Long Valley Caldera
Field Trip
June 14-16, 2014
Mammoth Pacific binary generating plants at Casa Diablo in Long Valley Caldera.
View is from the top of a peak within the caldera’s resurgent dome looking southwest
toward the southern topographic margin of the caldera. The caldera moat in the
background is the focus of recent seismicity during the last two decades of caldera
unrest. Paleozoic metamorphic rocks in the Sierra Nevada range to the south are the
source area for a landslide block that slid into the caldera on a gassy cushion of ash late
in the caldera’s collapse. Photo by R. Sullivan
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LONG VALLEY CALDERA
GEOTHERMAL AND MAGMATIC SYSTEMS
Gene A. Suemnicht
EGS Inc., Santa Rosa, CA 95404
INTRODUCTION
Long Valley Caldera in eastern California has been explored for geothermal resources
since the 1960s. Early shallow exploration wells (<300m) were located around Casa
Diablo near the most prominent hot springs and fumaroles on the southwest flank of the
Resurgent Dome (Figure 1). Later deep (±2000m) wells explored the southeastern
caldera moat and evaluated lease offerings in and around the caldera’s Resurgent Dome.
Data from these wells revealed that the principal geothermal reservoir in Long Valley is
not located directly beneath the Casa Diablo Hot Springs and is not currently related to
the Resurgent Dome. Instead, the hydrothermal system appeared to be more complex
with shallow production at Casa Diablo supplied by upflow and outflow from a more
extensive deeper geothermal source beneath the western caldera moat.
Geothermal development at Casa Diablo occurred in stages under various operating
companies. Ormat Nevada Inc., currently owns the project, operating it under the
Mammoth Pacific LP and generating 40 MWe (gross) from three power plants utilizing 6
production wells and 5 injection wells. Until 2005, production of ~630 L/sec (~10,000
gal/min) of moderate temperature (170ºC) fluids for the geothermal plants was limited to
~0.7 km2 (~165 ac) around Casa Diablo from shallow wells (<200m) completed in
permeable Early Rhyolite eruptive units on the southwestern edge of the Resurgent
Dome. MPLP drilled deeper (~450m) production wells in 2005 in the southwest caldera
moat around Shady Rest in a development area that Ormat refers to as Basalt Canyon.
The deeper west moat wells produce 185°C fluids from deep Early Rhyolite units and the
upper part of the Bishop Tuff. In 2006, an average of 126 L/sec (2000 gal/min) of higher
temperature brine from the Basalt Canyon wells began to be piped 2.9 km to sustain
production at the Casa Diablo plants allowing several older shallow wells to be shut in.
Ormat plans to expand production from Basalt Canyon, install new generation facilities
and increase generation to approximately 60 MWe (gross).
GEOLOGIC SETTING
Long Valley Caldera is a 17 X 32 km (10 X 20 mi) depression created by the eruption of
an estimated 600 km3 of Bishop Tuff 760,000 years ago (Bailey and others, 1976). The
caldera is the largest feature in the Mono-Long Valley volcanic field that includes
Pleistocene-Recent eruptive centers of Mammoth Mountain and the Mono–Inyo volcanic
chain. Volcanism associated with Long Valley began ~ 4 Ma ago with widespread
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eruptions of intermediate and basaltic lavas accompanying the onset of large-scale
transtensional faulting that formed the eastern front of the Sierra Nevada and the Owens
Valley graben (Figure 1). Discontinuous erosional remnants of precaldera extrusive
rocks are scattered over a 4000 km2 area around the caldera suggesting an extensive
mantle source region (Bailey and others, 1989). Rhyolitic eruptions began ~ 2 Ma ago
from multiple vents of the Glass Mountain eruptive complex along the northeast rim of
the present-day caldera (Metz and Mahood, 1985) (Figure 1). These more silicic
eruptions suggest the initiation of magma accumulation and differentiation in the shallow
crust (Bailey and others, 1989; Hildreth, 2004). The caldera-forming eruption partially
evacuated the underlying magma chamber and the floor of the caldera subsided up to 2
km in segments separated along pre-existing inherited basement faults (Suemnicht and
Varga, 1988). Approximately 350-400 km3 of the Bishop Tuff filled the caldera
depression. Post-collapse eruptions have continued to fill the caldera over the last
600,000 years (Bailey and others. 1976; Bailey 2004; Hildreth 2004). A series of rhyolite
flows and tuffs (Early Rhyolite) mark the onset of resurgence in the west-central part of
the caldera approximately 600,000 years ago (Figure 1). Coarsely porphyritic Moat
Rhyolites erupted later around the Resurgent Dome beginning in the north approximately
500,000 years ago progressing to the south around 300,000 years ago and approximately
100,000 years ago in the west (Bailey and others, 1976; 1989). The western Moat
Rhyolites erupted during a period of more voluminous basaltic and andesitic flows that
began approximately 200,000 years ago in the western caldera extending beyond the
caldera margins to the south and west. These more mafic eruptions include basalts in the
southwestern caldera moat (Casa Diablo flow of Bailey and others, 1976), eruptive events
such as the 80,000 year-old Devil’s Postpile basaltic andesite and more recently, the
8,000 year-old Red Cones vents south of the caldera. A series of rhyodacitic eruptions
also occurred in the western caldera moat 110,000 – 50,000 years ago (Mahood and
others, 2010); the most prominent of these is Mammoth Mountain on the southwestern
topographic rim of the caldera (Bailey 2004). Hildreth (2004) suggested that the mixed
mafic-rhyodacitic volcanism represented a separate magmatic system outside the
caldera’s ring-fracture system.
The most recent eruptions in the area occurred along the Mono-Inyo volcanic chain
extending from the western caldera moat northward to Mono Lake (Figure 2). Eruptions
along the chain began approximately 40,000 years ago and have continued to historic
times. Bursik & Sieh (1989) identified 20 small eruptions (erupted volumes <0.1 km3)
within the chain over the past 5000 years occurring at intervals ranging from 250 to 700
years. The most recent dome-forming eruptive events began at the north end of the Mono
Craters about 600 years ago and culminated with phreatic eruptions in the south at the
Inyo Craters and the north flank of Mammoth Mountain (Bursik & Sieh 1989; Mahood
and others, 2010). The magma source for these eruptions is an 8–10-km-long dike that
trends north out of the caldera. In 1987, the DOE funded several core holes that
penetrated the dike at Obsidian Dome north of the caldera and at the Inyo Craters in the
caldera’s west moat. Petrologic studies of the core indicate that the eruptions may have
tapped magmas of three different compositions venting at several places along strike
(Eichelberger, 2003). A shallow intrusion beneath Mono Lake approximately 200 years
ago lifted lake-bottom sediments to form Pahoa Island erupting a small volume of
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andesitic lava from vents on the island’s north side (Bursik & Sieh, 1989). The
progression of eruptions over the past 2 Ma from Glass Mountain on the eastern caldera
margin to Mammoth Mountain on the west and the Mono–Inyo volcanic chain to the
north suggests that the magmatic system that erupted to form Long Valley Caldera has
declined with time and has been supplanted by complex mixtures of magma erupting
contemporaneously from the active Mammoth Mountain–Inyo Domes magmatic system
at times triggered by the intrusion of more mafic magma into the shallow crust (Hildreth,
2004; Mahood and others, 2010).
Figure 2. Locations ages and types of eruptive events in the Mono-Inyo volcanic zone
(from USGS)
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Tectonics
Long Valley Caldera lies within an east-west embayment or offset in the northwest
trending Sierran escarpment (Mayo, 1934). The east-dipping normal faults that form the
escarpment mark the western edge of crustal extension in the Basin and Range Province.
The caldera is located at the northern end of the Owens Valley graben, part of the Eastern
California Shear Zone, a region of transtensional deformation along the western edge of
the Basin and Range that extends to the north along the Walker Lane in western Nevada
(Figure 2) (Hill, 2006). Transtensional deformation and active magmatism within the
Eastern California Shear Zone and the Walker Lane reflect the combined influence of
dextral slip along the San Andreas Fault transform boundary between the Pacific Plate
and North American Plates and the westward progression of crustal extension across the
Basin and Range Province. Right-lateral slip distributed across this transtensional zone
accounts for 15 - 25% of the relative motion between the Pacific and the North American
plates (Dixon and others. 2000, Faulds, 2004). Active magmatic processes in the region
are generally attributed to upwelling of the underlying asthenosphere as the crust
stretches, thins and fractures in response to transtensional extension across the Eastern
California Shear Zone - Walker Lane corridor (Hill, 2006).
The tectonic interaction between the Eastern California Shear Zone and the Walker Lane
system localizes volcanism at transtensional pull-apart sections in the Mono-Long Valley
region. Long Valley Caldera is located at the western end of the Mina Deflection, a
broad zone of northeast-trending left-lateral faults that form a right jog in the regional
right-lateral fault system of the Owens Valley providing a link to the Walker Lane to the
northeast (Figure 3). Crustal extension in the Mono Basin is the result of left lateral shear
across the Mina Deflection and the extension increases southward into Long Valley
Caldera. Extensional strain is complexly resolved within the caldera because of the
faulting and fracturing that defined the pre-existing Sierran range front embayment. The
western and central parts of the caldera are superimposed over several pre-existing fault
systems that defined the pre-caldera eastern Sierra range front. The floor of the caldera,
rather than being a simple planar feature, is segmented into a number of discrete blocks
with varying offsets that accommodated 1-2 km of subsidence as the caldera’s floor
foundered. The basement structure of Long Valley, buried under nearly 400 km3 of
Bishop Tuff and later volcanic fill, has been interpreted from detailed mapping, regional
geophysical surveys and drilling data (Bailey, 1976; Suemnicht and Varga, 1988; Bailey,
2000). The Discovery Fault Zone in the western caldera is a prominent surface feature
that is a direct extension of the Mina Deflection and an expression of the original range
front embayment before the caldera collapsed. The pre-existing tectonic framework
controlled the configuration of the caldera floor and, through faults and fault
intersections, controlled the location of postcollapse eruptive centers. The inherited deep
basement structures of the western caldera provide the high fracture density and deep
permeability for the source of the present geothermal system in Long Valley.
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Figure.3 Shaded relief map of east-central California and western Nevada, showing the
location of Long Valley Caldera (LVC) and the Mono–Inyo volcanic chain (MIVC)
outlined by red ovals. Increasing elevation is indicated by colors ranging from dark
green (at approximately sea-level) to light grey (>2400 m with maximum elevations
reaching c. 4300 m). The transition from yellow-green to brown marks the 1700 m
contour. Heavy black lines indicate major Quaternary faults. The red line in the Owens
Valley indicates surface rupture from the M ≈ 7.6 Owens Valley earthquake of 1872.
White arrows indicate the sense of displacement across the Eastern California Shear
Zone (ECSZ), the Mina Deflection (MD), and the Walker Lane (WL). The white outline in
the inset shows the map location with respect to the San Andreas Fault (SAF) in coastal
California (CA) and the Basin and Range Province in western Nevada (NV) with
opposing white arrows indicating the sense of extension. (After Hill, 2005)
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Caldera Unrest
Moderate to strong historical earthquakes occur regularly in the eastern Sierra Nevada
and the Owens Valley south of Long Valley Caldera (Ellsworth, 1990). The northern end
of the rupture zone of the M ~ 7.6 Owens Valley earthquake of 1872 extended to within
60 km of Long Valley Caldera (Figure 4) and earthquakes M>5 occurred outside the
caldera before 1970 (Cramer & Toppozada 1980; Ellsworth 1990).
Long Valley Caldera began a period of unrest in the late 1970s that included earthquake
swarms, approximately 80 cm of inflation within the resurgent dome, changes in the
outflow from hot springs and fumaroles and increased CO2 emissions around the flanks
of Mammoth Mountain (Figure 1). The gas emissions on Mammoth Mountain have been
accompanied by rising 3He/4He ratios interpreted as potential indicators of magma
moving to shallower crustal levels. The largest magnitude earthquakes occurred within
the Sierran block south of the caldera while caldera activity was marked by earthquake
swarms, long-period (LP) and very-long period (VLP) volcanic earthquakes. An intense
earthquake sequence that included four M>6 earthquakes within and around the caldera
on May 25, 1980 occurred within days of the May 18, 1980 eruption of Mount St Helens
and in that context, raised strong concerns about the eruptive potential of a large active
magma chamber beneath the caldera.
Volcanic hazard concepts related to the continuing unrest within the caldera evolved
rapidly as research progressed on the Mono-Long Valley magmatic system (Hill, 2006).
Based on Long Valley data and a better understanding of restless calderas worldwide,
large silicic calderas can go through sustained periods of episodic unrest, separated by
years to decades of relative quiescence, all without producing an eruption (Newhall &
Dzurisin 1988; Newhall, 2003). Caldera unrest can also be more intense and may extend
beyond the comparatively short restless periods associated with central vent volcanoes.
Volcanic earthquakes, increased magmatic gases and changes in geothermal
manifestations have all occurred in Long Valley Caldera without an eruption.
