VOLCANISM ALONG THE EASTERN SIERRA NEVADA FRONTAL SCARP

The Sierran frontal fault system marks the western boundary of the Basin and Range extensional domain. It is where the Sierran-Great Valley block is shearing away from Great Basin as it rotates counterclockwise along the Eastern California shear zone.

This province, particularly between the eastward extensions of the Murray and Mendocino fracture zones, is considered a "slab-free" zone where an upwelling of aesthenosphere replaced the descending lithosphere to produce late Tertiary and Quaternary magmatic and Volcanic cycles.

Fractures developed in response to the tensional forces throughout the Basin and Range. At the western margin, they facilitated the rise of the Sierra Nevada arc batholith and the rise of basaltic magmas. The ensuing volcanics represent activity that predated the formation of substantial relief of the Sierra Nevada and continued without significant lull until 200 years ago. Resurgent doming and seismic activity at the Long Valley caldera is a modern reminder that this region is still a potential volcanic threat.

This study explores major volcanism along the Sierran frontal fracture zone from the Coso Range northward to Mono Lake. These centers of volcanism roughly exhibit a northward progression of their inceptions, although none of them can be exempted from future activity. The most highly developed magma chamber exists at Long Valley, and most scientific research has been focused in that area.

COSO RANGE

Located along the southern Sierran fault scarp, the Coso Range is slightly older than the inception of the rise of the Sierra Nevada in this western edge of the Great Basin. Approximately 35 km3 of volcanic rocks were spread over roughly 400 km² of Mesozoic plutonic and older metamorphic rocks during two major eruptive episodes [Duffield, Bacon and Dalrymple, 1980] (refer to figure 1). Fumarole activity and intermittent hot springs has prompted interest in the Coso area as a potential hydrothermal resource region. Any residual magma chamber, however, does not exist as a large, shallow reservoir. The vapor-dominated geothermal system that does exist derives its heat primarily from the most recent volcanic episode.

The first period of volcanism began roughly 4 million years ago as lithospheric fractures developed as a result of regional extension. The ensuing 1.5 million years witnessed the extrusion of 3.1 km3 of volcanic rocks (roughly 90% of the Coso field total) which generally evolved continuously toward more silicic lavas [Duffield, Bacon and Dalrymple, 1980].

Basalt flows dominated the earliest volcanism, erupting from widely scattered vents marked now by eroded cinder cones. Great lateral extent of the flows suggests a subdued terrain existed until about 3 million years ago, when the Sierra Nevada batholith rose along the eastern faults. Several basalt flows of 2-5 m thicknesses were spread laterally for several kilometers. Local aggregate accumulations exceed 120 meters in some places. Most of the basalts contain a few percent of plagioclase and olivine phenocrysts, with some flows containing also clinopyroxene phenocrysts [Duffield, Bacon and Dalrymple, 1980].

Compositional evolution of the magma is evident from the chemically more evolved andesitic and dacitic extrusions which overly the basalts [Duffield, Bacon and Dalrymple, 1980]. The less voluminous and widespread andesites typically exhibit up to 15 % phenocrysts of plagioclase and green clinopyroxene with minor olivine and rare orthopyroxene. Andesites near Haiwee Ridge and Volcano Butte comprise the bulk of polygenetic volcanoes plugged by dacite. Dacite plugs commonly contain as much as 30 % andesitic and basaltic inclusions up to several meters in size. Intricately shaped boundaries between inclusions and the dacite suggest some of the rocks represent mechanical mixing of dacite magma with hotter, less viscous mafic magma [Duffield, Bacon and Dalrymple, 1980]. Dacite flows also exist near Haiwee Ridge, Volcano Butte, and as two flows 3 km northeast of Coso Hot Springs. These dacites contain 10-30 % phenocrysts of plagioclase with various combinations of quartz, cpx, opx, amphibole, and biotite.

Dating of the volcanic rocks substantiates Duffield et al.’s claim that the magma chamber was evolving toward more silicic compositions. Most basalts have been dated between 3.7 to 3.3 Ma. Andesites generally overly the more voluminous basalts while dacites are younger still with less areal extent than the andesites.

