Layout 6


Geochemical monitoring of volcanic lakes. A generalized box model
for active crater lakes

Dmitri Rouwet1,*and Franco Tassi2

1 Istituto Nazionale di Geofisica e Vulcanologia, sezione di Palermo, Palermo, Italy
2 Università degli Studi di Firenze, Dipartimento di Scienze della Terra, Firenze, Italy

ANNALS OF GEOPHYSICS, 54, 2, 2011; doi: 10.4401/ag-5035

ABSTRACT

In the past, variations in the chemical contents (SO4
2−, Cl−, cations) of

crater lake water have not systematically demonstrated any relationships
with eruptive activity. Intensive parameters (i.e., concentrations,
temperature, pH, salinity) should be converted into extensive parameters
(i.e., fluxes, changes with time of  mass and solutes), taking into account
all the internal and external chemical–physical factors that affect the
crater lake system. This study presents a generalized box model approach
that can be useful for geochemical monitoring of  active crater lakes, as
highly dynamic natural systems. The mass budget of  a lake is based on
observations of  physical variations over a certain period of  time: lake
volume (level, surface area), lake water temperature, meteorological
precipitation, air humidity, wind velocity, input of  spring water, and
overflow of  the lake. This first approach leads to quantification of  the
input and output fluxes that contribute to the actual crater lake volume.
Estimating the input flux of  the "volcanic" fluid (Qf - kg/s) –– an
unmeasurable subsurface parameter –– and tracing its variations with
time is the major focus during crater lake monitoring. Through expanding
the mass budget into an isotope and chemical budget of  the lake, the box
model helps to qualitatively characterize the fluids involved. The
(calculated) Cl− content and dD ratio of  the rising "volcanic" fluid defines
its origin. With reference to continuous monitoring of  crater lakes, the
present study provides tips that allow better calculation of  Qf in the future.
At present, this study offers the most comprehensive and up-to-date
literature review on active crater lakes.

1. Introduction
The crater lakes that top active volcanoes are windows

into the deeper parts of the volcano hydrothermal systems.
More than 80 post-Holocene volcanoes worldwide have a
lake in their craters [Simkin et al. 1981, Delmelle and Bernard
2000a]. A permanent crater lake should meet a combination
of physical constraints: (i) the bottom of the lake should be
protected against water seepage by physical sealing; (ii) the
meteorological precipitation should be abundant (snowfall

and its melt water, rainfall); (iii) the input of "volcanic" fluid
should be sustained; and (iv) the heat input from the active
volcano should be limited, to avoid drying out of the lake by
evaporation [Brown et al. 1989, Pasternack and Varekamp
1997]. The necessary coincidence of these physical factors
stresses the vulnerability of the accumulated water masses
inside active volcanic craters.

In historical times, only 8% of volcanic eruptions have
occurred beneath water bodies (crater lakes or shallow
seawater), although these eruptions have produced about
20% of the total fatalities [Mastin and Witter 2000]. Phreatic
or phreatomagmatic explosions through crater lakes can
eventually expulse water from the crater basin and generate
base surges, lahars, tsunamis or seiches [Sheridan and
Wohletz 1983, Mastin and Witter 2000]. Coincidences
between eruptive and seismic activities at volcanoes with
crater lake systems have been described in previous studies
[Hurst and McGinty 1999, Christenson 2000, Martínez et
al. 2000, Obha et al. 2008]. Recent modeling of eruptions
from beneath active crater lakes has stated that the eruption
style depends on the lake depth, the eruption pressure, and
the mass ratio of lake water to superheated vapor
[Morrissey et al. 2010]. Even during non-eruptive periods
or quiescent degassing, crater lakes can pose indirect
volcanic risks. The acidic crater lake brines can seep into
the volcanic edifice, dissolve the wall rock, and decrease the
volcano flank stability [Rowe and Brantley 1993, Pasternack
and Varekamp 1994, Rowe et al. 1995, Sanford et al. 1995,
Delmelle and Bernard 2000b, Kempter and Rowe 2000,
Taran et al. 2008, Varekamp et al. 2001, Varekamp 2008,
Varekamp et al. 2009]. These processes can lead to lahars,
or flank failure and sector collapses [López and Williams
1993, Kerle and van Wijck de Vries 2001, Reid 2004,
Procter et al. 2010], which can trigger major eruptions
[Voight et al. 1983]. The distribution of acidic brines into
the hydrological network and plume emission into the

Article history
Received February 1, 2011; accepted May 3, 2011.
Subject classification:
Geochemical monitoring, Active crater lakes, Box model, Mass budget, Isotope and chemical budget.

161



atmosphere can affect vegetation and agricultural activity,
leading to human health problems [Sriwana et al. 1998,
Delmelle and Bernard 2000b, Heikens et al. 2005, Löhr et al.
2005, van Rotterdam-Los et al. 2008]. In dormant volcanic
systems, the CO2-rich lake bottom waters can lead to
catastrophic turnovers (Nyos-type events), followed by the
dispersion of a toxic CO2 cloud in the area around the lake
[Kling et al. 1987, JVGR 1989, Evans et al. 1994, Kusakabe
et al. 2000a, Chiodini et al. 2010].

The presence of a water body (a hydrothermal system
or crater lake) along up-rising patterns of volcanic gases
causes water vapor and acidic gas species (CO2, SO2-H2S,
HCl, HF) to dissolve, which establishes particular P-T-pH-
redox conditions inside the volcano hydrothermal system. In
active crater lakes, the acidic species are abundant, for
creation of a pH lower than 3.8 to lose CO2 to the
atmosphere as a gas phase. The chemical–physical
parameters of crater lakes can reflect and preserve signals
related to eventual changes in volcanic activity. This
"memory" of crater lakes provides an extra tool, as
compared to open-conduit degassing volcanoes, where the
recording of variations of gas discharge rates related to deep
fluid inputs is directly lost to the atmosphere, often without
being detected. When using crater lake chemistry and
dynamics as monitoring tools, the challenge is to decipher
the degassing history of the volcano over a period of
observation.

Over the four decades of crater lake research, some
crater lakes that manifest eruptive activity have been
extensively studied (Ruapehu, New Zealand; Poás, Costa
Rica; Lake Yugama, Japan, Copahue, Argentina-Chile; Santa
Ana, El Salvador; Rincón de la Vieja, Costa Rica; Raoul,
Kermadec Islands; Voui, Vanuatu) [Giggenbach 1974,
Brown et al. 1989, Oppenheimer and Stevenson 1989,
Takano and Watanuki 1990, Hurst et al. 1991, Rowe et al.
1992a, Rowe et al. 1992b, Christenson and Wood 1993,
Christenson 1994, Ohba et al. 1994, Rowe 1994, Takano et
al. 1994, Rowe et al. 1995, Christenson 2000, Ohba et al.
2000, Varekamp et al. 2001, Bernard and Mazot 2004, Tassi
et al. 2005, Christenson et al. 2007, Ohba et al. 2008, Bani et
al. 2009, Christenson et al. 2010a, Ohsawa et al. 2010]. The
time series of datasets and the interpretations for these
events can be considered the first examples of
discontinuous volcanic monitoring using crater lake
chemistry and dynamics. In the present study, we propose
the box model approach for active crater lake systems, to
better understand how physical and chemical changes in a
crater lake can be related to possible variations in volcanic
activity. We give tips on how to improve each parameter
that contributes to the box model, to better constrain the
relationships between a deep fluid input and its effects on
crater lake chemical–physical properties. Moreover, this
report presents the most complete bibliographical

compilation on active crater lakes at present, as a helpful
guide for future studies.

