california

Daily Logs for Field Trip to Northern California

June 16 to June 22, 2003: The Coast Range and the Sierra Nevada



The trip will cover a portion of the Franciscan terrane north of San Francisco for three days, followed by a

three day transect across the Sierra Nevada to Mono Lake and Mono Craters. On the trip we will examine

classical sections through a continental margin, ranging from ophiolites and turbidites to blueschists and

eclogites. We will cross the world’s most famous transform fault zone, then the Great Valley to Yosemite.

We will see exposures of Sierran granite batholiths and their contact metamorphic rocks, glacial moraines,

Mono Lake tufas, Mono Craters, Bishop Tuff and Long Valley Caldera. The logs include map numbers

keyed to the Northern California Atlas and Gazetteer.



Day 1: Monday, June 16



Fly out of Detroit at 7 AM, arrive in Bay Area at Oakland Airport at 1 PM, pick up SUVs. [Map 104] Drive to

Tiburon Peninsula. Drive to I880 from airport, go N on I 880 to I 580 W across the San Rafael Bridge (ca.

30 mi.). Go past San Quentin (watch out for quail), turn S on Rt. 101, take second exit S on Paradise Dr.

(2 mi.) Go past turnoff to Ring Mountain Preserve, turn left on Trestle Glen Drive, left on Tiburon Blvd,

and left again on Reed Ranch Dr. (2 mi.) Stop at small parking lot (1 mi.).



Stop 1. Take trail toward Turtle Rock to see blocks of eclogite, blueschist and metagreywacke. Tiburon

Peninsula is the type locality of lawsonite, CaAl2Si2O7(OH)2.H2O, which was found as veins in blueschists

(Ransome, 1895). Lawsonite is stable above about 3 kbar and decomposes by a variety of reactions to

form clinozoisite above about 400°C at 5 kbar (Perkins et al., 1980); it is stable up to about 7 GPa along a

cold geotherm and remains a significant consituent of a cold subducted basalt to ca. 200 km depth (Poli &

Schmidt, 1997). Find store and shop for 1-2 days dinner, breakfast and lunch. Return to Tiburon Blvd.,

take it W and enter on S exit of US 101. If time permits, continue S to southern Sausalito exit (Alexander

Ave. exit, go W. under freeway on last exit before Golden Gate Bridge). Head W into Fort Baker. After

about 2 mi. turn SW onto Conzelman Rd. and drive to outcrops by Battery 129 (4 mi.)



Stop 2. Kink folds of Franciscan chert and some interlayered basalts and diabases. Radiolaria from

these cherts are Late Jurassic, with adjacent cherts to the E yielding Early Cretaceous microfossils

(Murchey & Jones, 1984). Then return to US 101 N and exit on Rt. 1 (8 mi.). If we are running late, skip

Stop 2, go S on US 101, exit on Rt. 1 (2 mi.). Drive up SR 1, following signs to Muir Woods (6 mi.).



Stop 3. Take hike along creek (no hammers). The poorly exposed rocks are Franciscan.

Return to cars, drive SW on Muir Woods Rd. to SR 1. Turn right N on SR 1 (4 mi.).

{Map 103] Note wave cut terraces and marine stacks of Franciscan rocks along SR 1 after the ocean

comes into view between Muir Beach and Stinson Beach.

SR 1 is in a valley along the San Andreas fault zone NW of Bolinas Bay. In this valley two streams, Pine

Gulch Creek on the W side (heading SE), and Olema Creek on the E (heading NW), run parallel to each

other but drain in opposite directions. Drive past Olema. The epicenter of the 1906 San Francisco

earthquake was located just N of here.

[Map 93] Turn W at Point Reyes Station on Sir Francis Drake Blvd. (26 mi.). Stop where the road leaves

Tomales Bay (ca. 6 mi.).



Stop 4. Road cut of granitic rocks on north side of road.

Head to Point Reyes (30 mi.). [Map 103] We will stay at the boathouse near the lighthouse (10 mi).



Stop 5. Examine granites and unconformity capped by Miocene conglomerates The granites are part of

the Salinan block extending from Santa Barbara to Bodega Bay. What are likely source rocks for most of

the cobbles? Total distance for day: 120 mi.

Stop 6. Visit the lighthouse.

Day 2: Tuesday, June 17

[Map 103] Leave lighthouse, returning NE on Sir Francis Drake Blvd. [Map 93]. Take turnoff SE toward

Drakes Beach (15 mi.).



Stop 1. Walk along shore to examine wave cut cliffs. The gently folded and faulted fossiliferous

sediments are of Pliocene to Miocene in age. Drakes Bay is one of several proposed localities N of San

Francisco where Sir Francis Drake beached his ship the Golden Hind in 1579. He named the place Nova

Albion for its white cliffs. Old ceramic pieces have been reported here, and the Bolinas Breastplate was

found about 70 years ago on Tomales Bay just N of Olema. The composition of the metal suggests it is

English and probably from the 16th Century (Appendix I).

Return to cars and head N to Sir Francis Drake Blvd back to SR 1 (25 mi.). Turn left and drive N on SR 1

to Valley Ford. (20 mi.)



Stop 2. Valley Ford is a classical eclogite, blueschist and amphibolite locality (Bloxam, 1961). High-grade

blocks litter contact of Franciscan greywackes with unconformably overlying Miocene sediments. The

Franciscan greywackes are jadeitized N of a mapped fault.

Continue N on SR 1 past Bodega Bay (5 mi.). [Map 92] Drive through Jenner (10 mi.) and turn E on SR

116 along the N side of the Russian River. Why does the Russian River have its name? Continue E to

Guerneville (12 mi.) and take Armstrong Woods Rd. N to Austin Creek State Recreation Area (3 mi.).



Stop 3. Reserve a campsite at Bullfrog Pond State CG (first come, first served, $12/vehicle + $2 entrance

fee, no showers, 707-869-2015). Lunch at the adjacent Armstrong Redwoods State Park. Take a walk

through the redwoods.

Return to Guerneville and Jenner on SR 116 (15 mi.). Proceed south on SR 1 to Goat Rock Beach (5 mi.)



Stop 4. Examine serpentinite, mélange and greenstone at Goat Rock. What is a mélange? Many high-

grade knockers (tectonic blocks) have a close association with serpentine bodies, which usually mark

major faults within and bounding the Franciscan terrane. The serpentinite and greenstone probably

represent disrupted portions of an ophiolite, and the knockers are from an older subduction complex (ca.

150-165 Ma) that are tectonically introduced along faults and in serpentinites.

Return to Jenner where we may need to stop and shop for dinner (5 mi.).



Stop 5. Examine eclogite and blueschist blocks on Jenner beach (Krogh et al., 1994). Do any of these

rocks have actinolite sheaths? What relationships are indicated among the lithologies?

Drive road up open slope of hill behind Jenner (2 mi.).



Stop 6. Examine jadeitized greywacke blocks. Crawford (1965) made a map of all the blocks.

Drive N. of Jenner on SR 1 (2 mi.).



Stop 7. Eclogite/blueschist occurrence on road if still there. Observe eclogite – blueschist transition.