By about 3 Ma, magmatic evolution had created a highly silicic melt which resulted in a major pyroclastic eruption of rhyodacite originating from the south end of Haiwee Ridge. Repeated tappings of the magma chamber blanketed the region with air-fall pumice typically several meters thick with a maximum thickness of 12 meters near the vent. Rhyodacite flows and pyroclastic lithic clasts exceeding 1.5 meters across occur near the vent. Ash flows from Haiwee Ridge deformed and incorporated underlying plastic lacustrine sediments. This and the development of a layered structure and contact zones enriched with vapor-phase minerals suggests eruption occurred partially within a lake [Duffield, Bacon and Dalrymple, 1980]. Volcanics ranging compositionally from andesite through rhyolite played only a relatively minor role after the Haiwee eruption for an additional 0.5 million years.

The second major volcanic episode is compositionally bimodal: consisting of alkali basalt and rhyolite volcanics spanning 1.1 to 0.4 Ma. Basalt is generally either older or younger than the rhyolite but local complexities make eruptive history difficult to decipher [Duffield, Bacon and Dalrymple, 1980]. Basalts erupted first from roughly 10 cinder cones typically contributing 1-5 flows, each averaging several meters in thickness. Basalt apparently began erupting again about 0.5 Ma from nine or more vents. All of the basalts contain as much as 30 % phenocrysts of variable ratios of olivine, cpx, opaque oxides and plagioclase. Plagioclase crystals have been found as large as 5 cm in some flows and others contain pebble to cobble-sized inclusions of granitic rock [Duffield, Bacon and Dalrymple, 1980].

The Coso Range is probably most noted for its Pleistocene rhyolite field [Wood, 1990]. This area of ~100 km2 is centered around Sugarloaf Mountain, and contains 38 rhyolite domes. Most domes and flows are younger than 0.15 Ma with the youngest dated at ~44,000 years old. Only 5 domes are older than 0.3 Ma according to Duffield et al. The entire rhyolite field is mantled by pyroclastic debris; and many domes are surrounded by explosion rings, recording the repeated explosions that occurred prior to dome formation. Some pyroclastic debris travelled as far as 20 km to the east suggesting some fairly violent activity although it pales in comparison to the Haiwee explosion [Duffield, Bacon and Dalrymple, 1980].

The petrography of the rhyolite is unusually homogeneous throughout the field. All the domes, except for one, consist of rhyolites containing less than 2 % phenocrysts of quartz, sanidine, oligoclase, opx, cpx, fayalite, ilmenite, horneblende, and biotite. A single dome contains ~7 % phenocrysts of the same minerals. Inclusions of vesicular basalt and granitic rocks are scattered throughout some domes. Chemical analyses have distinguished seven general sets of trace element variations between the rhyolitic extrusions which suggests to Bacon et al. [1979], that discrete bodies of magma erupted from a single magma reservoir. Volcanism ceased in the Coso Range with the final activity expiring at Sugarloaf Mountain 44,000 years ago.

The geothermal system at Coso is concentrated within the central Pleistocene rhyolite field, centered 3 km east of Sugarloaf Mountain (figure 2). High heat flow, low apparent resistivity, and significant fumarolic activity indicate an active geothermal system [Bacon, Duffield, and Nakamura, 1980].

Geothermal activity is localized along a north-northeast trending region of en echelon normal faults between Sugarloaf Mountain and Coso Hot Springs. These faults also parallel anamolous electrical resistivity readings and are subparallel to the strikes of normal faults associated with Basin and Range extension [Duffield et al., 1980]. Apparently the fractures are preferred conduits for geothermal fluids. Heat flow measurements reveal a maximum of 23 HFU based on temperature gradients in drill holes. By comparison, the heat flow at Long Valley Caldera surpasses 48 HFU. The rhyolite field is geophysically characterized by low resistivity [Fox, 1978] and low telluric J value [Jackson and O’Donnell, 1980].

Hot springs activity is intermittent at Coso Hot Springs where surface flow depends upon sufficient local precipitation. An anamolously low poisson ratio found by Combs and Rotstein [1976] suggested to them a highly fractured and vapor dominated basement rock. However, chloride-rich hot waters from wells suggest a hot water geothermal system. Thorough alteration of basement rocks adjacent to some domes has created fracture-filling amorphous silicas, sulfates, sulfur and cinnabar, adding evidence to at least a former hot water system [Duffield et al., 1980]. Altered rocks of fumarolic areas were onced mined for their mercury content [Ross and Yates, 1943].