2. Active crater lakes
Volcanic lakes are classified into various groups

according to their state of activity, based on their physical
and chemical characteristics, such as low-activity to peak-
activity lakes and no-activity lakes [Pasternack and Varekamp
1997], and CO2-dominated, quiescent and "active" crater
lakes [Varekamp et al. 2000]. Active crater lakes are generally
gas dominated, and the solute compounds in crater lake
water are thought to depend on the fluid dynamics in the
magmatic–hydrothermal systems. Variations in the absolute
concentrations of chemical compounds in crater lake waters
are only partly indicative of variations in the degassing
regime. In some cases, the cation (Al3+, Fe2/3+, Ca2+, Na+, K+,
Mg2+, Mn2+), SiO2, and SO4

2− and Cl− contents in lake waters
act as precursors of eruptive activities [Ohba et al. 2008].
However, many nonvolcanic processes (e.g. dilution,
evaporation, water mass loss, secondary mineral
precipitation) affect the absolute solute contents, which
prevents a systematic interpretation of the temporal patterns
of compositional parameters. Increases in the Mg2+/Cl− ratio
of crater lake water have been interpreted as eruption
precursors at Ruapehu volcano since the 1970's [Giggenbach
1974, Christenson 2000, Christenson 2010a], and in Lake
Yugama in the 1980-1990's [Ohba et al. 2008]. In volcano
hydrothermal systems, Cl− is generally considered a
conserved species [Pasternack and Varekamp 1994], while
Mg2+ results from water–rock interactions that involve an
unaltered rock (~fresh magma). Therefore, Mg2+/Cl−

increases in lake water are believed to indicate the intrusion
of a new magma batch at shallow depth [Giggenbach 1974].
However, for hyperacidic lakes (pH ≤0; e.g. Poás), Cl− can
escape at the lake surface as a gas phase, while Mg2+ remains
unaffected, making the Mg2+/Cl− indicator often useless for
the monitoring of set-ups [Rowe et al. 1992a]. Thiosulfate
concentrations in acidic lake water have been correlated to
pre-eruption activities of subaqueous fumaroles [Takano
1987, Takano and Watanuki 1990]. Thus, thiosulfates can
sometimes be used as volcanic precursors, but these species
can be detected in only a few lakes, depending on the
temperature regime and SO2/H2S ratio of the entering
volcanic fluid [Takano 1987]. The use of these geochemical
parameters can help to provide a better understanding of the
volcanic activity beneath crater lakes, although they are not
systematically applicable for each crater lake. As volcano
monitoring aims to quantify variations with time of
parameters sensitive to changes in the volcanic status, it is
recommended to convert the intensive parameters (solute
concentrations and their ratios, pH and temperature) into
extensive parameters (mass fluxes of water or chemical
compounds).

ROUWET AND TASSI

162



163

3. The box-model approach
In highly dynamic systems like active crater lakes, the

main concept is the residence time of the lake water, which
can be defined as [Varekamp 2003, Taran and Rouwet 2008]:

(1)

where R is the residence time (s) at a certain time of
observation, V is the lake volume (or mass) at a certain time
of observation (kg), and F is the difference between the input
and output water fluxes at a certain time of observation
(kg/s). This definition of the residence time assumes steady-
state for water (total input = total output, constant lake
water volume), which is almost never the case for highly
dynamic active crater lakes [Varekamp 2003]. On the other
hand, the volume of a crater lake arises from the difference
between the input and output fluxes of the water mass added
to the lake water volume at the previous time of observation
(Figure 1). If the observation or sampling frequency is higher
than the residence time of the lake water, the input and
output fluxes can be assumed to be constant for the
observation period, and the water mass budget of the lake
can be expressed in a differential form:

(2)

where VL is the resulting crater lake volume (kg), VL0 is the
crater lake volume at the previous time of observation (kg),
Qf is the input flux of "volcanic" fluid from beneath the lake
(kg/s), Qm is the input flux of meteorological water (kg/s),
Qsp is the input flux of water from springs located outside
the lake (kg/s), Qe is the evaporative flux of water from the
lake surface (kg/s), Qs is the seepage flux of water out of the
lake system (kg/s), Qo is the overflow flux of water out of
the lake basin (kg/s), Dt is the period (s) between two times
of observation of the crater lake (≈ sampling frequency). In
Equation (2), all of the fluxes are positive. The "volcanic"
fluid can be defined as a fumarolic condensate that enters the
lake bottom either as a vapor from the underlying volcano
hydrothermal system, or directly, as an often highly saline
and acidic liquid brine. To apply the mass balance to an active
crater lake, it is thus strictly necessary that the observation
period is shorter than the residence time of the crater lake
water. If not, possible indicators of the changes in the activity
of the volcano hydrothermal system will be "flushed out" of
the lake before being detected. For example, at Lake Yugama
(Kusatsu-Shirane volcano, Japan), R is less than one year, so
a monthly sampling frequency is adequate to detect possible
variations in activity [Ohba et al. 1994, 2000, 2008]; while the
El Chichón crater lake has a very low R of ~2 months, and
thus requires a sampling frequency of at least monthly
[Taran and Rouwet 2008].

3.1. Lake depth, surface area and volume
The first unknown parameter in Equation (2) is the

variation in the lake volume (VL - VL0). In general, the lake
volume is deduced empirically and independently for each
crater lake [Hurst et al. 1991, Ohba et al. 1994, Martínez et al.
2000, Rouwet et al. 2004, Tassi et al. 2009]. The lake surface
area can be obtained by global-positioning system (GPS)
measurements in the field [Rouwet et al. 2008], topographic
maps, satellite images [Zlotnicki et al. 2008], or photographic
methods [Rouwet et al. 2004, Rouwet 2011]. The lake depth
can be obtained by direct measurements or by echo-sounding
surveys along horizontal profiles from a raft [Takano et al.
2004, Christenson et al. 2010b]. On actively degassing lakes,
bottom profiles from echo-sounding can be disturbed by the
rising bubbles [Caudron and Bernard 2010]. The lake volume
can be highly variable due to the presence of sulfur pools at
the lake bottom [Oppenheimer and Stevenson 1989, Hurst et
al. 1991, Christenson and Wood 1993, Takano et al. 1994],
secondary mineral precipitation [Christenson et al. 2010a] and
changes in the sedimentary regime [Rouwet 2011], phreatic
eruptions breaching the lake bottom sediments [Bani et al.
2009, Rouwet et al. 2010], or dome intrusion that results in
uplift of the lake bottom [Bernard et al. 2010]. To improve
the accuracy and frequency of depth measurements, floating
buoys with pressure gauges have been recently installed on
lakes, to register variations in hydrostatic pressure (Y. Taran
and A. Bernard, person. comm.).

3.2. Input flux of  volcanic fluid
The input flux of volcanic fluids (Qf), i.e. the only real

"volcanic" factor that affects the crater lake mass balance, is
an unknown parameter. However, quantifying Qf and its
variations with time will provide insights into possible
variations in volcanic activity, which is within the scope of
volcano monitoring programs. Solving Equation (2) for this
"volcanic" parameter (Qf) is possible when the physical "non-

A BOX MODEL FOR ACTIVE CRATER LAKES

V

V Q t Q t Q t Q t Q t Q t
L

L0 f m sp e s o

=

+ + +D D D D D D- - -

Figure 1. Box model for active crater lakes, observed for a period of
time t0 to t.

