neutralisation
reduction
scrubing
atmospheric
gas-water
equilibrium
metabolic
products
two-phase
liquid-vapor
degassing
magma
single-phase
vapor
single-phase
liquid
convective
high
permeability
conduit/vent
circulation
convective
circulation
convective
circulation
CO , SO , H S, HCl, HF
2 2 2
CH , NH ,H S
4 3 2
meteoric
components
N , O , Ar, etc.
2 2
Fig. 1 Schematic model for component source gases in a
typical volcanic lake plumbing system. The scale is
purposefully not shown, as the configuration is generic to
both deep and shallow magmatic sources, with only the
topology of the two-phase liquid region adjusting verti-
cally to P-T-X constraints
Gases in Volcanic Lake Environments 127

phase liquid region (Fig. 1). Processes operating
therein, and their impact on the gas composition,
define the so-called hydrothermal source com-
ponent in these systems, as discussed below. For
modelling purposes (both conceptual and
numeric), the hydrothermal region is bound
hydrostatically on its outer margins by non-
thermal groundwater, and the nature of this
boundary is both dynamic and responsive to
changes in volcanic activity. As a result of the
convective processes operating in this zone,
hydrothermal fluids are typically composite
mixtures of meteoric and magmatic source fluids,
with chemical processes operating to generate
distinctly hydrothermal compositional signatures.
Another, potentially dominant component
signature in volcanic lakes is that derived from
biological activity in both lakes and sediments
associated with them (i.e., gases of biogenic
origins). A wide range of opportunistic microbial
organisms gain metabolic energy from oxidation-
reduction reactions involving principally mag-
matic-hydrothermal C, S and N, often leaving
distinctive isotopic signatures (discussed by
Mapelli et al. this volume), described briefly
below.
The gases which derive from these sources are
individually identified in this section, and therein
the following sections.
2.1 Magmatic Gases and Processes
Gases released from active volcanoes owe their
bulk compositions to a variety of complex pro-
cesses and interactions, both chemical and
physical, as they ascend from their source
regions in the mantle to the surface. Excellent
reviews of the compositional characteristics and
evolution of magmatic volatiles are available
from a number of sources, including Carrol and
Webster (1994), Giggenbach (1996), Wallace
(2005), Webster and Mandeville (2007), Aiuppa
et al. (2009), Oppenheimer et al. (2011) and the
many references therein. Only aspects most
commonly affecting chemical signatures in vol-
canic lake systems are introduced here.
2.1.1 Major Component Species
The principle components of magmatic gases are
H, O, C, S and Cl and these comprise >95 % of
magmatic volatiles, regardless of tectonic source
environment.
Hydrogen
Water is the most prevalent H-bearing volatile
species in magmas, followed by comparatively
minor amounts of H
2
,NH
3
,H
2
S, and trace
amounts of CH
4
. Arc magmas tend to have
higher water contents and larger ranges (0.2–6wt
%, Métrich and Wallace 2008) than their ridge
counterparts (0.12–0.5 wt%, Sobolev and Chau-
ssidon 1996;Saaletal.2002). Water enrichment
along arcs is the result of dehydration reactions
in the subducting oceanic crust and sediments,
whereas the large range of observed volatile
contents in arc environments is attributable to
heterogeneity of source materials, differentiation,
and assimilation (e.g. Fischer 2008). Water par-
tially dissociates in magma to OH, but Silver
et al. (1990) have shown that maximum solu-
bility is controlled by saturation equilibrium with
its vapour phase. Isotopic signatures of magmatic
water range from highly depleted values in ridge
and intraplate settings (δ
2
H=−100; Taylor and
Sheppard 1986) to ca. δ
2
H=−20 in subduction
zone environments (Taran et al. 1989; Giggen-
bach 1992).
Carbon
CO
2
is the predominant carbon species in volcanic
gases. It is the least soluble of the major constit-
uents, and is probably already saturated with
respect to a separate vapour phase at deep crustal
depths (Anderson 1975; Wallace 2005). Primary
basaltic MORB and hot spot melts are inferred to
have rather similar CO
2
contents of 0.6 and 0.7 wt
% respectively (Gerlach et al. 2002; Fischer et al.
2005), whereas arc basalts are comparatively CO
2
enriched, with contents ranging between 0.6 and
1.3 wt% CO
2
(Wallace 2005). As for water, this is
attributed to the very efficient recycling of sub-
ducted C in arc systems (e.g., Fischer 2008). CO
2
is present in the melt phase as CO
3
–
, and Fine and
Stolper (1986) have suggested that CO
2
degassing
128 B. Christenson and F. Tassi

has an oxidising effect on the remaining melt
phase, as described by the following reaction:
CO
3þHþþ0:5H
2¼CO2þH2O:ð1Þ
CO is a comparatively minor species charac-
terised by relatively fast reaction kinetics with
CO
2
. As such, it is particularly useful as both a
redox and temperature geoindicator in magmatic-
hydrothermal environments (e.g., Giggenbach
1987; Chiodini and Cioni 1989).
δ
13
C signatures of volcanic C are variably
influenced by three main source signatures,
depending on tectonic setting. Mantle signatures
of δ
13
C=−6.5 ±2.5 ‰are typically encountered
in rift margin and hot spot centres, whereas
organic material (δ
13
C≤−20 ‰) and carbonate
signatures (δ
13
C*0‰) typically contribute to
subduction zone derived magmatic CO
2
(Sano
and Marty 1995). In addition, δ
13
C signatures are
also affected by melt-vapour fractionation, where
the heavier isotope systematically favours the
vapour phase (Mattey 1991). While this process
has the potential to mask original isotopic source
signatures, such fractionation may also provide
insights into the nature and extent of batch
degassing processes, including potential recog-
nition of new pulses of magmatic degassing (as
at Mt Etna, Chiodini et al. 2011).
Sulfur
Sulfur is a particularly important component in
volcanic lake systems, where it typically plays a
major role in mediating the redox state of the
proximal magmatic-hydrothermal environments
(Giggenbach 1987), but its deposition in ele-
mental form can also become an important con-
trol on permeability therein (Hurst et al. 1991;
Oppenheimer 1992; Christenson 1994). Thor-
ough reviews of S chemistry in magmatic envi-
ronments have been provided by Wallace and
Edmonds (2011) and Oppenheimer et al. (2011),
for which only a few key aspects relevant to
volcanic lake environments are presented here.
S behaviour in magmas is complicated by the
fact that it may exist over a range of valence states
(−2, 0, +4 and +6). In the melt phase, therefore,
S may speciate across S
2
°,SO
3
2−
and SO
4
2−
,
according to the oxidation state imposed by Fe in
the melt-mineral system (Carmichael and Ghiorso
1986). This variation is reflected in the range of S
gas species which may coexist with the melt,
including H
2
S, S
2
,SO
2
,SO
3
and OCS.
S solubility in silicate melts is limited by the
occurrence of non-volatile S phases, including
Fe–S–O immiscible liquids and sulphide miner-
als under reducing conditions, and sulphate
minerals under more oxidising conditions (Wal-
lace and Edmonds 2011), with melt-phase solu-
bility generally increasing with oxidation state. S
solubility is also linked to total FeO content,
ranging from 800 to 1,200 ppm in MORB basalts
with 8–10 wt% total FeO, to 2,000–2,400 ppm at
15–16 wt% FeO (Wallace and Carmichael 1992).
In more evolved arc magmas, however, S shows
minor inverse solubility with silica contents
(Scaillet and Pichavant 2003).
Thermodynamic modelling of primitive
MORB melts at 1,280 °C shows that the total
abundance of S gas species varies directly with
redox state of the melt, with the sulphur content
of the gas phase ranging from ca. 0.05 to over
40 % for NNO values of −3 to +1 respectively
(Oppenheimer et al. 2011). These authors also
found that SO
2
is the predominant S-bearing gas
species in equilibrium with the melt above
NNO = −2.5, with the next most prevalent spe-
cies being S
2
.H
2
S becomes predominant only
below NNO *−2.5.
The δ
34
S signature of bulk mantle S is similar
to that of the internationally accepted standard
value derived from the Canyon Diablo Troilite
meteorite. MORB sulfides, for example, cluster
around −0.3 ±2.3 ‰VCTD, whereas OIB
sulphides are +1.0 ±1.9 ‰VCTD (Marini et al.
2011). However, there is strong, temperature-
dependent fractionation of
34
S between SO
2
and
H
2
S (+10 to +2 over the temperature range 300–
1,000 °C; (Friedman and O’Neil 1977), making
the δ
34
S isotope systematics a powerful tool for
understanding volcanic lake environments (e.g.,
Marini et al. 2011; Kusakabe et al. 2000;
Delmelle and Bernard this volume).
Gases in Volcanic Lake Environments 129