The heat which drives the convective geothermal system is believed to be derived from a residual magma chamber or region of partial melt detected by teleseismic data [Reasenberg et al., 1980]. This low-velocity zone is up to 10 kilometers in diameter and centered a few kilometers southeast of Sugarloaf Mountain. It has been suggested that the low velocity readings are more in line with a major crustal discontinuity that bisects the range. Seismic refraction studies by Walter and Weaver [1980] show that the crust behaves brittly to a depth of at least 8 km. Geophysical evidence does not provide obvious substantial evidence for the existence of a significant magma reservoir below the Pleistocene rhyolite field [Duffield et al., 1980].

Other evidence, however, may suggest that there exists a growing magmatic system below the Coso Range. High levels of local seismicity and fractured flows and a general trend toward greater extruded volumes of aphyric rhyolite may suggest that a magma chamber is developing from the partial melt of the basement rocks resulting from injections of basalt [Bacon, Duffield, and Nakamura, 1980].

BIG PINE VOLCANIC FIELD

Located between Independence and Big Pine, the Big Pine volcanic field exposes roughly 120 km2 of extruded material originating from more than 40 vents [Moore and Dodge, 1980]. Vents are localized around normal faults throughout the field. The greatest volumes of lava have seemingly emerged from between two abutting northwest-southeast oriented grabens that comprise the floor of the Owens Valley basin [Wood, 1990].

The Big Pine volcanics are chiefly olivine alkali basalts that began erupting ~1.2 Ma. The largest domes rise 250 m and the largest flows are 3 km wide by 9 km long. The basalts contain abundant granitic xenoliths derived from the plutons through which it passed. Some contain peridotite (lherzolite and wehrlite) xenoliths which were carried from the upper mantle and exhibit metamorphic fabrics [Wood and Kienle, 1990]. Rhyolites were also emplaced, but only a single rhyolite dome was created in the later stages of volcanism, and probably represents assimilation of country rock and differentiation of source material. A general trend toward more silicic magmas is evident in individual vents, but is not clearly evident for the volcanic field as a whole [Wood and Kienle, 1990]. Most flows were dated to be between 0.5 Ma and 0.1 Ma [Wood, 1990].

The volcanics of the Big Pine field have been interpreted as extrusions of rapidly rising magma plumes originating in the upper mantle [Darrow, 1972]. These melts must have risen fast enough to prevent the destruction of the ultramafic rocks containing metamorphic textures which would not have lasted long in a magma chamber equilibrating to the pressure-temperature realm of near surface conditions. The magma’s composition would change, particularly since the intruded plutons would easily enrich the hot melt with silicic minerals. The rapid rise required a long and continuous pathway which may best be facilitated by the reduced confining pressures and faulting of the Basin and Range regional extension.

LONG VALLEY

The most likely location for a large residual magma body along the eastern Sierran frontal scarp is at Long Valley. It is one of only a few places in the continental United States where magma comes intriguingly close to the Earth’s surface. It is one of only three locations in the conterminous U.S. where a major caldera-forming explosion occurred within the last million years (Yellowstone and Valles are the others). The Department of Energy is drilling at Long Valley caldera for the purpose of developing geothermal energy while scientific investigation monitors the region as a potential hazard.

Long Valley has recently reminded scientists of its danger when it shook the region with four 6+ magnitude quakes one week after the eruption of Mt. St. Helens. A crustal bulge had risen a foot during the two years prior to the St. Helens eruption and has continued to grow since at a slower rate [Hill, Bailey, and Ryall, 1985]. It is believed that the bulging is the result of magma injection into the chamber below the caldera. The caldera formed as a result of a catastrophic eruption (0.73 Ma) which emitted 600 km3 of pumice and ignimbrite (roughly 550 times the volume of pyroclastics that were spewn from Mt. St. Helens in 1980). An eruption of that scale would affect the lives of millions today and causes close watch of any changes in caldera activity.

