Volcanological perspectives on Long Valley, Mammoth Mountain, and Mono Craters: several contiguous but discrete systems

The inversion of local earthquake data (LED) for three-dimensional velocity structure requires the simultaneous solution of the coupled hypocenter-model problem. The Aki-Christoffersson-Husebye method (ACH) involves the inversion of large matrices, a task that is often performed by approximative solutions when the matrices become too big, as is the case for most LED, considering the coupled inverse problem. Such an approximate method (herein referred to as approximate geotomographic method) is used to perform tests with LED to obtain the best suited inversion parameters, such as velocity damping and number of iteration steps. The ACH method has been proposed for use of teleseismic data. Several adjustments to the original ACH method, which are necessary for use of LED, have been developed and are discussed. Such adjustments are the separation of the unknown hypocentral from the velocity model parameters for the inversion, the use of geometric weighting and step length weighting, the calculation of a minimum one-dimensional (1D) model as the starting three-dimensional (3D) model for the model inversion, and the display of an approximate resolution matrix (ray density tensors) before the inversion is performed. The ray density tensors allow the block cutting, e.g., the definition of the 3D velocity grid, to better correspond with the resolution capability of the specific data set. The adjustments to the method are tested by inversion of realistic LED of known variance. Synthetic LED are also used to demonstrate the effects of systematic errors, such as mislocations of seismic stations, on the resulting velocity field. Using the data sets from Long Valley, California, Yellowstone National Park, Wyoming, and Borah Peak, Idaho, the effects of improvements to the ACH method and of the data filtering process are shown. The use of the minimum 1D models for routine earthquake location improves this location procedure, as shown with the relocation of shots for the Long Valley and Yellowstone areas. The three-dimensional velocity fields obtained by the ACH method for the Long Valley and Yellowstone areas show local anomalies in the p velocity that can be correlated with tectonic and volcanic features. A pronounced anomaly of low p velocity below the Yellowstone caldera can be interpreted as a large magma chamber. However, the bulk of the paper addresses problems of the inversion method. The LED from the areas mentioned above are used to numerically and theoretically tune the inversion method for the defects that all real data contain. It is shown that one of the most important steps for any inversion of LED is the selection of the data for quality and for geometrical distribution.

Geotomography with local earthquake data

and the roofward decline in liquidus temperature of the zoned melt, prevented significant crystallization against the roof, consistent with dominance of crystal-poor magma early in the eruption and lack of any roof-rind fragments among the Bishop ejecta, before or after onset of caldera collapse. A model of secular incremental zoning is advanced wherein numerous batches of crystal-poor melt were released from a mush zone (many kilometers thick) that floored the accumulating rhyolitic melt-rich body. Each batch rose to its own appropriate level in the melt-buoyancy gradient, which was selfsustaining against wholesale convective re-homogenization, while the thick mush zone below buffered it against disruption by the deeper (non-rhyolitic) recharge that augmented the mush zone and thermally sustained the whole magma chamber. Crystal^melt fractionation was the dominant zoning process, but it took place not principally in the shallow melt-rich body but mostly in the pluton-scale mush zone before and during batchwise melt extraction.

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Compositional Zoning of the Bishop Tuff

Volcanic processes operate over a wide range of time scales that require a variety of instruments and techniques for their study. Short-period seismology typically covers the band 0.1–1 s, while broadband seismology can extend that band to 0.1–100 s. Borehole strainmeters may be used to cover the very wide band from 0.1 s to 100 days and Global Positioning System (GPS) surveys are useful to track deformation over time scales ranging from days to decades. Arrays of three-component broadband seismographs coupled with arrays of borehole strainmeters are required to monitor the dynamics of magmatic and hydrothermal fluids prior to and during eruptions. Inversions of broadband data may be performed to image the forces operating at a source and infer the fluid pathway geometry and mass transport balance in a volcano. The accuracy of such inversions depends on the degree of resolution achieved for the volcanic structure. High-resolution tomography based on iterative inversions of seismic travel-time data can image three-dimensional structures at a scale of a few hundred meters provided adequate local short-period earthquake data are available. Hence, forces in a volcano are potentially resolvable for periods longer than a few seconds. Studies of long-period events and tremor at periods ≤ 1s offer constraints on pressure fluctuations resulting from unsteady mass transport, and observations of volcano-tectonic activity tell us about hydraulic fracturing processes, as well as brittle response of the rock matrix to stress changes induced by rapid injection and/or withdrawal of fluids. Short-period networks are routinely used to monitor, locate, and assess the source properties of volcano-tectonic earthquakes and long-period events, while small-aperture seismic antennas are required to track tremor and separate source, path, and site effects in tremor wavefields. This chapter reviews quantitative methods used in the interpretation of source processes and structures in volcanoes and discusses some of the major challenges faced by volcano seismologists in their quest to understand eruptive behavior.