Long Valley remains on an active volcanic hazard alert status although the US
Geological Survey states that earthquake activity within and adjacent to the caldera has
remained at a comparatively low level since 1999. The caldera is probably not underlain
by a laterally extensive, upper-crustal magma body capable of feeding a major eruption
(Eichelberger, 2003), but none of the current data exclude the possibility of smaller (<1
km3) eruptions from smaller magma sources or a series of phreatic explosions similar to
the historical eruptions that occurred along the Inyo-Mono dike (Miller, 1985). The
1978-2004 period of unrest most likely was associated with the addition of ~0.3 km3 of
magma and hydrous fluids at a depth of 6-7 km beneath the resurgent dome. Seismic
tomography studies might resolve a 1–2 km diameter magma body but not the smaller
melt volumes noted above (Hill and Prejean, 2005). The dacitic magma chamber beneath
Mammoth Mountain has probably crystallized because the last eruption occurred 52,000
years ago (Hildreth, 2004). Hill & Prejean (2005) ascribed the 1989 earthquake swarm
beneath Mammoth Mountain that included mid-crustal long-period earthquakes and
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Figure 4. Shaded relief map illustrating the relation of Long Valley Caldera (LVC) to
major tectonic elements and M>5 earthquakes in eastern California (the red rectangle in
inset shows the map area). Heavy black and red lines are faults with Holocene slip and
historic (post-1870) slip, respectively. The surface trace of the M=7.8 Owens Valley
earthquake of 1872, labeled ‘1872’. Orange circles are post-1977 earthquakes (small for
M>5; large with decadal year included for M>6). Yellow circles are 1870–1978 M>6
earthquakes (not including aftershocks to the 1872 main shock with decadal year
indicated. WC is Wheeler Crest (from Hill, 2005).
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increased CO2 venting to a “mid-crustal plexus of basaltic magma (that) remains capable
of feeding future mafic eruptions.” This magma plexus presumably fed eruptions of the
mafic field surrounding Mammoth Mountain, including the 8,000 year-old Red Cones
vents, and it is the likely heat source for the 700 year-old phreatic explosion vents on the
northeast flank of Mammoth Mountain.” Some long period earthquakes have occurred
west of the Mono Domes (Pitt & Hill 1994) but the area has remained comparatively
quiet during the unrest in Long Valley. Based on a geologic history of 20 eruptions over
the last 5000 years and the eruption at Pahoa Island approximately 200 years ago (Bailey
2004), the young silicic domes of the Mono–Inyo volcanic chain still have the potential
to produce significant eruptive events.
GEOTHERMAL SYSTEM
The geothermal system in Long Valley has varied through time (Bailey and others, 1976,
Sorey and others, 1978; 1991). Different mineral assemblages in and around the
Resurgent Dome and differing age dates indicate that mineralization occurred in two
separate phases (Sorey and others, 1991). The caldera supported an intense hydrothermal
system from 300,000 to 130,000 years ago producing widespread hydrothermal alteration
in and around the Resurgent Dome. The current hydrothermal system has probably been
active for only the last 40,000 years, but prominent surface manifestations occur in many
of the older system’s established outflow zones at comparatively low elevations in the
south central portion of
the caldera around the
Resurgent Dome
(Figure 5). As much as
75% of the current
hydrothermal outflow
occurs in the thermal
springs at Hot Creek on
the southeastern edge
of the Resurgent Dome
and geochemical
estimates of source
reservoir temperatures
range from 200 oC –
280oC (Sorey and
others, 1978; 1991;
Mariner and Wiley,
1976).
Figure. 5. Generalized distribution of surface
manifestations in Long Valley Caldera
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Hydrothermal manifestations are notably absent in the western caldera moat (Bailey and
others, 1976); however, detailed mapping (Suemnicht and Varga, 1988) and remote
sensing studies of the western caldera (Martini, 2002) identify many high-temperature
alteration mineral assemblages <100,000 years old that are related to vigorous
hydrothermal outflow along penetrative faults in the western caldera. The alteration
mineralogy shows that significant surface manifestations occurred at higher elevations in
the western caldera in the early phases of the current hydrothermal system and the current
pattern of outflow to the southeast toward Casa Diablo may have resulted from active
fracturing and faulting opening older hydrothermal flow zones allowing outflow along
permeable zones at lower elevations.
GEOTHERMAL EXPLORATION
Early exploration drilling in Long Valley focused on the southern and central part of the
caldera. Production wells drilled in the 60s evaluated shallow production around Casa
Diablo’s fumaroles and hot springs. The first deep well (66-29) was drilled in 1976 to
evaluate the resource potential in the southeast moat (Figure 6). Numerous shallow
gradient holes evaluated the heat flow associated with the Resurgent Dome in the 1970s
and geothermal lease sale opportunities in the late 70s and early 80s prompted shallow
and intermediate drilling to assess lease blocks within the Resurgent Dome.
Clay Pit-1 and Mammoth-1 drilled in 1979 were the first deep wells drilled in the
Resurgent Dome of Long Valley and the first deep tests to penetrate the entire section of
the caldera fill (Figure 6). Mammoth-1 drilled through 390m of Early Rhyolite, 863m of
Bishop Tuff and 230m of metasedimentary rocks that correlate with the Mt. Morrison
roof pendant to the south, bottoming at 1605m. Mammoth-1 was also the first well
within the caldera to encounter a block of chaotically mixed metapelite and granite at
466m in the upper section of Bishop Tuff that was interpreted as a landslide block based
on cuttings alone (Suemnicht, 1987). Drilling to evaluate federal lease offerings during
the 1980s and scientific drilling to evaluate various eruptive processes expanded the
understanding of the western part of the caldera. Unocal’s deep well IDFU 44-16
penetrated the caldera fill, Tertiary volcanic rocks and metamorphic rocks to a depth of
2000m near the Inyo Craters (Figure 6). The well encountered temperatures of 218oC at
1100m, the highest yet measured in Long Valley, but proved unproductive because of a
limited thickness of reservoir rocks and the incursion of cold water beneath the
production zone (Suemnicht, 1987).
Later scientific drilling by Sandia National Labs west of 44-16 (Figure 6) established that
1000 m of vertical offset on the caldera’s western ring fracture system occurs within a
kilometer distance between the two wells along the western structural margin of the
caldera (Eichelberger and others, 1988). Additional scientific drilling in Long Valley
included a core hole at Shady Rest (Figure 6) (Wollenberg and others, 1989), and an
ultra-deep (3 km) Long Valley Exploratory Well (LVEW, Figure 6) intended to test the
presence of magma near the center of the Resurgent Dome (Finger and others, 1995).
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Figure 6. Caldera outline and locations of temperature gradient holes (T designation),
intermediate and deep wells drilled in Long Valley. Wells in orange drilled into
intracaldera Bishop Tuff and wells in red have penetrated the intracaldera fill and are
completed in the Paleozoic metamorphic rocks of the caldera basement. (Suemnicht and
others, 2006)
The results of deep drilling on the Resurgent Dome indicate that present-day thermal
conditions are controlled in places by vertical flow of relatively cold water in steeply
dipping faults that formerly provided channels for high-temperature fluid upflow (Sorey
and others, 2000). In the central caldera, current temperatures and gradients are relatively
low and show little evidence of possible magmatic temperatures at depths of 5-7 km
postulated on the basis of recent deformation, seismic interpretations and shearwave attenuation of teleseismic waves (Rundle and others, 1986; Rundle and Hill, 1988).
Drilling results establish that magmatic activity beneath the central part of the caldera has
waned over the past ~100,000 years while similar activity in the western caldera has
increased.
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RECENT DRILLING
Mammoth Pacific LP has continued to assess the productive potential of the caldera’s
western moat. Shallow test wells confirmed the general temperature distribution within
the shallow hydrothermal system and expanded the understanding of the intra-caldera
geology. A similar section of metapelite/granite was encountered at the top of the Bishop
Tuff in both the 38-32 corehole drilled to 353m less than a kilometer south of Casa
Diablo (Bailey, pers. com.) in 1992 and the BC 12-31 corehole drilled to 600m
approximately 2 km west of the production facilities in 2002 (Figure 7).
Figure 7. Distribution of the landslide block (blue) of Paleozoic metasedimentary rocks
from the southern rim of Long Valley Caldera (Suemnicht and others, 2006)
The intracaldera landside block is relatively coherent and covers approximately 3 km2 of
the southern moat between Mammoth-1 on the north, and BC 12-31 on the west. The
landslide block is absent in surrounding wells that penetrate into the Bishop Tuff on the
Resurgent Dome and at Shady Rest, approximately 1 km northwest of BC 12-31. The
landslide lies at or near the top of the Bishop Tuff. Lithologic interpretations of cuttings
from Mammoth-1 indicate that the well penetrated an upper ash-rich section of highly
altered Bishop Tuff before passing through a 43m thick section of landslide breccia and
then back into highly altered ash-rich Bishop Tuff. Core from 38-32 and BC 12-31
indicate the landslide block is a highly indurated relatively undisturbed meta-breccia or
conglomerate with a dense grey matrix of rock fragments and larger rounded or
subrounded clasts of granite, hornfels, metapelite and metaquartzite typical of lithologies
that crop out in the southern topographic wall of the caldera (Figure 8).
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Figure 8. Structural cross-section of the southern caldera margin showing the landslide
block encountered in exploration corehole 38-32 and the Mammoth-1 deep test well at
Casa Diablo (Suemnicht and others, 2006).
The landslide block does not contain pumice or phenocrysts from the Bishop Tuff and the
southern wall lithologies and indurated metamorphic matrix preclude the possibility that
the block might be glacial till. Based on the excellent core samples from 38-32 and BC
12-31, the Bishop Tuff below the landslide block is lithic rich with an intensely altered
glassy matrix permeated with gas void spaces. The landslide block lies entirely at the top
of the Bishop Tuff in both 38-32 and BC 12-31 and the underlying remarkably
undisturbed soft ash probably represents the upper ash cloud at the top of the Bishop Tuff
that acted as a cushion beneath the landslide block (Figure 7). Intense clay alteration has
obscured many of the primary features of the ash but much of the porous tuff matrix
appears to be a gassy froth that retains doubly terminated quartz crystals diagnostic of the
Bishop Tuff.
SHALLOW HYDROTHERMAL SYSTEM
Landslide blocks are common features that have been mapped in several eroded collapsed
calderas (Lipman, 2003). In Long Valley, a block of metasedimentary rock from the
southern caldera rim slid into the southern Long Valley moat on a gassy cushion of
Bishop ash late in the eruption and collapse sequence. The temperature distribution
within the caldera indicates that the present-day outflow of the deeper hydrothermal
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system occurs along penetrative NW-SE faults related to the Resurgent Dome and E-W
ring fracture faults that control the southern structural margin of the caldera. Upflow at
these fracture intersections occurs at comparatively low elevations at the base of the
Rhyolite Plateau west of Shady Rest. Within the corresponding shallow hydrothermal
system, the landslide block controls the flow of hydrothermal fluids southeast around the
southern part of the Resurgent Dome (Figure 9). Intense alteration of the over/underlying
ash froth, overlying clay-rich lacustrine sediments and the impermeable nature of the
landslide block isolates the shallow hydrothermal system from deeper flow and maintains
moderate-temperature outflow in shallower Early Rhyolite or lacustrine units east of
Shady Rest.
Figure. 9. Lithology and temperatures (oC) from wells in the southern moat of Long Valley
caldera. Wells extend from the projected outflow zone of he current geothermal system at Shady
Rest to existing production area at Casa Diablo represented by the deep Mammoth-1 well
(Suemnicht and others, 2006) .
Early proposals used the available temperature data to suggest that hotter outflow was
separated from colder recharge water because of density (Blackwell, 1985) and while
density separation might prevail for a time, it would be a transient condition. Shallow
hydrothermal circulation could only be sustained where hot water is physically separated
from the underlying permeable Bishop Tuff. All of the drilling results within the caldera
show that the strong head of cold recharging waters from the caldera rim has a significant
effect on the periphery of the hydrothermal system. Sharp temperature reversals of
nearly 100 oC are commonly found on the structural caldera margins where high
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temperature upflow is affected by cold recharge penetrating into the deeper fractured
Bishop Tuff (Suemnicht, 1987).
The physical separation between the Bishop Tuff and the overlying impermeable
landslide block in the southern part of the caldera allows the shallow moderate
temperature hydrothermal system to survive. Once thermal outflow has made it into the
shallow permeable Early Rhyolites east of Shady Rest, it is effectively separated from
underlying rocks and the pervasive influence of cold recharge waters around the rim of
the caldera (Figure 10). Production wells at Casa Diablo are completed in an upper
170oC zone in Early Rhyolites while injection wells are completed in the deeper (~700m)
permeable Bishop Tuff below the landslide block. The impermeable landslide block,
lacustrine clays or intensely altered clay-rich upper part of the Bishop Tuff control
shallow hydrothermal circulation in the southern caldera allowing sustained shallow
production at Casa Diablo by isolating warm shallow outflow from deep cold recharge
that might quench the system.