Continue N to Fort Ross Historical Park (13 mi.) and stop at if time permits ($2 fee). Continue N on SR 1

past Fort Ross. [Map 82] Take Ft. Ross Rd. E inland to Cazadero area (15 mi.).



Stop 8. Walk Ward Creek N of Cazadero to see metasediments, metavolcanic rocks (some of which are

eclogites: Shibakusa & Maekawa, 1997), blueschists, thin metacherts, as well as eclogite knockers

(Coleman & Lee, 1963). Thin marble layers and veins have aragonite partially inverted to calcite (Coleman

& Lee, 1962). The Franciscan is one of the few blueschist terranes with preserved aragonite, which

backreacts rapidly if retrogressed at > 180°C (Carlson, 1980).

Continue on Ft. Ross Rd. E to Cazadero Rd, turn right and take Cazadero Rd. SE to SR 116 (10 mi.), turn

left and drive through Guerneville (2 mi.). Return to Bullfrog Pond CG (3 mi.). Total distance for day: ca.

160 mi.

Day 3: Wednesday, June 18



[Map 82] Return to Russian River Rd. turn left (E), head to Healdsburg on Westside Rd. (N of river).

Turn left (W) on Mill Cr. Rd. (20 mi.) to Junction School (1 mi.).

Stop 1. Junction School eclogite body (Borg, 1956).

Return to Westside Rd., continue into Healdsburg (3 mi.). Gas up and shop for dinner if needed at

Healdsburg. This area is part of the Sonoma Valley, which together with the Napa Valley to the SE,

produce much of the best California wine. Drive N on US 101, turn right on Alexander Valley Rd. to SR

128 toward Calistoga [Map 84], continue SE to St. Helena, [Map 94], turn NE on SR 128 at Rutherford

(19 mi.), follow SR 128 to turnoff to left with Berryessa-Knoxville Rd. [Map 84] (10 mi.).



Stop 2. Metabasalts and rodingites in Capell Creek. Rodingites are Ca-metasomatites that usually form

on basalt dikes that cut serpentinites. The Ca metasomatism forms Ca-rich minerals such as hydrogarnet

and prehnite, and here pumpellyite as well. The metasomatism occurs during serpentinization of augite

and alteration of plagioclase, releasing Ca. The basalt loses Si to the serpentinite, and the rodingite may

only have 20-30 wt % SiO2.

Return S to SR 128, turn left on SR 128, head to Monticello Dam. Continue past dam to high outcrops on

the SW side of the road (14 mi.).



Stop 3. Examine the folded Lower Cretaceous Knoxville Fm. of the Great Valley sequence on outcrops

below dam along Putah Creek. Have lunch.

[Map 85] Take SR 128 east to Winters (5 mi), pass over I 805, pick up CO E6 E to Davis (20 mi.). Take

SR 113 S (1 mi.) to I 80 E toward Sacramento (7 mi.), [Map 86] take US 50 E past I 5 and enter US 99 S

(3 mi.) [Maps 96-97] past Stockton [Map 107] and Modesto to Merced (117 mi) [Map 118], and take SR

140 E. [Map 119] Stop six miles E of the village of Catheys Valley.



Stop 4. Road cuts of the Jurassic Mariposa Slate. Note relationship of slaty cleavage to bedding

features. The Mariposa terrane is a marine shale that was intruded further E by Sierran granites.

[Map 120] Continue E to Mariposa (37 mi.). Check status of gas. At 11.6 mi. E of the intersection of SR

49 and SR 140 is small bridge over Bear Creek continue 0.9 mi. N.



Stop 5. Pullout to examine road cuts. The rocks are greenstones in contact with slaty siltstones of the

Paleozoic Calaveras terrane. These are ocean floor deposits that formed in a convergent zone.

[Map 110] Continue E, 8 mi. past Briceburg, stop at sign “Oldest Rocks of the Yosemite Region” on west

side of highway away from Merced River, here flowing N.



Stop 6. Complexly folded Calaveras sediments.

Continue E to El Portal and stop at railroad exhibit on N side of road. (29 mi.)



Stop 7. Examine exposures of Calaveras terrane showing garnet-bearing marble contact

metamorphosed by granites outcropping to E.

Continue E. to Arch Rock Entrance Station of Yosemite Park ($20/vehicle fee). Note V-shaped Merced

Gorge even though it was glaciated. Proceed N on SR41 to Crane Flat and pick up SR 120 W (10 mi.).

Go W 1 mi beyond park boundary to Evergreen Rd. Turn right on Evergreen Rd. and proceed to Dimond

O CG (1 mi.). Site 33 and group site 34/35 are reserved ($60, NFS CG, reservations for 19 people).

Total distance for the day: 300 mi.

Day 4: Thursday, June 19



[Map 110] Return to SR 120 E and head S on SR 41 to Yosemite Park (cf. Appendix II for a summary of

the geology of the park). Stop at Tunnel View (mi.)



Stop 1. Views of El Capitan, Halfdome, Bridalveil Fall and its hanging valley. What causes hanging

valleys?

Continue S on SR 41, turn left at Chinquipin Junction up Glacier Point Rd. to Glacier Point (23 mi.).



Stop 2. Spectacular views of the Valley, Vernal and Nevada Falls, and High Sierras.

Return three miles to down road.



Stop 3. walk out to septum on way to Cathedral.



Stop 4. Pegmatites are exposed on the glacial pavement on trail to xx

Return to Yosemite Valley on SR 41, head to Visitor Center (30 mi.).



Stop 5. Yosemite Visitor Center. Examine geology displays.

Return to cars, return back down valley to road sign V-7



Stop 6. Stop at base of El Capitan. Walk 1/4 mi. toward base of El Cap to view relationships among the

diorite bodies, pegmatite and aplite dikes, and the host El Capitan granite.

Return to cars, continue W to road sign V-11.



Stop 7. Valley View. Final view of Yosemite Valley

Take SR 41 N, and turn E. on SR 120. Gas up. Drive toward Tioga Pass on SR 120, and stop 8.9 miles

east of intersection with SR 41 1.2 mi. E. of small stream, South Fork of the Tuolumne River.



Stop 8. Pegmatite dike. Examine dike, note wide variation of grain size and mafic inclusions in El

Capitan granite. Are they cognate inclusions, entrained mafic magma or xenoliths?

Continue on SR 120 E, note exfoliation domes and glacial features along road; stop at Tenaya Lake.



Stop 9. Take short walk north of highway across from entrance to walk-in campground to examine glacial

features in more detail.

Continue to Tioga Pass Entrance Station.



Stop 10. Tioga Pass Entrance Station. Walk east from station for a mile on unmarked trail to Dana Peak

to examine contact between metamorphic and igneous rocks. Some of us may be able to scale Dana

Peak if the snow is not too deep. The country rocks are Triassic volcanic, volcanoclastic and sedimentary

rocks that are interpreted as a caldera by Schweickert and Lahren (1993b, 1999).

Drive E off Tioga Pass. Turn N at Saddlebag Lake Rd. and drive into campground (5 mi.).



Stop 11. Examine at a distance the visible contact relations exposed at Saddlebag Lake. Depending on

time, walk down trail to lake at far western edge of campground to seen calcsilicate hornfelses. At this

locality Mine Creek discharges into the meadows from a small canyon. May be able to inspect

relationships of hornfels units on trail in canyon unless there is too much snow.