R V F=



volcanic" factors Qm, Qsp, Qe, and Qs are fully understood
and quantified.

3.3. Input flux of  meteorological water
The input flux of meteorological water (Qm) to a crater

lake system depends on the precipitation rate, which
commonly refers to a certain period (e.g. daily), and on the
surface of the catchment area for meteorological
precipitation. To measure the precipitation rate, a rain gauge
should be installed as near as possible to the lake. For lakes in
tropical or temperate areas, the meteorological precipitation
is a real-time parameter, whereas for lakes in cold areas (high
altitudes and/or latitudes), meteorological precipitation is
often snow, which will be released into the lake with a time
delay, after the snow melts in spring and summer. The
catchment area of meteorological precipitation is the crater
surface area, and considered in this way, it is the total amount
of meteorological water that finally contributes to the crater
lake water volume by direct entrance, superficial run off or
after infiltration. In reality, part of the meteorological water
is dissipated by permeation through the volcanic edifice, and

by evaporation from the ground of the emerged catchment
area before contributing to the lake water volume. Such
losses strongly depend on the subsurface structure and
hydrogeology of the summit system, which is generally a
poorly constrained factor. To deduce the "effective
catchment area" for meteorological precipitation [Ohba et
al. 1994], it is best to observe the direct effects of short-term
major rain events on the (sudden) lake volume increase.
Nevertheless, it is not always possible to observe such events.
For this reason, taking the crater area as the catchment area
for meteorological precipitation is often the best alternative,
although this can cause overestimation of the Qm value.

3.4. Input flux of  spring water
Thermal or cold springs can significantly affect the mass

balance of a crater lake [Ohba et al. 1994, Delmelle et al.
1998, Stimac et al. 2004, Rouwet et al. 2008]. For example, a
geyser-like spring is a major fluid source for the El Chichón
lake (Chiapas, Mexico) [Taran et al. 1998, Rouwet et al. 2004,
Rouwet et al. 2008, Taran and Rouwet 2008] (Figure 2). The
input flux of springs (Qsp) located outside the catchment area

ROUWET AND TASSI

164

Figure 2. Discharge of a thermal spring (Qsp) towards an active crater lake (El Chichón, Chiapas, Mexico). (a) The white arrow indicates the discharge
of thermal spring water towards the lake. The people indicate the scale (photograph: D. Rouwet). (b) A detail of the stream flow (photograph: Y. Taran),
and (c) Geyser-like thermal spring, with a shoe to indicate scale (photograph: D. Rouwet). The white box indicates the location of the geyser-like thermal
spring, enlarged in (c).



165

of a lake can be directly measured in the field, by means of
a flow-meter device.

3.5. Evaporative output flux
The most prominent feature that proves that a crater

lake acts as a calorimeter and window into the volcano
hydrothermal system beneath the lake is the evaporative
heat and mass loss from the lake surface (Figures 3, 4). The
evaporative output (Qe) strongly depends on the difference
between the temperatures of the lake surface and the
ambient air above the lake, and the air humidity at the lake
water temperature. The Brown et al. [1989] equation, which
was applied for the first time at Poás Laguna Caliente (Costa
Rica), is the most widely accepted method to calculate the
evaporative heat loss (HL - W/m

2) at the lake surface from
the latent heat of vaporization of water (L-MJ/kg):

(3)

where E is the average particle flux, which is dependent on
temperature and wind speed; u is the friction velocity (m/s),
which is dependent on wind speed; qw is the saturation vapor
concentration of water in air at the lake temperature

(kg/m3); and qa is the vapor concentration in air at the
ambient temperature (kg/m3). Generally, a low wind speed
of 2 m/s is accepted to calculate the evaporative mass losses,
considering that a crater lake surface is often screened from
the dominant strong winds at high elevations of volcano
summits by the crater rim [Brown et al. 1989, Taran and
Rouwet 2008] (Figure 3). Oppenheimer [1997] suggested that
the bulk temperature measured a few centimeters beneath
the lake surface is tenths to a few degrees higher than the
real surface water temperatures (skin effect). Often, the
measured temperature thus erroneously results in
overestimations of calculated evaporative fluxes. However,
as broadly discussed in the literature, the wind speed
immediately above the lake surface probably dwarfs the DT
and humidity effects [Brown et al. 1989, Adams et al. 1990,
Hurst et al. 1991, Ohba et al. 1994]. Hurst et al. [1991] stated
that «the calculation of evaporative losses is not
straightforward with experimentally based formulae in
apparent conflict». With the present instrumental means, it
is impossible to measure the water (vapor) lost from a crater
lake surface by evaporation. To date, no direct and consistent
measurements of wind speed above crater lakes have been
produced. It is strongly recommended to install permanent

A BOX MODEL FOR ACTIVE CRATER LAKES

Figure 3. Evaporation (Qe) from Poás crater lake, Costa Rica (December 2009; photograph: D. Rouwet). The lake temperature was ~45 ˚C, while the
ambient temperature was ~10 ˚C.

H LE q q uL w a= -^ h



meteorological stations to measure wind speed and
direction, air humidity and temperature, and atmospheric
pressure near the lake shore, coupled with a fixed
temperature probe in the lake, to enhance the confidence of
the Qe parameter in Equation (3).

3.6. Seepage output flux
The seepage loss (Qs) through the crater lake floor is an

important subsurface process. A reliable estimation of the
seepage flux can be carried out through solving the mass
balance for Qs during quiescent periods of short duration,
assuming a zero Qf [Rouwet et al. 2004]. Although it is often
considered as constant afterwards, Qs is probably highly
variable with time, and depends on: (i) the hydrostatic
loading of the lake water mass [Terada et al. 2010]; (ii) the
lake bottom surface area [Taran and Rouwet 2008]; (iii) the
physical sealing of the lake bottom by clays [Rouwet 2011],
secondary minerals or molten sulfur pools [Oppenheimer
and Stevenson 1989, Hurst et al. 1991, Christenson and Wood
1993, Takano et al. 1994]; and (iv) the opening and sealing of
fractures. During periods of peak activity, upward steam
pressures probably inhibit direct seepage at the lake bottom,
or force the lake water to infiltrate through alternative

routes. Moreover, phreatic eruptions can breach the lake
bottom sediments or sulfur pools, strongly modifying the
morphology and permeability of the lake bottom, and
inevitably affecting the lake seepage. Understanding the local
tectonics inside and around craters can help to locate possible
seepage routes. Structural mapping using CO2 flux surveys
on and around crater lakes can give qualitative insights into
the fluid migration regimes [Mazot and Taran 2009, Mazot
et al. 2011]. Geophysical surveys (self potential, resistivity,
magnetometry, microgravimetry) can increase the
knowledge of the seepage regime of crater fluids [Fournier
et al. 2004, Rymer et al. 2009, Fournier et al. 2010, Jutzeler et
al. 2011]. These indirect approaches, however, do not help to
quantify the seepage output flux. The results of hydrological
and geophysical findings can be subjected to
hydrogeochemical modeling to verify if the deduced fluid
circulations are physically possible [Christenson et al. 2010a].
By trial-and-error, this can allow a semi-quantitative
estimation of the permeability that controls the seepage.

3.7. Overflow output flux
Some lakes can reach the crater rim and overflow

towards the volcano flank (e.g. Lake Nyos, Cameroon;

ROUWET AND TASSI

166

Figure 4. Evaporation (Qe) from Paulina Lake, Newberry Caldera, Oregon, USA (August 2010; photograph: D. Rouwet). The lake temperature was ~18 ˚C,
while the ambient temperature was ~5 ˚C.