ACKNOWLEDGEMENTS
Ormat Nevada Inc. graciously allowed publication of their proprietary information and
John Bernardy of MPLP arranged our tour of the Casa Diablo production facilities.
Thanks to Dave Hill, Bill Evans and Jim Howle of the USGS for providing updates on
caldera seismicity and unrest and recent data . This field guide is dedicated to Roy Bailey
of the USGS whose pioneering work in Long Valley provided vital insight into the
workings of caldera eruptive systems worldwide.
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REFERENCES
Bailey, R.A., Dalrymple, G.B. and Lanphere, M.A., 1976, Volcanism, structure and
geochronology of the Long Valley caldera, Mono County, California. Journal of
Geophysical Res., Vol. 81, No. 5, pp. 725-744.
Bailey, R.A., Dalrymple, G.B. and Lanphere, M.A., 1989, Quaternary Volcanism, of the
Long Valley Caldera and Mono-Inyo Craters Eastern, Field Trip Guidebook. 28th
International Geological Congress, 36 pp.
Bailey, R.A., 1989, Geologic map of Long Valley caldera, Mono-Inyo Craters Volcanic
Chain, and Vicinity, Eastern California. US Geol. Survey Map I-1933.
Bailey, R.A., 2004, Eruptive History and Chemical Evolution of the Precaldera and
Postcaldera Basalt-Dacite Sequences, Long Valley, California: Implications for
Magma Sources, Current Seismic Unrest and Future Volcanism. US Geol.
Survey Prof. Paper 1692 75pp
Benioff, H. and Gutenberg, B., 1939, the Mammoth “Earthquake Fault” and Related
Features in Mono County, California. Seismological Soc. of America, Bull.
29/333-340.
Blackwell, D.D., A transient model of the geothermal system of Long Valley caldera,
California. Journal of Geophysical Res., Vol. 90, No. B13, pp. 11,229-11,242.
Brophy, P., 2008, Geothermal reservoir evaluation. GRC Annual Meeting workshop,
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Long Valley Caldera (USA) and El Chichón (México). Geofísica Internacional 48
(1), 171-183 (2009)
Bursik, M., 2003, Overview: tectonic-magmatic interactions at Mono-Inyo Craters. in:
Understanding a Large Silicic Volcanic System: An Interdisciplinary Workshop
on Volcanic Processes in Long Valley Caldera-Mono Craters.
Bursik, M. and K. Sieh,1989. Range Front Faulting and Volcanism in the Mono Basin,
Eastern California, J. Geophys. Res., 94, 15,587-15,609.
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Long Valley Caldera
Field Trip Log
June 14-16, 2014
Prepared by:
Gene Suemnicht
Bastien Poux
Waterfall Towers Suite A-112
2455 Bennett Valley Road
Santa Rosa, CA 95404
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DAY 1: Casa Diablo and Long Valley Caldera
The trip leaves from University of Nevada campus in Reno, Nevada. We drive south on
Highway 395 passing the Steamboat Springs geothermal field and hot spring areas
around Bridgeport before driving through the Mono Lake basin and entering Long Valley
Caldera at Deadman Summit. Continue south on Highway 395 to the Mammoth
Lakes/Devils Postpile exit at Highway 203. At the bottom of the exit ramp, turn right
(west) toward Mammoth Lakes. Continue west and uphill through the Town of
Mammoth Lakes. Turn right (north) on Minaret Summit Rd. Drive to the Main Lodge at
Mammoth Mountain and park at the base of the gondola facilities. Ride the gondola to
the top of the mountain and walk out of the gondola building uphill to the northeast to the
Inyo National Forest summit sign: Mammoth Mountain. Elevation 11, 060 feet.
Cumulative distance (to the nearest mi/km) – 170 mi. / 274 km
Stop 1 – Caldera Overview, Mammoth Mountain Summit
The view from the summit of Mammoth Mountain encompasses Long Valley Caldera,
the northern Owens Valley and the Sierra crest. The topographic margin of the caldera
extends from the Palisades below San Joaquin Ridge on the west, Bald Mountain on the
north, the Glass Mountain eruptive complex on the northeast, Lake Crowley to the east
and Laurel Mountain and the Sierra front to the south. The White Mountains, on the
eastern skyline beyond the Glass Mountain complex represent the eastern side of the
Owens Valley rift with elevations of 4373m (14347ft).
Long Valley Caldera formed when the roof of the caldera magma chamber ruptured
760,000 years ago erupting 600 km3 of Bishop Tuff. The roof of the partially emptied
magma chamber collapsed to form the 17 x 32 km (10x20 mi) depression of Long Valley
Caldera. Drilling confirms that the structural margin of the caldera and the inherited
faults that controlled the caldera collapse can be as much as 2 km (1.2 mi) inboard of the
current topographic rim of the caldera. As pressures recovered within the Bishop Tuff
magma chamber, the roof began rising and lifting the overlying caldera fill. The low hills
in the west central part of the caldera are the composite Early Rhyolite eruptive centers of
the Resurgent Dome that began erupting at ~600,000 years ago. The caldera moat is the
space between the caldera’s topographic rim and the Resurgent Dome. Smaller Moat
Rhyolite eruptions began in the north around 500,000 years ago, southeast near the
airport approximately 300,000 years ago and in the western moat around 100,000 years
forming a rhyolite plateau above the Town of Mammoth Lakes. Most eruptive events
over the last 160,000 years have occurred in the western caldera varying from mafic to
silicic compositions suggesting a complex mixture of contemporaneous magma sources
(Hildreth, 2004; Mahood and others, 2010). Mammoth Mountain on the western caldera
rim is a composite of a dozen dacite/rhyolite centers that erupted between 120,000 and
58,000 years ago. Mammoth Mountain and Lincoln Peak were erupted between 68,00058,000 years ago. The summit dome where we stand is dated at 68,000 years and
detailed age dating indicates that the main bulk of the mountain was formed by dacite
eruptions in less than 2,000 years (Mahood and others, 2010). The Mono-Inyo volcanic
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chain extends from the Inyo Domes in the northwestern caldera moat to the Mono Craters
and Mono Lake to the north. Eruptions began in the Mono Craters approximately 40,000
yrs. ago. Based on tree ring data, The Deadman, Glass Creek and Obsidian domes all
erupted in the late summer in 1350 AD (Millar and others, 2006); an exceptionally well
constrained pre-historic eruption age. The Mono-Inyo chain is the most prominent
surface evidence of current magmatic activity in Long Valley. Phreatic explosions
related to the Mono-Inyo dike intrusion excavated phreatic explosion pits on the north
face of Mammoth Mountain 659–737 yrs. ago (Sorey and others, 1998). Many of these
features have been modified by construction of the mountain’s recreational facilities.
Only a tephra hill remains of what used to be an explosion crater that formed the base of
Chair 19 (now Discovery Chair) west of the lodge. Stratified eruption deposits are found
in drainages eroded between ski runs.
Areas of dead trees began to appear where high levels of CO 2 soil gas were noted
following a six-month period of seismic swarms beneath the mountain in 1989 (Sorey
and others, 1998)(Figure 4). Elevated soil gas flux and tree-kill areas occur along faults
or fault intersections around all
but the eastern flank of
Mammoth Mountain, most
notably along the northern flank
near Chair 12 and along the
southern flank at Horseshoe
Lake just outside the
southwestern margin of the
caldera. The cold CO 2
discharge does not contain 14C
but is accompanied by elevated
levels of magmatic helium
(3He/4He ratios ~ 5 times the
atmospheric ratio [R/Ra]).
Groundwater with the same
inorganic carbon and helium
isotopic compositions flows off
the flanks of the mountain. The
source of these inorganic carbon
and helium emissions appears to
be a long-lived gas reservoir
formed above a region of
magmatic intrusion, sealed at
Figure 1. Tree kill areas around Mammoth Mountain and
the top by hydrothermally
Horseshoe Lake to the south (USGS).
altered volcanic and
metasedimentary rocks that are
periodically penetrated by
intrusions of basaltic magma
(Sorey and others, 1998).
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Ride down the gondola and return to the vehicles. Exit the ski area parking lot and head
east toward Mammoth Lakes on Minaret Summit Road. Turn left (east) at the stoplight
intersection of Minaret Summit Road and Highway 203. Continue east through the Town
of Mammoth Lakes. At the eastern side of town, turn left (north) at the Shady Rest
Campground road. Drive north past the camping area and bear right through the parking
lot. Turn right (east) onto a dirt road at a gap in the logs on the east side of the parking
lot. Continue a short distance to the Mammoth Pacific well pads on the left and park on
the western side of the pad.
Incremental Distance –6.1 mi / 9.8 km
Cumulative distance – 176.1 mi. / 283.1km
Stop 2 – Production Wells, Shady Rest
The USGS and DOE drilled core hole
RDO-8 to 732 m (2401 ft.) near the Shady
Rest campground in the southern caldera
moat in 1984 (Wollenberg and others,
1989). The lower portion of the well was
lost because of wellbore instability but
measured temperatures at 400m(1312 ft.)
were 203oC (397°F) and appeared to persist
at depth based on extrapolated survey
temperatures. This scientific effort and
exploration wells drilled in the western
Figure 2. Drilling BC 12-31 exploration
corehole in Basalt Canyon (photo by R.
caldera added to understanding the Long
Sullivan)
Valley geothermal system. Drilling
established that the principal geothermal reservoir in Long Valley was not located
directly beneath Casa Diablo and did not appear to be related to the Resurgent Dome.
Instead, the hydrothermal system appeared to be more complex and the shallow
production at Casa Diablo is supplied by upflow from a western geothermal source and
outflow at or near Shady Rest. Mammoth Pacific LP has continued to evaluate the
geothermal potential of the southern caldera moat, drilling intermediate-depth exploration
wells 66-31 and 38-32 in 1992, BC 12-31 in 2002 and production wells 57-25,66-25, 1425 and 12-25 east and north of the Shady Rest core hole. The production wells are nearly
500 m (1640 ft.) deep and produce from the upper fractured portion of the Bishop Tuff.
The current production of ~2000gpm is piped to Casa Diablo and commingled with
production from earlier wells.
Continue east on Sawmill Flat Road to Highway 203 (BEWARE OF TRAFFIC).
Continue east on Highway 203 crossing beneath the Highway 395 overpass. Continue
east toward Casa Diablo. Turn left (north) on Casa Diablo Road and enter the power
plant headquarters at the control gate at the top of the hill. Access to the geothermal
development is by permission of Mammoth Pacific LP.
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Incremental Distance –4.0 mi/ 6.4 km
Cumulative distance – 180.1 mi. / 289 km
Stop 3 – Mammoth Pacific Production Facilities, Casa Diablo
Casa Diablo Hot Springs is an active thermal area on the southwestern side of the
caldera’s Resurgent Dome (Figure 3). Most of the prominent hot springs within the
caldera including Hot Creek (Stop 5) occur in the southern caldera moat or are localized
along faults around the southern edge of
the Resurgent Dome. The springs at Casa
Diablo were an early gathering place for
local Piute Indians, settlers and
prospectors. During the late 1800s, the
site was a way station on the BodieMammoth City stage line and later an
automobile rest stop on Old Highway 395.
The old route survives as a paved roadway
dividing the east and west portions of the
geothermal field (Figure 4). Historical
records and photographs recount the
Figure 3. Generalized distribution of surface
erratic discharge of the thermal features
manifestations in Long Valley Caldera.
that varied from relatively quiet outflow to
sporadic geysering.
Active and relict fumaroles,
mudpots and hot springs at Casa
Diablo are localized along a
major northwest trending
normal fault system that forms a
graben within the Resurgent
Dome. Hydrothermal alteration
marks the trace of a fault that
cuts 600,000 year-old Early
Rhyolite of the Resurgent Dome
on the northeastern side of the
field. Younger 129,000-62,000
year-old postcaldera mafic lavas
flood the southwestern caldera
moat and lap against the
Resurgent Dome. Active
Figure 4. Wells and temperature gradient holes in and
fumaroles on the western side
around
the
Casa
Diablo
development
area
(modified
from
of
the geothermal field are aligned
Calif. Div. of Oil, Gas and Geothermal Resources).
along a fault scarp that uplifts and
exposes these moat basalts. Casa Diablo was the location of several seismic swarms that
occurred during the initial phases of caldera unrest and Resurgent Dome uplift from 1980
to 1983. Some of these events were marked by ground cracking, changes in spring
discharge, vigorous reactivation of a previously dormant fumarole immediately east of
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the old highway and new or reactivated hot springs around the edge of the Resurgent
Dome.