Halogens
The halogens constitute the last major compo-
nent group in volcanic emissions affecting vol-
canic lakes, and include in decreasing order of
abundance, Cl, F, Br and I. While both Cl and F
are relatively abundant in crustal rocks with
mean concentrations of 550 and 240 ppm
respectively, Br and I are typically present at
levels less than 1 ppm. Cl is the dominant halo-
gen in volcanic emissions and crater lakes, and it
is an important conservative constituent in
hydrothermal systems generally. As such, our
very cursory discussion here will be limited to
this single specie, although its general behaviour
remains applicable to the halogens as a whole,
and the reader is referred to the vast literature on
halogen behaviour in magmatic systems (e.g.,
Aiuppa et al. 2009; Pyle and Mather 2009, and
references therein).
Cl concentrations in magma show consider-
able variation between differing tectonic regimes
and melt compositions. Cl in basalts from ridges
and oceanic islands have Cl contents gener-
ally <800 ppm. However, Cl is essentially an
incompatible component in silicate melts, tend-
ing to fractionate into the remaining melt frac-
tions during crystallisation (e.g., Anderson
1975), as evident in the occurrence of melt
inclusions in calc-alkaline volcanics with Cl
contents exceeding 7,500 ppm, and high silica
peralkaline magmas up to 1.2 wt% (Aiuppa et al.
2009). Once magmas become saturated with
respect to H
2
O, however, Cl is strongly frac-
tionated into the fluid phase at supercritical
conditions (Kilinc and Burnham 1972). With
transition to sub-critical conditions, this fluid
becomes immiscible, with Cl strongly fraction-
ating into the liquid fraction to form NaCl-rich
brine and a low-density vapour phase (Shinohara
et al. 1989). Further pressure decline leads to the
brine becoming increasingly hypersaline (see
Henley this volume), ultimately entering the
halite + vapour region in PTX space (e.g.,
Shinohara and Fujimoto 1994). It is typically this
associated aqueous vapour, in conjunction with
the aforementioned volatiles species, which
constitutes the magmatic volatile phase on active
volcanoes. The predominant Cl species in vol-
canic emissions is HCl (e.g., Symonds et al.
1988), with high temperature fumarolic emis-
sions from andesitic volcanoes (>500 °C) typi-
cally ranging between 1,000 and 14,000 mmol/
mol (Giggenbach 1996).
2.1.2 Minor and Trace Gases
In volcanic systems, N
2
is typically a non-reactive
component (Giggenbach 1996) derived from one
of three main component sources. These include:
primordial mantle gas, N derived from subducted
sedimentary/organic material along convergent
margins (Hilton et al. 2002;Fischer2008) and also
that contained in meteoric waters convectively
swept into volcanic hydrothermal environments
(Giggenbach 1987,1996).
15
Nsignaturesofthe
upper mantle, as represented by MORB analyses,
have δ
15
N
2
values uniformly in the vicinity of −5.0
(i.e., depleted in
15
N relative to air, Marty and
Humbert 1997), whereas sediment-derived N
2
is
distinctly heavier, ranging from +5 to +7 (Cartigny
and Ader 2003). Interestingly, deep mantle signa-
tures from mantle plumes have been shown to be
quite variable (−5 to +5), raising questions about
deep mantle convection and mixing of subducted
sediments (Marty and Dauphas 2003), although
others argue that N is efficiently recycled at sub-
duction zones (i.e., volatilised) from subducted
slabs and related sediments (Fischer et al. 2002).
Numerous other trace-level gases are found in
volcanic emissions, including most notably, the
noble gases which serve as useful tracers (Hilton
and Porcelli 2013), metal species (e.g., Symonds
et al. 1987,1992; Henley this issue), and a
multitude of trace and ultra-trace species derived
from interactions between major component
sources (e.g., Symonds and Reed 1993).
2.1.3 Gas Solubility in Magma
A general theoretical treatment relating solubility
of the major gases to total confining pressure in
andesitic melts was presented by Giggenbach
(1996), and his summary diagram is reproduced
in Fig. 2. In its formulation, Henry’s constants
130 B. Christenson and F. Tassi

were derived from Ostwald coefficients presented
by Zhang and Zindler (1989), and were used to
assess the solubility of the predominant gases in
White Island (New Zealand) magmas. Total gas
pressure is derived from summing the partial
pressures of component species via:
Pt¼Rci;o=RvQi;Tþ1

Ki;T;ð2Þ
where c
i,o
is the initial concentration of species i,
R
v
is the vapour-melt volume ratio, Q
i,T
is the
Ostwald coefficient for species i, and K
i,T
is the
temperature dependent Henry’s law constant.
Solute gas contents are plotted in Fig. 2as
functions of vesicularity, V
v
, (volume %) which
is related to R
v
by:
Vv¼100 Rv=ð1þRvÞ:ð3Þ
Assuming single step vapour separation, the
curves in Fig. 2show that CO
2
has low solubility
in the melt phase, comparable to the noble gases,
and is >90 % evolved by the time vesicularity
approaches 1 %. At 50 % vesicularity, all of the
noble gases, CO
2
and some 70 % of the water are
exsolved, but just 20 % of total S and virtually
none of the HCl has entered the vapour phase.
From these relations it is clear that depth of
degassing and decoupling of gas from magma are
critically important controls on the composition
of gases reaching volcanic lake environments.
Distal magma sources will result in gases com-
prised predominantly of CO
2
, as in the case of
maar lake systems, whereas shallow degassing
will provide the full complement of magmatic
volatiles and heat in the lake systems.
Another interesting insight to be gained from
Eq. 3is also portrayed in Fig. 2. Here, considering
CO
2
and H
2
O to be the predominant pressure
generating species in the melts allows the total
vapour pressure to be calculated as a function of
R
v
, and thereby the depths at which vapour-melt
separation may begin. For melts with CO
2
con-
tents of 10,000 mg/kg, the calculated depth
exceeds 80 km, suggesting that phase separation is
probably already occurring close to the depths of
magma generation. It is clear from Fig. 2that the
depth of gas-melt decoupling places a first order
constraint on the relative compositions of the
principle gas species reaching the surface, and
thereby the chemical characteristics of volcanic
lakes into which they discharge. Evidence sug-
gests that vapour segregation readily occurs in the
plumbing systems of erupting basaltic volcanoes
(Aiuppa et al. 2010; Allard 2010), through the
generation of volatile-rich foams (Vergniolle and
Jaupart 1990; Sparks 2003), or through separation
processes invoked by branching conduit geome-
tries as proposed for Etna (Burton et al. 2003), or
Mount Cameroon (Suh et al. 2003). What this
means for volcanic lakes, is that deeply decoupled
volatile streams will consist largely of CO
2
(plus
noble gases), whereas shallow decoupling will
contribute CO
2
, S species and halogens directly
into the lake environment.
It is well recognised that many volcanoes emit
large quantities of gas in the absence of significant
eruption activity, including Etna (Aiuppa et al.
2007); Popocatépetl (Delgado-Granados et al.
2001), Ambrym (Bani et al. 2009), Miyakejima
(Shinohara et al. 2003) and a wide range of less
prodigeous emitters, some including volcanic
lakes (e.g., Aso, Shinohara et al. this issue), White
Fig. 2 Gas solubility in magma as a function of
vesicularity (after Giggenbach 1996). CO
2
is relatively
insoluble in melts, having a bulk solubility similar to that
of the noble gases
Gases in Volcanic Lake Environments 131