Return to cars and to SR 120 (5 mi.), proceed E to the S end of Ellery Lake and stop if time permits.



Stop 12. Examine contact relations at Ellery Lake.

Drive down Tioga Pass, viewing lateral and terminal moraines in Lee Vining Canyon.

Continue on SR 120 E [Map 112] At Lee Vining (60 mi.), take US 395 N [Map 101] and turn W on

Virginia Lake Rd. just past Conway Summit 10 mi.). A group campsite (site 40) is reserved (NFS CG,

$42, ca. 7600’, no showers, fishing) at Trumbull Lake CG just below Virginia Lakes (7 mi.). Total distance

for day: 90 mi.

Day 5: Friday, June 20



[Map 101] Return to US 395 (7 mi.), take US 395 10 mi. S, [Map 112] turn E on gravel road (Picnic

Ground Rd., first road NE off US 395) drive 0.7 mi. E along lake.

Stop 1. Tufa deposits, algal mounds formed from springs along edge of Mono Lake.

Return to US 395 and drive 6 mi. S to SR 120 E. Take 120 E a mile or so to Mono Craters.

Stop 2. Pumice and obsidian. Mono Craters volcanic rocks are eroded domes that formed in cinder

cones.

Return to US 395 and continue south to Drive S (20 mi.) past the Owens River Rd. and turn left on dirt

road to Lookout Mtn.

Stop 3. View of Long Lake caldera at Lookout Mtn. Bailey (1987) described the Plio-Pleistocene volcanic

history of the area, from which the following has been extracted. Eruptions of basalt and andesite lavas

began about 2.8 Ga ago in the Long Valley area, followed by emplacement of quartz latite and rhyolite

domes, flows and tuffs until 0.8 Ma. A catastrophic collapse of the magma chamber occurred 0.73 Ma

ago, producing the Long Valley caldera and 600 km 3 of ash falls and flows that deposited the Bishop Tuff,

and which reached up to 3000 km E to Kansas and Nebraska. The Long Valley caldera was 2-3 km deep

and 20 x 30 km across. Postcaldera rhyolite eruptions and domes formed after the collapse, partially

filling the caldera and continuing episodically up to 0.1 Ma ago. Geophysical data indicate that a partly

molten magma chamber still is present. Younger lavas formed near Mammoth Mountain and continued

until the last 40 ka, followed by eruptions in Mono Craters and Inyo Craters until the last 500 a, forming

pumice deposits and the obsidian domes.

Although the earthquake swarm of 1980 led to incautious predictions by geologists of a new eruption

in Long Valley, causing property values to plummet and geologists to become hated in the area, the recent

history suggests that future eruptions should occur in the Mono-Inyo craters area to the N and W.

Following an earthquake swarm with up to magnitude 5 and 6 in the Mammoth Lakes area in 1980, a hole

was drilled 3 km deep into Long Valley caldera by the USGS (Appendix III).

Return to US 395, turn N to Owens River Road. Turn E and head to outcrops of the Bishop Tuff

Stop 4. On way past Hot Creek Geyser, check out if swimming is possible. If so, park at Hot Creek

Geyser, take plunge in hot springs.

Return to US 395 and drive past Mammoth Mountain turnoff to Mammoth Scenic Loop Rd. at turnoff

opposite road to Lookout Mtn. (24 mi.). Drive to Inyo Craters sign (3.3 mi.), turn right and follow dirt road

to parking area for Inyo Craters

Stop 5. Inyo Craters. Take trail W to the craters. Inyo Craters comprise three phreatic craters, two of

which contain small crater lakes. The stratigraphy exposed on the N side of the southern crater is the

following: dark trachyandesite flow, cinder (2 m), glacial till (8 m), pumice (1 m), and bedded rhyolite with

blocks of volcanic, granitic and metamorphic rocks (8 m). What causes phreatic eruptions?

Drive S. from craters to SR 203, turn right and stop at outcrops at Minaret Summit.

Stop 6. Metamorphic rocks.

Stop 7. Outcrops down to Agnew Meadow.

Drive into Devils Postpile National Monument (17 mi.).

Stop 8. Take walk to basalt columns (Stop 8, Appendix IV).

Return to US 395 (17 mi.) and drive N to June Lake turnoff (15 mi.), W to June Lake CG (3 mi). W e have

walk-in sites 3, 4 and 5 reserved (NFS CG, $72, no showers, fishing, 760-647-3000). Total distance for

day: 100 mi.

Day 6: Saturday, June 21



[Map 112] Return to US 395 (3 mi.). Drive N on US 395 past Lee Vining [Map 101], Bridgepport, and

along the W Walker River [Map 91] past Topaz Lake [Map 90], where we briefly enter Nevada. Just past

Minden, take SR 88 S to SR 207 W and pick up US 50 W at W South Lake Tahoe (114 mi.). Drive on US

50 W to SR 89 [Map 89] and drive out to Emerald Bay State Park. (mi.)



Stop 1. Look at Lake Tahoe.

Return to US 50 W and head over Echo Summit. Stop at road cuts by Echo Summit (



Stop 2. Continue W. on US 50 W along the W. Fork of the American River [Map 88].



Stop 3. Furrowed conglomerate deposits – what is their origin? (

Proceed to Placerville (83 mi.). Gas up and get food for dinner. Continue W. on US 50 [Map 87] to

Sacramento (29 mi.) [Map 86] Pick up I 80 W on W side of Sacramento. Take I 80 W [Map 95] to I 680

(38 mi.), take 680 S across the Martinez Bridge past Walnut Creek (25 mi.) [Map 105] to Danville (7 mi.).

Turn E. and head into Mt. Diablo SP (5 mi.). We have a 30 person group campsite reserved (Buckeye,

BG, $47.50). Set up camp.

Drive towards the top of Mt. Diablo. Appendix V is a detailed road log of the drive.



Stop 4.



Stop 5. Views and outcrops of Franciscan rocks. The Franciscan sediments on the flanks of Mt. Diablo

are cherts and greywackes. Little is known about the metamorphism of these rocks, although further S at

Mt. Hamilton and Pacheco Pass in the Diablo Range are subregional belts of jadeite metagreywackes

(Ernst, 1971). Because both the Bay Area and the Sierran peaks are visible from this point, Mt. Diablo

was an important base for a topographic survey across the United States extending from San Francisco to

New York at the end of the 19th Century. Subsequent releveling in the 20th century using Mt. Diablo

(elevation 3,850’) was flawed because it continues uplifting at 3 mm/yr. Total distance for day: 280 mi.

Call 925-837-0904 one day prior to arrival date to confirm status of park re fire danger.



Day 7: Sunday, June 22

[Map 105] Get up early, return to Danville (5 mi.). Take I 680 S to I 580 W (9 mi), take I 580 W to I 880 (2

mi.), N on I880 to Oakland Airport (7 mi.) for flight at 11:30 AM. Total distance for day: 20 mi.

References



General field guides

Alt, D.D. & Hyndman, D.W. (1975) Roadside Geology of Northern California. Mountain Press, Missoula,

MT, 243 p.

