5.1. Magnetic Data
 The distribution and dip of faults in the study area were estimated from ground-magnetic data using Geosoft Inc. Oasis Montaj software. The total intensity of magnetization is the sum of the intensity of induced and remnant magnetization [Plouff, 1975]. We assumed constant remnant magnetization because all of the potential magnetic sources at Masaya are much younger than the last reversal of the Earth's magnetic field (∼0.78 Ma) [Cande and Kent, 1995], and tectonic rotations are thought to be minimal during this time. Therefore we were able to isolate the induced magnetization, creating a model using bulk values as we are only interested in the total vector of magnetization.
 A peak-to-peak amplitude of over 1000 nT (Figure 3b and Table 1) is consistent with vertical separation of basalt that has a strong magnetization [e.g., La Femina et al., 2002]. We consider the magnetic source to be a lava flow unit of high bulk magnetization beneath scoria of low bulk magnetization with some faulting. Other potential sources of the magnetic anomaly, such as a dike, would produce anomalies with a vector of magnetization inconsistent with those observed. Modeling is used to estimate the depth to the top of the lava beneath a veneer of scoria. We used vectors of magnetization of magnitude 2 × 10−2 for the lava flow and 1 × 10−6 for the scoria. This produces a plausible lava flow structure with a depth <20 m and a variable depth to the top of the lava of +/−10 m (Figure 4). This structural model and the observed topographic offset indicate the presence of SE-dipping normal faults in the map area.
5.3. Vadose Zone
 Our various geophysical data and models provide constraints on the shallow hydrothermal system that can be used in a numerical model. Transient electromagnetic soundings throughout the caldera indicate that the groundwater table is approximately 250 m beneath the surface in the study area [MacNeil et al., 2007], although there is also some evidence of shallower groundwater [Mauri et al., 2012]. Magnetic data and models are consistent with a faulted terrain, dominated in this area by steeply dipping (60°) normal faults. SP and CO2-flux data indicate that these faults serve as barriers to flow, rather than enhancing flow along the fault plane, and that CO2 flux anomalies average around 2 × 10−5 kg/m2s (Table 1). Shallow temperature measurements recorded over a period of several years at five different depths (0.15 to 1.5 m) at Comalito cinder cone show average temperatures of between 46 and 74°C, and maximum temperatures of between 61 and 82°C (Figure 6) [see Pearson et al., 2008]. Thus the hydrothermal system should be locally dominated by upward flow of fluids through the footwalls of steeply dipping faults, with near-surface temperatures of ∼70°C.
Figure 6. Example time series of temperature recorded in the diffuse degassing zone at Comalito cinder cone, measured at 66 cm depth. The average near-surface temperature of 70°C is used as a constraint in the numerical modeling. Temperature anomalies correspond to episodes of minor volcanic activity in Santiago crater [seePearson et al., 2008].
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 We created a rectangular, 2-dimensional grid 1000 m wide and 250 m deep (Figure 7c). In the x-direction the grid had 74 elements; 60 cells of 5 m in the center to represent the 100-m wide fracture zone and 100 m on either side, and 7 elements of 50 m adjacent to each vertical boundary (Figure 7). In this way, the vertical boundary conditions should not affect the area of interest within the fracture zone. The vadose zone above the approximately 250-m-deep water table [MacNeil et al., 2007] was represented by 25 layers of 10 m. Within the grid, less-permeable faults dipped at approximately 60° (Figure 7c).
Figure 7. (a) SP measurements recorded along profile c-c′ (Figure 3b) in 2006. (b) Modeled vertical fluid flux along a profile just below the surface, using the geometry in Figure 7c. (c) Geometry used in a TOUGH2 model to represent vadose-zone conditions. Other parameters used in the model can be seen inTable 2. The green square represents the cell where the maximum temperature was located. (d) Output of the TOUGH2 model after injecting equal parts air-water mixture at 2 × 10−4 kg/m2s for 2500 years. Black arrows represent fluid flow, and colored contours show temperature. Heat and fluid flux increase toward the footwall of the left fault.