Island (Werner et al. 2008), Poás (Rowe et al.
1992), and Ruapehu (Christenson 2000, Chris-
tenson et al. 2010). In these cases, the amounts of
emitted gas well exceed those which could be
derived from associated erupted magmas, if any.
So the question is, by what mechanism(s) do the
gases reach the surface?
Bubble accumulation and coalescence to form
foams, with subsequent shearing and collapse to
form permeable channel-ways for gas migration,
have been proposed to explain differential gas
transfer and observed slug-flow and lava foun-
taining phenomena (Jaupart and Vergniolle 1988;
Vergniolle and Jaupart 1990). Sparks (1978)
modelled the diffusive transfer of water vapour
through ascending basaltic melts and found that
bubble diameters ranged between 0.1 and 1.0 cm,
consistent with the majority of observed scoria
bubble diameters. However, for a typical basalt
melt viscosity range of between 30 and 3 Pa.s,
bubble diameters of between 14.5 and 4 cm,
respectively, are required for bubbles to move
significantly faster than the magma (Sparks 2003),
suggesting that other transfer mechanisms are at
play in the formation of foams.
Carrigan (1983) proposed forced convection
within magma conduits of radii as small as 5 m
diameter, as a means of offsetting heat loss and
attendant solidification of magma in feeder con-
duits for active volcanoes. The model incorpo-
rates buoyant rise of vesiculated melt to some
level in the volcano conduit, where decoupling
takes place and gases either accumulate, or are
released into the surrounding hydrothermal
environment, after which the denser, degassed
magma convectively sinks back to the source.
Since then, compelling evidence has emerged for
magmatic convection being the principle driver
for magmatic degassing at numerous volcanoes
(e.g., Kazahaya et al. 1994; Allard 1997; Ste-
venson and Blake 1998; Shinohara et al. 2003;
Oppenheimer et al. 2004). By this means, we
have at least one mechanism for continuous, or
near continuous delivery of magmatic volatiles to
the hydrothermal environments associated with
volcanic lakes.
2.2 Hydrothermal Gases
and Processes
As previously described, the hydrothermal envi-
ronment in magmatic-hydrothermal systems
effectively begins at the transition from single
phase vapour to the 2-phase vapour-liquid
envelopes surrounding a magma conduit (Fig. 1).
The gases involved in this environment consist of
the aforementioned magmatic inputs, the mete-
oric gases (of atmospheric origin, primarily N
2
,
O
2
, Ar), and gases formed through heteroge-
neous reaction of these species with the host
rocks.
2.2.1 Redox State
A most remarkable feature of proximal mag-
matic-hydrothermal environments is the wide
range of redox conditions found within, a vari-
ability which owes its existence to the interplay
between the two principle redox buffers. A
measure of redox state, as proposed by Giggen-
bach (1987), is the readily measurable ratio of
fH2=fH2o(with parameter R
H
defined = log
(fH2=fH2o)). Gases leaving magmas typically have
their redox state buffered by the temperature- and
pressure-dependent equilibrium between H
2
S
and SO
2
, the so-called “magmatic gas buffer”of
Giggenbach (1987):
H2Sþ2H2O¼SO2þ3H2:ð4Þ
At temperatures close to 1,000 °C, this buffer
closely approaches the FMQ (i.e., fayalite-mag-
netite-quartz) buffer, which is generally accepted
as being the primary redox control in silicate melts
Carmichael and Ghiorso (1986). At this temper-
ature R
H
is approximately −2.0. On departure
from the magma, and with decreasing tempera-
ture, however, the gas buffer drives the redox state
of the gas stream to ever higher oxidation potential
(decreasing R
H
), where at temperatures
approaching ca. 300 °C, the redox potential of the
gas stream is very close to that governed by
magnetite-hematite equilibrium at 1 bar pressure.
132 B. Christenson and F. Tassi

As described by Giggenbach (1987), observed
deviations of gas compositions away from the
magmatic gas buffer (i.e., toward lower oxidation
potentials with decreasing temperature) is nor-
mally attributable to interaction with Fe
2+
and Fe
3
+
in silicate rocks, whose redox potential is well
represented by the reaction between fayalite (a
proxy for FeO) and hematite according to:
Fayalite þH2O¼Hematite þQuartz þH2;
ð5Þ
which has a nearly temperature independent R
H
value of *−2.8. Buffering by rocks in this
environment effectively converts SO
2
in the
hydrothermal environment to H
2
S.
As we shall see below, another reaction which
indirectly affects the redox state of the magmatic-
hydrothermal environment is that governing the
formation of the elemental sulphur from the
magmatic gas stream:
2H2SþSO2¼3Seþ2H2Oð6Þ
Here, the forward reaction depletes the vola-
tile stream of H
2
S, driving Eq. (4) to the left,
effectively lowering fH
2
.
As in the magmatic realm, CO
2
is also the
predominant C-gas in the hydrothermal envi-
ronment. Here it enters into temperature- and
redox-sensitive equilibria with both CO and CH
4
(Giggenbach 1980,1987; Taran 1986; Chiodini
et al. 1993). The reaction kinetics for:
CO þH2O¼CO2þH2ð7Þ
are demonstrably fast, and have been shown to
carry the redox signature of either the gas or rock
buffers within the magmatic-hydrothermal envi-
ronment (Giggenbach 1987; Chiodini et al.
1993), and the reaction readily equilibrates in
either the liquid or vapour phase. The reaction:
CH4þ2H2O¼CO2þ4H2;ð8Þ
on the other hand, has much slower kinetics, and is
typically buffered by the FeO–FeO
1.5
rock buffer
within the liquid-phase environment marginal to
high temperature conduits (Taran and Giggenbach
2003). As such, CH
4
may be regarded as both a
tracer and component species of the hydrothermal
equilibrium environment (Chiodini 2009).
A good example of the interplay between
hydrothermal and magmatic source environments
is found on White Island (Fig. 3) where the rel-
ative variations amongst N
2
,CH
4
and CO
2
in a
single fumarole show a progressive shift towards
magmatic signatures since the late 1970s through
to the present time. Three end-member compo-
nents can be described in terms of their relative
N
2
/CH
4
ratios, including a N
2
-rich meteoric
component, a CH
4
-rich hydrothermal gas and a
relatively CH
4
-depleted magmatic endmember.
N
2
/CH
4
ratios close to 0.5 in the late 1970s trend
towards a CH
4
-depleted signature of N
2
/
CH
4
> 30 by 2004, during the magmatically
active period of the volcano.
Heavier hydrocarbons of thermogenic deriva-
tion are also typically present as either trace or
ultra-trace species in hydrothermal environments
associated with magmatic systems such as White
Island (e.g., Taran and Giggenbach 2003), and
Vulcano Island (Tassi et al. 2012), with alkanes
Fig. 3 Variability in hydrothermal and magmatic end-
member components in a low temperature fumarole from
White Island, New Zealand. CO
2
/CH
4
ratios for the
magmatic and hydrothermal endmembers are ca. 50
and >50,000 respectively
Gases in Volcanic Lake Environments 133