Harbaugh, J.W. (1974) Field Guide Northern California. Kendall/Hunt, Dubuque IA, 123 p.



Franciscan rocks

Bailey, E.H. & Blake, M.C. (1969) Late Mesozoic tectonic development of western California.

Geotectonics 3, 148-154.

Bailey, E.H., Irwin, W.P. & Jones, D.L. (1964) Franciscan and related rocks, and their significance in the

geology of western California. Calif. Div. Mines Geol. Bull. 177 p.

Bloxam, T.W. (1959) Glaucophane-schists and associated rocks near Valley Ford, California. Am. J. Sci.

257, 95-112.

Borg, I.Y. (1956) Glaucophane schists and eclogites near Healdsburg, California. Geol. Soc. Am. Bull.

67, 1563-1583.

Broecker, M. & Day, H.W. (1995) Low-grade blueschist facies metamorphism of metagreywackes,

Franciscan complex, northern California. J. Metam. Geol. 13, 61-78

Brothers, R.N. & Grapes, R H. (1989) Clastic lawsonite, glaucophane, and jadeitic pyroxene in

Franciscan metagraywackes from the Diablo Range, California. Geol. Soc. Am. Bull. 101, 14-26

Cloos, M. (1985) Thermal evolution of convergent plate margins: thermal modeling and reevaluation of

isotopic Ar-ages for blueschists in the Franciscan Complex of California. Tectonics 4, 421-433.

Cloos, M. (1986) Blueschists in the Franciscan Complex of California: petrotectonic constraints on uplift

mechanisms. In Blueschists and Eclogites, Geol. Soc. Am. Mem. 164, 77-93.

Coleman, R.G. & Lee, D.E. (1962) Metamorphic aragonite in the glaucophane schists of Cazadero,

California. Am. J. Sci. 260, 577-595.

Abstract. Metamorphic aragonite is the dominant polymorph of CaCO3 within concordant and discordant

carbonate-rich lenses and veins in the glaucophane schist belt exposed along Ward Creek, Cazadero,

California. This schist sequence is part of a larger metamorphic belt within the Franciscan Formation of

Jurassic and Cretaceous age. Petrofabric study shows that the fabric element of the aragonite marbles is

compatible with the lineation of the enclosing schists. A careful chemical, optical, and X-ray study of 4

purified aragonite specimens shows that this metamorphic aragonite conforms in every respect with the

physical properties recorded from aragonite. The mineral is biaxial negative, 2V ca. 18°,  = 1.530±0.001,

 = 1.680±0.002, and  = 1.685±0.001. A maximum of 1.26 mol % SrCO3 was found by chemical analysis.

The presence of aragonite as a metamorphic mineral within these glaucophane schists suggests that the

P-T conditions for this facies may be unique. Experimental work on the inversion of calcite to aragonite

has shown that aragonite is the stable high pressure polymorph of CaCO 3; the formation of aragonite in

the glaucophane schists indicates that P greater than 4 kbar prevailed during metamorphism. Tectonic

stresses developed during the formation of the glaucophane schists may have increased the pressure

considerably above the lithostatic load pressure and promoted the stable formation of aragonite marbles.

Coleman, R.G. & Lee, D.E. (1963) Glaucophane-bearing metamorphic rock types of the Cazadero area,

California. J. Petrol. 4, 260-301.

Coleman, R.G., Lee, D.E., Beatty, L.B. & Brannock, W.W. (1965) Eclogites and eclogites: their

differences and similarities. Geol. Soc. Am. Bull. 76, 483-508.

Abstract. Eclogites are divisible into three groups based on mode of occurrence; garnets and pyroxenes

both reflect compositional changes related to occurrence. Garnets in eclogite inclusions in kimberlites,

basalts, or ultramafic rocks are greater than 55 % Pyp, in bands or lenses of eclogite within migmatite

gneissic terrains, 30 to 55 % Pyp, and in bands or lenses of eclogite within alpine-type metamorphic rocks,

less than 30 % Pyp. Jd content progressively increases in the first two, whereas Di decreases. Difference

in Ca-Mg partition between coexisting garnet-pyroxene in eclogites of the same bulk composition indicates

a broad range of P-T conditions during crystallization. Field evidence strongly suggests elimination of the

single eclogite metamorphic facies concept.

Courtillot, V., Feinberg, H., Ragaru, J.P., Kerguelen, R., McWilliams, M. & Cox, A. (1985) Franciscan

Complex limestone deposited at 24°N. Geology 13, 107-110.

Crawford, W. A. (1965) Studies in Franciscan metamorphism near Jenner, California. PhD thesis, Univ.

California, Berkeley, 142 p.

Dalla Torre, M., De Capitani, C., Frey, M., Underwood, M.B., Mullis, J. & Cox, R. (1996) Very low-

temperature metamorphism of shales from the Diablo Range, Franciscan Complex, California: new

constraints on the exhumation path. Geol. Soc. Am. Bull. 108, 578-601.

Ernst, W.G. (1971) Petrologic reconnaissance of Franciscan metagreywackes from the Diablo Range,

Central California Coast Ranges. J. Petrol. 12, 413-437.

Jones, D.L., Blake, M.C. Jr., Bailey, E.H. & McLaughlin, R.J. (1978) Distribution and character of upper

Mesozoic subduction complexes along the West Coast of North America. Tectonophys. 47, 207-222.

Kerrick, D.M. & Cotton, W.R. (1971) Stability relations of jadeite pyroxene in Franciscan metagraywackes

near San Jose, California. Am. J.Sci. 271, 350-369.

Krogh, E.J., Oh, C.-W. & Liou, J.G. (1994) Polyphase and anticlockwise P-T evolution for Franciscan

eclogites and blueschists from Jenner, California, USA. J. Metam. Geol. 12, 121-134.

Krueger, S.W. & Jones, D.L. (1989) Extensional fault uplift of regional Franciscan blueschists due to

subduction shallowing during the Laramide Orogeny. Geology 17, 1157-1159.

Lee, D.E., Coleman, R.G. & Erd, R C. (1963) Garnet types from the Cazadero area, California. J. Petrol.

4, 460-492.

Abstract. The purpose of this study is to investigate the garnets present in metamorphic types III (bedrock

schists) and IV (tectonic blocks). 16 garnet analyses are from (aragonite-bearing) type III glaucophane

schists, and 8 are from type IV glaucophane schists. Type IV rocks include California eclogites. Type III

rocks include metabasalt, metachert, metashale, meta-ironstone, and metacarbonate that were formed

under high P/T. The garnets display a wide range of composition, dominantly Alm, Sps, and Grs. Type IV

rocks are mainly metabasalts that were probably formed under higher P-T than type III rocks. There is a

difference between garnets from type III rocks and those from type IV (including eclogites); the latter

contain less Sps and more Pyp, and the dominant molecules are Alm and Grs. The 4 garnets from

California eclogites have an average Pyp content of about 10 mol %, and they extend the range of

composition reported for eclogite garnets. Quantitative spectrographic determinations of minor elements

are listed for each of the garnets described. The values determined for some of the minor elements have

a wide range and a capricious distribution over a few feet of outcrop area. As a group, both the garnets

from type III rocks and those from type IV are pyralspites with large contents (as much as 35 mol %) of

ugrandite. This unusual admixture of the pyralspite and ugrandite series may have resulted in part from

the conditions (high P/T) under which the enclosing rocks were recrystallized.