Magma Power Co drilled the first geothermal exploration wells in Long Valley at Casa
Diablo between 1959 and 1962. A series of nine wells were drilled to total depths
ranging from 125m to 324m (410-1062ft) adjacent to active fumaroles west of Old
Highway 395. The wells encountered reservoir temperatures of ~170°C (338°F) and
Magma planned to construct a 15 Mw generation facility in 1962; however, flash/binary
production technologies were still in the early stages of development and the project was
shelved. Initial geothermal development in 1984 included four production wells and 2
injection wells supplying a 10 Mw (gross) pilot binary MP-1 production facility within
the original Magma well field. Successful generation led to additional development
drilling and in 1990 the installation of two 15 Mw binary MP-II and PLES-1 generating
plants. The project currently uses 6 production wells and 5 injection wells in the Casa
Diablo area plus production from 2 wells in the western caldera (Stop 2) to produce ~757
L/sec (~12,000 gal/min) of moderate temperature (170ºC) fluids to generate a combined
40 Mw (gross) of electricity. The project is managed and operated by Mammoth Pacific
LP, a partnership between Ormat Nevada Inc. and Constellation Power supplying base
load electricity under contract to Southern California Edison (SCE). Ormat has
modernized the older MP-1 plant and plans additional production and injection in Basalt
Canyon to increase production to ~1135 L/sec (~18,000 gal/min) and 60MWe (gross).
Mammoth-1, drilled by Unocal east of Old Highway 395 in 1979, was the first well at
Casa Diablo to evaluate potential deep production from the intracaldera Bishop Tuff and
the caldera basement. Mammoth-1 drilled through 390m (1280ft) of Early Rhyolite,
863m (2831ft) of Bishop Tuff and 230m (755ft) of metasedimentary rocks that correlate
with the Convict Lake roof pendant to the south, bottoming at 1605m (5265 ft.).
Mammoth-1 was also the first well within the caldera to encounter a landslide block of
chaotically mixed metapelite and granite at 466m in the upper section of Bishop Tuff
(Suemnicht, 1987). The well is currently used as a backup injector.
The results from Mammoth-1 and other deep tests indicate that production at Casa Diablo
is limited to shallow (<200m; 650 ft.) permeable Early Rhyolite flows and that the source
of the current system is not directly beneath the Resurgent Dome (Suemnicht, 1987).
Mineralization ages (Sorey and others, 1991) indicate that the Bishop Tuff magma
chamber supported a vigorous geothermal system from 300, 00 to 130,000 years ago
producing widespread hydrothermal alteration in and around the Resurgent Dome.
Drilling results establish that magmatic activity beneath the central part of the caldera has
waned over the past 100,000 years while activity in the western caldera has increased.
Shallow wells at Casa Diablo produce 170oC (338°F) outflow that is supplied by upflow
from an active geothermal system to the west. Injection wells return the produced fluid
to deeper (750m; 2460 ft.) permeable zones in the underlying Bishop Tuff. The shallow
Early Rhyolite reservoir is stratigraphically separated from the underlying Bishop Tuff by
an impermeable landslide block that controls the vertical distribution of shallow
hydrothermal circulation in the southern caldera allowing sustained production and
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injection at Casa Diablo by isolating warm shallow outflow from deeper cold natural
recharge and injection fluids that might quench the system.
Exit the geothermal facility and turn left on Casa Diablo Road. Turn right (west) on
Highway 203 and cross under the freeway and continue toward the Town of Mammoth
Lakes. Bear right (south) onto Highway 395 and drive approximately 1.5 miles (2.4 km)
to the intersection with Hot Creek Fish Hatchery Road. Turn left (north) on Hot Creek
Fish Hatchery Road and continue to the intersection with Hot Creek Rd. Park to the side
of the road at the entrance to the hatchery.
Incremental Distance –3.9 mi/ 6.3 km
Cumulative distance – 185 mi. / 295.3 km
Stop 4 – Hydrothermal Outflow Monitoring, Hot Bubbling Pool
Hot Bubbling Pool is part of a network of monitoring points (Figure 5) that includes hot
springs, fumaroles, cold springs, and stream flow and precipitation measurements in and
around the geothermal development at Casa Diablo. The US Forest Service, the US
Bureau of Land Management
and Mono County require
hydrologic monitoring of
geothermal operations in the
caldera. Monitoring is carried
out by the U.S. Geological
Survey and Mammoth Pacific
and is overseen by the Long
Valley Hydrologic Advisory
Committee, a cooperative
effort of both public and
private sector stakeholders.
Hydrologic data have been
collected and interpreted by the
USGS since the program’s
Figure 5. Hydrologic monitoring sites in Long Valley (USGS)
inception in 1985.
Monitoring data establish that the Long Valley hydrologic system is affected by
variations in precipitation, recharge, geothermal production, non-thermal groundwater
withdrawals, earthquakes and crustal deformation. Recharge, dominated by snowmelt
and precipitation at high elevations around the caldera’s topographic rim, appears to be
the primary influence on the caldera’s hydrologic system (http://lvo.wr.usgs.gov).
Thermal springs occur along faults in the southern and eastern parts of the caldera in and
around the Resurgent Dome. Approximately 75% of the thermal water discharge in the
caldera occurs at Hot Creek (Stop 5) while minor springs contribute comparatively
smaller but important thermal outflow. For example, warm springs at the Hot Creek Fish
Hatchery operated by the California Department of Fish and Game immediately to the
east of Hot Bubbling pool account for only 2-5% of the caldera’s total thermal outflow
but their thermal water contribution raises water temperatures an average of 5oC (41°F)
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above background. Hatchery fish are planted in many surrounding Sierra lakes and
streams and are an important part of regional recreation and the local tourist industry. An
example of the influence of geothermal development comes from production changes in
1991 that resulted in reservoir pressure declines of approximately 30 psi (~70 ft.) at Casa
Diablo (Sorey, 2005). Water levels in Hot Bubbling Pool, a 90oC (194°F) hot spring ~ 5
km (3mi) from the production facilities at Casa Diablo declined about 5 feet in 1991
when injection wells were deepened and production rates were increased at Casa Diablo.
In contrast, springs in Hot Creek Gorge (Stop 5) some 10 km (6mi) east of Casa Diablo
were unaffected by the 1991 pressure declines at Casa Diablo. As production has shifted
to Basalt Canyon, water levels in Hot Bubbling Pool have nearly recovered to previous
levels.
Recent studies of spring flow, temperature and water chemistry at the Fish Hatchery have
shown that no significant temperature changes have occurred in the mixed thermal and
non-thermal warm springs in response to geothermal development at Casa Diablo. This
appears to reflect the fact that Hatchery spring temperatures are buffered by heat transfer
between reservoir rock and fluid in the flow zone carrying mixed water from the Casa
Diablo area to the Hatchery (Sorey and Sullivan, 2006).
Continue east approximately on Hot Creek Rd. and park at public access area at the top
of Hot Creek gorge.
Incremental Distance – 3.6 mi/ 5.8 km
Cumulative distance – 188.6 mi. / 301.1 km
Stop 5 – Surface Manifestations, Hot Creek
The most prominent hot springs and
fumaroles in Long Valley are located at
Hot Creek, Little Hot Creek, and the
Alkali Lakes around the southern and
eastern edge of the Resurgent Dome.
The hot springs and fumaroles in Hot
Creek Gorge discharge ~250L/sec (3968
gpm) of near-boiling water representing
70-75% of the geothermal system’s
outflow. The creek cuts through a
288,000 year-old rhyolite flow that
Figure 6. Hot Creek Gorge view west (USGS photo).
originated from a Moat Rhyolite dome 4
km to the south (Bailey, 1976; Bailey
and others, 1989). The lower portion of the flow erupted into Pleistocene Long Valley
Lake burying existing lake sediments and pervasively hydrating the rhyolite. Brecciated
and perlitically altered flow units are exposed along the paved trail that leads down to the
Hot Creek. Bailey (1976) described the Hot Creek flow as sparsely phorphyritic,
sanidine-augite bearing rhyolite; a composition slightly different from other hornblendebiotite rhyolites of the moat eruptive sequences. He interpreted the mineralogic
difference as evidence that new mafic magma injections rejuvenated the Long Valley
magma chamber, inflating the Resurgent Dome and initiating basaltic eruptions in the
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western caldera moat. Age dates on the hot spring mineralization indicate that part of
the hydrothermal alteration common in Hot Creek gorge resulted from a vigorous
geothermal outflow from 300,000 – 100,000 years ago (Sorey and others, 1991).
Intense hydrothermal alteration and boiling (93°C;199°F) springs at Hot Creek are
localized along two north-striking faults that cut the rhyolite flow and form a small
graben that contains the Hot Creek swimming hole. Smaller or cooler springs are
localized along the contact between the flow and underlying lake sediments. Numerous
earthquakes that occurred during caldera unrest that began in 1980 have often affected
the flow of the springs. Additional boiling springs developed or were reinvigorated in
May 2006 expanding beyond the protective fencing on the north side of the gorge.
Changes in spring discharge were accompanied by temperature increases in nearby
monitoring wells and pressure increases in adjacent cold-water aquifers in response to
above normal precipitation in the preceding winter (Farrar and others, 2007). The
swimming area remains closed as a precautionary measure.
Exit the Hot Creek parking lot and head southeast on Hot Creek Road. Turn left (north)
on Owens River Road and immediately left onto Little Antelope Valley Road. The road
parallels Little Hot Creek for approximately 3 miles. Pull of to the right (north) side of
the road at the Hundley Clay Pit.
Incremental Distance – 9.6 mi/15.5 km
Cumulative distance – 198.2 mi. / 316.6 km
Stop 6 – Clay Pit
The Hundley Clay Pit is a large area of intense hydrothermal alteration in the Little
Antelope Valley in the northern part of the Resurgent Dome. Alteration is controlled by a
series of north-striking normal faults that are northern splays of the Quaternary/Recent
Hilton Creek fault system that Bailey (1976) considered the principal post-collapse
structural element in the central caldera. Hilton Creek fault splays also control
hydrothermal outflow at Hot Creek (Stop 5) to the south and east. Similar hydrothermal
alteration at Casa Diablo (Stop 3) occurs along northwest fault splays that define a medial
graben in the west-central part of the Resurgent Dome roughly along the route of
Highway 395. Clay Pit alteration is predominantly kaolinite with lesser amounts of
alunite and opal veins in Pleistocene tuffaceous lakebeds or older 600,000-year-old Early
Rhyolite units. The pit faces expose relict primary bedding in the lakebeds and relict flow
banding in altered Early Rhyolite units. The deposit is capped by an opaline claystone that
ranges in thickness from 3 in. (8 cm) to 20 ft. (6 m). Alteration occurred before deposition of
the youngest lakebeds suggesting that Clay Pit is related to the vigorous older
hydrothermal system developed within the central caldera 300,000 – 130,000 years ago.
Older or reactivated faults still control circulation in the current hydrothermal system that
has been active for the last 40,000 years. The Little Hot Creek springs occur along fault
splays that define the limits of alteration approximately 1 mile (1.6 km) east of Clay Pit.
The Hundley mine has operated intermittently since 1952. Standard Industrial Minerals,
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the current owner, trucks kaolinite from the Hundley pit to the company mill north of Bishop.
Kaolinite uses include paint filler, plastic, rubber, paper processing, Portland cement,
ceramics, insecticides, pharmaceuticals and stucco (Wilkerson and others, 2007; Lipshie,
2001). Prior to the federal Geothermal Steam Act that defined geothermal as a mineral
resource, Magma Power Company maintained claims on this and other minor kaolinite
prospects during the 1970’s to maintain grandfather mineral/geothermal rights and retain
leasing rights for federal geothermal lease sales in the 1980’s.
Clay Pit-1, on the southern side of the Hundley mine, drilled by Unocal in 1979, was
another deep test to 1847m (6060 ft) that evaluated potential production from the most
prominent hydrothermally altered area within the caldera. Clay Pit-1 drilled through
335m (1280ft) of Early Rhyolite, 1085m (3559 ft) of Bishop Tuff and 230m (755ft)
bottoming in 128m (420 ft) rhyodacite intrusive. Pervasive alteration of the rhyodacite
precluding age dating; however, major oxide chemistry closely matches moat rhyodacites
that erupted 500,000 years ago less that a mile from the Clay Pit mine. Clay Pit-1 was the
first well within the caldera to encounter younger intrusions into or beneath the Bishop
Tuff that could be a heat source for a hydrothermal system. The bottomhole temperature
of 148°C (298°F) was less than shallow production wells at Casa Diablo and a
conductive geothermal gradient and low permeability confirmed a lack of convective
circulation (Suemnicht, 1987). The results from Clay Pit-1 and Mammoth-1 confirmed
that the Resurgent Dome is not the source of the current active hydrothermal system in
Long Valley.
Continue west on Antelope Valley Rd approximately 2.5 miles (4 km). Turn right (north)
at the intersection of USFS Rd. 3S13 Stay to the right and continue approximately .8 mi
(1.3 km) to the Long Valley Exploratory Well site. Turn left (west) onto the wellsites
and park on the south side of the wellpad.