167

Ruapehu, New Zealand; Boiling Lake, Dominica) [Tuttle et
al. 1987, Christenson 2000, Schaefer et al. 2008, Fournier et al.
2009, Christenson et al. 2010a]. In some cases, this overflow
can have disastrous effects, leading to destructive lahars
descending from volcano flanks (Chiginagak, Alaska;
Ruapehu, New Zealand) [Schaefer et al. 2008, Procter et al.
2010]. Such overflows (Qo) can also be induced intentionally
by human engineering, as drainage tunnels, to control the
lake water volume, to avoid overflows leading to lahars or
dissolved gas accumulation in the bottom waters of "Nyos-
type" lakes (e.g. Kelud, Indonesia; Lago Albano, Italy)
[Bernard and Mazot 2004, Carapezza et al. 2008]. As is the
case for the input flux of spring waters, the overflow output
flux of lake water can be generally measured directly in the
field. Most crater lakes, however, do not show overflow
(Qo.=.0 kg/s).

4. Isotope and chemical budget analysis
So far, the lake water chemistry has not been necessary

to calculate the input and output fluxes that affect the mass
budget, which is based only on the physical variations of the
crater lake (lake level, surface area, volume, temperature,
atmospheric parameters, meteorological precipitation).
However, when lake water sampling is possible, the box-
model approach can be expanded to the isotope and
chemical budget of the crater lake. This has the great
advantage that sources can be better characterized, physical
processes can be detailed, and fluxes of chemical compounds
can be calculated. Considering that active crater lakes do not
have a chemocline, due to lake convection [Christenson and
Wood 1993, Christenson 1994], a water sample from the
surface is generally representative of the vertical water
column. Sampling and analytical procedures should remain
identical during a period of monitoring, to avoid data
inconsistencies. When crater walls are nearly vertical, a
rappel descent is necessary (e.g. Irazú, Costa Rica), or specific
sampling devices have to be lowered into the lake from the
crater rim (Rincón de la Vieja, Costa Rica and Aso, Japan)
[Tassi et al. 2005, Miyabuchi and Terada 2009]. At erupting
volcanoes, sampling can become extremely dangerous. 

4.1. Isotope budget
As all of the terms in Equation (2) represent water

masses, by multiplying each term with the corresponding
values of the water isotopic composition (dD or d18O), the
isotope budget of a crater lake is given by (Figure 5):

(4)

where dD is the water isotopic composition, where the L,
L0, f, m, sp, e, s, and o subscripts refer to the lake water, the
lake water at the previous time of observation, the volcanic

fluid entering from beneath the lake, the meteorological
water, the spring water discharging into the lake, the vapor
coming off the lake surface, the seeping lake water, and the
overflowing lake water, respectively. dD is conveniently used
because: (i) the d18O is sensitive to O-exchange between
oxidized S-species and the water in variable thermal regimes
[Kusakabe et al. 2000b], which is a common situation for
active crater lakes; and (ii) the d18O shows a positive isotope
shift due to water–rock interaction processes, which are
abundant in volcano hydrothermal systems [Craig 1963,
Giggenbach and Stewart 1982]. The isotope balance
approach goes beyond a pure mass balance of crater lakes, as
the isotopic composition of the contributing fluids provides
qualitative insight into the chemical and physical processes
that control isotopic fractionation. Moreover, with the dD
budget, we consider only the water, without considering the
chemical compounds entering through the gas phase.

The only fluids that are impossible to sample are the
entering volcanic fluid (f) and the steam coming off the lake
by evaporation (e). Therefore, solving Equation (4) for dDf or
dDe is most adequate, although one of the two parameters
has to be deduced or calculated separately. If the rising
volcanic fluid is assumed to be pure magmatic (andesitic)
water, the dDf is expected to range from −30‰ to −10‰
[Taran et al. 1989, Giggenbach 1992]. However, this
assumption is not straightforward, as the condensing
volcanic fluid does not necessarily directly originate from the
magma, but can be influenced by boiling and steam
separation processes, lake recycling, or mixing processes. The
isotopic composition of local meteorological water can be
determined for rain water samples (dDm). The dDsp is directly
obtained from the analysis of the spring water sample. The
dDsp of each specific period of observation is often preferred
to a long-term average dDsp. The isotopic composition of
steam that originates from heated water bodies has been
discussed in the literature [Giggenbach 1978, Giggenbach

A BOX MODEL FOR ACTIVE CRATER LAKES

Figure 5. Box model for the isotope budget, observed for a period of
time t0 to t.

V D Q D t Q D t Q D t+ + +d d d dD D D -=

Q D t Q D t Q D td d dD D D- - -

DdVL L L0 L0 f f m m sp sp

e e s s o o



and Stewart 1982]. The dDe of the vapor released at the lake
surface is subject to kinetic or equilibrium fractionation, and
strongly depends on: (i) the differences between the
temperature of the top layer of the lake water and the
ambient temperature above the lake [Ohba et al. 2000, Taran
and Rouwet 2008]; (ii) the humidity of the boundary layer
(the air layer immediately above the lake surface) [Varekamp
and Kreulen 2000, and references herein]; (iii) the pH; and
(iv) the salinity of the lake water [Horita et al. 1993a, 1993b].
The kinetics and equilibrium isotopic fractionation for dDe
under a large variety of clearly controllable conditions is
quantified in experimental set-ups [Horita and Wesolowski
1994, Cappa et al. 2003]. Nevertheless, the natural system,
i.e. the evaporating crater lake, is far from an experimentally
stable situation, which forces assumptions to be adopted
regarding the calculated dDe introduced in Equation (4).
Different equations to calculate dDe can be applied, although
the exact isotopic composition of the evaporative steam
from a crater lake at the moment of observation will hardly
ever coincide with the calculated dDe. The dDs and dDo
values introduced in Equation (4) are simply taken to be the
average of dDL for the period of observation ((dDL +dDL0)/2).
Among the parameters in Equation (4), dDe is the one with
the largest uncertainty. A challenge for the future is to find an
adequate technique to sample the evaporative steam.

4.2. Chemical budget
As already emphasized by the mass balance approach,

increased release of volcanic fluids related to changes in
volcanic activity are masked by nonvolcanic fluid
contributions. The only certainty is that the main anionic
species (Cl−, SO4

2−, F−) and acidity originate from a
magmatic-related fluid entering the lake, i.e. volcanic fluid
and/or thermal springs. Dilution by meteorological and/or
spring water, salinity increase by evaporation, and element
loss through seepage or lake overflow are inevitable.

Following the above statements, it is not safe to suggest that
an increase in the anionic content of crater lake waters is
only due to the enhanced input of the volcanic fluid.
Chloride is considered be chemically conserved in acidic
hydrothermal environments, while sulfur species are subject
to precipitation as secondary minerals (anhydrite, alunite,
native sulfur). For this reason, past studies have basically used
Cl− as the guiding species in chemical balances [Ohba et al.
1994, Rouwet et al. 2004, Rouwet et al. 2008, Taran and
Rouwet 2008]. Based on Equation (2), we can now multiply
each term with the corresponding Cl− content (Figure 6):

(5)

where Cl is the Cl− content, and the L, L0, f, m, sp, e, s, and
o subscripts are those used in Equation (4). Equations (2), (4)
and (5) can be solved mathematically for the desired
unknown parameter. For El Chichón crater lake, Qsp was
calculated [Rouwet et al. 2004, 2008], while Clf was estimated
for Yugama Lake following a similar approach [Ohba et al.
2000]. It is clear that the expression "Q Cl Dt" represents a
Cl− mass, and Equation (5) respects the relative proportions
of Cl− given by the various Q Dt.