The precaldera volcanic history begins with the widespread venting of basalts and andesites at 3.2 Ma. This venting occurred over much of the region at the inception of the rise of the Sierra Nevada. This period of volcanism lasted until 2.6 Ma and extruded trachybasaltic, andesitic and quartz latitic lavas over an area greater than 4000 km2 stretching from the Adobe Hills to the northeast up into the High Sierra of the southwest [Bailey, 1984]. The quartz latites were extruded somewhat after the more mafic lavas were and occur primarily within the vicinity of the present day caldera (specifically along the north rim and along San Joaquin Ridge). The quartz latites are perhaps representative of the differentiation of magma indicative of a magma chamber developing in a large and shallow region beneath the present caldera site [Bailey, 1984].

The earliest volcanics that can be attributed directly to the magma chamber are the rhyolites of Glass Mountain [Bailey, Dalrymple and Lanphere, 1976] (refer to figure 3). These comprise a complex of domes, flows, and tephras emplaced during the period 1.9 to 0.8 Ma.

Fig. 3. Geologic map of the Long Valley and Mono Basin regions showing distribution of volcanic rocks. HSF, Hartley Springs Fault; HCF, Hilton Creek Fault; SLF, Silver Lake Fault; WCF, Wheeler Crest Fault; CD, Casa Diablo; HC, Hot Creek. From Hill, Bailey, and Ryall, [1985].

The accumulated thickness of Glass Mountain exceeds 1000 meters and is surrounded by pumice falls, ash falls, ash flows, blocky avalanche deposits, and epiclastic conglomeratic sediments [Rinehart and Ross, 1957]. Glass Mountain rises along the northeast rim of the modern caldera and its southern slopes and flanks are now residing within the caldera beneath 1500 meters of Bishop Tuff. The collapse of Glass Mountain’s southern side allows vents and intrusive centers to be well exposed along its southern face. The vents are arranged in an arc which subparallels the caldera implying an incipient ring fracture system may have leaked the rhyolites from the magma chamber [Bailey, Dalrymple and Lanphere, 1976].

The Glass Mountain rhyolites are very high in silica (77 %) and generally are aphyric to sparsely porphyritic with sanidine and biotite. According to investigations by Metz and Mahood [1983], the Glass Mountain volcanics are chemically more evolved than the Bishop Tuff which was ejected later. They offer two long-term trends which explain the tendency toward less-evolved compositions based on geochemical studies. First, an apparent introduction of mafic magma into the chamber enriched the volcanics in Manganese and Rubidium around 1.74 Ma. Also, successively younger lavas exhibit mineral assemblages indicative of progressively higher temperatures. These lines of evidence along with the relatively continuous period of eruptions between 1.9 and 0.8 Ma suggests repeated additions of mafic magma into the precaldera chamber [Metz and Mahood, 1983].

The roof of the magma chamber ruptured catastrophically at 0.73 Ma ejecting 600 km3 of magma which became to be known as the Bishop Tuff. The Mt. St. Helens eruption ejected approximately 0.18 % of this volume. Ash flows from Long Valley inundated over 2200 km2 [Hildreth and Mahood, 1986], primarily within upper Owens Valley, Adobe Valley, and Mono Basin, as well as surmounting the Sierra crest through the gorge of the Middle Fork San Joaquin River and possibly reached the edge of the San Joaquin Valley [Bailey, 1984]. Another ash flow travelled fifty miles southward through the Owens Valley and past the present site of Bishop, which gave it its name. Fallout from the ash clouds left a recognizable ash layer as far away as central Nebraska (see figure 4) [Harris, 1988]. The Bishop Tuff itself is crystal rich and high in silica (75-77%). It may contain as much as 30 % phenocrysts of quartz, sanidine, plagioclase, biotite, and Fe-Ti oxides.

The volume of the cavity left after magma ejection caused the roof to collapse while a 1.5 km depth of Bishop Tuff buried the remains. Detailed analysis of the Bishop Tuff by Hildreth [1974] revealed that the ash deposits represent the stratigraphically inverted contents of the magma chamber which emptied from the top down. The Bishop Tuff ash falls erupted at a temperature of 745 degrees Celcius and a pressure of at most 2 kbar while the ash flows erupted at 800 degrees Celcius and a pressure of at least 3 kbar [Hildreth, 1974]. This implies that the magma chamber’s roof was at ~6 km depth when the eruption began and the final ash flows came from ~10 km deep. The seismic refraction study of Hill [1976] and the gravity model of Kane, Mabey and Brace [1976], suggest that the low-density fill within the caldera is approximately 3 km thick which correlates nicely with the geochemical information of Hildreth.