Incremental Distance – 4.8 mi/ 7.7 km
Cumulative distance – 203 mi. / 324.3 km
Stop 7 – LVEW DOE well
The Long Valley Exploration
Well was drilled to 2.3 km
(7545 ft) in 1991 and then
deepened to its current total
depth of 3 km (9842 ft) in
1998. The well was drilled in
the approximate center of 80
cm (2.6 ft) of uplift within the
Resurgent Dome to evaluate
the potential for extracting
energy from magmatic or
near-magmatic conditions in
Figure 7. Long Valley Exploration Well in 1991. (USGS photo)
Long Valley. Measurable
uplift and continued seismicity were interpreted as evidence for potential intrusion of new
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magma that in 1980 initiated a period of unrest in the caldera. The original funding for
the drilling came from the DOE and the well is currently part of the International
Continental Scientific Drilling Program (www.icdp-online.org).
The well penetrated 662m (2171 ft) of Early Rhyolite, 1178m (3865 ft) of Bishop Tuff
and entered metamorphic roof pendant rocks at 1891m (6204 ft). The metamorphic rocks
in the lower portion of the well correlate with the Mt. Morrison metamorphic roof
pendant exposed in the southern caldera rim and at Convict Lake to the south
(McConnell, 1995). Although permeable fluid zones were encountered in the crystalline
metamorphic rocks at depths of 2-3 km, the maximum temperatures in the well are only
104°C (219°F) (Figure 8). Significant data from the well were summarized in a special
issue of the Journal of
Volcanology and Geothermal
Research (v. 127, no. 3-4,
2003).
The well remains as an
observation point for
studying caldera unrest. The
water level response to rock
strain within the well is
significantly greater than
strain sensitivities for
shallower observation wells
(Roeloffs, 2001). The
LVEW well is one of the best
sites within the caldera for
Figure 8. Measured (solid) and extrapolated (dashed) LVEW
monitoring rock strain related to
temperature profile (Sorey and McConnell, 2000).
seismicity and potential magmatic
Generalized lithologies: post-caldera Rhyolites (PCR),
intrusion. A downhole seismometer Bishop Tuff (BT), Mesozoic metavolcanic rocks (Mzmv), and
was installed in the well in 2003 and Paleozoic metasediments (Pzms). The fracture zone
is currently the deepest observation encountered by the well and projected ductile conditions are
point for seismic monitoring within included.
the caldera (Peter Malin, pers. comm.)
Return to Antelope Valley Rd. and turn right (southwest). Return back past the Casa
Diablo production facilities and Casa Diablo Rd. Cross under Highway 395, drive
through the Town of Mammoth Lakes and turn right (north) on Minaret Summit Rd.
Drive to the Alpenhof Lodge and park.
Incremental Distance – 8.4 mi/ 13.5 km
Cumulative distance – 211.4 mi. / 337.8 km
- END OF THE FIRST DAY –
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DAY 2: Western Caldera
The second day of the trip leaves from The Alpenhof Lodge and all road mileage is based
on distance from that point.
Exit the Alpenhof parking lot and turn right (north) on Minaret Summit Rd. Continue
west through the Mammoth Mountain parking lot to the Forest Service entry hut at
Minaret Summit. Continue 5.5 mi (8.9 km) west and down the winding road towards
Red’s Meadow and the San Joaquin River valley. Smell the brakes burning. Slow and
turn right (west) into the Devil’s Postpile parking lot and park. Hike approximately 0.3mi
(0.5 km) to the base of the lava flow.
Cumulative distance – 13.5 mi. / 21.7 km
Stop 1 – Devils Postpile National Monument
Devils Postpile is a porphyritic basalt
flow that erupted ~ 80,000 years ago
from a cinder cone(s) ~ 2 mi (3.2 km)
to the north around Soda Springs,
flowed down the San Joaquin river
drainage and ponded in a lowland
against a Tahoe age recessional
moraine (Bailey, 04; Mahood and
others, 2010). The Devils Postpile
flow is one spectacular example of a
series of mafic lavas that erupted from
several widely scattered vents both
inside and outside the caldera around
the base of Mammoth Mountain. The
mafic units range in age from 120,000
to 57,000 years old and were erupted
during the same time that smaller
volume high silica dacites and
rhyolites were erupting in the western
caldera moat and massive volumes of
dacites and rhyodacites erupted to
form the bulk of Mammoth Mountain
(Mahood and others, 2010).
Figure 9. National Pak Service map of Devils
Postpile National Monument
Relatively uniform composition and
the ~400 ft. (122m) thickness of the
ponded lava resulted in slow cooling
between the granite at the base and
the upper surface of the flow exposed
to the atmosphere. The hexagonal
joint pattern that defines the columns
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developed as contraction
cracks as a cooling front
moved from the edges of the
flow inward toward the
center (Figure 10). Kilauea
lava lakes form surface
cracks at ~900°C (1690°F)
but at lower temperatures of
~750°C (1382°F) in the
interior bulk of flows
(Goehring and Morris, 2008).
The 120° angles of the joints
are common in many natural
settings where volume loss
Figure 10. Columnar jointing in the Devil’s Postpile
during shrinkage results in a
flow (from USGS).
uniform distribution of
horizontal stress. The
hexagonal shape provides the greatest stress relief and allows the largest surface area for
dissipating heat. As might be expected, the fractures propagate intermittently and branch
to join other cracks after reaching a critical length (Huber and Eckhardt, 2002). The
columnar jointed section of the Devils Postpile flow is smaller compared to jointed
sections in massive flood basalts or the famous Fingal’s Cave in Scotland or the Giant’s
Causeway in Ireland but the 12- to 18-m (40- to 60-ft) length of the columns and high
percentage (55%) of perfect hexagonal columns makes Devils Postpile unique.
Standing on top of the outcrop,
the current surface of the flow is
polished and striated by
glaciation (Figure 11). The
Middle Fork of the San Joaquin
drainage was established by the
Pleistocene and the glaciers most
likely flowed from north to south
within the canyon. Other
evidence of glaciation includes
glacial erratics (boulders or
Figure 11. Glacial striae on the surface of Devil’s
cobbles carried by a glacier a
Postpile flow (from USGS).
distance from their point of
origin) distributed within the monument and along the river drainage. Glaciation removed
much of the edge of the Devils Postpile flow exposing the flow interior. Since the flow
erupted 80,000 years ago near the end of Tahoe age glacial advances, the current surface
of the flow is the result of younger Tioga (12,000-25,000 years) glaciers. Estimates based
on theoretical cooling/thickness relationships suggest that glaciation may have removed
as much as 50% of the original thickness of the flow. (NPS, 2009).
National Park Service brochures recount the struggle to establish Devils Postpile as a
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national monument. The boundaries of Yosemite National Park included Devils Postpile
in 1890 but timber and mining interests pressured Congress to reduce the park’s size. The
first park Superintendents were military officers who agreed that the southeast part of
Yosemite should be returned to the public domain and Congress removed 500 mi2 (1,300
km2) from Yosemite in 1905 including Devils Postpile. Walter Huber, the district
engineer for the U.S. Forest Service in 1910, received an application from a mining
company to dynamite the Postpile and use the rock as the core of a hydroelectric dam on
the Middle Fork of the San Joaquin River. Huber, John Muir, and other members of the
(then) new Sierra Club opposed the plan and led a letter-writing appeal to Washington.
President Taft designated the area as Devils Postpile National Monument on July 6, 1911.
Exit the parking lot and turn right (northeast) onto the Red’s Meadow/Minaret Summit
Road. Drive back up road to Minaret Summit (optional stop). Continue east through the
ski area parking lot and head east on Minaret Summit Rd. toward the Town of Mammoth
Lakes. Turn left (north) at the small side entrance road to the Earthquake Fault parking
lot.
Incremental Distance – 9.5 mi/ 15.3 km Cumulative distance – 23 mi. / 37 km
Stop 2 – Earthquake Fault
Figure 12. Earthquake Fault
view north.
The Earthquake Fault is the best-known dilational ground
crack in the western part of Long Valley Caldera (Figure
12). A more appropriate name for the fracture would be
Earthquake Fissure because there is little evidence that the
rocks on either side of the fracture have moved vertically
or laterally relative to one another. Irregularities in the
joint faces across the fissure nearly match perfectly with
some suggestion of right-lateral pull-apart separation. The
fissure is up to 3m (10 ft) wide and 18m (60 ft) deep and
cuts through 113,000 year old dacites or 81,000 year old
rhyodacites of Mammoth Mountain (Mahood and others,
2010). Snow from the winter months commonly lasts all
year round in the deep bottom portions of the fissure and
local Indians stored their food in the fissure during warmer
summer months.
Earthquake Fault is one of a series of western caldera fissures and faults (see stop #3)
aligned along the Inyo-Mono Craters volcanic chain that, in a regional sense are part of
the east-west stretching that is gradually widening the entire Basin and Range. Early
studies (Benioff and Guttenberg, 1939; Rinehart and Ross, 1964; Bailey and others,
1976) concluded that these features are tectonic in origin and suggested that the cracks
and faults might represent a southern extension of the Harley Springs Fault from the
north. Mastin and Pollard (1988) concluded that the faults and fissures in the western
part of Long Valley are related to crustal dilation above the southern portion of 8–10-kmlong dike that strikes north out of the caldera supplying dome-forming eruptive events in
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the north Mono Craters about 600 years ago (Bursik & Sieh 1989) and here along the
south end of the Inyo Domes (Sorey and others, 1998). Both interpretations are
compatible because the Inyo-Mono dike is probably controlled by preexisting faults that
allowed magma into the shallower crust similar to previous intrusive events in the
western caldera while the cracks and fractures are related to shallow uplift and
extensional stresses immediately above a rising Mono-Inyo dike.
The age of Earthquake Fault is unkown or at least poorly constrained. Tree rings indicate
that the trees growing in the crack are ~160 years old. The crack lacks pumice fill so the
fissure formed after the eruption of the Inyo Craters 600 years ago. Indian legends
recount a massive earthquake in the area ~ 200 years ago causing the Earthquake Fault
rift. Separating fact from myth is difficult but the age and cause are tantilizing since the
legendary event roughly corresponds to a shallow intrusion beneath Mono Lake ~200
years ago that pushed up lake-bottom sediments to form Pahoa Island and produced a
small volume of andesitic lavas from vents on the island’s north side (Lajoie, 1968;
Bursik & Sieh 1989) (Day 3, Stop #3).
Benioff and Guttenberg (1939) placed stainless steel pins in opposite walls of the fissure
and they or CalTech colleagues periodically measured the distance between the pins with
a steel rod to determine strain over time. Through the 1940s, the fissure widened by
approximately a millimeter but after a 20-year break, 1967 measurements found the crack
had closed enough that the steel caliper bar would no longer fit between the pins. No firm
evidence was ever presented to substantiate rumors that Earthquake Fault was reactivated
during the 1980s Mammoth Lakes earthquakes. Soil settling noted at the bottom of the
rift was probably related to consolidation of loose material during the series of M6
earthquakes in May 1980 and M5 events in January 1983. A number of collapse or
settling pits up to 5 m (16 ft) in diameter also formed along this and other ground cracks
in the western caldera during the Mammoth Lakes earthquake sequence. The US Forest
Service restricted access to the fissure during the seismic unrest of the 1980s and
removed damaged ladders that had previously allowed tourist access to deeper levels of
the rift.
Exit the Earthquake Fault parking lot and turn left (east) on Minaret Summit Road toward
the Town of Mammoth Lakes. BEWARE OF TRAFFIC! Turn left (north) at the
intersection with Mammoth Scenic Loop. The road crosses a plateau of 100,000 year-old
Moat Rhyolites that underlie Mammoth Knolls on the north side of Mammoth Lakes,
turns northeast and drops down into the Dry Creek basin. Turn left (north) on Inyo
Craters Road and follow the road north to the Inyo Craters parking lot. The trail to the
craters is on the northwest side of the parking lot and wanders approximately 0.3 mi
/500m northwest through the forest to the base of the craters.
Incremental Distance – 7.3 mi/ 11.7 km Cumulative distance – 30.3 mi. / 48.7 km
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Stop 3 – Inyo Craters
Phreatic explosions created
the Inyo Craters
approximately 600 years ago
(Bursik & Sieh 1989)
(Figure 13). The magma
source for these eruptions is
an 8–10km (5-6 mi)-long
dike that trends north out of
the caldera. Phreatic craters
occurred where the dike
encountered groundwater
and extend from the north
flank of Mammoth
Figure 13. Inyo Craters, view north to Deadman Dome (USGS
Mountain (across the
photo)
parking lot south of the
Mammoth Mountain Inn) to the Mono Basin. Dome-forming eruptive events occurred
where magma made it to the surface at the north end of the Mono Craters (Day 3, Stop 2)
and south within Long Valley Caldera at the Inyo Domes immediately north of the
craters.
The Inyo Craters are aligned north south along
the presumed strike of the Inyo dike (Figure 14).