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 Rock properties were measured by Chiodini et al. , MacNeil et al. , or are typical of fractured basalt (Table 2). Bulk permeability within the fracture zone was determined by comparison of model outputs with observed surface temperatures (see Tables S1 and S2 in the auxiliary material). Permeability of the fault was set at the same value as the wall rock for simplicity, as both are relatively impermeable compared to the fracture-zone rock. The model was run for 2500 years, approximately the length of time since the caldera formed [Williams, 1983].
Table 2. Parameters Used in Numerical Modelsa
|Permeability of wall rock||m2||1 × 10−16|
|Permeability of fault rock||m2||1 × 10−16|
|Wet heat conductivity||W/mC||1.45|
 The model was initially at atmospheric pressure and temperature (101 kPa and 25°C) with 80% air and 20% water. The upper boundary of the model domain had the same pressure and temperature conditions, but with 99% porosity and a very large volume to represent the atmosphere. This effectively fixed the temperature and pressure at atmospheric values. Both of the vertical boundaries were impermeable to flow. The lower boundary condition was set to represent the water table, and so was 100% water.
 The lower boundary condition is the major constraint of the model as it dictates how, at what temperature, and at what rate, fluids flow through the fracture zone. As the fracture zone appears to host all heat and fluid flow, this was the only area where different boundary conditions were used. Initially the pressure was atmospheric, and then water and air were injected at a temperature and rate determined by comparison of model outputs with observed surface temperatures. Either enthalpy or temperature could be prescribed; by making the rock density (and therefore specific heat capacity) arbitrarily large (2 × 1040 kg/m3) we fixed the injection temperature, thereby avoiding problems with enthalpy during phase changes. Temperatures of 70 and 120°C were used to ensure that injection of both water and water vapor were modeled (see Tables S1 and S2). Air and water were injected in equal parts as suggested by gas geochemistry [Chiodini et al., 2005]. Air was used to represent the gases in the vadose zone for a number of reasons: the TOUGH2 model can only work with water and either air or CO2, gas geochemistry suggests that at Masaya the gases are ∼45% air, 45% water, and 10% CO2 [Chiodini et al., 2005]; and we have found that the air module is more computationally stable. As the thermal properties of air and CO2 are quite similar compared to water or water vapor, we believe that the differences between the two models would not be significant.
 A range of simulations were run with fracture-rock permeability of between 1 × 10−14 and 1 × 10−10 m2and air-water injection rates of between 1 × 10−7 and 1 × 10−3 kg/m2s (Figure 8). After 2500 years of fluid injection, the air, water and water vapor mixture rise steadily, primarily controlled by the geometry of faults (Figure 7d). Most fluid gets redirected shallowly to travel along the footwall of the west fault, but some also travels between the faults (Figures 7b and 7d). Temperatures correlate with gas flux, increasing near the footwall of the west fault. Simulated fluid flux is in good agreement with fluid flux inferred from SP (Figures 7a and 7b).
Figure 8. Maximum near-surface temperature as a function of fluid injection rate. Numbers labeling the curves are permeabilities in m2. (a) Injection of equal mixture of 120°C water and air. (b) Injection of equal mixture of 70°C water and air. The observed near-surface temperature of 70°C can only be reached in the former case, when hotter fluid is injected at depth.