Fig. 4. Distribution of ashfall from the eruption at Long Valley Caldera (LV). It thins to about half an inch at its eastern extremity.

Probably greater than 20 km3 of accidental lithic fragrments accompanied the magma upon ejection [Wildreth and Mahood, 1986]. These rocks, derived from the local lithological differences within the caldera, can then be used to approximate the vent locations as the eruption(s) progressed. Wildreth and Mahood [1986] used this idea to determine the chronological development of the Long Valley caldera eruption. Lithics within the initial fallout did not consist of any of the Glass Mountain obsidian or northern rim basalts, but did have a significant volume of metasediments. This places the initial vent along an intracaldera projection of the Mt. Morrison roof pendant. This, in combination with a quartzite fraction and lithics of the Wheeler Crest quartz monzonite, places the vent location where these lithologies coexist: a narrow zone immediately east of the Hilton Creek Fault. Correlation of ash flows with the caldera lithologies presented a problem due to substantial weathering and welding of the pyroclastics. It became apparent, however, that vents began to propagate around the ring fracture system that became accessible due to subsidence. The propagation proceeded counterclockwise along the ring fracture system until it reached the northern margin where the hottest ash flows were ejected last (some up into the San Joaquin canyon). Further, the transition from a single vent to multiple vents within the ring fracture system took place after the initial ash fall was ejected and after ~100 km3 of ash flows had erupted. Another 100 or so km3 erupted as the vents propagated counterclockwise around the rim [Wildreth and Mahood, 1986].

Eruptions resumed shortly after the eruption of the Bishop Tuff, but were generally confined within the caldera walls [Bailey, Dalrymple and Lanphere, 1976]. Crystal-poor domes, flows and tuffs accumulated to a depth of 500 meters in places, some of which exhibit low amplitude cross-bedding and sorting suggestive of deposition within a caldera lake. Jet black obsidian was a typical component of these early rhyolites. The relative lack of phenocrysts indicates that the lavas erupted at near liquidus temperatures. Three mineralogic facies were mapped by Bailey [1974] and exhibit increases in phenocryst percentages as the rhyolites decrease in age. Silica content was typically 75% for all the lavas and they are subsequently presumed to have undergone convective mixing after the depressurization of the Bishop Tuff magma [Bailey, Dalrymple and Lanphere, 1976]. The texture and the relative enrichment of the magma with the elements Ca, Ba, Zn, Mn, Sr, and P has caused Bailey [1983] to suspect contamination by deeper and more mafic magma.

The early rhyolites represent a time span of 0.1 million years (0.73 to 0.63 Ma). During this time, there also existed resurgent doming in the middle of the caldera. Outwardly tilted lake terraces and the deposition of beach cobbles 80 meters higher on the dome than the caldera walls, indicate that the resurgent dome spent its early life as an island [Bailey, Dalrymple and Lanphere, 1976]. Decrease in the amount of tilting of successively younger flows indicates that resurgent doming was waning during the emplacement of the youngest early rhyolites.

Around 0.51 Ma, coarsely porphyritic hornblende-biotite rhyolite began erupting from a region between the dome and the caldera walls. These moat rhyolites contain as much as 20 % phenocrysts of hornblende, biotite, quartz, sanidine, and plagioclase. A second phase erupted ~0.3 Ma, and a third erupted 0.1 Ma. They were emplaced in a clockwise fashion around the resurgent dome and the 0.2 My periodicity may represent the time required for sufficient pressure to build in order for the magma to penetrate the resurgent dome’s ring fracture system [Bailey, Dalrymple and Lanphere, 1976]. This second period represents relative cooling and crystallization within the magma chamber according to Bailey, [1983]. A third phase occurred at about 0.3 Ma and resulted in thermal and chemical rejuvenation of the melt by basaltic intrusions [Bailey, 1983]. This third phase is coincident with activity in the Mono-Inyo system. Subsequent activity involving the caldera, including modern activity, is also related to magmatic injection along the Mono-Inyo Craters system which is connected to the chamber underlying Long Valley [Hill and Bailey, 1985]. Lava from the Mono-Inyo system is less evolved than the Long Valley magma [Bailey, 1983].