Lee, D.E., Coleman, R.G., Bastron, H. & Smith, V.C. (1966) A two-amphibole glaucophane schist in the

Franciscan Formation, Cazadero area, Sonoma County, California. USGS Pap. Rept. P 0550-C,

C148-C157.

Abstract. A detailed study of a large glaucophane schist tectonic block has shown that the assemblage

actinolite-glaucophane-garnet-epidote has been partly replaced by retrograde chlorite and pumpellyite.

The whole rock and each of the major mineral phases have been analyzed for major and minor elements.

Results of this study are consistent with the hypothesis that the primary metamorphic assemblage

represents a high-grade schist that has been dislocated from its bedrock location by upward tectonic

movement. The retrograde assemblage may have formed during the period of tectonic dislocation.

Lee, D.E., Thomas, H.H., Marvin, R.F. & Coleman, R.G. (1964) Isotopic ages of glaucophane schists

from the area of Cazadero, California. USGS Prof. Pap., Rept. P 0475-D, D105-D107.

Abstract. Glaucophane-bearing tectonic blocks are stratigraphically out of place with respect to the

glaucophane-bearing bedrock terrain on which they rest. Five isotope age determinations ranging from

130 to 150 m.y. indicate that both the tectonic blocks and the bedrock schists were recrystallized as part

of a Late Jurassic and Early Cretaceous metamorphic event.

Mattinson, J.M. (1986) Geochronology of high-pressure--low-temperature Franciscan metabasite: a new

approach using the U-Pb system. In Blueschists and eclogites, Evans, B.W. & Brown, E.H., eds.,

Geol. Soc. Am. Mem.164, 95-105.

Mattinson, J.M. & Echeverria, L.M. (1980) Ortigalita Peak gabbro, Franciscan Complex: U-Pb dates of

intrusion and high-pressure-low-temperature metamorphism. Geology 8, 589-593.

McDowell, F.W., Lehman, D.H., Gucwa, P.R., Fritz, D. & Maxwell, J.C. (1984) Glaucophane schists and

ophiolites of the Northern California Coast Ranges: isotopic ages and their tectonic implications.

Geol. Soc. Am. Bull. 95, 1373-1382.

McLaughlin, R.J., Kling, S.A., Poore, R.Z., McDougall, K. & Beutner, E.C. (1982) Post-middle Miocene

accretion of Franciscan rocks, northwestern California. Geol. Soc. Am. Bull. 90, 595-605.

Moore, D.E. (1984) Metamorphic history of a high-grade blueschist exotic block from the Franciscan

complex, California. J. Petrol. 25, 126-150.

Moore, D.E., Liou, J.G. & King, B.S. (1981) Chemical modifications accompanying blueschist facies

metamorphism of Franciscan conglomerates, Diablo Range, California. Chem. Geol. 33, 237-263.

Moore, D.E. & Liou, J.G. (1980) Detrital glaucophane schist pebbles from Franciscan metaconglomerates

of the Northeast Diablo Range, California. Am. J.Sci. 280, 249-264.

Moore, D.E. & Liou, J.G. (1979a) Mineral chemistry of some Franciscan blueschist facies

metasedimentary rocks from the Diablo Range, California. Geol. Soc. Am. Bull. 90, I 1089-1091, II.

1737-1781.

Moore, D.E. & Liou, J.G. (1979b) Chessboard-twinned albite from Franciscan metaconglomerates of the

Diablo Range, California. Am. Mineral. 64, 329-336.

Murchey, B.L. & Jones, D.L. (1984) Age and significance of chert in the Franciscan Complex, in the San

Franciscan Bay region. In Franciscan geology of Northern California, Blake, M.C., Jr., ed., Field Trip

Gdbk – Pac. Sect., Soc. Econ. Paleontologists Mineralogists, 43, 23-30.

Nelson, B.K. (1991) Sediment-derived fluids in subduction zones; isotopic evidence from veins in

blueschist and eclogite of the Franciscan Complex, California. Geology 19, 1033-1036.

Nelson, B.K. & DePaolo, D.J. (1982) Sr and Nd isotopic composition of Franciscan eclogite and

blueschist: a sampling of subducted crust? Eos 63, 1133.

Oh, C.-W., Liou, J.G. & Maruyama, S. (1991) Low-temperature eclogites and eclogitic schists in Mn-rich

metabasites in Ward Creek, California: Mn and Fe effects on the transition between blueschist and

eclogite. J. Petrol. 32, 275-302.

Oh, C.-W. & Liou, J.G. (1990) Metamorphic evolution of two different eclogites in the Franciscan

Complex, California, USA. Lithos 25, 41-53.

Pampeyan, E.H. (1963) Geology and mineral deposits of Mount Diablo, Contra Costa Co., California.

Calif. Div. Mines Spec. Rept. 80, 31 p.

Patrick, B.E. & Day, H.W. (1989) Controls on the first appearance of jadeitic pyroxene, northern Diablo

Range, California. J. Metam. Geol. 7, 629-639.

Platt, J.P. (1975) Metamorphic and deformational processes in the Franciscan Complex, California; some

insights from the Catalina Schist terrane. Geol. Soc. Am. Bull. 86, 1337-1347.

Radvanec, M., Banno, S. & Ernst, W.G .(1988) Chemical microstructure of Franciscan jadeite from

Pacheco Pass, California. Am. Mineral. 83, 273-279.

Ransome, F.L. (1895) On lawsonite, a new rock-forming mineral from the Tiburon Peninsula, Marin

County, California. Univ. Calif. Pub. Geol. Sci .1, 301-312.

Ross, J.A. & Sharp, W.D. (1988) The effects of sub-blocking temperature metamorphism on the K/Ar

systematics of hornblendes: Ar-40/Ar-39 dating of polymetamorphic garnet amphibolite from the

Franciscan Complex, California. Contr. Mineral. Petrol. 100, 213-221.

Shibakusa H. & Maekawa H (1997) Lawsonite-bearing eclogitic metabasites in the Cazadero area,

northern California. Mineral. Petrol. 61, 163-180.

Abstract.: In the Cazadero area, northern California, lawsonite-bearing eclogitic metabasites occur in

association with glaucophane schists. Lawsonite-bearing eclogitic metabasites are coarse-grained and

usually lack albite. Typical mineral assemblages are garnet-omphacite-lawsonite-quartz-epidote-

glaucophane-chlorite, garnet–omphacite-lawsonite-pumpellyite-epidote-glaucophane-quartz and garnet-

omphacite-lawsonite-pumpellyite-actinolite-quartz-glaucophane. They can be represented on an Al2O3-

Fe2O3-MgO-Na2O diagram in which all minerals are projected from quartz, lawsonite, almandine garnet,

and H2O-predominant fluid. On the basis of the garnet-clinopyroxene geothermometry and phase

relations, the metamorphic conditions for the formation of lawsonite-bearing eclogitic metabasites are

estimated at 360-445°C and > 9±1 kbar. Lawsonite-bearing eclogitic metabasites formed near the

univariant curve albite = jadeite + quartz. A petrogenetic grid constructed by Schreinemakers' method

shows that the lawsonite-bearing eclogitic metabasites in the Cazadero area formed under transitional P-T

conditions between those of the garnet-bearing glaucophane schists in New Caledonia and lawsonite-

bearing eclogitic metabasites in Corsica.

