Distribution and Transport of Thermal Energy within Magma–Hydrothermal Systems

Proximity to magma bodies is generally acknowledged as providing the energy source for hot hydrothermal reservoirs. Hence, it is appropriate to think of a “magma–hydrothermal system” as an entity, rather than as separate systems. Repeated coring of Kilauea Iki lava lake on Kilauea Volcano, Hawaii, has provided evidence of an impermeable, conductive layer, or magma–hydrothermal boundary (MHB), between a hydrothermal system and molten rock. Crystallization on the lower face of the MHB and cracking by cooling on the upper face drive the zone downward while maintaining constant thickness, a Stefan problem of moving thermal boundaries with a phase change. Use of the observed thermal gradient in MHB of 84 °C/m yields a heat flux of 130 W/m2. Equating this with the heat flux produced by crystallization and cooling of molten lava successfully predicts the growth rate of lava lake crust of 2 m/a, which is faster than simple conduction where crust thickens at t and heat flux declines with 1 / t . However, a lava lake is not a magma chamber. Compared to erupted and degassed lava, magma at depth contains a significant amount of dissolved water that influences the magma’s thermal, chemical, and mechanical behaviors. Also, a lava lake is rootless; it has no source of heat and mass, whereas there are probably few shallow, active magma bodies that are isolated from deeper sources. Drilling at Krafla Caldera, Iceland, showed the existence of a near-liquidus rhyolite magma body at 2.1 km depth capped by an MHB with a heat flux of ≥16 W/m2. This would predict a crystallization rate of 0.6 m/a, yet no evidence of crystallization and the development of a mush zone at the base of MHB is observed. Instead, the lower face of MHB is undergoing partial melting. The explanation would appear to lie in vigorous convection of the hot rhyolite magma, delivering both heat and H2O but not crystals to its ceiling. This challenges existing concepts of magma chambers and has important implications for use of magma as the ultimate geothermal power source. It also illuminates the possibility of directly monitoring magma beneath active volcanoes for eruption forecasting.

Imaging the magmatic system beneath the Krafla geothermal field, Iceland: A new 3-D electrical resistivity model from inversion of magnetotelluric data

SUMMARY Krafla is an active volcanic field and a high-temperature geothermal system in northeast Iceland. As part of a program to produce more energy from higher temperature wells, the IDDP-1 well was drilled in 2009 to reach supercritical fluid conditions below the Krafla geothermal field. However, drilling ended prematurely when the well unexpectedly encountered rhyolite magma at a depth of 2.1 km. In this paper we re-examine the magnetotelluric (MT) data that were used to model the electrical resistivity structure at Krafla. We present a new 3-D resistivity model that differs from previous inversions due to (1) using the full impedance tensor data and (2) a finely discretized mesh with horizontal cell dimensions of 100 m by 100 m. We obtained similar resistivity models from using two different prior models: a uniform half-space, and a previously published 1-D resistivity model. Our model contains a near-surface resistive layer of unaltered basalt and a low resistivity layer of hydrothermal alteration (C1). A resistive region (R1) at 1 to 2 km depth corresponds to chlorite-epidote alteration minerals that are stable at temperatures of about 220 to 500 °C. A low resistivity feature (C2) coincides with the Hveragil fault system, a zone of increased permeability allowing interaction of aquifer fluids with magmatic fluids and gases. Our model contains a large, low resistivity zone (C3) below the northern half of the Krafla volcanic field that domes upward to a depth of about 1.6 km b.s.l. C3 is partially coincident with reported low S-wave velocity zones which could be due to partial melt or aqueous fluids. The low resistivity could also be attributed to dehydration and decomposition of chlorite and epidote that occurs above 500 °C. As opposed to previously published resistivity models, our resistivity model shows that IDDP-1 encountered rhyolite magma near the upper edge of C3, where it intersects C2. In order to assess the sensitivity of the MT data to melt at the bottom of IDDP-1, we added hypothetical magma bodies with resistivities of 0.1 to 30 Ωm to our resistivity model and compared the synthetic MT data to the original inversion response. We used two methods to compare the MT data fit: (1) the change in r.m.s. misfit and (2) an asymptotic p-value obtained from the Kolmogorov–Smirnov (K–S) statistical test on the two sets of data residuals. We determined that the MT data can only detect sills that are unrealistically large (2.25 km3) with very low resistivities (0.1 or 0.3 Ωm). Smaller magma bodies (0.125 and 1 km3) were not detected; thus the MT data are not sensitive to small rhyolite magma bodies near the bottom of IDDP-1. Our tests gave similar results when evaluating the changes in r.m.s. misfit and the K–S test p-values, but the K–S test is a more objective method than appraising a relative change in r.m.s. misfit. Our resistivity model and resolution tests are consistent with the idea of rhyolite melt forming by re-melting of hydrothermally altered basalt on the edges of a deeper magma body.