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 Permeability is unknown, but can be partially constrained by the known near-surface temperature. Modeling shows that a range of conditions can approximate the temperature, but that some heat is lost in transport and the base temperature must therefore be above 70°C (Figure 8b). A permeability range of between 1 × 10−11 and 1 × 10−14 m2is possible for near-surface maximum temperatures of 70°C (Figure 8a), although the low injection rate of between 1 × 10−7 and 1 × 10−6 kg/m2s required with a permeability of 1 × 10−11 m2 is not supported by surface gas flux measurements. The higher injection rates of >1 × 10−5 kg/m2s that result in 70°C maximum near-surface temperature with other permeabilities are however in good agreement. For a low permeability of 1 × 10−13 m2 or less, fluid flow (particularly water) is inhibited and the system only heats up with a high injection rate (Figure 8a). In the extreme case of a permeability of 1 × 10−14 m2 and gas flux of 1 × 10−3 kg/m2s, the fluid is trapped and pressure builds up beyond physically realistic values so that the model crashes (Figure 8b). The unusually high gas fluxes observed at Masaya volcano [Lewicki et al., 2003] suggest that fluid can flow freely through this fracture zone, and therefore we believe that a permeability of 1 × 10−12 m2 is the most appropriate. This is typical of basalt, which is in the range of 10−14 to 10−9 m2 [Freeze and Cherry, 1979; Saar and Manga, 1999; Ingebritsen and Manning, 2010].
 The injection rate at depth also needs to be determined from modeling, and comparisons with surface flux values allow us to refine the likely range. Field measurements suggest a surface CO2 flux of 2 × 10−5 kg/m2s, and as 10% of the measured gases are CO2 [Chiodini et al., 2005], this would imply a total fluid flux of 2 × 10−4 kg/m2s. Given a permeability of 1 × 10−12 m2, the models suggest that an injection rate of ∼1 × 10−4 kg/m2s would give rise to 70°C surface temperatures (Figure 8a), in good agreement with these surface measurements.
 Using these models we can estimate a number of surface parameters, and compare them with measured values where available. Using an injection rate of 2 × 10−4 kg/m2s, we find that the optimum temperature at the water table is 96°C (Figure 9a), which would allow some boiling to occur creating the water vapor that is visible at the surface. The maximum gas flux in the layer below the atmosphere is 1.7 × 10−4 kg/m2s (Figure 9b), close to the 2 × 10−4 kg/m2s measured but suggesting that the injection rate at depth may be higher than 2 × 10−4 kg/m2s. The maximum modeled surface heat flux is 162 W/m2 (Figure 9c), higher than the 91 W/m2 measured by Chiodini et al. , but feasible given the uncertainty associated with heat flux measurements.
Figure 9. Model results as a function of injection temperature given a permeability of 1 × 10−12 m2 and an injection rate of 2 × 10−4 kg/m2s. Surface (a) temperature, (b) gas flux and (c) heat flux are all of the same order of magnitude as measured surface values.
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 Fluid velocity is also affected by permeability and heat input, and these models allow us to estimate travel times. Maximum gas velocities vary from 5 × 10−5 m/s in the low-permeability models to 5 × 10−3 m/s in a system with a high permeability of 10−10 m2 and a fluid input of 1 × 10−3 kg/m2s. This implies travel times ranging from 60 days to 14 hours between the water table (250 m depth) and the surface. The velocity for our optimum model (permeability 1 × 10−12 m2, fluid injection 2 × 10−4 kg/m2s) is 6.7 × 10−4 m/s, resulting in a minimum travel time of ∼4 days for fluids to travel through the system.
 For these models, the permeability of the wall and fault rock have been assumed to be relatively impermeable at 1 × 10−16 m2, but this is an arbitrary value. For a permeability of 1 × 10−14 m2 or less the models appear to be relatively insensitive to permeability (Figure 10). Unfractured basalt is typically on the order of 1 × 10−14 m2 or less [Freeze and Cherry, 1979]. For fault permeability close to the permeability of the fracture, the near-surface temperature of 70°C is not reached (Figure 10b) because circulation increases and the fluid flow distribution is not focused, supporting our theory that relatively impermeable faults are creating barriers to fluid flow.
Figure 10. Variations given a fracture permeability of 1 × 10−12 m2, fluid injection of 2 × 10−4 kg/m2s and 96°C base temperature as determined by modeling. (a) Model geometry. (b) Maximum temperature as a function of varying either wall rock permeability (blue line with diamonds) or fault rock permeability (red dashed line with crosses). They must both be significantly less than the 1 × 10−12 m2fractured rock for model temperatures to match the measured surface temperature of ∼70°C. With a more- or less-permeable layer for the top 50 m (green dashed line with triangles) the temperature becomes too cold or too hot respectively.