The entering volcanic fluid cannot be sampled;
therefore, the most common procedure is to calculate Clf
using Equation (5), when the other parameters are known.
Unless acidic rain is formed due to the absorption of plume
gases into the clouds, the entrance of Cl− from
meteorological water can be neglected (Clm =0mg/l). Spring
waters can be sampled and Clsp should thus be known. A
significant HCl loss by evaporation occurs from the surface of
hot hyper-acidic lakes (e.g. Poás, Costa Rica; Aso, Japan). In
this case, Cle is the unknown term because the Cl2-rich vapour
is still impossible to sample. A promising tool is to determine
the chemical species in evaporation plumes at active crater
lakes by direct plume measurements [Christenson et al.
2010a, Christenson et al. 2010b, Shinohara et al. 2010]. For
lower temperature, lower acidity lakes, Cl2 is not lost from
the lake surface (Cle =0mg/l). For the Cl

− lost by seepage, it
is most correct to use the average Cl− content of the lake
water for the period of observation (Cls = (ClL + ClL0)/2). If
overflow occurs, the same is true for Clo (Clo =(ClL +ClL0)/2).

At El Chichón, no volcanic Cl− is thought to enter the
lake system, and the geyser-like springs were found to be the
only Cl− source for the crater lake during the period 1995-
2007 [Rouwet et al. 2004, Rouwet et al. 2008, Taran and
Rouwet 2008]. Following similar approaches, Ohba et al.
[1994, 2000] calculated a Clf even higher than the Cl

−

contents in high-temperature volcanic gases, which
suggested that volcanic fluids entering beneath crater lakes
are Cl−-enriched due to multi-step steam separation
processes, and probably to lake water recirculation. By

ROUWET AND TASSI

168

Figure 6. Box model for the chlorine budget, observed for a period of
time t0 to t.

V Cl V Cl Q Cl t Q Cl t Q Cl t

Q Cl t Q Cl t Q Cl t
L L L0 L0 f f m m sp sp

e e s s o o

= + + +D D D

D D D

-

- - -



169

multiplying Clf by Qf, Cl
− fluxes can be calculated. In our

opinion, monitoring Cl− fluxes of the volcanic fluid is the
most direct way to detect changes in volcanic input.

5. Conclusions
Based on a literature review, this study offers a

generalized box-model approach that can be applied in
continuous or discontinuous geochemical monitoring of
active crater lakes. Intensive geochemical parameters (solute
concentrations, pH) do not systematically relate to variations
in eruptive activities at volcanoes with crater lakes. On the
contrary, the box model allows the calculation of extensive
parameters (mass fluxes), providing a better representation
of the chemical–physical variations with time. By
quantification of the nonvolcanic factors that affect the
physical–chemical state of an active crater lake, the box
model allows calculation of the volcanic input flux (Qf),
which cannot be measured directly. A challenge for the
future is to better constrain these nonvolcanic factors, in
order to better define Qf. As well as giving insights into the
changes in volcanic activity (Qf), an understanding of the
seepage (Qs) and overflow (Qo) regime of crater lake systems
can be useful in volcanic surveillance of indirect risks related
to crater lakes (impact of contaminants on the environment,
flank failure, lahars).

The error of discontinuous monitoring methods is
impossible to quantify, as it largely depends on: (i) the error
in the single contributing parameters; (ii) the residence time
of the lake water specific for each lake; (iii) the related
sampling/observation frequency, which depends on access
to the crater lake (~safety limits during eruptive cycles,
logistics); and (iv) the averaging of the values for observation
periods longer than the residence time of the lake. Evolution
towards continuous monitoring (visual-thermal cameras,
and meteorological stations with data transmission systems,
P-T-pH-conductivity probes) will increase the frequency,
quality and safety of the data gathering. The quantitative and
qualitative characterization of temporal variations in the
volcanic and nonvolcanic parameters of the box model offers
a means to better constrain the physical and chemical
processes that occur in crater lakes with volcano
hydrothermal systems. Beyond the concept of continuous
versus discontinuous monitoring, this box model is applicable
for low-cost (mass budget), over medium-cost (chemical and
isotope budget) to high-cost (continuous) monitoring set-
ups. If volcano monitoring is a too ambitious goal for some
crater lakes, within the limits of the available means, the box
model proposed here will definitely help to develop basic
conceptual models of crater lakes with volcano
hydrothermal systems.

Acknowledgements. The authors wish to thank the Editor, S.
Inguaggiato, for his insightful handling of this study, and reviewer A.
Caracausi for fruitful discussions that helped us to improve the manuscript.

We are indebted to all of the previous workers on active crater lakes for
their indirect contributions to this review paper.

References
Adams, E., D.J. Cosler and K.R. Helfrich (1990). Evaporation

from heated water bodies: predicting combined forced
plus free convection, Water Resour. Res., 26, 425-435.

Bani, P., C. Oppenheimer, J.C. Varekamp, T. Quinou, M. Lardy
and S. Carn (2009). Remarkable geochemical changes and
degassing at Voui crater lake, Ambae volcano, Vanuatu, J.
Volcanol. Geotherm. Res., 188, 347-357.

Bernard, A. and A. Mazot (2004). Geochemical evolution of
the young crater lake of Kelud volcano in Indonesia,
Water-Rock Interaction (WRI-11), Wanty & Seal II eds.,
A.A. Balkema Publishers, 87-90.

Bernard, A., B. Barbier, C. Caudron, D. Syahbana, A. Solikhin,
S. Kunrat and V Hallet (2010). The monitoring of volca-
nic lakes in Indonesia, 7th Workshop of the Commission
of Volcanic Lakes, Costa Rica.

Brown, G., H. Rymer, J. Dowden, P. Kapadia, D. Stevenson,
J. Barquero and L.D. Morales (1989). Energy budget
analysis for Poás crater lake: implications for predicting
volcanic activity, Lett. Nature, 339, 370-373.

Cappa, C.D., M.B. Hendricks, D.J. DePaolo and R.C. Cohen
(2003). Isotopic fractionation of water during evapor-
ation, J. Geophys. Res., 108, 4525-4534.

Carapezza, M.L., M. Lelli and L. Tarchini (2008). Geochemis-
try of the Albani and Nemi crater lake in the volcanic
district of Alban Hills, J. Volcanol. Geotherm. Res., 178,
297-304.

Caudron, C. and A. Bernard (2010). Hydroacoustic quantifi-
cations of CO2 bubbles in volcanic lakes, 7th Workshop
of the Commission of Volcanic Lakes, Costa Rica.

Chiodini, G., A. Costa, D. Rouwet and F. Tassi (2010). Model-
ing CO2 air dispersion from gas driven lake eruptions,
Abstract V23A-2384, AGU Fall Meeting.

Christenson, B.W. and C.P. Wood (1993). Evolution of the
vent-hosted hydrothermal system beneath Ruapehu
Crater Lake, New Zealand, Bull .Volcanol., 55, 547-565.

Christenson, B.W. (1994). Convection and stratification in
Ruapehu crater lake, New Zealand: implications for Lake
Nyos-type gas release eruptions, Geochem. J., 28, 185-197.

Christenson, B.W. (2000). Geochemistry of fluids associated
with the 1995-1996 eruption of Mt. Ruapehu, New
Zealand: signatures and processes in the magmatic-
hydrothermal system, J. Volcanol. Geotherm. Res., 97, 1-30.