Two southern craters are approximately 200m
(650 ft) across and 60m (200 ft) deep and
contain small cold (11°C;52°F) lakes that show
no evidence of thermal influence. The color of
the water in each lake is apparently the result of
differing types of algae or organic matter. A
phreatic eruption also excavated a small
northern crater at the summit of Deer Mountain,
Figure 14. Conceptual diagram of the Inyo
a 115,000 year-old moat rhyolite dome similar
dike (USGS adaptation after J. Fink)
in age composition to the rhyolite plateau north
of Mammoth Lakes. The light colored phreatic
explosion debris at the summit (Figure 13) is fragmented hornblende rhyolite of the dome
and the darker debris around the two lower craters is predominantly andesite derived
from underlying flows exposed in the walls of the southern crater (Bailey and others,
1989). The stratigraphy exposed in the craters includes andesite flows overlain by a thin
(1m) layer of Inyo Craters pumiceous rhyolite tephra overlain by a succession of poorly
sorted phreatic explosion debris. Based on detailed stratigraphic studies Mastin and
Pollard (1988) determined that the craters erupted from north to south with light-colored
explosion deposits from Deer Mountain summit overlain by dark debris from the middle
crater in turn overlain by darker debris from fragmented andesites underlying the
southern crater. Phreatic debris from the southern crater extends as far as 0.6 mi/1km to
the northeast and includes granite and metamorphic fragments probably from underlying
till or basement rocks.
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The Inyo Craters lie within a small .6 km wide
graben defined by numerous faults and
fissures with as much as 20 m of displacement
extending 2.5 km (1.5 mi) north and south
(Figure 15). Similar fissure and fracture zones
are common in the western moat of Long
Valley Caldera and control similar phreatic
eruptive centers. Mastin and Pollard (1988)
attributed these fissure zones to uplift and
extension above a rising Mono-Inyo dike
although the intrusion probably followed some
existing inherited structures similar to
previous intrusive events in the western
caldera over the last 200,000 years (Suemnicht
and Varga, 1988). Phreatic eruptions were
widespread along the Mono-Inyo dike and
numerous equally impressive but less
accessible craters occur to the north and west
of the Inyo Craters that do not conform to a
uniform north-south orientation.
Figure 15. Generalized map of faults and ground
cracks around the Inyo Craters (After Mastin and
Pollard, 1988)
A series of DOE -funded core holes
penetrated and sampled the Mono-Inyo dike
inside and outside the caldera (See also Stop 5). The RDO-4 corehole was drilled from
the top of the fault scarp on the western side of the south Inyo Crater at an angle of 68°,
and penetrated 320m (1050 ft) of post caldera andesites and basalts, 50m of gravel, 360m
of Early Rhyolite tuffs and flows, 63m of glacial till bottoming in 30m of Sierran
metamorphic rocks (Eichelberger and others, 1988) (Figure 16). The highest temperature
in the hole was 81°C (178°F at 500m (1640 ft). The hole encountered limited amounts of
juvenile high-silica rhyolite and vent breccias composed of Early Rhyolite and andesite
directly under the southern crater. Petrologic studies of the core indicate that the
eruptions may have tapped magmas of three different compositions venting at several
places along strike (Eichelberger, 2003).
Return to Inyo Craters Road. Turn left (northeast) at the first Forest Service road on the
left. Park on the northwest side of the 44-16 wellpad.
Incremental Distance – 0.6 mi/ 1 km
Cumulative distance – 30.9 mi. / 49.7 km
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Stop 4 – Inyo Domes 44-16
Unocal drilled well 44-16 approximately 1 km (0.6 mi.) east of the Inyo Craters to test
the geothermal potential of the western caldera moat. While 44-16 is the hottest well
drilled in Long Valley (218°C at 1100m)(424°F at 3608 ft.), it is unproductive because a
strong incursion of cold water penetrates beneath a limited in thickness of Bishop Tuff
reservoir rocks (Suemnicht, 1987). The Inyo-4 corehole and 44-16 establish the location
of the western structural margin of Long Valley Caldera. Between the two wells, the
caldera’s ring fracture system offsets the metamorphic rocks of the caldera floor by
approximately 1km in less than 1km lateral distance (Figure 17). As with many calderas,
the structural margin of the ring fracture system can be considerably different from the
topographic margin or rim of the caldera that we see today.
Figure 16. Inyo-4 core hole (from
Eichelberger and others, 1988)
Figure 17. Stratigraphic correlations
between the Inyo-4 corehole and IDFU 4416 adjacent to the western ring fracture
system of Long Valley (From Eichelberger
and others, 1988)
Return on Inyo Craters Road to Mammoth Scenic Loop. Turn left (northeast) and drive to
Highway 395. Cross the southbound lane and turn left (northwest) toward June Lake.
Turn left at Obsidian Dome Road. Turn left (south) on USFS road 2S10 and park at a
locked gate on the west flank of the dome. Hike past the gate up the quarry road into the
dome interior to a flat space with some sparse trees to the southeast.
Incremental Distance – 12.2 mi/ 19.6 km
Cumulative distance – 43.1 mi. / 69.4 km
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Stop 5 – Obsidian Dome
Obsidian, Glass Creek, and Deadman domes are
aligned north-south along the central-southern
part of the Inyo dike. The domes and numerous
phreatic craters cross the topographic and
structural boundaries of Long Valley Caldera.
Obsidian Dome erupted just outside the
topographic margin of the caldera. Dome
formation is the end product of an eruptive
sequence that begins with phreatic (steam)
explosions excavating a crater (like the Inyo
Craters at Stop 3) followed by pyroclastic
eruptions of ash, lapilli and pumice progressing
to extrusion of lava at the surface. Each eruptive
phase can be found along the length of the Inyo
dike. Initial explosive phases are related to dike
intrusion into the groundwater table or to the
high volatile content of the magma
(predominantly gasses and water). The gas
content of the magma declines during the
explosive phase decreasing the drive for
fragmenting and ejecting magma. Eventually the
eruption
progresses to a comparatively passive
Figure 18. Obsidian and Glass Creek
extrusion of thick pasty rhyolite lava to form a
domes with plan views of the DOE
dome. All three domes erupted in the late summer
coreholes penetrating the Inyo dike.
of 1350 AD based on detailed dating studies on
(From Eichelberger and others, 1985)
trees that died in the resulting ash fall that covers
White Wing Mountain to the west (Millar, et al.2006). The USGS refers to the event(s) as
one of the best-constrained prehistoric eruption dates in the world.
The DOE drilled coreholes in
Obsidian Dome in 1984 as part of
a project to intersect and sample
the Inyo dike (Eichelberger and
others, 1985). A 55° inclined hole
in the dome interior was drilled to
the northeast to a slant depth of
624m (2047 ft) in the approximate
location where we are standing
(Figure 19). The corehole entered a
51.1m (167.5 ft) thick magma
conduit of mixed rhyolite and
brecciated granite starting at 482m
(1581 ft, true vertical depth) below
the surface. A second hole inclined
at 54° was drilled almost due west
Figure 19. Obsidian Dome corehole 1984 (view SE).
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from a phreatic explosion crater located between Obsidian Dome and Glass Creek Dome.
The corehole reached 829m (2720 ft) slant depth and penetrated 7.3m (29.5 ft) of a
rhyolite obsidian dike starting at 640m (1581 ft, true vertical depth).
The DOE coring project was remarkably
successful in penetrating the Inyo dike
directly beneath obvious surface features
and in sampling the quenched rhyolite
magma supply for Obsidian Dome.
Temperatures in the dike hole outside the
dome (Figure 20) are generally in the
range of background heat flow for the
Sierran block indicating a limited amount
of heat from the Inyo dike event with
little broad regional affect. The actual
conduit beneath Obsidian Dome proved
to be more complex than originally
expected. The chemical composition of
the intrusive phases span the entire
compositional range of Inyo chain but, as
discussed at the Inyo Craters (Stop 3), is
Figure 20. Temperature-depth relationships for
coreholes penetrating the Inyo dike. (From
still simpler than the range of basalt to
Eichelberger and others, 1985)
high silica rhyolite compositions
penetrated in the RDO-4 corehole inside
the caldera. The conduit hole core shows the dike is complexly brecciated beneath the
dome interior with abundant wall rock xenoliths in roughly ⅓ of the penetrated volume.
In contrast, the southern dike hole between Obsidian and Glass Creek domes is
comparatively intact and free of xenoliths implying that a large quantity of volatiles must
have escaped during the early eruptive phases because a volatile rich rhyolite magma at
less than 640m depth should be explosively fractured. Glass is scarce in the dike intrusion
compared to the distinctly glassy extrusive dome. The coring confirmed that pre-existing
faults were a principal control in guiding magma to the surface and that mapping the
internal structure of domes and surrounding surface features can be important clues to the
underlying structural controls beneath a thick blanket of otherwise featureless young
volcanic rocks.
Return on Obsidian Dome/Glass Flow Road to Highway 395. Turn right (southeast)
toward Mammoth Lakes. Turn right (southwest) at Mammoth Scenic Loop road and head
toward Mammoth Mountain and return to Alpenhof Lodge.
Incremental Distance – 14.3 mi/ 23 km
Cumulative distance – 57.4 mi. / 92.3 km
- END OF THE SECOND DAY –
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DAY 3: Mono Basin
The final day of the trip leaves from The Alpenhof Lodge and all road mileage is based
on distance from that point.
Exit the Alpenhof parking lot and head west on Highway 203 through the Town of
Mammoth Lakes. Turn right (north) on Minaret Summit Rd. Continue west to the
intersection with the Mammoth Scenic Loop and turn right (north) to cross a plateau of
100,000 year-old Moat Rhyolites and continue northeast through the Dry Creek basin.
BEWARE OF TRAFFIC at the intersection with Highway 395. and cross the highway.
Turn left (northwest) toward June Lakes and Mono Lake. Exit the caldera at Deadman
Summit. Continue and approximately 1.5 mi. /2.4 km from the summit and pull over to
the eastern side of the road near the sign noting Obsidian Dome Road. BEWARE OF
TRAFFIC!
Incremental Distance – 16.3 mi/ 26.2 km
Stop 1 - Bishop Tuff, Northern Lobe
The Bishop Tuff eruption produced approximately 600 km3 of rhyolite ash evacuating the
upper part of the Long Valley magma chamber causing the chamber roof to subside
approximately 1-2 km creating Long Valley Caldera. Ashfall from the eruption is
recognized east as far as Kansas and Nebraska (Izett, 1982) and in deep sea cores from
the eastern Pacific (Sarna-Wojcicki and others, 1987). Collapse of the eruption column
produced as many as seven ash flow lobes that blanketed 1500 km2 surrounding the
caldera (Hildreth, 1979). The ash flow completely covered precaldera topography and is
thickest near the margin of the caldera thinning at the northern and southern distal edges.
The ash flow sheet at this stop is the moderately welded upper part of the northern Mono
Basin lobe of the eruption (Hildreth, 1979). At least 75% of the original expanse of
northern Bishop Tuff is covered by post-caldera deposits of glacial outwash, Lake
Russell/Mono Lake sediments and Mono Craters tephra. The remnant outcrops are
referred to as the Aeolian Buttes in the Mono Basin. The mineral assembly of
quartz+sanidine+ biotite is typical of the Bishop Tuff. Hildreth (1979) noted that the
mineral compositions within the Tuff changed as the eruption progressively tapped
deeper and hotter levels of the magma chamber. Calculated Fe-Ti oxide eruption
temperatures of 756-790oC (1393-1454°F) establish the Mono Basin lobe of the Bishop
Tuff as the highest temperature phase of the eruption. The pumice fragments within the
Tuff are still relatively coherent at this location with some evidence of collapse or
welding in the upper part of the ash flow sheet. Some xenoliths can be found in northern
and eastern lobes of the Tuff but they are much more prevalent in southern ash flow
lobes. Hildreth and Mahood (1986) interpreted rock types and lithic fragment amounts to
indicate that the caldera forming eruption progressed from the south central margin near
Lake Crowley counterclockwise to the north and west culminating on the northern
caldera margin. Detailed stratigraphic studies and eruption modeling suggest that the
Bishop Tuff eruption probably took a week or less (Wilson and Hildreth, 1997).
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Most volcanic activity over the last 200,000 years has occurred in the western moat of
Long Valley and over the last 50,000 years volcanism has occurred north of the caldera in
the Mono Basin. Obsidian Dome across Highway 395 to the west is one of a series of
eruptions that occurred 600 yrs. ago along the Mono-Inyo volcanic chain. Pumice outfall
from eruptions over the last 2,000 years mantle the outcrops at this location and much of
the area including the interior of the Mono Basin.
Rejoin Highway 395 and continue north past June Lake Junction toward the town of Lee
Vining. Turn right (east) on Highway 120. Drive approximately 3 mi. / 5 km. and turn
left (north) on a dirt road poorly marked as the access to Panum Crater. Continue north
approximately .5 mi. /.8 km. to the parking lot at the southeastern edge of Panum Crater.