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 The magnetic models suggest a shallow scoria layer underlain by lava. However, a uniform permeability distribution appears to give the best fit to the data. Adding a more permeable shallow layer results in cooler temperatures near the fault, while a less permeable layer results in hotter temperatures (Figure 10). We infer that the variation in permeability between fractured lava and scoria is less significant than their very different magnetic properties.
5.4. Saturated Zone
 The degassing system at Masaya volcano is unusually stable and open, and the distribution of diffuse degassing can be used to infer variations in fluid flux at depth. The three distinct zones of diffuse degassing along the fracture zone on the flank of Masaya volcano (Figure 1) are all characterized by stable and consistently elevated temperature and CO2 flux. Soil temperatures at the area of diffuse degassing at Comalito cinder cone range between 55 and 80°C [Pearson et al., 2008], and are similar to those measured at the degassing area closest to the crater (38 to 83°C). Measured CO2 fluxes are similar at all three degassing zones, varying between 2.3 and 2.6 × 10−5 kg/m2s. At Comalito cinder cone, we measured gas fluxes that were stable over remarkably long periods of time, varying by less than 10% from average over 5 months. The uniformity in temperature and CO2 flux across the three degassing zones suggests that they share a common source. Gases at Comalito cinder cone, 3.5 km from the active crater, retain a magmatic component [Lewicki et al., 2003], but as deep groundwater flow is from Comalito toward the active crater [MacNeil et al., 2007] there must either be separate magmatic sources, or one extensive source at depth. Lava ponding in the active crater in 2006 corresponded to changes in temperature at the Comalito cinder cone diffuse degassing zone [Pearson et al., 2008], suggesting that there may in fact be an aerially extensive magmatic source at depth beneath the caldera.
 For three distinct degassing zones to occur with one magmatic source, the source must either be focused by variations in permeability or from convection of groundwater. To determine whether the consistent spacing, temperatures and gas fluxes in the distinct degassing zones can be explained by groundwater convection, we created a TOUGH2 model of the saturated zone. We used a grid of 35 by 110 cells extended over 3500 m depth and 5500 m length respectively (Figure 11a). The model was 2D; the third dimension was one cell of 100 m to represent the thickness of the fracture zone. Horizontally the fracture was modeled as the 3500 m-long high-permeability zone, plus 1000 m of lower permeability on either side (Table 2 and Figure 11a) to ensure that the boundary conditions did not affect the model. A depth of 3500 m was assumed for simplicity, as there is no information to constrain depth to the source. The entire model was fully saturated at atmospheric temperature and pressure initially. The top boundary was fixed at atmospheric conditions, and the vertical boundaries were no flow. Heat or fluid was injected at various rates along the bottom boundary to replicate the volcanic source. This model was also run for 2500 years.
Figure 11. (a) Geometry used for numerical model, representing a cross-section of the flank fracture at Masaya (Figure 1b) within the saturated zone. Other parameters can be found in Table 2. (b) Model results with heat injection of 5 W/m2. (c) Model results with hot fluid injected at 5 × 10−5 kg/m2s with an enthalpy of 632 kJ/kg.
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 TOUGH2 modeling shows that a range of conditions can give rise to convection along the length of the fracture system (Figure 11 and Table S3). A permeability of between 1 × 10−14 and 1 × 10−11 m2 was modeled with either heat or fluid injected at a variety of rates (Figure 12). Heat was injected at between 0.1 and 50 W/m2. Fluid was injected at an enthalpy of 632 or 1300 kJ/kg (∼150°C or 300°C respectively), and with an injection rate of between 1 × 10−6 and 1 × 10−4 kg/m2s. With a permeability of 10−14 m2 the system was generally conductive, but for higher permeability convection readily occurred.