The Long Valley magma chamber drew much interest during the period surrounding the earthquakes of 1980, and concern has continued since. The resurgent dome grew 0.5 meter during the period 1975-1983.  This has been interpreted as reinflation of the magma chamber [Goldstein and Stein, 1988]. Prior to 1980, all earthquake activity was focused at depths greater than 7 km. Immediately afterward, there were numerous seismic events shallower than 5 km which proceeded to migrate northward into the caldera [Ryall and Ryall, 1983]. Additionally, there were earthquake swarm events of extended duration which accompanied an increase of seismic activity after the four 6+ magnitude quakes of May 25, 1980 (see figure 6). These spasmodic tremor events are potentially suggestive of magma injection [Ryall and Ryall, 1983] and they accompanied the shallowing and northward migration of the earthquake foci. New fumarole activity also developed northeast of the swarm area within the caldera. All these factors indicated the growing (or shallowing) activity of the magma chamber. Today, magmatic gases have pervaded the soils and are disturbing biologic activity in the root zones, killing lodgepole pine and red fir trees in the vicinity of Mammoth Lakes [Reich, 1994].

Fig. 5. from Ryall and Ryall, 1983

The interest in determining the size, shape, and character of the Long Valley magma chamber produced a resurgence in scientific endeavor after the four 6+ magnitude earthquakes struck the southern edge of the caldera on May 25, 1980. There had been a general agreement upon the volume of melt and depth of the magmatic body, but no consensus had developed in regards to its size and shape. A model developed mainly from seismic results [Hill and Bailey, 1985] indicates that the modern magma chamber could have a volume between 500 and 1000 km3. If a large body of melt exists at the depths indicated by seismic and deformation information, then an obvious thermal anomaly should be detected...but is not [ Lachenbruch, Sass, Munroe, and Moses, 1976].

A three-dimensional gravity anomaly was detected by Carle [1988] who used integrated detailed geologic structure and a high degree of accuracy to produce the model seen in figures 7a, 7b, and 8. This model is little different than the seismic refraction interpretations of Hill and Bailey [1985], in depicting a large and anomalously low density mass beneath the caldera, but Carle’s model shows more definition in the basement and fill lithologies. An anomalous low-density body was also detected by Carle to exist beneath the Devils Postpile. Lachenbruch et al., [1976] detected a high heat flow over Devils Postpile, as well, suggesting that a zone of partial melt may be underfoot.

The seismic interpretations of Hill and Bailey [1985] indicate that the roof of the magma chamber lies roughly 6-7 km below the northwest margin of the resurgent dome. Attenuations of P and S waves from nearly 300 local earthquakes by Sanders and Ryall [1983] support the roof reflection at this depth as well as a fingerlike protrusion extending to within 4.5 km of the surface below the spasmodic tremor confluence of figure 6. Ryall and Ryall [1984] suggest smaller magmatic bodies exist along the Hilton Creek Fault near the south rim at depths of 5-6 km. The search for magma through the interpretation of magnetotelluric data by Park and Torres-Verdin [1988] precluded the existence of a large (~400 km3), conductive magma body at a depth between 6-13 km. Their data does not eliminate the existence of smaller (<40 km3) or larger, more resistive bodies (like a wet, granitic magma:5 ohm m, or a partial melt) because their technique cannot detect in this range. High quality borehole and surface seismograms were analyzed by Hauksson [1988] and showed a constant compressional-to-shear wave velocity with depth in the suspected location of the shallowest fingerlike protrusion of magma. This discourages the existence of a magma chamber, but the data cannot resolve a small melt zone with a diameter smaller than 3 km.

The model that is emerging from the onslaught of information is the existence of a midcrustal zone below the western part of the caldera that consists of separate, small batches of melt [Goldstein and Stein, 1988]. The following color figure illustrates this synthesis of opinions in developing a realistic model of the Long Valley magma chamber.