New Methods and Future Trends in Seismological Volcano Monitoring

Mars’s seismic activity and noise have been monitored since January 2019 by the seismometer of the InSight (Interior Exploration using Seismic Investigations, Geodesy and Heat Transport) lander. At night, Mars is extremely quiet; seismic noise is about 500 times lower than Earth’s microseismic noise at periods between 4 s and 30 s. The recorded seismic noise increases during the day due to ground deformations induced by convective atmospheric vortices and ground-transferred wind-generated lander noise. Here we constrain properties of the crust beneath InSight, using signals from atmospheric vortices and from the hammering of InSight’s Heat Flow and Physical Properties (HP3) instrument, as well as the three largest Marsquakes detected as of September 2019. From receiver function analysis, we infer that the uppermost 8–11 km of the crust is highly altered and/or fractured. We measure the crustal diffusivity and intrinsic attenuation using multiscattering analysis and find that seismic attenuation is about three times larger than on the Moon, which suggests that the crust contains small amounts of volatiles. The crust beneath the InSight lander on Mars is altered or fractured to 8–11 km depth and may bear volatiles, according to an analysis of seismic noise and wave scattering recorded by InSight’s seismometer.

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Constraints on the shallow elastic and anelastic structure of Mars from InSight seismic data

During the summers of 1982 and 1983, field work was completed on 10 new seismic refraction profiles in the Long Valley-Mono Craters region located within the technically active zone between the eastern Sierra Nevada and the Basin and Range province. The four profiles presented here sample the upper 7–10 km of the crust within the western half of Long Valley cladera, the Mono craters ring fracture system, and the “normal” crust north and northeast of Long Valley. The entire area shares a common crystalline basement consisting primarily of Mesozoic granitic rocks. At a depth of 2 km beneath the topographic surface, the P wave velocity is uniformly 5.6 km/s beneath which the velocity increases with a gradient of about 0.1 s−1. Clear secondary arrivals observed on a profile crossing the west part of the caldera support evidence from 1972 seismic refraction profiles for a reflecting boundary at a depth of 7–8 km beneath the west margin of the resurgent dome, which may represent the top of a magma chamber. Structural details resolved in the upper 2 km of the crust lead to the following conclusions: (1) P wave velocities in the crystalline basement at depths shallower than 2 km vary from 3.6 to 5.0 km/s. This variation reflects differences in the degree of fracturing of the upper portion of the basement. (2) An apparently homogeneous basement in the upper 7–10 km beneath the Mono Craters ring fracture system indicates that any extensive magma chamber must be at least 10 km deep there. (3) The downdropped crystalline basement beneath the western half of Long Valley caldera dips genty (5°–10°) to the northeast. Apparent offsets along steeply dipping caldera-bounding faults are 1.0–1.7 km along the northern, western, and southern sides of the caldera with largest offsets in the north. (4) Rocks forming the caldera fill fall into essentially three P wave velocity groups. From the surface downward these velocity groups are 100–400 of 1.2–1.8 km/s material corresponding to unconsolidated fluvial or lacustrine deposits or highly fractured rhyolite flows, 200–400 m of 2.8–3.1 km/s material corresponding to sparsely jointed postcaldera rhyolite, rhyodacite and basalt flows, and approximately 1000 m of 3.9–4.4 km/s material corresponding to deposits of Bishop Tuff.

J. H. Luetgert

Papers

Constraints on the upper crustal structure of the Long Valley‐Mono Craters Volcanic Complex, eastern California, from seismic refraction measurements