generally predominating in the rock-buffered
environments, although varying mixtures of
alkenes and alkanes typically occur in the more
oxidising (i.e., H
2
S–SO
2
gas-buffered) environ-
ments. The species are typically thermogenic in
origin, derived from kerogen cracking in the
hydrothermal environment (Taran and Giggenbach
2003).
NH
3
bears a similar equilibrium relationship
to N
2
as CH
4
and CO have to CO
2
(Eq. 8)in
hydrothermal environments, i.e. it is present in
subordinate concentrations relative to the main
species and therefore behaves more as an indi-
cator rather than an iso-molar redox buffer Gig-
genbach (1980). The equilibrium reaction is:
2NH3¼N2þ3H2ð9Þ
The two equilibria share other characteristics
as well, including equilibration under reducing
conditions, and similarity in their relative rates of
equilibration which are relatively slow in com-
parison to H
2
–H
2
O, CO–CO
2
and H
2
S–SO
2
(Giggenbach 1987).
2.2.2 Aqueous Solubility of Gases
in the Hydrothermal
Environment
All gases are soluble in water to varying degree.
At relatively low gas pressures, the aqueous
concentration of any gas are conveniently
described by Henry’s Law,
xi¼Pi=KH;ið10Þ
where x
i
is the molality, P
i
is the partial pressure
(bar), and K
H,i
is the temperature dependent
Henry’s Law constant (with units of bar/molal)
for species i. Where the total pressure of volatile
species exceeds the confining pressure (typically
hydrostatic in hydrothermal environments), gases
will partition between vapour and liquid phases
according to:
Bi¼xi;v=xi;l¼zvKH;i=PH2Oð11Þ
where B
i
is the vapour-liquid distribution coeffi-
cient, and z
v
and PH2Oare the compressibility and
pressure of water vapour, respectively.
Henry’s Law constants for the major gas
species found in volcanic hydrothermal envi-
ronments are shown as functions of temperature
in Fig. 4(as bar·molal
−1
, using regressions of
Fernandez-Prini et al. (2003), Naumov et al.
(1974) and Truesdell et al. (1989). Two broad
solubility groups are apparent in these data. The
smaller, lighter and/or mono-atomic gas species
(e.g., He, Ar, H
2
,N
2
,O
2
, CO and CH
4
) are less
soluble than their larger and/or heavier molecular
counterparts (CO
2
,H
2
S, HCl, SO
2
and NH
3
), and
the former typically shows aqueous solubility
minima between 0 and 100 °C.
The more soluble protonated species typically
exhibit dissociative solvent-ion interaction,
leading to pH-dependent speciation, and this
somewhat complicates their behaviour as vola-
tiles in the magmatic-hydrothermal environment.
For example, CO
2
undergoes temperature- and
pH-dependent speciation across H
2
CO
3
°, HCO
3
−
and CO
3
2−
.At20°C and pH < 4, for example, it
exists almost entirely as hydrated molecular CO
2
and H
2
CO
3
(e.g., Stumm and Morgan 1981), and
behaves effectively as an inert species. Where pH
may be controlled by mineral buffers at levels >4,
dissociation of H
2
CO
3
provides a source of
protons for hydrolysis reactions between fluid
and rock Giggenbach (1981). H
2
S is similar in
behaviour, with a pK
1
of ca. 7.5, whereas
Fig. 4. Henry’s law constants for typical gases in
volcanic lakes. Units are bar/molal. Reference data are
from: Fernandez-Prini et al. (2003), Naumov et al. (1974),
and Truesdell et al. (1989)
134 B. Christenson and F. Tassi

ammonia predominates over ammonium only at
pH levels > ca. 9.
HCl is a gas with bulk solubility similar to
that of H
2
S. It is, however, a strong electrolyte
which becomes significantly associated only at
very low pH and/or temperatures approaching
the critical point Ruaya and Seward (1987).
Therefore, contact between high temperature
HCl-bearing vapours and liquid water generally
leads to dissolution and dissociation of the HCl,
potentially leading to the formation of highly
acidic solutions (e.g., Truesdell et al. 1989).
These processes are discussed further below.
2.3 Meteoric Gases
As already mentioned, the meteoric constituents
are composed of atmospheric gases introduced
into the hydrothermal environment by air satu-
rated rain water entering into groundwater sys-
tems, or in the case of volcanic lakes, by direct
equilibration with air which is in contact with the
lake water. The principle components are, not
surprisingly, those of highest partial pressure in
the atmosphere (i.e., N
2
,O
2
and Ar), and their
equilibrium concentrations in groundwater are
proportional to their temperature-dependent
Henry’s Law solubilities. As an example, two
component mixing between atmospheric and
magmatically derived end-members is clearly
evident in solute N
2
, Ar and CO
2
compositions
from Ruapehu Crater Lake (Fig. 5; data from
Christenson et al. submitted). Here fumarolic
discharges from both Ruapehu and White Island
represent the magmatic-hydrothermal end mem-
ber, whereas the theoretical meteoric component
signatures are calculated for temperature ranging
between 0 and 40 °C using the Henry’s constants
of Fernandez-Prini et al. (2003). CO
2
/N
2
ratios
vary between ca. 40 and 3 over the 3 year-long
period represented by the data, and reflect vary-
ing emissions from the magmatic system on the
volcano. Interestingly, projection back to the N
2
–
Ar axis reveals a N
2
/Ar ratio of ca. 55, well
above the meteoric range of 36–40. This higher
value most probably represents the N
2
-enriched
magmatic component which is prevalent on NZ
arc-type hydrothermal environments (e.g.,
Christenson et al. 2002), with the N
2
/Ar ratio
of *1,000 for Ruapehu (Christenson 2000).
2.4 Biogenic Gases
The distribution of biomass along the vertical
water column of a lake is controlled by water
chemical-physical conditions, i.e. temperature,
pH and redox. Biomass is relatively abundant in
the epilimnion, decreases to a minimum in the
metalimnion and increases in the lower anoxic
hypolimnion (Niewolak 1974; Jones 1978; Kato
and Sakamoto 1981; Coveney and Wetzel 1995;
Simon 1998). Irrespective of their primary origin,
CO
2
,CH
4
,N
2
,O
2
, S gases and H
2
dissolved in
volcanic lakes are involved in biogeochemical
processes (Mapelli et al. this issue). As shown in
the following sections, lakes hosted in quiescent
volcanoes, characterized by relatively low tem-
perature, almost-neutral pH and establishment of
meromictic conditions (i.e., a permanent thermal
and chemical stratification), the vertical distri-
bution of these gases and that of microbial pop-
ulations are intimately related.
2.4.1 CO
2
Although dissolved CO
2
in volcanic lakes are
typically fed by sub-lacustrine gas discharges
(Sigurdsson et al. 1987; Tazieff 1989; Aeschbach-
Hertig et al. 1999; Tassi et al. 2009), the fate of
dissolved CO
2
strongly depends on addition-
consumption processes related to biotic respira-
tion, anaerobic decomposition of organic matter
and microbial oxidation of CH
4
(Rudd et al. 1974;
Rich 1975,1980), especially in lakes hosted in
quiescent volcanoes. In an aerobic environment,
i.e. in the shallower portions of crater lakes, mic-
roalgae and cyanobacteria along with higher
plants, are capable of CO
2
-consuming oxygenic
photosynthesis Nelson and Ben-Shem (2004),
which proceeds following two main pathways:
light energy conversion to biochemical energy by
a photochemical reaction, and CO
2
reduction to
Gases in Volcanic Lake Environments 135

organic compounds such as sugar phosphates,
through the use of this biochemical energy by
Calvin-cycle enzymes. In the hypolimnion, CO
2
biogenic and geogenic inputs are counteracted by
microbial reduction processes mainly by metha-
nogens (Schoell 1988; Whiticar 1999), a group of
microorganisms phylogenetically affiliated to the
kingdom Euarchaeota of the domain Archaea
Woese et al. (1990). Anaerobic methanotrophy
coupled to Fe or Mn reduction Valentine (2002)
and using nitrates as substrates Raghoebarsing
et al. (2006) can also occur, although most of CH
4
oxidation is carried out in the epilimnion (Hanson
and Hanson 1996; Lopes et al. 2011). The com-
bination of loss and addition of CO
2
related to the
various biogeochemical processes occurring at
different depth along the water vertical column
maintain stable CO
2
-rich reservoirs typically
characterizing Nyos-type meromictic volcanic
lakes.
2.4.2 CH
4
Methanogenic processes are active within sedi-
mentsofanoxichypolimniaofmeromicticlakes
Rudd and Taylor (1980). In the presence of free
oxygen, particularly at oxic/anoxic boundaries,
CH
4
is easily consumed by methanotrophs phylo-
genetically affiliated to the a,hand gsubdivisions
of kindom Proteobacteria in the domain Eubacteria
(Hanson and Hanson 1996). Methanogenic pro-
cesses proceeds through carbonate-reduction and
acetate fermentation pathways Schoell et al. (1988),
which can be described by the following reactions:
CO2þ8Hþþ8e!CH4þH2Oð12Þ
and
CH3COOH !CH4þCO2ð13Þ
where the * indicates the intact transfer of the
methyl position to CH
4
. Reaction (12) prevalently
Fig. 5 Relative solute CO
2
–N
2
–Ar contents in Ruapehu
Crater lake water, showing mixing relations between
magmatic and meteoric lake water components. Ruapehu
source gas composition is taken from Christenson (2000),
and bears close similarities to gases to values observed
from White Island (Giggenbach and Sheppard 1989), both
having N
2
/Ar ratios ≥1,000. The meteoric endmember
component is N
2
-enriched over the range of theoretical
equilibrated air-saturated water compositions (calculated
for temperatures of 0–40 °C), reflecting excess N
2
in the
source gas. Open circles reflect time series data collected
over the period 2007–2012, and open triangles represent
multiple samples collected by boat over a single day in
2010
136 B. Christenson and F. Tassi