Incremental Distance – 12.3 mi/19.8 km Cumulative distance – 28.6 mi. / 77.4 km
Stop 2 – Panum Crater
Panum Crater is the farthest northern
vent along the Mono-Inyo dike that
erupted between AD 1325-1365 (Bursik
& Sieh 1989). The vent is a textbook
example of a rhyolite plug-dome volcano
and this stop serves as the prime
illustration of eruptive progressions
common to the Mono-Inyo dike event
600 years ago (Figure 18). The tephra
ring is approximately 0.8 km in diameter Figure 21. Panum Crater view SW(from USGS)
and is composed of pumice, ash, lapilli,
obsidian fragments and minor amounts of granitic or metamorphic fragments derived
from underlying glacial deposits. Detailed studies by Sieh & Bursik (1986) indicate that
phreatic explosions cleared the throat of the Panum eruptive center, progressed to
pyroclastic flow and surge deposits, constructed the outer tephra ring and eventually
extruded a high-silica rhyolite dome as magma made it to the surface. Stratigraphy of the
surrounding pyroclastic deposits indicates two periods of explosive activity from the
same eruptive center. The northern part of an early rhyolite dome apparently exploded or
collapsed producing a block avalanche deposit to the northwest into Mono Lake (Wood,
1977; Sieh & Bursik, 1986) (Figure 19). Later pyroclastic ashfalls healed the breach of
the tephra ring and a composite dome was built by further intrusions. Panum rhyolite has
a silica content of 75% and the rocks in the dome interior grade from highly pumiceous to
dense obsidian. The flow textures within the dome illustrate the complex composite
nature of dome formation. Breadcrust textures are common where viscous rhyolite
surfaces cooled and degassed.
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Eruptive styles along the Mono-Inyo dike are
apparently controlled by magma type, the depth
of penetration into the shallow crust, variations
in country rock and groundwater (Figure 14).
Portions of the dike that intruded local
groundwater may have fragmented producing
phreatic explosion craters like the Inyo Craters
and the phreatic craters on the north face of
Mammoth Mountain without surface extrusions.
Early phreatic eruptions at Panum Crater cleared
the path for later magma to make it to the
surface and eventually build the rhyolite dome
in the center of the tephra ring. The eruptive
progression continued at other places along the
dike producing high-silica rhyolite flows that
overrode tephra deposits producing the Inyo
Domes about 550 years ago (Miller 1985).
Figure 22. Generalized map and cross
sections of Panum Crater (after Bursik
and Sieh, 1989)
Return to Highway 120 and turn right (west). At the intersection with Highway 395, turn
right (north) toward the town of Lee Vining. Approximately .25 mi (400m) north of
town, turn right on the access road to the US Forest Service National Scenic Area
Visitors Center.
Incremental Distance – 12.7 mi/ 20.4 km Cumulative distance – 41.3 mi / 97.8 km
Stop 3 – Mono Craters Visitor Center
The Mono Craters National Scenic Area is operated by the Inyo National Forest and
provides an excellent overview of Mono Lake and the Mono-Inyo volcanic chain. The
Mono Craters are the result of several thousand years of volcanic activity that began
~40,000 years ago. At least 20 small volume eruptions have occurred along the chain
over the past 5000 years at intervals of 250 to 700 years (Bursik & Sieh, 1989). The
progression of eruptions over the past 2 Ma from Glass Mountain on the eastern margin
of Long Valley Caldera to Mammoth Mountain on the west and the Mono–Inyo volcanic
chain in the north suggests that the caldera’s magmatic system has declined with time and
has been supplanted by mixed composition eruptions from the active Mammoth
Mountain–Inyo Domes magmatic system (Hildreth, 2004). The progression of eruptions
in the area did not end 600 years ago with Panum Crater or with the extrusion of the Inyo
Domes. A shallow intrusion beneath Mono Lake ~250 years ago pushed up lake-bottom
sediments to form Pahoa Island and produced a small volume of andesitic lavas from
vents on the island’s north side (Lajoie, 1968; Bursik & Sieh 1989). Mixed magma
compositions are consistent with regional crustal extension in the Eastern California
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Shear Zone and the Mina Deflection connection to the Walker Lane to the northeast.
Bursik (2003) suggested that dikes beneath the domes intrude a pull-apart structure to
accommodate strain otherwise resolved by purely extensional faulting.
The eastern Sierra is also a key location for understanding regional glacial events over the
last 2-3 Ma. Prominent lateral and terminal moraines along the Sierran range front west
of the visitor’s center are the result of the latest Tahoe (75,000 – 60,000 years) and Tioga
(25,000 – 12,000 years) glacial advances and retreats during the late Pleistocene. The
Casa Diablo till intercalated between ~200,000 year-old moat basalts in the southern
caldera moat in Long Valley provide well-dated evidence of a Mid-Late Pleistocene
200,000 year-old glacial event that may also correlate with Mono Basin glaciation of
similar age. The Sherwin till underlies the Bishop Tuff at Sherwin Creek and marks a
glacial retreat approximately 800,000 years ago. Tills on McGee Mountain south of
Long Valley are evidence of an earlier 2.5 Ma glaciation (Table 1).
Table 1. Correlation of glacial sequences in the Eastern Sierra (adapted from Lipshie, 2001)
North American Continental
Glacial Sequences
Post Wisconsin
Late Wisconsin
Early Wisconsin
Illinoian
Kansan
Nebraskan
Sierra
Glacial Sequences
Matthes
Recess Peak
Hilgard
Tioga
Tenaya
Tahoe
Mono Basin
Casa Diablo
Sherwin
McGee
Ages
(years before present)
0-700
2000-4000
9000-10,500
12,000-25,000
~ 45,000
60,000 – 75,000
~130,000
200,000(?)
~800,000
1.5 – 2.5 Ma
Mono Lake is the saline remnant of Lake Russell, an extensive freshwater lake that
formed in the Mono Basin during Pleistocene glacial retreats. The glacial lake is named
for Israel Russell who mapped and studied the region in 1881. He interpreted the
moraines along the range front as evidence of at least two major glacial events in the
Sierra, mapped the glacial lake shorelines and correlated the glacial advances with high
water levels in Lake Lahontan and Lake Bonneville elsewhere in the Great Basin
(Russell, 1889). The shoreline terraces of Lake Russell extend to an elevation of 2188m
(6948 ft.) (Putnam, 1950) or roughly 302m (990 ft.) above the present 1886m (6187ft)
level of Mono Lake and are evident along the shore of the lake west of the visitor’s
center. The scarp along the western shoreline marks the Lee Vining Fault part of the
Sierran range front fault system. As we drive north along the lakeshore, note the uplifted
shoreline terraces and deformed young lakebeds along the fault scarp.
The predominant water source for Mono Lake is snowmelt in the creeks that flow into the
Mono Basin from the Sierran range front. The Los Angeles Department of Water and
Power (LADWP) acquired water rights throughout the basin and in 1941 began exporting
water from Walker, Parker, Lee Vining and Rush creeks through an 18.5km (11.5mi)
tunnel under the Mono Craters to the Owens River in Long Valley, and eventually to the
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Owens Valley aqueduct that extends 359 km (223mi) from Lake Crowley in Long Valley
to the northern part of the Los Angeles basin. Other minor streams that drain into the
closed basin contribute relatively small volumes of water. Without continued freshwater
supply evaporation eventually exceeded inflow resulting in declining lake levels and
rising salinities. The total dissolved solid content (TDS) of Mono Lake water was 50,
000mg/ L in 1941 and at the lake’s lowest level in 1982 the salinity was 99,000mg/L.
The declining water levels eventually formed a peninsula on the north side of Negit
Island exposing migratory bird nesting areas to predators. The expropriation of water was
eventually restricted by court decisions and in 1994 the California State Water Resources
Control Board ordered LADWP to release enough water to maintain the elevation of the
lake at 1945m (6381 ft). The current lake level is ~2m below that elevation but Negit is
once again an island and salinities are anticipated to stabilize around 70, 000 mg/L.
The young siliceous Mono Craters eruptive complex is an obvious geothermal
exploration target enhanced by the occurrence of several hot springs within the Mono
Basin (Figure 23). Records from the LADWP Mono Craters tunneling project document
large volumes of 36°C (97°F) water and CO 2 (Jaques,1939). Some recent publications
cite “boiling hot water” where the tunnel crossed beneath the central axis of the crater
chain; a claim unsubstantiated in LADWP records that cite vertical shafts sunk through
“…sands and silts which, due to the water pressures behind or under them, as they were
approached, would cause them to boil up and fill the shaft for many feet,” (emphasis
mine) (Jaques, 1939). The tunnel passes through 884 m (2,900 ft) of glassy or fractured
rhyolite, “large crevasses” (fault zones) and heavy flows of CO 2 . The maximum water
flow from all the tunnel headings was 1262 L/s (20,000 gpm) under substantial head that
would certainly account for water “boiling” into the tunnel shafts.
Wildcat exploration wells on the north and south shores of Mono Lake did not encounter
encouraging temperatures to depths of 744m (2440 ft)(Figure 23). A Unocal temperature
gradient hole at the southeast end of the volcanic chain encountered 28°C (82°F) at 893m
(2929 ft). The Mono Craters are certainly a spectacular example of comparatively recent
silicic volcanism but the available drilling data indicate that neither the Inyo section of
the dike (Stops 3 and 5) or the longer-lived Mono eruptive complex provide regionally
extensive heat for a substantial hydrothermal system in the shallow crust.
Return to Highway 395. Turn right (north) heading toward Reno. Leave the Mono Basin
at Conway Summit and continue north to the intersection with State Route 270. Turn
quickly right (east) and continue approximately 9 miles (14.5 km) to Bodie State Park.
Incremental Distance – 16.2 mi/ 26 km
Cumulative distance – 57.5 mi. / 123.8 km
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Figure 23 – Generalized topographic map of the Mono Basin showing warm springs and
geothermal exploration wells (from CDOGGR well files)
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Stop 4 (optional) – Bodie
The gold and silver mining town of
Bodie is one of the best preserved
ghost towns in the old west and its
designation as a National Historic
Landmark and a California State Park
has kept it from being carted away
piece by piece over time. Most of the
information about the town and the
geology of the mineral deposit are
taken from Chesterman and others,
1986; Hollister and Silverman, 1995;
Figure 24. Bodie, CA.
John and others, 2011. The main
purpose of this (optional) stop is to review some of the extensive and well-documented
hydrothermal processes responsible for the mineralization at Bodie. Epithermal mineral
deposits like Bodie provide excellent conceptual models and analogs for modern
geothermal systems particularly in intermediate-composition volcanic environments that
host more than 50% of the world’s known and developed geothermal systems (Brophy,
2008).
The Middle-Late Miocene Bodie Hills volcanic field is a large ~700 km2 (270 mi2) longlived (9my) but episodic andesite-dacite eruptive center that is part of the southern
ancestral Cascades arc (John and others, 2011). Peak periods of eruptive activity include
trachyandesite stratovolcanoes between14.7-12.9 my ago and trachyandesite-dacite dome
fields between ~9.2-8.0 my ago. Tertiary pyroclastic rocks, flows and intrusive plugs
(Silver Hill series, Murphy Spring Tuff Breccia, Potato Hill Formation) around Bodie are
equivalents of a regional suite of calc-alkaline basalt, andesite, dacite, and rhyolite that
erupted in the Eastern Sierra during the Miocene. Diffuse regional low-grade
hydrothermal alteration over 30 km2 (11 mi2) is related to several earlier volcanic centers
but not necessarily to economic mineralization. The precious metal mineralization at
Bodie is closely related to the structurally controlled emplacement of late stage 9.2-8.0
my old silicic (dacite-rhyolite) plugs late in the Silver Hill Volcanic Series (Chesterman
and others, 1986; John and others, 2011). The largest plugs of Bodie Bluff and Standard
Hill are the host rock for mineralized quartz veins that account for more than 90 percent
of the total production of gold and silver. Equivalent plugs at Red Cloud Mine, Queen
Bee Hill, and Sugarloaf to the south are smaller but have fewer quartz veins. The major
NNE striking Moyle Footwall and the Standard Vein faults form a graben between Bodie
Bluff and Standard Hill and were the principal focus of fluid circulation, intense
alteration and mineralization. Later vein deposition occurred on other pre-mineralization
faults. Post-mineralization ENE and WNW faults (Mono and Tioga) offset many ore
veins (Chesterman and others, 1986).
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Hydrothermal alteration is widespread in the district. For reference the general economic
geology terms for low to high-grade mineralization are:
• Propylitic – chlorite, epidote, various clay minerals, albite, pyrite, and minor
quartz (usually characterized as greenish in color)
• Argillic - montmorillonite, illite, sericite, quartz, and pyrite (progression to
micaceous mineralization)
• Potassic – adularia, quartz, sericite, (progressive silicification)
These assemblages are analogous to the sequence of hydrothermal alteration minerals
that are commonly logged at progressively greater depths and higher temperatures in
geothermal wells and would be characteristic of an alteration “cap” around many
developed geothermal systems. The margins of the Bodie district are characterized by
initial propylitic (greenish) alteration followed by later by more pervasive light-colored
argillic and potassic alteration sequences. Potassically altered rocks are lighter colored,
and tend to resemble the original rock in color and texture. Potassic alteration is
accompanied by extensive and pervasive development of irregular quartz-adularia veins
with or without calcite. Silicic alteration generally occurred last at Bodie and the silicaaltered rocks are light colored, hard, and form a cap on argillicly-altered rocks.
