In a reconnaissance study of Owens Valley, Lee (1906) described "dark colored ... andesitic tuffs and breccias" along the Owens River. The Bishop Tuff was first investigated (and named) by Gilbert (1938), who not only recognized the deposit as the product of pyroclastic flows (similar to nuée ardentes), but also described the post-emplacement features produced by welding and secondary crystallization. 

Rinehart and Ross (1957) mapped the Bishop Tuff within the Casa Diablo Mountain Quadrangle. Putnam (1960) discussed the evolution of the Owens River and Rock Creek Gorges. Sheridan (1965) described the mineralogy and development of the ash-flow zones. Dalrymple et al. (1965) established a 0.71 Ma potassium-argon age of the Bishop Tuff. Bateman (1965) mapped Bishop Tuff of the southern Volcanic Tableland and discussed its structural and erosional modifications. Sheridan (1970) reported "fossil" fumaroles and relationships to ash-flow zones. Crowder and Sheridan (1972) mapped ash-flow sub-units of the Bishop Tuff in the White Mountain Peak Quadrangle. Ragan and Sheridan (1972) analyzed the compaction by welding of the tuff. Bailey et al. (1976) presented the relationship of the Bishop Tuff to regional volcanism and Long Valley caldera. Hildreth (1979) showed that the Bishop Tuff successively tapped deeper, less evolved levels of a zoned magma chamber, and identified several flow "lobes". Horn and Sheridan (1983) calculated pyroclastic flow emplacement temperatures modelled from density profiles. Hildreth and Mahood (1986) analyzed the lithic inclusions to infer an eruption sequence and possible source vents of the Bishop Tuff. 

 

Regional Geologic Setting

The Bishop Tuff is located in a geologically young and historically active area of volcanism and tectonism. This area in eastern California, adjacent to Nevada (Fig. 1), lies at the west edge of the Basin and Range Province. Extensional tectonics created topography that not only controlled the paths of pyroclastic flows that deposited the Bishop Tuff, but also faulted and warped the ignimbrite after it had cooled. 

Pre-Tertiary rocks of the region include folded late Pre-Cambrian or early Cambrian to Triassic sedimentary and Mesozoic volcanic rocks of the White Mountains. Metamorphosed equivalents of the earlier sedimentary sequence are preserved as roof pendants and isolated bodies in the Sierra Nevada. In Mesozoic time the area was intruded by felsic quartz-bearing granitic plutons, ranging from granodiorite to alaskite (Bateman 1965), of the Sierra Nevada batholith, the root of a magmatic arc. The batholith was uplifted in Tertiary time and stripped of its cover, producing a range with relief of a few thousand feet. Final uplift was achieved before the onset of Pliocene glaciation (Christensen 1966).

By approximately 3 Ma the White Mountains and Sierra Nevada stood high over the intervening Owens Valley. Long Valley formed an embayment into the north-south trend of the Sierra Nevada, an "encroachment of Basin and Range extension into thick and stable Sierra Nevada crust" (Hill et al. 1985). A magma body below Long Valley would be the source of later volcanism.

Late-Tertiary volcanism was largely basaltic, but also included andesitic and rhyodacitic compositions. Though not direct products from the Long Valley magma chamber, these eruptions represent early differentiation of a mafic parent (Bailey 1987). The most extensive exposures of basalt, a lava plain covering some 770 km2, are north of Adobe Valley (Gilbert et al. 1968). In the Owens River Gorge basaltic lavas and cinder layers blanket a granitic basement ridge. The basalts had flowed upon an eroded granitic terrain of moderate relief.

The rhyolite complex of Glass Mountain represents the first volcanic products from the Long Valley magma chamber (Fig. 1). Two periods of eruption, 2.13 to 1.20 Ma and 1.10 to 0.79 Ma (Metz and Mahood 1985), produced more than 15 km3 of domes, lava flows, breccias, pumice, and ash. High-silica rhyolite compositions, 76.8 to 77.9% SiO2 (Metz and Mahood 1985), include black crystal-poor obsidian, pink and grey lava, and white pumice. The eruptions were not sufficient to trigger caldera collapse (Metz and Mahood 1985).

Pleistocene to Recent age glacial deposits are found in the vicinity of the Sierra Nevada. Pleistocene Sherwin Till forms a thick deposit near the head of Rock Creek Gorge, where it overlies the granitic and metamorphic basement. It also occurs below the Bishop Tuff, as observed in the roadcut of US 395 known as the "Big Pumice Cut" (Sharp 1968) and in a tunnel excavated for the Los Angeles Department of Water and Power aqueduct (Putnam 1960). Weathering of granitic boulders in the till indicates a gap of tens of thousands, but less than 100,000 years, between the glaciation and the Bishop Tuff eruption (Sharp 1968). Sherwin deposits also are found between the basalts and the Bishop Tuff in the Owens River Gorge. The sandy, sorted matrix indicates that these are outwash deposits rather than till (Sharp 1968).

The Bishop Tuff was deposited more than 700,000 years ago, starting with a Plinian eruption followed by a series of pyroclastic flows from multiple source vents within Long Valley. The caldera, a 17 by 32 km elliptical structure, formed as the roof foundered above and collapsed into a partially depleted magma chamber.