occurs in sulfate-free marine sediments, whereas
in freshwater environments the two processes are
competitive (e.g., Takai 1970; Winfrey et al.
1977). Methane produced by microbial methyl-
type fermentation is typically
2
H-depleted (δ
2
Η–
CH
4
<−250 ‰V-SMOW) and
13
C-enriched
(δ
13
C–CH
4
>−70 ‰V-PDB) with respect to that
produced by bacterial CO
2
reduction Whiticar
et al. (1986), although these boundary values may
vary depending on the maturity and type of the
organic source Whiticar (1999).
2.4.3 N
2
The triple–bounded N
2
molecule is involved in
biological processes to produce reactive N
2
-
bearing compounds, such as NO
x
and NH
x
.
Therefore, the distribution of N
2
concentrations
in stratified lakes is controlled biological pro-
cesses able to fix or produce N
2
. Biological N
2
fixation, which depends on light (Tison et al.
1977), and presence of bioavailable trace metals
(Hysenstrand et al. 1988) is carried out in water,
on the sediment surface and in sediment pores by
heterocyst-forming species such as cyanobacteria
(Loeb and Reuter 1981; Valelia 1991) and
methane-oxidizing bacteria (Rudd and Taylor
1980). Denitrification, i.e. the reduction of NO
3
−
to NO
2
−
and then to N
2
(Ahlgren et al. 1994)
commonly occurs at anaerobic conditions. In this
process, NO
3
−
is used as an electron acceptor
during respiration of bacterial species such as
Pseudomonas and Clostridium. Bacterial N
2
production may also occur through NH
4
oxida-
tion with NO
2
−
as electron acceptor, a process
termed as anammox (Anaerobic Ammonium
Oxidation) that is carried out by the bacteria
phylum Planctomycetes (Jetten et al. 1998).
2.4.4 O
2
Oxygen concentrations in permanently stratified
volcanic lakes typically decrease with depth, since
this gas is depleted rapidly by oxidative processes.
This produces a vertical O
2
profile termed clino-
grade, where aerobic epilimnion and anaerobic
hypolimnion can be clearly distinguished. Oxy-
gen consumption, which is primarily due to
biological oxidation of organic matter, intensively
occurs at all the depths, but in the epilimnion is
frequently offset by water circulation and photo-
synthesis, two O
2
renewal mechanisms that are
not active in the hypolimnion. Pure chemical
oxidation and photochemical oxidation induced
by ultraviolet light may also significant contribute
to oxygen depletion of deep lake waters (Laane
et al. 1985; Wetzel et al. 1995).
2.4.5 H
2
Cyanobacteria produce molecular H
2
through
both photosynthesis and anaerobic fermentation
processes (Asada and Kawamura 1986; Asada
and Miyake 1999). Hydrogen biogenesis and
consumption is related to metabolic activity of
several enzymes: nitrogenase catalyzes H
2
pro-
duction concomitantly with the reduction of N
2
to NH
4
, whereas hydrogenase is able to both take
up and produce H
2
(Bergman et al. 1997; Tam-
agnini et al. 2002). At anaerobic conditions, H
2
is
involved in mineralization processes of organic
matter, being used as electron donor by metha-
nogenic and sulfate-reducing bacteria (Zehnder
1978; Thaurer and Badziong 1980). These pro-
cesses mainly occur at the water-sediment inter-
face, thus H
2
slowly diffuses from the lake
bottom sediment toward the surface is efficiently
consumed before arriving at the lake surface.
3 Chemical and Physical Models
of Sub-lake Environments
As the foregoing suggests, our understanding of
the chemical interaction between ascending
magmatic heat and volatiles and Earth’s hydro-
sphere is reasonably well advanced, despite the
fact that much of what goes on in these envi-
ronments is inaccessible to us at the Earth’s
surface. Much of our early understanding stem-
med from time series studies of crater lake sys-
tems during volcanically active periods (e.g.,
Giggenbach and Glover 1975; Delmelle et al.
2000; Rowe et al. 1992). Such time series data
have been used to constrain numerical simula-
tions of these systems, including reaction path
Gases in Volcanic Lake Environments 137

models for equilibrium gas and water chemistry
(Christenson and Wood 1993), and more
recently, heat-mass transport and reactive trans-
port models of these same systems (Chiodini
et al. 2003; Todesco et al. 2003; Christenson
et al. 2010; Christenson and Young 2010). In the
next section we briefly review results from two
such models.
3.1 Gas-Groundwater Interaction
The absorption of a rising magmatic vapour into
groundwater (Fig. 6) can be readily simulated
with the reaction pathway algorithm REACT
Bethke (1996). In this example, 10 mol of a high
temperature fumarolic gas (White Island fuma-
role F3; Table 1) are numerically titrated over
100 steps (i.e., 0.1 mol total gas added per step)
into 1 kg of the relatively dilute Silica Rapids
spring water which issues from the flank of Mt
Ruapehu (Table 2). Since REACT has no
enthalpy constraint, temperatures are monotoni-
cally increased from 20 to 300 °C during the run
(2.8 °C per step) to simulate heating that might
occur through a combination of condensation of
the hot vapour into the groundwater and con-
ductive heating in the conduit region. All gases
are taken into aqueous solution during the titra-
tion, i.e., there is no vapour-liquid partitioning.
System Eh and pH are plotted in Fig. 6a. The
spring water has a near-neutral pH initially, and
oxygen levels were set to be in equilibrium with
atmosphere (*6 ppm O
2
). There are rapid neg-
ative shifts in both Eh and pH in the first reaction
steps which, not surprisingly, demonstrates the
low buffering capacity of the dilute spring water.
pH decreases to a minimum of 1.3 at ca. 250 °C,
but then climbs back up to ca. 1.5 with continued
heating to 300 °C. Conversely, the abrupt decline
of system Eh to a minimum value of *+0.1 V in
the first step is followed by a progressive
increase to a maximum *+0.28 V at 270 °C,
before declining to +0.25 V at 300 °C. These
changes correspond closely to the appearance
and temperature-controlled stability of elemental
sulphur in the system (Fig. 6b). Elemental sul-
phur precipitation initiates at the outset, and
continues through to approximately 240 °C, after
which it begins to redissolve. Temperature
appears to be the controlling factor in sulphur
stability in this system, which in turn controls the
redox state of the fluid.
Fig. 6 Reaction path simulation for the titration of
10 mol of White Island fumarolic gas into 1 kg of Silica
Rapids spring water (Mt Ruapehu), with monotonic
heating from 20 to 300 °C. apH and Eh variations with
reaction progress (each symbol represents 1/100 of the
reaction progress; bMinerals formed during the reaction
sequence (moles); cAqueous S speciation; dAqueous C,
F and Cl speciation with reaction progress
138 B. Christenson and F. Tassi