Christenson, B.W., C.A. Werner, A.G. Reyes, S. Sherburn, B.J.
Scott, C. Miller, M.J. Rosenburg, A.W. Hurst, and K. Britten
(2007). Hazards from hydrothermally sealed volcanic
conduits, EOS, 88 (50), 53-55.

Christenson, B.W., A.G. Reyes, R. Young, A. Moebis, S. Sher-
burn, J. Cole-Baker and K. Britten (2010a). Cyclic processes
and factors leading to phreatic eruption events: insights

A BOX MODEL FOR ACTIVE CRATER LAKES



from the 25 September 2007 eruption through Ruapehu
Crater Lake, New Zealand, J. Volcanol. Geotherm. Res.,
191, 15-32.

Christenson, B.W., A. Mazot, and K. Britten (2010b). Gas trans-
fer through Ruapehu Crater Lake: Insights gained from a
recent water-borne survey, Abstract V23A-2388, AGU Fall
Meeting.

Craig, H. (1963). The isotopic geochemistry of water and
carbon in geothermal areas, In: Nuclear Geology in
Geothermal Areas, Conference Proc. (Spoleto, Italy),
edited by E. Tongiorgio, 17-53.

Delmelle, P., M. Kusakabe, A. Bernard, T. Fischer, S. de Brouwer
and E. del Mundo (1998). Geochemical and isotopic
evidence for seawater contamination of the hydrothermal
system of Taal Volcano, Luzon, the Philippines, Bull.
Volcanol., 59, 562-576.

Delmelle, P. and A. Bernard (2000a). Volcanic lakes, In: Ency-
clopedia of volcanoes, Academic, New York, 877-895.

Delmelle, P. and A. Bernard (2000b). Downstream composition
changes of acidic volcanic waters discharged into the
Banyupahit stream, Ijen caldera, Indonesia, J. Volcanol.
Geotherm. Res., 97, 55-75.

Evans, W.C., L.D. White, M.L. Tuttle, G.W. Kling, G. Tanyileke
and R.L. Michel (1994). Six years of change at Lake Nyos,
Cameroon, yield clues to the past and cautions for the
future, Geochem. J., 28, 139-162.

Fournier, N., H. Rymer, G. Williams-Jones, and J. Brenes (2004).
High-resolution gravity survey: Investigation of subsurface
structures at Poás volcano, Costa Rica, Geophys. Res. Lett.,
31; doi: 10.1029/2004GL020563.

Fournier, N., F. Witham, M. Moreau-Fournier, and L. Bardou
(2009). Boiling Lake of Dominica, West Indies: High-
temperature volcanic crater lake dynamics, J. Geophys.
Res., 114; doi: 10.1029/2008JB005773.

Fournier, N., M. Moreau and R. Robertson (2010). Disappear-
ance of a crater lake: implications for potential explosivity
at Soufrière volcano, St Vincent, Lesser Antilles, Bull.
Volcanol.; doi: 10.1007/s00445-010-0422-3.

Giggenbach, W.F. (1974). The chemistry of Crater Lake, Mt
Ruapehu (New Zealand) during and after the 1971 active
period, New Zeal. J. Sci., 17, 33-45.

Giggenbach, W.F. (1978). The isotopic composition of waters
from the El Tatio geothermal field, Northern Chile,
Geochim. Cosmochim. Acta, 42, 979-988.

Giggenbach, W.F. and M.K. Stewart (1982). Processes control-
ling the isotopic composition of steam and water
discharges from steam vents and steam-heated pools in
geothermal areas, Geothermics, 11 (2), 71-80.

Giggenbach, W.F. (1992). Isotopic shift in waters from geo-
thermal and volcanic systems along convergent plate
boundaries and their origin, Earth Planet. Sci. Lett., 113,
495-510.

Heikens, A., S. Sumarti, M. van Bergen, B. Widianarko, L.

Fokkert, K. van Leeuwen and W. Seinen (2005). The
impact of hyperacid Ijen Crater Lake, risks of excess
fluoride to human health, Sci. Tot. Environm., 346, 56-69.

Horita, J., D.J. Wesolowski and D.R. Cole (1993a). The activity-
composition relationship of oxygen and hydrogen
isotopes in aqueous salt solutions: I. Vapor-liquid water
equilibration of single salt solutions from 50 to 100˚C,
Geochim. Cosmochim. Acta, 57, 2797-2817.

Horita, J., D.R. Cole and D.J. Wesolowski (1993b). The activity-
composition relationship of oxygen and hydrogen
isotopes in aqueous salt solutions: II. Vapor-liquid water
equilibration mixed salt solutions from 50 to 100˚C and
geochemical implications, Geochim. Cosmochim. Acta,
57, 4703-4711.

Horita, J. and D.J. Wesolowski (1994). Liquid-vapor fraction-
ation of oxygen and hydrogen isotopes of water from
the freezing to the critical temperature, Geochim.
Cosmochim. Acta, 58 (16), 3425-3437.

Hurst, A.W., H.M. Bibby, B.J. Scott and M.J. McGuinness
(1991). The heat source of Ruapehu Crater Lake;
deductions from the energy and mass balances, J.
Volcanol. Geotherm. Res., 6, 1-21.

Hurst, A.W. and P.J. McGinty (1999). Earthquake swarms to
the west of Mt Ruapehu preceding its 1995 eruption, J.
Volcanol. Geotherm. Res., 90, 19-28.

Jutzeler, M., N. Varley and M. Roach (2011). Geophysical
characterization of hydrothermal systems and intrusive
bodies, El Chichón volcano (Mexico), J. Geophys. Res.;
doi: 10.1029/2010JB007992.

JVGR (1989). The Lake Nyos Event and Natural CO2 Degas-
sing I, Special issue edited by F. Le Guern and G.E.
Sigvaldason, J. Volcanol. Geotherm. Res., 39 (2-3), 182 pp.

Kempter, K.A. and G.L. Rowe (2000). Leakage of Active Crater
lake brine through the north flank at Rincón de la Vieja
volcano, northwest Costa Rica, and implications for
crater collapse, J. Volcanol. Geotherm. Res., 97, 143-159.

Kerle, N. and B. van Wijck de Vries (2001). The 1998 debris
avalanche at Casita volcano, Nicaragua-Investigation of
structural deformation as the cause of slope instability
using remote sensing, J. Volcanol. Geotherm. Res., 105,
49-63.

Kling, G.W., M.A. Clark, H.R Compton, J.D. Devine, W.C.
Evans, A.M. Humphrey, E.J. Koenigsberg, J.P. Lockwood,
M.L. Tuttle and G.N. Wagner (1987). The 1986 Lake Nyos
gas disaster in Cameroon, West Africa, Science, 236,
169-175.

Kusakabe, M., G.Z. Tanyileke, S.A. McCord and S.G. Schla-
dow (2000a). Recent pH and CO2 profiles at Lakes Nyos
and Monoun, Cameroon: implications for the degassing
strategy and its numerical simulation, J. Volcanol.
Geotherm. Res., 97, 241-260.

Kusakabe, M., Komoda, Y., Takano, B. and T. Abiko (2000b).
Sulfur isotopic effects in the disproportionation reaction

ROUWET AND TASSI

170



171

of sulfur dioxide in hydrothermal fluids: implications for
the d34S variations of dissolved bisulfate and elemental
sulfur from active crater lakes, J. Volcanol. Geotherm.
Res., 97, 287-307.