The geothermal sytem of the Long Valley caldera is already being harnessed by the Dept. of Energy. It must have been in existence, at its current level, for at least 30,000 to 40,000 years to produce the the saline deposits at Searles Lake which were derived from it [Hill and Bailey, 1985]. Bailey, Dalrymple, and Lanphere [1976] believe that the geothermal system was much more active in the past, perhaps on the level of some of Yellowstone’s geyser basins. They reason that the decline in surficial activity is due primarily to the "self-sealing" processes that restrict the permeability of the sediments. Silicification, argillization, and zeolitization performed by the circulating fluids are effectively clogging the pore spaces and reducing the permeability as well as the areal extent of the system.

Most of the surface activity (hot springs and fumaroles) are along direct extensions of the Hilton Creek Fault [Bailey, Dalrymple and Lanphere, 1976] (see figures 9). This relationship may be a direct result of the high activity of this normal fault as a primary player in the Sierran frontal scarp system. The modern activity, in effect, provides a weak link in the self-sealing processes that occur, allowing fluids to permeate. Besides being dictated by north-northwest trending faults like the Hilton Creek, the hot springs also tend to be somewhat controlled by the arcuate ring fracture system of the resurgent dome [Bailey, Dalrymple and Lanphere, 1976].

MONO-INYO CRATERS

The magmatic model by Bailey [1982], (refer to figure 10), illustrates a popular belief that the magmatic systems of Long Valley, Mono Lake, and the Mono-Inyo Craters utilize the Sierran frontal fault system to remain connected to each other and therefore exhibit some influence over one another. Periodic invasions of magma, for example, from the Mono-Inyo system into Long Valley caldera, during the period 0.2-0.05 Ma, are more mafic and less evolved than the Long Valley lavas [Bailey, 1983]. The three systems are connected by the Hartly Springs and Mono Lake faults which also facilitate a more-or-less linear arrangement to the volcanics from Mammoth Mountain to Black Point (refer to figure 3).

Mammoth Mountain, which rests atop the western rim of Long Valley caldera, is a complex of over 20 overlapping domes and silicic lava flows [Harris, 1988]. These quartz latite and rhyolite domes and flows were emplaced during five eruptive sequences as defined by Koeppen [1983]. Initially the flows were of sparsely phyric biotite and trachyandesite which centrally erupted ~0.2 Ma. The middle three sequences formed the bulk of the mountain by extruding coarsely phyric biotite, hornblende, or pyroxene rich quartz latites. The fifth phase extruded pumiceous coulees of coarsely phyric biotite and hornblende rhyolite from the summit. Quartz latites and basalts erupted around the perifery of Mammoth Mountain during these sequences. Geochemical studies by Koeppen suggest that the lavas came from a progressively evolving magma chamber that is unrelated to Long Valley magma. The Mono Craters consist of a 12 km long arc of 30 silicic domes and explosion craters between the Long Valley caldera and Mono Lake. The arc represents the eastern side of an 18 km diameter proposed ring fracture system [Wood, 1990] or the border of a mylonitized Cretaceous pluton [Kistler, 1967]. The ring fracture system has subsided some 200 meters since the eruption of the Bishop Tuff [Archauer, Greene, Evans, and Iyer, 1986]. This implies that a magma system below the Mono Craters has developed sufficient size to allow for downsagging to occur [Wood, 1990]. Seismic data does imply the existence of a low-velocity feature below the southern part of the craters. A magma chamber between 200 and 600 km3 positioned between 10-20 km depth would explain the seismic data gathered [Archauer, Greene, Evans, and Iyer, 1986]. Archauer et al. proposed an explanation that the low-velocity feature is the presence of a partial melt (20 %) in a magma chamber below the craters and partly controlled by the Sierran frontal fault system or a mylonitic border zone of a Cretaceous pluton exposed in the June Lakes area.

The Mono Craters volcanics are fed by a ring dike from a deep magma source [Bailey, 1984]. Except for one old dome of rhyodacite on the northwest side of the craters, all the activity has occurred within the last 40,000 years with the majority being younger than 10,000 years old [Wood, 1990], and the youngest eruptions occurred along the northernmost 5 km of the chain as recently as 1400 AD [Sieh, 1983]. The chemistries of the Mono Craters (excluding the old dome) are surprisingly homogeneously rhyolitic, despite significant variations of phenocrysts. This suggests that they are derived from a magmatic source that is large enough to support convective homogenization [Bailey, 1984]. This is not supported by seismic data, however, unless the chamber is in excess of 20 km deep.