The paragensis or progression of alteration and mineralization is very important in
mineral exploration and no less significant for identifying and evaluating a potential
geothermal system. Age dating and the progression of mineralization at Bodie provide
important constraints on the viability of active geothermal systems. Focused intense
alteration occurs within the larger area of regional lower grade argillic alteration.
Recent 40Ar/39Ar dating of adularia in the gold-bearing quartz veins and in surrounding
hydrothermally altered rocks indicates that alteration and mineralization occurred over a
period of ~1.5 my and was directly related to younger siliceous dacite intrusives at Bodie
Bluff and Standard Hill (John and others, 2011). Epithermal mineralization accompanied
repeated normal faulting resulting in vein development along the fault planes. Individual
mineralized structures commonly formed along new fracture planes during separate
repeated fault movements with resulting broad zones of veinlets in the walls of the major
veined faults.
For more than a century geologists have recognized that epithermal precious metal
deposits are actually fossil hot spring systems. Bodie and many other well-studied
epithermal mineral deposits are important analogs for exploring and evaluating current
igneous-related geothermal systems. Vigorous hydrothermal systems are related to longlived, repeatedly intruded, tectonically active settings and share common characteristics:
•
Heat sources
Young shallow intrusions are fundamentally important for conventional igneousrelated geothermal systems. Highly siliceous systems are priority exploration
targets because they are part of a larger differentiating intrusive system at depth
and siliceous intrusions are emplaced at relatively shallow levels in the crust
resulting in high heat flow and fluid circulation at economically accessible depths.
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The epithermal system at Bodie is an example of a long-term series of regional
andesitic systems culminating with the emplacement of more siliceous daciterhyolite intrusions at Bodie Bluff- and Standard Hill that resulted in a vigorous
focused hydrothermal system. Shallow, easily accessible geothermal systems
significantly enhance the economics of geothermal development and epithermal
systems related to shallow silicic intrusions like Bodie generally occur within
~ 1km of the surface.
•
Timing and duration
Hydrothermal systems are ephemeral and short-lived. The pervasive permeability
that is critical for a productive geothermal system is also an efficient means of
convecting heat away from a crustal heat source. Simple conductive heat flow
modeling for the Long Valley Caldera suggests that a 2 my history of eruptions
from an upper crustal magma chamber could only be sustained by repeated
magma replenishment (Lachenbruch, and others 1976). Numerical modeling
suggests that an intrusion emplaced in a crustal setting with a minimal amount of
convective heat loss would remain a viable heat source for only 800, 000 yrs.
(Cathles and others, 1997). More precise 40Ar/39Ar dates for well studied mineral
systems suggest a narrow range of ages for single intrusion systems (Marsh, 1995;
attached) and some short-lived systems are younger than the lowest 10,000 yr
limit of current 40Ar/39Ar dating techniques of (Stein and Cathles, 1997). Still,
geologic evidence indicates many geothermal systems are the current or youngest
phase of much longer-lived hydrothermal systems. The current Long Valley
geothermal system succeeds an older system that peaked ~300,000 yrs. ago. The
Steamboat geothermal system has been active over a span of nearly 3 my. The
Geysers geothermal system has apparently been active for more that 1.5 my. The
implication is that repeated replenishment from multiple intrusive events is
important for the development of ore deposits and active geothermal systems.
•
Structural control
Permeability is a critical element for any viable geothermal system. Based on
detailed studies, epithermal mineral deposits and current active hydrothermal
systems are directly controlled by faults, fault intersections, fault complexities
(transfer and accommodation zones) and repeated fault movement creating
permeability to allow circulation of large volumes of fluid in otherwise
impermeable crystalline rocks. The tectonic setting and therefore the style of
faulting is also a critical element. Transtensional and dilatant settings provide the
vertical permeability that is often crucial in allowing vertical convection of
geothermal fluids. Bodie is a prime example of how pre-existing or repeated
complex faulting focused the emplacement of shallow crustal heat sources and
generated permeable pathways for hydrothermal circulation that resulted in vein
development along fault planes.
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Geothermal exploration is the search for heat. Drilling is the only way of directly sensing
heat but it is generally a much more advanced, intrusive and expensive part of an
exploration program. Geology, geochemistry and geophysics are the core geothermal
exploration methods before any holes are drilled but they evaluate the effects of heat.
Interpreting geological, geochemical or geophysical data depends on drawing
comparative analogies with known settings to develop a conceptual model of a potential
resource. Unfortunately, most geothermal systems are only partially explored or poorly
understood and none are exposed in three dimensions like a mine. The best exploration
tool is a well-constrained conceptual model. The best quantified model constraints come
from detailed evaluations of epithermal systems because ore bodies provide a chance to
examine what a geothermal system looked like before, during and after it evolved.
The Bodie area remains a potential Walker Lane exploration target. The lode mines
operated productively between 1877 and 1942 but the lower-grade disseminated goldsilver part of the deposit continues to attract attention. The main geothermal attraction is
the renewed post 5 my volcanism and rapid extension. The nearest (very) young eruptive
enter is Aurora Crater, a 250,000 yr. old basaltic cinder cone on the southwest side of the
town of Aurora. Gradient drilling partially compiled from mineral exploration holes
suggest heat flow above Basin and Range background of 1.7-2.8°F/100 (31-51 °C/km).
Precious Metal Mineralization
Gold and silver production at Bodie
came from the pervasive mineralization
of the several systems of quartz veins in
and around the intrusive dacite plugs of
Bodie Bluff-Standard Hill. Including
Silver Hill area farther south, gold and
silver were recovered from quartz veins
and irregular silicified zones in dacite
flows and tuff breccia. Ore minerals
include native Au, pyrargyrite
(Ag 3 SbS 3 ), tetrahedrite (Cu,Fe,Zn,
Figure 25. Standard Mill ~1932
Ag) 12 Sb 4 S 13 stephanite (Ag 5 SbS 4 ), and
pyrite. The ratio of Au to Ag by weight
in the northern part of the district was about 1:12, and in the southern part about 1:40.
(Chesterman and others, 1986; Hollister and Silverman, 1994).
The total value of production from Bodie through 1942, including the boom years 18761888, was $34m (Chesterman and others, 1986) or roughly one year of peak Comstock
production valued at $36m (Smith, 1943). The total documented production of 1.458 M
oz. Au and 7.28 M oz. Ag combined would yield roughly $2.5B at 2012 prices
($1500/oz. Au and $27/oz. Ag). Fantastic high-grade production early in the district’s
history yielded incredible financial rewards. The Syndicate mill crushed 1,000 tons of
bonanza ore from the Bodie Mine in 1878, yielding $601,103 in six weeks (Smith, 1943).
An average 100-pound bar of the district’s bullion contained about 73 pounds Ag and 27
pounds of Au. The term “Bodie gold” assumes a certain silver content. Bodie’s bullion
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was very pale in color because of the high silver content that decreased the value because
silver is less valuable than gold. Early Nevada placer miners at Washoe noted their gold
flakes and nuggets were lighter in color than the California gold they were accustomed to
panning. Dan De Quille (William Wright) in The Big Bonanza, his 1876 account of the
Comstock Lode, explained that: “The gold dug in the placer-mines of California is worth
from $16 to $19 per ounce, whereas the gold taken from the croppings of the Comstock
was worth no more than $11 or $12 per ounce.” Smith (1943) recognized the effect
noting: “All of the gold in the (Bodie) district was heavily alloyed with silver, and was
rarely worth more than $12 per ounce. In some of the shallow placer diggings, at the
south end of the ridge, the gold was worth only from $3 to $8 per ounce.”
Silver/gold alloy with was also the principal source of silver at Bodie but disseminated
silver mineralization was a secondary source. White quartz veins were variably
discolored by sulfides that 19th century miners referred to as “sulphurets”; an old mining
term for Au/Ag bearing sulfide ore. Sulphurets at Bodie included argentite (Ag 2 S;
known on the Comstock as the “blasted blue stuff”), chalcopyrite (CuFeS 2 ), galena
(PbS), pyrargyrite (Ag 3 SbS 3 ; a crimson mineral called “ruby silver”), pyrite (FeS 2 ; fool’s
gold), stephanite (Ag 5 SbS 4 ) and tetrahedrite (Cu,Fe,Zn,Ag) 12 Sb 4 S 13 ). Sulphurets
required specialized treatment beyond the milling and ore processing methods of the
time. Microscopic particles of free gold often represented the primary value of sulfide
ores. Scarcer forms of silver were also found at Bodie. Horn silver (AgCl) was especially
valued because it could be easily processed and recovered by ordinary milling. Rare
metallic or native Ag occurred deep underground beyond near-surface oxidation. Openpit mining has been considered at Bodie when precious metal prices are high. Drilling by
Galactic Resources during the 1980’s found economically viable veinlet swarms around
the lode orebodies. The USGS (Long and others,1998) estimated the disseminated Bodie
resource at 24.6 million tons averaging 0.07oz Au/ton and 0.42oz Ag/ton. Galactic
shelved the project in the late 1980’s when their open-pit mining proposal proved bitterly
controversial and precious metal prices declined. Cleanup liabilities at Summitville,
Colorado eventually forced Galactic into bankruptcy and the Desert Protection Act of 1994
withdrew 5,000 acres of existing Bodie claims appraised at $65 million. California State
Parks acquired the acreage in 1997 for $6 million. (Wilkerson and others, 2007).
History
Bodie was one of the roughest, toughest, rip-roaringest, rootin-tootinest mining camps in
the west. The town’s name is supposedly a misspelling of the last name of William (a.k.a.
Waterman or Wakeman) S. Bodey who discovered placer gold in1859 near what is now
called Bodie Bluff. Bodey died in a blizzard the following November while making a
supply trip to Monoville and never saw the development of the town that supposedly
bears his name. The discovery of gold and silver at Bodie coincided with other
discoveries at nearby Aurora, Nevada and the Comstock Lode in Virginia City, Nevada.
While those towns boomed, interest in Bodie remained lackluster. By 1868 only two
companies had built stamp mills at Bodie, and both had failed. A collapse at the Bunker
Hill mine in 1875 exposed a rich quartz lode vein on Bodie Bluff that quickly attracted
speculators and investors who eventually purchased the claim and organized the Standard
Company. By 1880 the town had an estimated population of 10,000 people bustling with
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robbers, miners, storekeepers, gunfighters and prostitutes. Town records list 65 saloons
stretching for a mile along Main Street that included or were part of gambling halls,
opium dens and houses of ill repute. Rosa May, a local madam, is credited with giving
life-saving care to many during a serious epidemic at the height of the town’s boom but,
like many soiled doves of the day, she was not regarded as a “respectable woman” and is
buried outside the town cemetery fence.
Legend and fiction often masqueraded as fact in the 19th century Wild West and at least
some of Bodie’s notoriety was based on exaggeration. Embellishing the potential value of
a mineral deposit was an integral part of soothing speculators’ nerves and assuring
continued investment. A notorious town legend held that a man was killed every day of
the week in Bodie. That might be an exaggeration from a particularly violent week but,
based on court records for murder convictions, the homicide rate was at least 30 times the
current US average (Roth, 2010). An unattributed quote in one local newspaper
commented that, “Bodie is becoming a summer resort—no one killed here last week.”
Presumably, the undertaker never had a slow day. Mark Twain tried his hand at mining in
nearby Aurora and got his first newspaper job writing for the Esmeralda Star. He
chronicled life in the Esmeralda Mining District for the Virginia City Territorial
Enterprise and in his book Roughing It. Twain said that he never had occasion to kill
anybody with the Colt Navy revolver he carried, but he had "…worn the thing in
deference to popular sentiment, and in order that I might not, by its absence, be
offensively conspicuous, and a subject of remark." One story (sometimes wrongly
attributed to Twain) is about a little girl whose family was moving to Bodie and wrote in
her diary "Goodbye God, I'm going to Bodie." In one of the finest of early spin jobs, the
ever-enterprising local newspaper editor “Lying” Jim Thompson denied the slight
claiming that the girl was misquoted or the sentence mis-punctuated and that she had
actually said “Good! By God I’m going to Bodie.”
Clean out your pockets before leaving. Bodie is a historic park and removing any artifact
is strictly prohibited. More importantly, among the town’s many ghost stories and
legends of haunting the most infamous is the Bodie Curse. Anyone who removes
something from Bodie is cursed with bad luck and misfortune until the item or items are
returned. Every year park rangers receive objects mailed back to Bodie with letters
(preserved in the Museum) apologizing to the town and imploring its guardian spirits to
forgive their offense. If you have carried anything off with you, this is the first of your
bad luck. We can’t risk riding with you and you’ll have to find our own way back.
Return west on Highway 270 to Highway 395.Turn right (north) heading toward Reno.
BEWARE OF TRAFFIC! Continue north through Bridgeport, the county seat of Mono
County. Continue north of Highway 395 to the UNR campus.
Incremental Distance – 137 mi/ 220 km
Cumulative distance – 194.5 mi. / 343.8 km
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Figure 26. Generalized geologic map
of the Bodie Hills (after John and
others, 2011)
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