Löhr, A.J., T.A. Bogaard, A. Heikens, M.R. Hendriks, S. Sumarti,
M.J. van Bergen, C.A.M. Van Gestel, N.M. Van Straalen,
P.Z. Vroon and B. Widianarko (2005). Natural pollution
caused by the extremely acidic crater lake Kawah Ijen,
East Java, Indonesia, Environ. Sci. & Pollut. Res., 12 (2),
89-95.

López, D.L. and S.N. Williams (1993). Catastrophic volcanic
collapse: relation to hydrothermal processes, Science, 260,
1794-1796.

Martínez, M., E. Fernández, J. Valdés, V. Barboza, R. Van der
Laat, E. Duarte, E. Malavassi, L. Sandoval, J. Barquero
and T. Marino (2000). Chemical evolution and activity of
the active crater lake of Poás volcano, Costa Rica, 1993-
1997, J. Volcanol. Geotherm. Res., 97, 127-141.

Mastin, L.G. and J.B. Witter (2000). The hazards of eruptions
through lakes and seawater, J. Volcanol. Geotherm. Res.,
97, 195-214.

Mazot, A. and Y. Taran (2009). CO2 flux from the volcanic lake
of El Chichón (Mexico). Geofís. Int., 48 (1), 73-83.

Mazot, A., D. Rouwet, Y. Taran, S. Inguaggiato and N. Varley
(2011). CO2 and He degassing at El Chichón volcano,
Chiapas, Mexico: gas flux, origin and relationship with
local and regional tectonics, In: S. Inguaggiato, H.
Shinohara and T. Fischer (eds), Geochemistry of volcanic
fluids: a special issue in honor of Yuri A. Taran, Bull.
Volcanol., 73 (4), 423-442.

Miyabuchi, Y. and A. Terada (2009). Subaqueous geothermal
activity revealed by lacustrine sediments of the acidic
Nakadake crater lake, Aso Volcano, Japan, J. Volcanol.
Geotherm. Res., 187, 140-145.

Morrissey, M., G. Gisler, R. Weaver and M. Gittings (2010).
Numerical model of crater lake eruptions, Bull. Volcanol.;
doi: 10.1007/s00445-010-0392-5.

Ohba, T., J. Hirabayashi and K. Nogami (1994). Water, heat
and chloride budgets of the crater lake, Yugama at
Kusatsu-Shirane volcano, Japan., Geochem. J., 28, 217-231.

Ohba, T., J. Hirabayashi and K. Nogami (2000). D/H and
18O/16O ratios of water in the crater lake at Kusatsu-
Shirane volcano, Japan, J. Volcanol. Geotherm. Res., 97,
329-346.

Ohba, T., J. Hirabayashi and K. Nogami (2008). Temporal
changes in the chemistry of lake water within Yugama
Crater, Kusatsu-Shirane Volcano, Japan: Implications for
the evolution of the magmatic hydrothermal system, J.
Volcanol. Geotherm. Res., 178, 131-144.

Ohsawa, S., T. Saito, S. Yoshikawa, H. Mawatari, M. Yamada,
K. Amita, N. Takamatsu, Y. Sudo and T. Kagiyama (2010).
Color change of lake water at the active crater lake of
Aso volcano, Yudamari, Japan: is it in response to change

in water quality induced by volcanic activity?, Limnology,
11, 207-215.

Oppenheimer, C. and D. Stevenson (1989). Liquid sulphur
lakes at Poás volcano, Nature, 342, 790-793.

Oppenheimer, C. (1997). Ramifications of the skin effect for
crater lake heat budget analysis, J. Volcanol. Geotherm.
Res., 75, 159-165.

Pasternack, G.B. and J.C. Varekamp (1994). The geochemistry
of the Keli Mutu crater lakes, Flores, Indonesia, Geochem.
J., 28, 243-262.

Pasternack, G.B. and J.C. Varekamp (1997). Volcanic lake sys-
tematic I. Physical constraints, Bull. Volcanol., 58, 528-538.

Procter, J.N., S.J. Cronin, I.C. Fuller, M. Sheridan, V.E. Neall
and H. Keys (2010). Lahar hazard assessment using Titan2D
for an alluvial fan with rapidly changing geomorphology:
Whangaehu River, Mt. Ruapehu, Geomorphology, 116,
162-174.

Reid, M.E. (2004). Massive collapse of volcanic edifices trig-
gered by hydrothermal pressurization, Geology, 32, 373-376.

Rouwet, D., Y. Taran and N.R. Varley (2004). Dynamics and
mass balance of El Chichón crater lake, Mexico, Geofís.
Int., 43, 427-434.

Rouwet, D., Y. Taran, S. Inguaggiato, N. Varley and J.A. San-
tiago Santiago (2008). Hydrochemical dynamics of the
"lake-spring" system in the crater of El Chichón volcano
(Chiapas, Mexico), J. Volcanol. Geotherm. Res., 178,
237-248.

Rouwet, D., R.A. Mora-Amador and C.J. Ramírez-Umaña
(2010). A combined Cl and isotope balance approach for
Laguna Caliente, Poás: precursory signals for single
phreatic eruptions during the ongoing eruptive cycle?,
7th Workshop of the Commission of Volcanic Lakes,
Costa Rica.

Rouwet, D. (2011). A photographic method for detailing the
morphology of the floor of a dynamic crater lake: the El
Chichón case (Chiapas, Mexico). Limnology; doi:
10.1007/s10201-011-0343-7.

Rowe, G.L., S. Ohsawa, B. Takano, S.L. Brantley, J.F. Fernández
and J. Barquero (1992a). Using Crater Lake chemistry to
predict volcanic activity at Poás Volcano, Costa Rica, Bull.
Volcanol. 54, 494-503.

Rowe, G.L., S.L. Brantley, M. Fernández, J.F. Fernández, A.
Borgia and J. Barquero (1992b). Fluid-volcano interaction
in an active stratovolcano: the Crater Lake system of
Poás Volcano, Costa Rica, J. Volcanol. Geotherm. Res.
64, 233-267.

Rowe, G.L. and S.L. Brantley (1993). Estimation of the dis-
solution rates of andesitic glass, plagioclase and pyroxene
in a flank aquifer of Poás Volcano, Costa Rica, Chem.
Geol., 105, 71-87.

Rowe, G.L. (1994). Oxygen, hydrogen, and sulphur isotope
systematics of the crater lake system of Poás Volcano,
Costa Rica, Geochem. J., 28, 263-287.

A BOX MODEL FOR ACTIVE CRATER LAKES



Rowe, G.L., S.L. Brantley, J.F. Fernández and A. Borgia (1995).
The chemical and hydrologic structure of Poás Volcano,
Costa Rica, J. Volcanol. Geotherm. Res. 64, 233-267.

Rymer, H., C.A. Locke, A. Borgia, M. Martínez, J. Brenes, R.
Van der Laat and G. Williams-Jones (2009). Long-term
fluctuations in volcanic activity: implications for future
environmental impact, Terra Nova; doi: 10.111/j.1365-
3121.2009.00555.x. 

Sanford, W.E., L.F. Konikow, G.L. Rowe Jr and S.L. Brantley
(1995). Groundwater transport of crater-lake brine at
Poás Volcano, Costa Rica, J. Volcanol. Geotherm. Res, 64,
269-293.