Tagami, T. & Dumitru, T.A. (1996) Provenance and thermal history of the Franciscan accretionary

complex: constraints from zircon fission track thermochronology. J. Geophys. Res. B 101, 11,353-

11,364.

Terabayashi, M. & Maruyama, S. (1998) Large pressure gap between the coastal and central Franciscan

belts, Northern and Central California. Tectonophys. 285, 87-101.

Wakabayashi, J. & Deino, A. (1989) Laser-probe 40Ar/39Ar ages from high grade blocks and coherent

blueschists, Franciscan Complex, California: preliminary results and implications for Franciscan

tectonics. Geol. Soc. Am. Abstr. Prog. 21, 267.



Phase equilibria

Carlson, W.D. & Rosenfeld, J. (1981) Optical determination of topotactic aragonite-calcite growth kinetics;

metamorphic implications. J. Geol. 89, 615-638.

Evans, B.W. (1990) Phase relations of epidote-blueschists. Lithos 25, 3-23

Frey, M., de Capitani, C. & Liou, J.G. (1991) A new petrogenetic grid for low-grade metabasites. J.

Metam. Geol. 9, 497-511.

Perkins, D., Westrum, E.F. Jr. & Essene, E.J. (1980) Thermodynamic properties and phase relations of

some minerals in the system CaO-Al2O3-SiO2-H2O. Geochim. Cosmochim. Acta 44, 61-84.

Poli S. & Schmidt M.W. (1997) The high-pressure stability of hydrous phases in orogenic belts: an

experimental approach on eclogite-forming processes. Tectonophys. 273, 169-184.

Redfern, S.A.T., Salje, E. & Navrotsky, A. (1989) High-temperature enthalpy at the orientational order-

disorder transition in calcite: implications for the calcite/aragonite phase equilibrium. Contr. Mineral.

Petrol. 101, 479-484.



Older metamorphic, volcanic and plutonic rocks in the Sierra Nevada

Brook, C A. (1977) Stratigraphy and structure of the Saddlebag Lake roof pendant, Sierra Nevada,

California. Geol. Soc. Am. Bull. 88, 321-334.

Schweickert, R.A. & Lahren, M.M. (1987) Continuation of Antler and Sonoma orogenic belts to the

eastern Sierra Nevada, California, and Late Triassic thrusting in a compressional arch. Geology 15,

270-273.

Schweickert, R.A. & Lahren, M.M. (1993) Triassic-Jurassic magmatic arc in eastern California and

western Nevada; arc evolution, cryptic tectonic breaks, and significance of the Mojave-Snow Lake

Fault. In Mesozoic paleogeography of the Western United States. II. Dunne, G.C. and McDougall,

K.A., eds., Field Trip Gdbk. - Pacific Sect., Soc. Econ. Paleo. Mineral. 71, 227-246.

Schweickert, R.A. and Lahren, M.M. (1999) Triassic caldera at Tioga Pass, Yosemite National Park,

California: structural relationships and significance. Geol. Soc. Am. Bull. 111, 1714-1722.

Abstract. A Middle or Late Triassic volcanic vent structure, named the Tioga Pass caldera, is exposed

near the eastern boundary of Yosemite National Park, California. The caldera and related volcanic and

plutonic rocks--part of an early Mesozoic continental-margin magmatic arc in east-central California--

formed prior to or during an episode of contractional deformation in the arc. Field relationships show that a

widespread 222 Ma rhyolitic ash-flow tuff was erupted as an extensive outflow sheet during the formation

of the caldera. The Late Triassic Lee Vining Canyon pluton may represent the subvolcanic magma

chamber that was partially evacuated during the eruption of the ash-flow tuff. The caldera wall is now

exposed as a highly irregular boundary between prevolcanic basement and intracaldera rocks that formed

by a combination of initial caldera collapse and subsequent intracaldera intrusive and extrusive events.

Intracaldera rocks include a thick section of Triassic metasedimentary and metavolcanic rocks on Gaylor

Peak, together with the Dana sequence on Mount Dana. All of the Triassic rocks of the Saddlebag Lake

pendant later underwent strong deformation and metamorphism involving folding and thrusting during

Middle Jurassic time. The caldera fill is now exposed in the lower plate of an east-vergent Jurassic thrust,

which emplaced lower Paleozoic through Jurassic(?) metasedimentary and metavolcanic rocks

structurally above the caldera fill. The results of this study indicate that caldera formation may occur in a

contractional arc setting. Structural and stratigraphic relationships described here may also provide clues

to recognition of other caldera and vent complexes in highly deformed metavolcanic sequences in the

western United States and elsewhere.

Stevens, C.H. & Greene, D.C. (2000) Geology of Paleozoic rocks in eastern Sierra Nevada roof pendants,

California. In Great Basin and Sierra Nevada, Lageson, D.R., Peters, S.G. & Lahren, M.M., eds., GSA

Field Guide, 2, 237-254.

Abstract. Rocks in the major roof pendants of the eastern Sierra Nevada have been mapped in various

degrees of detail to better understand their stratigraphy, internal structure, and geologic history, and their

relationships to other rock assemblages in the region. Ten formations ranging in age from Middle(?)

Cambrian to Middle(?) Permian are recognized in these pendants, which along with other minor pendants,

constitute a tectonostratigraphic unit called the Morrison block. Rocks of the Morrison block were first

deformed by north-northwest-trending thrust faults and footwall synclines involving strata as young as

Early or Middle Permian. We designate this event, which correlates with a similar pre-middle Early Triassic

event recognized in rocks near Tinemaha Reservoir, the Morrison orogeny. Structures produced during

this orogeny include a probable cryptic thrust fault separating rocks assigned to the Morrison block from

those in the Big Pine Creek pendant, which may belong to the White-Inyo block, and the Nevahbe thrust,

which separates lower from upper Paleozoic rocks in the eastern part of the Mt. Morrison pendant and

may separate the Pine Creek and Bishop Creek pendants. In the Mt. Morrison pendant structures

produced during the Morrison orogeny apparently were later refolded twice prior to sinistral displacement

on the Laurel-Convict fault, which cross-cuts older structures and is intruded by a pre-latest Late Triassic

dike. Other thrust faults in the eastern Sierra Nevada include the Golconda thrust of early Middle Triassic

age and the Lundy Canyon thrust of Late Triassic age. The Golconda thrust system apparently overprints

the Roberts Mountains thrust and separates rocks of the Morrison block from those of the Golconda and

Roberts Mountains allochthons in the Saddlebag Lake pendant, and perhaps from those of the Roberts

Mountains allochthon in the Northern Ritter Range and Log Cabin Mine pendants. After thrust-faulting, but

prior to intrusion of the Late Triassic Wheeler Crest Granodiorite, dextral movement on the Tinemaha fault

displaced Paleozoic facies and structural belts in the Sierra Nevada northward, producing most of the

present complicated paleogeographic patterns apparent in the region. Other less important structures

such as the Laurel-Convict fault have further complicated the geology of the Morrison block.