The generation of a strongly acidic liquid phase
is perhaps the most notable aspect of this gas-
water interaction. As mentioned previously, and
shown in Fig. 6c, d, HCl and H
2
SO
4
are the
principle (strong) acids, and these are the primary
hydrolysis-promoting agents involved in ensuing
water-rock interaction in these environments. It is
evident that temperature is ultimately a limiting
factor on acidity, owing to the temperature-
dependence of solvent-ion interaction where, as
the dielectric constant for water decreases with
temperature (Uematsu and Franck 1980), acids
become increasingly associated (e.g., HCl; Ruaya
and Seward 1987). It is this effect alone that is
responsible for observed increases in pH above
250 °C in this comparatively simple model, and it
is a trend which continues to where acids become
fully associated at the point. Owing to the fact that
H
2
CO
3
is a weaker acid than either HCl, H
2
SO
4
or
even HF, it remains almost completely associated
in this model environment owing to the very low
pH. Whereas CO
2
is the predominant hydrolysis-
promoting agent in near-neutral pH hydrothermal
systems, and may in fact become buffered by the
presence of calcite in such systems (Giggenbach
1981; Arnórsson 1985; Arnórsson and Gunnl-
augsson 1985), in highly acidic environments
such as those found in active crater lakes, CO
2
behaves as a conservative constituent, similar to
any other non-reactive gas.
3.2 Gas-Lake Water-Rock Interaction
Christenson et al. (2010) simulated gas-water-
rock interactions in the Ruapehu vent environ-
ment using the 1D advective reactive transport
model X1t Bethke (1996). Here, we have further
refined the model to examine the effects of water-
rock interaction on key indicator gases in the vent
hydrothermal environment. The conceptual model
for this condensed vapour-water-rock interaction
is shown in Fig. 7, where a column of permeable
substrate (35 % porosity) measuring 10 cm
2
, and
10 m in length is initially filled with Ruapehu
Crater Lake water (Table 2), and is allowed to
equilibrate at 50 with the reactant minerals listed
in Fig. 7(as volume %). This scenario is similar to
what might be found in the vent region during
periods of quiescence. We then numerically inject
a 300 °C magmatic condensate proxy consisting
of 0.5 molal concentrations of both H
2
S(g) and
SO
2
(g), and 1 molal concentrations of HCl(g) and
CO
2
(g) under a head pressure of 0.2 MPa for 24 h
into the column. A fundamental assumption of
this model is that full equilibrium is established at
each time step. Whereas the reality is that reaction
kinetics will play an important, if not dominant,
role in governing which chemical reactions pre-
vail in these environments (see Henley this issue),
our knowledge of reaction rates is not well enough
advanced to rigorously account for kinetics in the
modelling of such complicated systems as pre-
sented here. In our case, we consider the concept
of full equilibrium as perhaps a theoretical end-
point in the process. The relatively low volume
percentage of the six reacting mineral phases is
chosen here to simulate the actual amount of
reactive surface that permeating liquid phase
would interact with, taking into account the effects
of armouring of pore and fracture surfaces by
precipitated minerals in the reaction column.
The initial alteration mineral assemblage is of
argillic nature, owing to the acidity of the starting
fluid, but is uniform along the length of the
Table 1 White Island fumarolic #3 composition from
White Island (Giggenbach and Sheppard 1989)
Date 31-8-78
T
m
(°C) 540
H
2
O 880,000
CO
2
84,320
S
t
18,600
N
2
1.9
HCl 4,080
HF 304
NH
3
20
He 0.68
H
2
2,100
Ar 0.86
O
2
<0.5
N
2
458
CH
4
2.2
CO 48.4
Units are μmol/mol
Gases in Volcanic Lake Environments 139

Table 2 Water compositions used in the numerical simulations
Date Feature Tm °C pH HCO
3
Li Na K Ca Mg Cl SO
4
B SiO
2
H
2
SFe Al As Br F NH
3
20-May-11 Silica rapids 8.2 6.82 47 <0.01 13.8 3.1 26 13.0 8.0 98 <0.3 35 <0.01 <0.08 0.34 <0.015 <0.03 <0.1 0.02
15-Aug-07 Crater lake 13.5 0.95 0.54 410 65 794 692 6,304 9,074 18 284 0.04 379 562 0.53 8.8 99 15.0
Units are mg/L
140 B. Christenson and F. Tassi

column, and consists of (in order of decreasing
abundance) jarosite, illite, pyrite, S, hematite,
siderite, K-nontronite, anhydrite, kaolinite,
dolomite, mesolite, Mg-saponite and quartz.
Results of the advective flow model are sum-
marised in Figs. 8and 9. Specific discharge
through the system attains a maximum at ca.
2.5 h, after which discharge declines monotoni-
cally to ca. 15 % of the peak flow by the end of
the simulation (24 h; Fig. 8a). The advancing
front of condensate is clearly delineated by Cl
concentration (Fig. 8b), which approaches the
end of the column after 24 h. Peak Cl concen-
trations broaden with time owing to dispersion
along the pathway, whereas total C in the fluid
(not shown) has a much steeper advancing front,
which is governed by carbonate mineral equi-
libria (discussed below). Porosity abruptly
declines from 35 to ca. 10 % in the first few
elements (Fig. 8c). Naturally, this is reflected in
permeability (Fig. 9a) which is derived as a
function of porosity in X1t (the default relation
for sandstone was adopted here, where log
k=15ϕ–5, with kbeing permeability (darcy), and
ϕthe porosity of the medium). A further but
smaller decrease in porosity and permeability
occurs at ca. 7 m after 24 h. Interestingly, this
perturbation is observed to migrate along the
column during the course of the simulation,
pointing to a progressing mineral front (dis-
cussed below). In contrast to the relatively rapid
advance of Cl through the system, the thermal
front associated with the hot condensate injection
advances less than half way along the column
within 24 h (Fig. 9a), pointing to decoupling of
heat and mass along the flow path as heat is
transferred into the host rock matrix. System Eh
and pH are plotted in Fig. 9b, and reveal further
evidence of chemical fronts advancing along the
Fig. 8 Reactive transport model results. aSpecific
discharge at 9.95 m. bAdvance of Cl front along flow
path with time. cPorosity along flow path after 24 h
10 cm
10 m
mc01
10 cm
X1t Reactive Transport Model
Single-phase liquid
Porosity:
Pressure Gradient:
Pore Fluid:
T (initial):
Reactant Fluid:
T (Reactant fluid):
Reactant Minerals:
Duration:
35%
0.2 Mpa
RCL (August 4, 2007)
50 oC
1 molal SO2, H2S, HCl, CO2
300 oC
Diop(.1%), Anor(.1%), Alb(.1%),
Ksp(.1%), Mt(.01%)
1 day
Fig. 7 Reactive transport conceptual model. Flow path is
10 m long, with axis-normal dimensions of 10 cm by
10 cm. Adapted from Christenson et al. (2010)
Gases in Volcanic Lake Environments 141

Fig. 9 Reactive transport
modelling results plotted as
a functions of position
along flowpath.
aTemperature and
permeability. Significant
drop in permeability occurs
at the inlet, and another at
approximately 7–8 m after
24 h, whereas the thermal
front has advanced to ca.
3 m after 24 h. bAn abrupt
increase in Eh corresponds
to changes in S species
concentrations. Low pH
front advances to ca. 7 m
after 24 h. cS
concentrations decline
abruptly between ca. 4 and
7.5 m, with H
2
S most
strongly depleted.
dElemental S precipitation
extends to ca. 4 m, after
which its absence
corresponds to the region
of abruptly increased in Eh,
reflecting its control over
the oxidation state of the
system. Carbonate and clay
minerals characterise the
interval from 7.5 to 10 m,
and correspond to localised
low permeability
142 B. Christenson and F. Tassi

flow path. The initial fluid redox state and pH are
preserved only in the last metre of the column (at
−0.15 and 6.4 V respectively, and are controlled
by the equilibrium mineral assemblage formed
from the initial interaction between lake water
and reactant phases described in Fig. 7. At the
inlet end, pH is strongly acidic (<1); here the
injected condensate has expended the pH-buf-
fering capacity provided by the initial equilib-
rium mineral assemblage.
Elemental S is the single most abundant pre-
cipitated mineral phase, and in conjunction with
quartz, alunite and pyrite, is responsible for the
aforementioned decline in permeability close to
the source (Fig. 9d; and Christenson et al. 2010).
However, consumption of H
2
S and SO
2
accord-
ing to Eq. (6) ultimately leads to under-saturation
with respect to S along the flow path (at dis-
tances >4 m after 24 h), with an abrupt increase
in system Eh thereafter (Fig. 9b). This is reflected
in the increased abundance of anhydrite, and the
appearance of hematite over the interval of 4–
7 m. Here, pH reaches its minimum values (<1)
owing primarily to declining temperature. The
formation of elemental S along the flow path, and
the large positive shift in redox potential in the
absence of this phase are also reflected in the
absolute and relative fugacities of H
2
S and SO
2
(Fig. 9c). In the presence of elemental sulphur
these species remain relatively constant.
Downstream of the elemental sulphur, how-
ever, H
2
S is effectively absent, as reflected in log
(fSO
2
/fH
2
S) varying from 0.4 to >30 between the
two zones. Similarly, the fugacities of H
2
and CO
also change abruptly across this interface, with
log (fH
2
/H
2
O) decreasing from −5.9 to −18.2
and log (fCO/fCO
2
) decreasing from −7.6 to
−21.3 over the interval between 0.8 and 4.5 m.
While the temperature gradient across this zone
has a direct role to play in both ratios, the effect
of temperature is minor compared to the pres-
ence/absence of elemental sulphur. It is note-
worthy that the changes across this interval occur
in the absence of any effective rock buffering
capacity as described previously (Giggenbach
1987; Chiodini et al. 2001), owing to the fact that
all Fe in the system occurs as Fe
3+
.
Interestingly, gases in equilibrium with the
alteration mineral assemblage occurring near the
end of the flow path have log (fH
2
/fH
2
O) ratios of
approximately −6.4. Although lower than those
encountered in either of the zones upstream, the
redox state of the fluid here is still considerably
more oxidising than the rock buffer value of
−2.8, suggesting that rocks in these environ-
ments have varying, if not fleeting capacities to
buffer redox potential of the fluids moving
through them. From this model, it is apparent that
the presence of elemental sulphur in these sys-
tems has an important bearing on the redox state
of the system, and the gases derived from them.
Reference was made earlier to the decrease in
permeability observed between 7 and 8 m after
24 h, and the fact that this appears to migrate along
the flow path close to the advancing Eh and pH
fronts. We can see from Fig. 9dthatthisperme-
ability perturbation corresponds to the presence of
kaolinite, siderite and dolomite, and as mentioned
previously, this corresponds to the advancing front
of CO
2
through the column. Carbonate minerals
have been previously identified as the main seal-
forming phases at Raoul Island Christenson et al.
(2007), contributing to gas over-pressuring and
gas-driven, phreatic eruption activity in that sys-
tem. This modelling shows that their formation is
the natural consequence of acid gas injection into
the more distal magmatic-hydrothermal environ-
ments, leading to the formation of highly reactive
acidic fluids which interact with their enclosing
rocks via hydrolysis reactions, thereby consuming
hydrogen ions and leading to the downstream
stabilisation and precipitation of permeability-
reducing mineral phases, including carbonates.
4 Hazards Associated with Gases
in Volcanic Lake Systems
4.1 Eruptions Induced by Selfsealing
of Sublacustrine Gas Conduits
In volcanic lake systems fed by magmatic sour-
ces lying within hundreds of metres of the lake
floor, the interaction between magmatic vapour,
Gases in Volcanic Lake Environments 143