Schaefer, J.R., W.E. Scott, W.C. Evans, J. Jorgenson, R.G.
McGimsey and B. Wang (2008). The 2005 catastrophic
acid crater lake drainage, lahar, and acidic aerosol
formation at Mount Chiginagak volcano, Alaska, USA:
field observations and preliminary water and vegetation
chemistry results, Geochem. Geophys. Geosyst., 9 (7);
doi: 10.1029/2007GC001900.

Sheridan, M.F. and K.H. Wohletz (1983). Hydrovolcanism:
basic considerations and review, J. Volcanol. Geotherm.
Res., 17, 1-29.

Shinohara, H., S. Yoshikawa and Y. Miyabuchi (2010). Degas-
sing of Aso Volcano, Japan through an acid crater lake:
differentiation of volcanic gas-hydrothermal fluids
deduced from volcanic plume chemistry, Abstract V23A-
2387, AGU Fall Meeting.

Simkin, T., L. Siebert, L. McClelland, D. Bridge, C. Newhall
and J.H. Latter (1981). Volcanoes of the World,
Hutchinson Ross, Pa. Usa, 233 pp.

Sriwana, T., M.J. Van Bergen, S. Sumarti, J.C.M. de Hoog,
B.J.H. van Os, R. Wahyuningsih and M.A.C. Dam (1998).
Volcanogenic pollution by acid water discharges along
Ciwidey River, West Java (Indonesia). J. Geochem.
Explor., 62, 161-182.

Stimac, J.A., F. Goff, D. Counce, A.C.L. Larocque and D.R.
Hilton (2004). The crater lake and hydrothermal system
of Mount Pinatubo, Philippines: evolution in the decade
after eruption, Bull. Volcanol., 66, 149-167.

Takano, B. (1987). Correlation of volcanic activity with sulfur
oxy-anion speciation in a crater lake, Science, 235, 1633-
1635.

Takano, B. and K. Watanuki (1990). Monitoring of volcanic
eruptions at Yugama crater lake by aqueous sulphur
oxyanions, J. Volcanol. Geotherm. Res., 40, 71-87.

Takano, B., H. Saitoh and E. Takano (1994). Geochemical
implications of subaqueous molten sulfur at Yugama
crater lake, Kusatsu-Shirane volcano, Japan, Geochem. J.
28, 199-216.

Takano, B., K. Suzuki, K. Sugimori, T. Ohba, S.M. Fazlullin,
A. Bernard, S. Sumarti, R. Sukhyar and M. Hirabayashi
(2004). Bathymetric and geochemical investigation of
Kawah Ijen Crater Lake, East Java, Indonesia, J. Volcanol.

Geotherm. Res., 135, 299-329.
Taran, Y.A., B.G. Pokrovsky and Y.M. Dubik (1989). Isotopic

composition and origin of water from andesitic magmas,
Dokl. (Trans.) Ac. Sci. USSR, 304, 440-443.

Taran, Y., T.P. Fischer, B. Pokrovsky, Y. Sano, M.A. Armienta
and J.L. Macías (1998). Geochemistry of the volcano-
hydrothermal system of El Chichón Volcano, Chiapas,
Mexico, Bull. Volcanol., 59, 436-449.

Taran, Y. and D. Rouwet (2008). Estimating thermal inflow
to El Chichón crater lake using the energy-budget,
chemical and isotope balance approaches, J. Volcanol.
Geotherm. Res., 175, 472-481.

Taran, Y., D. Rouwet, S. Inguaggiato and A. Aiuppa (2008).
Major and trace element geochemistry of neutral and
acidic thermal springs at El Chichón volcano, Mexico.
Implications for monitoring of the volcanic activity, J.
Volcanol. Geotherm. Res., 178, 224-236.

Tassi, F., O. Vaselli, B. Capaccioni, C. Giolito, E. Duarte, E.
Fernández, A. Minissale and G. Magro (2005). The
hydrothermal-volcanic system of Rincón de la Vieja
volcano (Costa Rica): A combined (inorganic and organic)
geochemical approach to understanding the origin of the
fluid discharges and its possible application to volcanic
surveillance, J. Volcanol. Geotherm. Res., 148, 315-333.

Tassi, F., O. Vaselli, E. Fernández, E. Duarte, M. Martínez,
A. Delgado Huertas and F. Bergamaschi (2009).
Morphological and geochemical features of crater lakes
in Costa Rica: an overview. J. Limnol., 68 (2). 193-205.

Terada, A., T. Hashimoto and T. Kagiyama (2010). Volcanic
lake system at Aso Volcano, Japan: fluctuations in the
supply of volcanic fluid from the hydrothermal system
beneath the crater lake, Abstract V23A-2386, AGU Fall
Meeting.

Tuttle, M.L., M.A. Clark, H.R. Compton, J.D. Devine, W.C.
Evans, A.M. Humphrey, G. Kling, E.J. Koenigsberg, J.P.
Lockwood and G.N. Wagner (1987). The 21 August 1986
Lake Nyos gas disaster, Cameroon, Final Report Dep. Int.
USGS.

van Rotterdam-Los, A.M.D., S.P. Vriend, M.J. van Bergen and
P.F.M. van Gaans (2008). The effect of naturally acidified
irrigation water on agricultural volcanic soils. The case
of Asembagis, Java, Indonesia, J. Geochem. Explor., 96,
53-68.

Varekamp, J.C. and R. Kreulen (2000). The stable isotope geo-
chemistry of volcanic lakes, with examples from
Indonesia, J. Volcanol. Geotherm. Res., 97, 309-327.

Varekamp, J.C., G.B. Pasternack and G.L. Rowe Jr (2000).
Volcanic lake systematics II. Chemical constraints, J.
Volcanol. Geotherm. Res., 97, 161-179.

Varekamp, J.C., A.P. Ouimette, S.W. Herman, A. Bermúdez
and D. Delpino (2001). Hydrothermal element fluxes from
Copahue, Argentina; a "bee-hive" volcano in turmoil,
Geology, 29, 1059-1062.

ROUWET AND TASSI

172



173

Varekamp, J.C. (2003). Lake contamination models for evolu-
tion towards steady state, J. Limnol., 62 (1). 67-72.

Varekamp, J.C. (2008). The volcanic acidification of glacial
Lake Caviahue, Province of Neuquen, Argentina, J.
Volcanol. Geotherm. Res., 178, 184-196.

Varekamp, J.C., A.P. Ouimette, S.W. Herman, K. S. Flynn, A.
Bermúdez and D. Delpino (2009). Naturally acid waters
from Copahue volcano, Argentina, Appl. Geochem., 24,
208-220.

Voight, B., R.J. Janda, H. Glicken and P.M. Douglass (1983).
Nature and mechanics of the Mount St. Helens rockslide-
avalanche of 18 May 1980, Geotechnique, 33, 243-273.

Zlotnicki, J., Y. Sasai, J.P. Toutain, E.U. Villacorte, A. Bernard,
J.P. Sabit, J.M. Gordon Jr., E.G. Corpuz, M. Harada, J.T.
Puningbayan, H. Hase and T. Nagao (2008). Combined
electromagnetic, geochemical and thermal surveys of
Taal Volcano (Philippines) during the period 2005-2006,
Bull. Volcanol.; doi: 10.1007/s00445-008-0205-2.

*Corresponding author: Dmitri Rouwet,
Istituto Nazionale di Geofisica e Vulcanologia, sezione di Palermo,
Palermo, Italy; email: dmitrirouwet@gmail.com.

© 2011 by the Istituto Nazionale di Geofisica e Vulcanologia. All rights
reserved.

A BOX MODEL FOR ACTIVE CRATER LAKES