Venting along an 11 km long Sierran frontal fault connecting the Long Valley caldera and the Mono Craters produced the Inyo Crater chain. Activity spanned the period from 4000 BC to 1350 AD, but activity was concentrated about 600 years ago [Wood, 1990]. Drillings discovered that the chain was fed by a 6 meter wide dike to the northern domes while the southern craters received magma from what is now a 20 km wide brecciated zone. The proposed interpretation [Miller?, 1985], is that the southern dike rose into wet fill in or near the caldera, created several phreatic explosion craters, and quenching itself. To the north, the drier environment allowed the dike to create domes and flows. Miller [1983] has claimed that 0.8 km3 of rhyolitic magma erupted along the chain in three episodes, the most recent being the intrusion by the dike.

A coarsely porphyritic rhyodacite apparently mixed with a finely porphyritic rhyolite in some of the flows with a north-to-south decrease in the abundance of rhyolite. Furthermore, the Mono Craters lavas are less enriched in large rare earth elements and the two chains do not seem to be related by crystal fractionation. This suggests that the two lavas from the Inyo chain and the one lava type from the Mono chain lava are derived from seperate batches of magma [Sampson, 1983].

The volcanics of the Mono Lakes region are, by contrast to Long Valley and the Mono chain, at the earliest stage of dike intrusion of the crust to produce volcanoes. The volcanic rocks are chemically distinct from all the other volcanics in the vicinity [Bailey, 1984], suggesting another separate magma source which may involve a relationship akin to the one illustrated by figure 10. The Mono Lake volcanics include a palagonite cone on the north shore of Mono Lake at Black Point. This feature is ~ 13.5 thousand years old and is a flat-topped pile of horizontally bedded ash. This formed as the ash was erupted underwater and allowed to settle slowly. The youngest volcanics are the dacitic lavas and ashes on Paohoa Island’s northern shore and are only ~ 200 years old [Wood, 1990].

MAGMA CHAMBER EVOLUTION

The magma systems of Long Valley, Mammoth Mountain, the Mono chain, the Inyo chain, and Mono Lake all are distinctive enough chemically from one another to be considered products of separate magma bodies. They may all be connected by a central root, but they seem to have been sufficiently isolated from one another to create individual chemistries. The magma systems also represent various stages of the development of a magma chamber.

The Coso Range, for example, has experienced a higher volume of phenocryst-free silica-rich lavas in the more recent periods of its activity. This suggests that it is experiencing a recharge of high-temperature (mafic) magmas which are keeping more mineral assemblages nearer their liquidus temperatures. This may indicate that Coso is in the process of redeveloping a magma chamber.

Long Valley has reached the caldera stage and has had a long history of a large mid-crustal magma chamber. Recent resurgence and seismic activity may suggest that it, too, may be experiencing recharge from hotter magmas. Its potential interconnectedness with the Mono-Inyo and Mono Lake systems via the Sierran frontal fault system only increases the possibility of activity in al of the active systems.

The arcuate Mono chain may be developing a ring fracture system which implies that enough mid-crustal magma has existed for downsagging to occur aboveit. Most of the volcanism is quite young here like it is at Mono Lake and the Inyo Craters. Mono Lake may be in the initial stages of dike intrusions where only isolated volcanoes occur.

If Mono Lake were to continue receiving hot magmas from below, the increased temperature would cause silicic materials to melt, if they were allowed enough time to equilibrate to their new environment. A magma chamber may then develop, and repeated recharges of mafic magmas would foster its growth. Long periods of quiescence would allow partial melting and magmatic differentiation the opportunity to increase the silica content and the danger of the magma chamber.

PARTING SHOT

The western boundary of Basin and Range lithospheric extension is marked by the steep Sierran frontal fault system. These normal faults have not only facilitated the rise of the Sierra batholith, but also the rise of magma from the deep crust. The volcanism permitted by the interaction of tectonic forces has produced frequent, recent, and catastrophic activity. How do you feel about encountering such an intimate glimpse of nature’s powerful lassitude?