ground water and rocks may lead to the forma-
tion of hydrothermal mineral seals which in
themselves create significant hazards. In these
cases, the very gases driving seal formation may
also serve as pressure-transmitting media behind
those blockages, and ultimately lead to their
failure, (e.g., Phillips 1972), with consequent
explosive phreatic and/or phreato-magmatic
eruptions as recently observed on Raoul Island
Christenson et al. (2007) and Ruapehu, Chris-
tenson et al. (2010).
Ruapehu Crater Lake, with a diameter of
some 500 m and a volume of approximately
9 million m
3
, sits atop the two active vents of this
andesitic massif situated in Tongariro National
Park. The volcano has maintained varying states
of activity since Europeans first observed it the
1860s, including two lake-expelling eruptions in
1945 and 1995. A most interesting characteristic
of the volcano, and one that is seemingly unaf-
fected by large eruptive events, is its cyclic dis-
charge of heat and volatiles through the vent-lake
system (e.g., Werner et al. 2006; Christenson
et al. 2010). Cycles vary between 3 and 9 months
duration, with lake temperatures historically
ranging between 10 and 60 °C. Phreatic eruption
activity on Ruapehu most typically occurs when
the lake is at or near the peak of its temperature
cycles (i.e., during open vent conditions). How-
ever, three significant events have occurred in the
last 25 years when the lake temperatures were at
minimum values (9 °C, December, 1988; 15 °C,
October, 2006, and 13 °C, September, 2007).
These eruptions occur with little or no pre-
cursory activity, and have so far proven difficult,
if not impossible to forecast. The largest and
most recent of these eruptions occurred in Sep-
tember of 2007 when some 160,000 m
3
of vent
materials were ejected through the lake (Kilgour
et al. 2010). The ejecta consisted of blocks
comprising a variety of vent-fill lithologies (la-
vas, sediments and breccias), spanning a range of
hydrothermal alteration intensities from fresh to
intensely altered, and ranging up to 2 m diameter.
Elemental sulphur was commonly observed
having exuded from pore space in some of the
ballistics, indicative of S having existed in a
molten state (T > 119 °C) in a portion of the vent
at the time of the eruption. Petrographic analysis
of the intensely altered ejecta show dense, pore
space accumulations of predominantly elemental
S, with lesser amounts of alunite, pyrite, anhy-
drite and trace segregations of halite (Christenson
et al. 2010), an assemblage noted previously in
ejecta from earlier eruptions (Christenson and
Wood 1993). This assemblage is nearly identical
to that predicted from the aforementioned reac-
tive transport modelling, and it appears to con-
stitute an impermeable seal formed through
repeated injection of magmatic volatiles through
the vapour saturated conduit environment.
Insights into the processes leading to these events
are gained from monitoring gas emissions from
the pre- and post-eruption systems.
Emissions of CO
2
,SO
2
and H
2
S from Ru-
apehu have been measured by airborne platform
approximately monthly since 2003 (methodology
discussed by Werner et al. 2006). Data for the
period January 2006 to July 2012 are shown
along with lake temperatures in Fig. 10, and this
time span encompasses the two phreatic eruption
events of October 2006 and September, 2007.
There is a striking coherence between heat and
volatile discharge in this system, indicating that
gas-charged batches of magma periodically cir-
culate into the conduit system, degas and thereby
drive the thermal cycling in the lake. TOUGH2
modelling of the passage of heat and CO
2
through the vent system provides insights into
the processes involved, showing early decou-
pling of gas from heat in the vent, with gas
moving ahead of the thermal pulse, followed by a
convective sweep of heat through the conduit
(Christenson et al. 2010). A total of 8 heating
cycles occurred between 1 January 2006 and 1
July 2012, and each has associated CO
2
maxima
of varying magnitudes. C/[SO
2
+H
2
S] mole
ratios (hereafter C/ΣS) vary widely during this
period, ranging from 3 to infinity, and are com-
pared to the proposed so-called magmatic value
of 5 for White Island (Giggenbach and Sheppard
1989). Ratios above this value may result to
some extent from variability in the magmatic gas
supply (i.e., varying depth of separation, as
described earlier), but sulphur-budgeting argu-
ments presented by Christenson (2000) point to
144 B. Christenson and F. Tassi

scrubbing of S in the hydrothermal system as the
predominant process regulating S emission from
this volcano. Hence, the period leading up to the
October 2006 eruption where ratios ranged from
40 to infinity, also point to a period of extensive
S scrubbing and seal formation beneath the lake.
Conditions in the vent plumbing apparently
changed after the minor vent-clearing eruption in
2006, after which small but measurable levels of
SO
2
were detected, suggesting that the gas
stream entering the lake had equilibrated with
elemental sulphur lying close to the lake floor in
the hydrothermal system. CO
2
emissions
throughout 2007 were erratic compared to those
measured during 2006, a feature which we
interpret to indicate growth and consolidation of
Fig. 10 Time series gas
emissions data for Mt
Ruapehu. Six heating
cycles are portrayed (lake
temperature in red), along
with airborne platform data
for CO
2
,SO
2
and H
2
S.
Periods of elevated C/ΣS
point to periods of S
deposition (i.e., scrubbing)
in the vent region. Adapted
from Christenson et al.
2010
Gases in Volcanic Lake Environments 145

the seal beneath the lake floor, with consequent
pressurisation of the sub-seal hydrothermal
environment. CO
2
emission measured immedi-
ately after the 2007 eruption was the highest yet
observed for this volcano (>1,000 T/d), and the
C/ΣS ratio was ca. 20 at this time.
A conceptual model for the state of the system
immediately prior to the 2007 eruption is shown
in Fig. 11. The main feature here is the accu-
mulation of a pressurised gas column behind a
seal of elemental sulphur, alunite and anhydrite,
infilling the interstitial pore space of what would
otherwise be permeable vent breccia material.
Gas emission through the northern vent of Ru-
apehu was very strong immediately after the
eruption. This is consistent with the eruption
Fig. 11 Conceptual model
for the sealed state of the
Ruapehu vent system prior
to the September 2007
eruption. Seal development
reduced gas transfer
through the vent, and led to
increased pore pressures
via formation of a
compressible gas column
beneath the seal. Printed
with permission from
Christenson et al. (2010)
146 B. Christenson and F. Tassi