Following the collapse of the caldera, Long Valley was filled by a Pleistocene lake (Mayo 1934). The waters reached a maximum elevation indicated by gravel capped terraces east of the caldera, and was emptied by 50,000 to 100,000 years B.P. (Bailey 1987)†. The evacuation of Long Valley Lake caused or contributed to the incision of Owens River Gorge (245 m deep) and Rock Creek Gorge (185 m deep) into the plateau of Bishop Tuff to the southeast. The mechanism by which lake drained has not been resolved; it has been attributed to both headward growth of the Owens River (Mayo 1934; Putnam 1960) and lake overflow (Rinehart and Ross 1957; Bailey et al. 1976).

Following caldera collapse volcanism continued in Long Valley with smaller volume rhyolite eruptions. Resurgent doming of the western part of the caldera occurred 40,000 to 100,000 years after the caldera formed (Bailey et al. 1976). Later activity included rhyolitic and rhyodacitic magmas (including the 0.18 Ma Mammoth Mountain complex), and basalts of the Devil's Postpile and June Lake area (Bailey et al. 1976). Holocene eruptions of the Inyo and Mono chain of craters and domes occur in an arc that intersects the northwest margin of Long Valley caldera, possibly the initiation of a new caldera cycle.



Welding and Crystallization Zones in Ash-Flow Tuffs

Pyroclastic flows originate from explosive fragmentation of volatile-rich magma. These gravity driven flows owe their mobility to rapidly expanding gases, and are capable of traveling hundreds of kilometers. The height at which the flows originate is probably the dominant factor controlling flow distance (Malin and Sheridan 1982). Once emplaced, cooling and compaction of the hot deposits can alter an initial homogeneous mass of glass, ash, and crystals into a layered unit of rock indistinguishable from one deposited as a lava flow. Such textural changes might be considered metamorphic (Sheridan 1965). Smith (1960a, 1960b) and Ross and Smith (1961), based on years of field observations, outlined the basis for modern concepts of the processes which produce zonation in welded tuffs. Riehle (1973) mathematically modelled the cooling and compaction of rhyolitic ash-flows.

Tuff is an excellent insulator, retaining heat in thick deposits that compact, or weld, under its own weight or overburden. The high temperatures also promote secondary crystallization within the sheet-like deposits. The ultimate product is a layered formation with varying degrees of compaction and crystallization. The Bishop Tuff (Fig. 3) displays all of the ash-flow zones described by Smith (1960b).



Zones of Welding. Welding of ash-flow tuffs is the cohesion of the hot glassy fragments, ranging from simple sticking together to plastic deformation that effectively reduces loose ash and pumice to a rock of nearly zero porosity (Smith 1960a). While a number of variables influence welding, the observed textural changes can be interpreted in terms of two: emplacement temperature and thickness of the unit (Smith 1960a). A thin sheet, despite a small load, may weld if it is emplaced at a high temperature. A thick sheet, however, may weld at lower temperatures due to insulation and greater load stress. Smith's (1960b) diagrams of zonation in sheets of various thickness and emplacement temperatures have served as excellent models for observed ignimbrites.

Compaction by uniaxial load stress is the mechanism of welding (Ragan and Sheridan 1972). The "Y"-shaped glass shards are compressed, aligned in the plane perpendicular to the load stress, and may bend around more rigid objects such as crystals or lithic inclusions. Carr (1981) used scanning electron microscopy to observe incipient welding between shards as lines of imperfections that develop near dust and gas at the juncture. Increasing degrees of welding reduces porosity, producing a denser rock. The pumice lapilli flatten and darken in color, their vesicles are squeeze closed, and ultimately reduce to obsidian-like lenses or "fiamme".

Three welding zones are described; zones of dense welding, partial welding, and no welding (Smith 1960b). The idealized cooling of a single ash-flow contains a central densely welded zone that grades to partly welded and non-welded zones toward the top and bottom surfaces and at distal margins. A "cooling unit" (Smith 1960b), represents one or more pyroclastic flows that cool as a package, with the zonation pattern just described. Smith  (1960b) demonstrated how the zones pinch out with distance from the source, as functions of both decreasing thickness and temperature. Typically, zone boundaries are gradational, and not all zones are always present.

The non-welded zone most closely represents the original texture of the deposit. Incipient welding may be present but only recognizable microscopically. The density ranges from that of pumice (0.8 g/cm3) to weakly indurated tuff (1.4 g/cm3), greater than 40% porosity assuming bulk density of 2.45 g/cm3 (Ragan and Sheridan 1972). An entire ash-flow may be non-welded, or the zone may only be found at the top and bottom of the sheet and its thin margins. Generally unconsolidated and subject to rapid erosion, the non-welded zone is usually only found in young ash-flow tuffs. The densely welded zone is defined by porosities of less than 10% (densities from 2.2 to 2.45 g/cm3) and represents compaction of unconsolidated material by a factor of greater than two (Ragan and Sheridan 1972). Its position, the hottest part of the sheet, is displaced below the center due to the greater rate of heat loss at its upper surface (exposed to the atmosphere) relative to the heat loss at its lower surface. The partly welded zone is transitional from non-welded tuff to densely welded tuff with a range of densities from 1.4 to 2.2 g/cm3.