Stevens, C.H. and & Stone, P. (2002) Correlation of Permian and Triassic deformations in the western

Great Basin and eastern Sierra Nevada: evidence from the northern Inyo Mountains near Tinemaha

Reservoir, east-central California. Geol. Soc. Am. Bull. 114, 1210-1221.

Abstract. Geologic relationships exposed near Tinemaha Reservoir southeast of Big Pine, east-central

California, provide key chronological and structural constraints for linking Permian and Triassic

deformational events recognized in the White and Inyo Mountains in the western Great Basin with those in

the eastern Sierra Nevada, particularly the Mount Morrison and Saddlebag Lake pendants. Permian to

earliest Triassic deformation in the Tinemaha Reservoir area produced a large north- to northwest-

trending, east-vergent, originally recumbent syncline (Mule Spring syncline) cut by an overriding thrust

(Strange Hill thrust). We correlate this deformation with a middle Permian to earliest Triassic contractional

deformation recognized in the southern Inyo Mountains, and with a major episode of folding and thrust

faulting in the eastern Sierra Nevada. After a period of tectonic quiescence and marine sedimentation in

the Early Triassic, rocks in the Tinemaha Reservoir area were refolded twice, producing distinctive sets of

steeply plunging folds. Similar structures in the Mount Morrison pendant that formed prior to intrusion of a

225±16 Ma dike along the Laurel-Convict fault are correlated with those in the Tinemaha Reservoir area.

The timing of these folding events may be similar to that of displacement on the Golconda and/or Lundy

Canyon thrusts in the Saddlebag Lake pendant. Close structural and stratigraphic ties suggest that rocks

in the northeastern part of the Mount Morrison pendant and Tinemaha Reservoir area, now separated by

approximately 65 km of dextral displacement along the cryptic Tinemaha fault, originally lay adjacent to

one another. This offset postdates Triassic folding and is inferred to predate emplacement of the latest

Triassic Wheeler Crest Granodiorite, which crops out across the projected fault trace.



Younger volcanic rocks east of the Sierra Nevada

Bailey, R.A. (1987) Long Valley Caldera, eastern California. In Centennial Field Guide, DNAG, Geol. Soc.

Am., Boulder, CO, trip 36, 163-169. QE 77.C461 Sci Lib

Bailey, R.A., Miller, C.D. & Sieh, K. (1989) Excursion 13B: Long Valley caldera and Mono-Inyo Craters

volcanic chain, eastern California. In Field Excursions to Volcanic Terranes in the W US, Chapin, C.

& Zidek, J.. eds., New Mexico Bur. Mines Mineral Res. Mem. 47, 227-254.

Sheridan, M.F. (1975) Tectonic displacement of the Bishop Tuff. In A Field Guide to Cenozoic

deformation along the Sierra Nevada Province and Basin and Range Province boundary. Calif. Geol.

28, 107-110.

Appendix VI. Field Trip to Long Valley Caldera, Mono Craters and vicinity



by Jon Davidson, taken largely from Bailey (19xx)

Geological setting

Long Valley caldera (Fig. 1) is located at the western edge of the Basin and Range province

straddling the eastern frontal fault escarpment of the Sierra Nevada, in which it forms a reentrant or offset

commonly referred to as the "Mammoth embayment." The floor of the caldera ranges in elevation from

2000 m in its eastern half, where it is dominated by Lake Crowley and sage and grass-covered Long

Valley, to 2600 m in its westerm half which is hillier and heavily forested. The caldera walls rise steeply to

elevations of 3000-3500 m on all sides except the east and southeast where the floor rises only 150 m

before merging with the Volcanic Tableland at 2300 m elevation. The Mono-Inyo Craters volcanic chain

extends from the western part of Long Valley caldera northward from Mammoth Mountain to Mono Lake,

a distance of 50 km. Although commonly described as subparallel to the Sierran front, the chain trends

nearly due north at a noticeable angle to the northwest-trending Sierran faults (Figs. 1, 2). The

prevolcanic basement in the area is mainly Mesozoic granitic rocks of the Sierra Nevada batholith plus

Paleozoic metasedimentary rocks and Mesozoic metavolcanic rocks of the Mt. Morrison and Ritter Range

roof pendants. The late Tertiary terrain upon which Long Valley volcanism was initiated was a maturely

eroded upland drained by westward-flowing streams.

Volcanism

Volcanism associated with Long Valley caldera (Bailey et al., 1976; Bailey, 1989) began with

widespread eruption of trachybasaltic-trachandesitic lavas between 3.6 and 2.2 Ma. Erosional remnants

of these precaldera lavas are scattered discontinuously over a 4000 km 2 area around the caldera (Fig. 2),

a distribution that suggests an extensive mantle source region for these initial mafic eruptions. Slightly

younger rhyodacite domes and flows associated with these mafic lavas erupted near the north and

northwest rims of the present caldera between 3.2 and 2.6 Ma. They probably represent the onset of

deep-crustal magmatic accumulation and differentiation that eventually culminated in formation of the

large, shallow Long Valley magma chamber from which subsequent more silicic eruptions originated. The

first eruptions from this silicic chamber were on the northeast rim of the present caldera at Glass

Mountain, where 1000 m of high-silica rhyolite domes, flows, and tuffs accumulated between 2.1 and 0.8

Ma (Metz & Mahood, 1985).

Catastrophic rupturing of the roof of the silicic magma chamber of 0.73 Ma expelled at least 600 km 3

of rhyolite magma as plinian ash falls and incandescent ash flows. This partial emptying of the chamber

caused collapse of its roof to form the 2-3 km deep oval depression of Long Valley caldera. The resulting

ash-flow deposits, the Bishop Tuff (Gilbert, 1938; Hildreth, 1979), inundated 1500 km 2 around the caldera

and accumulated locally to thicknesses approaching 200 m on the Volcanic Tableland and to lesser

thicknesses in upper Owens Valley, Adobe Valley, and Mono Basin. A large volume of Bishop Tuff also

ponded within the caldera as it collapsed; although the tuff is not exposed at the surface within the

caldera, drill holes have confirmed that as much as 1500 m of Bishop Tuff is buried beneath younger

volcanic and sedimentary caldera fill. During this climactic eruption, associated plinian ash clouds drifted

thousands of kilometers downwind and deposited an ash layer (informally termed the Bishop ash) as far

east as Kansas and Nebraska (Izett, 1982), as well as in southern California (Merriam & Bischoff, 1975).