Figure 12. Maximum near-surface temperature as a function of injection rate. Numbers labeling the curves are permeabilities in m2. (a) Heat injection. (b) Water injection with an enthalpy of 1300 kJ/kg (∼300°C). (c) Water vapor injection with an enthalpy of 632 kJ/kg (∼150°C). The modeled water table temperature of 96°C can be approximated with any of these scenarios.
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 The water table temperature estimated in the vadose-zone model provides a constraint when determining the permeability and injection rate within the saturated-zone system. When injecting heat, the permeability must be between 1 × 10−13 m2 and 1 × 10−12 m2for the near-surface temperature of 96°C to be reached (Figure 11a). The heat injection rate must be around 5 W/m2 for the lower permeability, and greater than 20 W/m2 for the higher permeability (Figure 12a). When fluid is injected, 96°C can be reached with any permeability, but with an injection rate of 1 × 10−5 kg/m2s or higher for 300°C fluid (Figure 12b), and closer to 1 × 10−4 kg/m2s for 150°C fluid (Figure 12c). Therefore any of these scenarios are theoretically possible.
 Three distinct diffuse degassing zones are observed at Masaya volcano (Figure 1), and therefore the model needs to test if three distinct plumes can be produced. This can be done with a range of heat injection, although the number of plumes does not vary directly with injection rate (Figure 13a). The higher the heat injection rate, the larger the number of plumes for a permeability of 1 × 10−11 m2 but it is extremely unlikely that the saturated zone is more permeable than the vadose zone. There are only two heat injection rates that result in three plumes for a permeability of 1 × 10−12 m2; 1 W/m2 and 20 W/m2 (Figure 13a). Both of these result in surface temperatures that are too cold (Figure 12a). For a permeability of 1 × 10−13 m2, injection rates of 5 W/m2 or higher result in plumes that reach the surface, and in some cases with 3 plumes (Figure 13a). An injection rate of 5 W/m2 gives a groundwater table fluid temperature of ∼90°C (Figure 12a) and therefore this could be a possible source of the groundwater convection. This results in three regularly spaced, consistent convective plumes (Figure 11b).
Figure 13. Number of plumes as a function of injection rate. Numbers labeling the curves are permeabilities in m2. (a) Heat injection. (b) Water injection with an enthalpy of 1300 kJ/kg (∼300°C). (c) Water vapor injection with an enthalpy of 632 kJ/kg (∼150°C). Three plumes are observed at Masaya volcano, and can be replicated with any of these models.
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 It appears that hot water entering the groundwater system at 3500 m depth also results in groundwater convection with three plumes. For a hotter fluid at around 300°C, it requires a fairly high permeability of 1 × 10−4 kg/m2s for three plumes to develop (Figure 13b). However, with a permeability of 1 × 10−12 m2 and injection of ∼150°C fluid, three plumes develop irrespective of fluid injection rate (Figure 13c). A groundwater table fluid temperature of 96°C results from a fluid injection rate of 5 × 10−5 kg/m2s (Figure 12c). This results in a more random pattern of convection than with simple heat injection (Figure 11c), but it is stable over 1000 years.
 Although modeling shows that both a uniform heat source and a uniform fluid source at depth can produce three zones of diffuse degassing, comparison of heat and fluid flow rates allow us to determine which is more likely. When heat is injected at 5 W/m2 into rock with a permeability of 1 × 10−13 m2, the maximum heat flux in the layer below the upper boundary layer is 17 W/m2 and the maximum fluid flux is 5 × 10−5 kg/m2s; significantly lower than the vadose zone model would suggest. However, when water is injected at 5 × 10−5 kg/m2s with an enthalpy of 632 kJ/kg into rock with a permeability of 1 × 10−12 m2, the maximum heat flux is 195 W/m2 and the maximum fluid flux is 5 × 10−4 kg/m2s. These are in good agreement with vadose zone values of 2 × 10−4 kg/m2s and heat flux of 162 W/m2. Therefore it seems entirely possible that hot fluid entering the groundwater system at around 3500 m is driving groundwater convection that results in the three distinct zones of diffuse degassing observed at Masaya volcano.