having been a gas-driven event, but the strong
post-eruptive degassing also points to the scale of
the gas accumulation within the hydrothermal
environment prior to the eruption. The elevated
C/ΣS ratio immediately after the eruption is also
consistent with storage of a magmatic gas in the
hydrothermal environment for some time behind
the seal, facilitating extensive scrubbing of H
2
S
and SO
2
via Eq. (6). SO
2
emission increased to
ca. 500 T/d ca. one week later, and C/ΣS fell
to <5, pointing to the degassing of a deeper and/
or less evolved magmatic gas source through an
open vent environment. A second, stronger pulse
of degassing occurred in early 2008, again
through the open vent, with C/ΣS ratios again
close to ca. 5. Since then, C/ΣS ratios have
progressively shifted upward, even during peri-
ods of unprecedented high CO
2
emission, sug-
gesting that S scrubbing and seal formation is
again occurring.
A possible mechanism by which accumulated
gas pressure behind such a seal may lead to its
failure is shown schematically in Fig. 12, adapted
from Phillips (1972). Here the Mohr stress circles
are plotted as functions of effective normal stress
(typically measured in units of tensile rock
strength), and are compared to a hypothetical
shear failure envelope. With maximum principle
stress being nearly vertical in the vicinity of the
vent seal, failure of the seal will occur when the
combination of effective normal stress and shear
stress matches the shear strength of the of the rock,
or in terms of the graph in Fig. 12, when the Mohr
stress circle (denoting differential stress) intersects
the shear failure envelope. This may occur
through either increasing the differential stress
(i.e., increasing the diameter of the stress circle),
or by increasing the pore pressure within the seal
which has the net effect of decreasing effective
normal stress, thereby shifting the stress circle to
the left towards the failure envelope. A state of
critical stress is achieved when the stress circle
intersects the failure envelope, and any minor
perturbation of the system may trigger its failure.
Christenson et al. (2010) referred to the 2007
eruption of Ruapehu as an “accidental”phreato-
magmatic eruption, noting that only a very minor
amount of juvenile material was present in the
ejecta. Central to their model of events leading to
the eruption was that a compressed gas cap had
accumulated behind the seal, raising pressures in
the vent to levels required for seal failure. Once
decompression was initiated at the seal, a pres-
sure transient migrated downward into the vent,
eventually reaching the top of the magma con-
duit, thereby explaining the small amount of
juvenile material observed in the ejecta.
4.2 Limnic Eruptions: Constraints
from the Lake Nyos Experience
Lakes hosted in quiescent volcanic systems, i.e.
those showing moderate degassing activity, are
typically affected by significant inputs of salt-
and CO
2
-rich fluids which favour the develop-
ment of vertically stratified water columns.
Under these conditions, mass transfer is mostly
controlled by slow diffusive mechanisms, leading
to accumulation of dissolved gases at depth. As
exhaustively described in other chapters of the
present book, meromictic volcanic lakes having
great depth and a large volume are able to store
large amounts of dissolved gases, which repre-
sent a tremendous hazard in case of destabiliza-
tion of water stratification, leading to a limnic
eruption (Sabroux et al. 1987), i.e. a catastrophic
release of those gases into neighbouring
Effective Normal Stress σN'
Shear Stress σ'S
MOHR FAILURE ENVELOPE
Δp
Fig. 12 Conceptual plot of effective normal stress versus
shear stress. Increasing pore pressure (Δp) in the seal
leads to displacement of the Mohr stress circle to the left,
towards the failure envelope. The intersection of the stress
circle and the failure envelope signifies a state of critical
stress for the system. Adapted from Phillips (1972)
Gases in Volcanic Lake Environments 147

environments, as occurred at Monoun and Lake
Nyos in 1984 and 1986, respectively (e.g., Kling
et al. 1987; Sigurdsson et al. 1987), resulting in
thousands of deaths.
There is still debate on whether the initial
destabilization of a stratified lake is related to
external triggering, such as an earthquake, a
landslide slumping into deep lake water or
extreme weather conditions (Sigurdsson et al.
1987; Kling et al. 1987,1989; Giggenbach 1990;
Evans et al. 1994), or a lake rollover rather occurs
“spontaneously”(Zhang 1996; Woods and Phil-
lips 1999). Considering that the CO
2
concentra-
tion at 58 m depth in Lake Monoun in January
2003 was very close to saturation (Kusakabe et al.
2008), the second hypothesis seems plausible.
Independent of the mechanism for the origin of
this natural phenomenon, the hazard for limnic
eruptions in a Nyos-type lake can only be mitigated
by artificial and controlled degassing of the bottom
water layer. However, this solution may severely
affect the shallow lacustrine environment, espe-
cially in lakes intensely populated by living
organisms such as Lake Kivu (DCO). The CH
4
and
CO
2
-rich gas reservoir hosted in this lake Schoell
et al. (1988), represent both a hazard for local
population and a possible huge energy resource.
According to the increase of gas accumulation rate
observed over the past decades in this lake Pasche
et al. (2010), artificial degassing has been consid-
ered necessary to prevent the progressive saturation
of deep water layers (Hirslund 2012). Nevertheless,
such an intervention may have dramatic conse-
quences for the *2 million people living in the
lake surroundings whose survival strongly depends
on lake water for personal use and fishery (Pasche
et al. 2009). This suggests that the hazard mitiga-
tion approaches at Lake Kivu should be carefully
evaluated on the basis of further studies to evaluate
different strategies aimed to mitigate the limnic
eruption hazard in this area.
5 Summary and Conclusions
Clearly, volcanic lakes are dynamic environ-
ments which are highly responsive to the inputs
of magmatic heat and/or volatiles. We have
shown that the depth of magmatic degassing is a
primary control on the compositional character-
istics of the volatile stream entering the limnic
environment, and this effectively controls the
chemical nature of the associated lakes. Shallow
magmatic degassing (with magma residing at 10s
to 100s of m depth) in volcanically active sys-
tems releases heat, CO
2
,SO
2
,H
2
S, HCl and HF
gases into overlying groundwater and/or lake
environments. This leads to development of a
magmatic hydrothermal system which envelopes
the magma conduits, and is characterised by
highly acidic solutions in its core which drive
hydrolysis reactions with their enclosing rocks.
Lakes sitting atop such conduit systems are
likewise acidic in nature, hot and they typically
carry high total dissolved solids. Modelling
results show that the precipitation of elemental
sulphur occurs close to the conduit within the
two phase liquid-vapour region, along with a
suite of advanced argillic alteration minerals,
which collectively fill pore space and reduce
permeability. Modelling also shows that acid as
fluids flow laterally away from the conduit,
neutralisation (i.e., water-rock) reactions lead to
the development of propyllitic mineral assem-
blages (including carbonates) in the presence of
fluids with near-neutral pH.
Deep magmatic degassing (i.e., sourced at
perhaps 10s of kms depth), on the other hand,
releases primarily CO
2
,N
2
and noble gases into
conduits, with sulphur and halogen gases
remaining largely or totally within the melt phase.
With low emission rates and/or long transport
pathways, such volatile streams typically decou-
ple from their associated heat to reach the surface
as cold gas emissions. CO
2
is the primary agent
driving hydrolysis in these systems, and model-
ling shows that equilibrium is rapidly achieved
with carbonate minerals, and in the presence of
argillic to propylitic alteration mineral assem-
blages. Of course, these examples represent the
extreme end members of a continuous spectrum of
possible intermediate degassing depths, and
potential intermediate crater lake characteristics.
Two principle hazards are associated with
gases in volcanic lake environments. In the acidic
systems, mineralogic seal formation resulting
148 B. Christenson and F. Tassi

from hot gas-water-rock interaction may lead to
accumulation of gas behind the seal, and raise pore
pressures to the point where the seal becomes
critically stressed, leading to its failure and con-
sequent phreatic eruptions. While seal formation
and consequent phreatic eruptions cannot be pre-
cluded in crater lakes which receive only cold,
deeply derived CO
2
discharge, an additional threat
in these lakes is that the gases may become strat-
ified in the otherwise stagnant water column. Once
the water column is saturated, experience from
Lake Nyos and Monoun show that even minor
perturbance of the lake may lead to explosive
exsolution and release of voluminous clouds of
heavy, suffocating CO
2
into adjacent environs.
Tropical regions, where seasonal limnic overturns
typically do not occur are most at risk from these
events, whereas lakes with significant heat inputs
are generally safe from stratification owing to
thermal instability, and consequent convection of
their water columns.
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