Bishop ash is found in deep-sea cores from the East Pacific Ocean (Sarna-Wojeicki et al., 1987). The

stratigraphy of the Bishop Tuff eruption has recently been recognized to be more complex than previously

described, and many of the pyroclastic flows can be shown to be co-plinian, so that flows occurred

simultaneously with fall, with distribution varying largely with topography (Wilson & Hildreth, 1997). The

volume of differentiated magma represented by the Bishop Tuff represents a real problem in terms of

origin. Whether it is derived by fractional crystallization of basalt, or by crustal melting, an enormous

volume of mantle-derived magma is required to provide the needed heat and/or mass (Knesel &

Davidson, 1997). Isotopic studies of the Bishop Tuff and earlier Glass Mountain rhyolites indicate that

magma chambers might exist in isolation for a million years or so (e.g. Halliday et al., 1989; Christiansen

et al., 1993, 1996; Davies et al., 1994). This is also a serious physical problem as it is very difficult to

prevent the magma from cooling and crystallizing over such a long period of time.

After collapse of the roof of the magma chamber, rhyolitic volcanism continued on the caldera floor.

Pyroclastic eruptions followed by extrusion of thin, hot, fluid rhyolite flows produced a 100-500 m thick

sequence of intracaldera tephra and lavas informally designated as the early rhyolite. They are typically

aphyric to sparsely porphyritic, containing less than 5% crystals of plagioclase, hypersthene, biotite, and

Fe-Ti oxides. They contrast strikingly with the preceding crystal-rich Bishop Tuff, suggesting a marked

change in magmatic conditions after caldera collapse. Simultaneous renewal of magma pressure

accompanying this early rhyolite episode uplifted, arched, and faulted the early rhyolite flows and tephra,

forming a resurgent dome with a northwest-trending medial graben (Figs. 2, 8). The resurgent dome

formed within 100,000 yrs after caldera collapse, between 730 and 650 ka as constrained by K-Ar ages of

the contemporaneous early rhyolite flows (Bailey et al., 1976; Mankinen et al., 1986).

After a quiescent interlude of about 100,000 yrs, crystal-rich rhyolite again began erupting within the

caldera, mainly in the moat--probably from ring fractures peripheral to the resurgent dome. This coarsely

porphyritic rhyolite, informally designated as the moat rhyolite, typically contains up to 20% phenocrysts of

plagioclase, quartz, sandidine, biotite, hornblende, and Fe-Ti oxides. It forms thick, steep-sided domes

and flows suggesting higher viscosity and lower temperature than the early rhyolite, and it probably

signaled the onset of cooling and crystallization of the magma chamber. The moat rhyolite erupted at

about 200,000 yr intervals at 500, 300, and 100 ka in clockwise succession around the resurgent dome, in

the northern, southeastern, and western sectors of the moat, respectively. The 100 ka western moat

rhyolites appear to be the youngest extrusions so far derived from the Long Valley magma chamber.

However, recent seismological and geodetic studies suggest that a body of partially molten magma still

underlies the resurgent dome and is a potential source for future eruptions (see below: Recent seismicity

and ground deformation).

This 3.6 Ma to 100 ka mafic-to-silicic sequence of volcanism centered on Long Valley caldera is

overlapped spatially and temporally by the Mono-Inyo Craters volcanic chain, a younger mafic-to-silicic

sequence localized along a 50 km north-trending fissure system which extends from Mammoth Mountain

on the southewstern caldera rim through the western caldera moat to Mono Lake. This younger sequence

began between 300 and 200 ka with the eruption of trachybasaltic-trachyandesitic lavas in and near the

west moat, where they accumulated to at least 250 m thickness and poured around the resurgent dome

sending lava tongues eastward into both the north and south moat (Fig. 2). Younger mafic lavas vented

successively farther north near June Lake (40-20 ka) and at Black Point on Mono Lake (13,300 yrs B.P.),

suggesting localization along a northward-propagating fissure system. During these mafic eruptions,

rhyodacite domes and flows erupted sporadically in the western part of the caldera as well as farther north

in and near Mono Lake. The greatest accumulations of rhyodacite are on the northwest and southwest

caldera margins where the fissure system apparently intersected caldera ring fractures. On the southwest

caldera rim, between 200 and 50 ka, repeated extrusion of rhyodacite domes and flows built the imposing

cumulovolcano of Mammoth Mountain.

Rhyolites began erupting along the Mono-Inyo fissure system about 35 ka--first at the Mono Craters

chain, NW of the caldera, and more recently at the Inyo Craters chain, which spans the NW caldera rim

and extends into the west moat. The Mono Craters form a 17 km long, arcuate chain of 30 or more

coalesced rhyolite domes, flows, and craters ranging in age from about 35 ka to 600 yrs ago (Wood,

1977b; Sieh & Bursik, 1986). They are composed of high-silica rhyolite and consist predominantly of

aphyric to sparsely porphyritic obsidian, pumiceous glass, and tephra (Kelleher & Cameron, 1989). The

most recent of these eruptions occurred at the northern end of the chain from a 6 km long line of vents

that ejected about 1 km3 of magma as plinian and subplinian pyroclastic falls and pyroclastic flows and

surges followed by vent-filling domes and flows (Sieh & Bursik, 1986).

The Inyo Craters chain forms a 12 km long, discontinuous line of mainly low-silica rhyolite domes,

flows, and craters ranging in age from about 6000 to 500 yrs B.P. The youngest Inyo eruptions (650-550

yrs B.P.) began explosively (Figs. 10-12) and culminated with extrusion of Obsidian, Glass Creek, and

Deadman Creek domes (Fig. 8), which apparently were fed by a shallow, 8 km long rhyolite dike (Miller,

1985). The two southernmost of these youngest domes consist of two distinctly different rhyolites--a light-

colored, coarsely porphyritic, pumiceous rhyolite and a more silicic, sparsely porphyritic, rhyolite--which

are locally commingled in spectacular "marbled-cake" fashion. Petrologic and geochemical studies

(Sampson & Cameron, 1987) indicate that the two types probably came from separate magma chambers

and commingled within the conduits during erupton. The youngest Inyo eruptions succeeded the

youngest Mono eruption by only a few years (Sieh & Bursick, 1986), so that the northern and southern

ends of the Mono-Inyo chain were active almost simultaneously about 600 yrs B.P.

Phreatic eruptions both preceded and succeeded these youngest magmatic eruptions along the Inyo

chain. Shortly before the Inyo tephra eruptions, phreatic explosions broke out on the north face of

Mammoth Mountain, forming six small craters subparallel to the caldera wall. Shortly after the tephra

eruptions, several phreatic eruptions occurred in the west moat near Deer Mountain, a 115 ka dome of

moat rhyolite, forming several large craters, two of which contain the Inyo Craters Lakes. The latter

eruptions, according to radiocarbon dating and dendrochronological evidence (Wood, 1977a), occurred

between 1340 and 1460 A.D. The most recent eruptive activity in the region, however, occurred in Mono

Lake between 1720 and 1850 A.D., with the emergence of Paoha Island during intrusion of a rhyolite

cryptodome (Stine, 1984).

Structure

The caldera ring fracture is not exposed, but its general location is suggested by the distribution of

the moat rhyolite vents, which are within the annular caldera moat and peripheral to the resurgent dome.

The dominant structural trend in the Long Valley region is northwest, subparallel to the Sierran frontal fault

zone, which parallels the elongation of the Sierran plutons and the associated roof pendants. Within the

caldera, this northwest trend is reflected by the medial graben on the resurgent dome and by the

alignment of the vents of the early and moat rhyolites.

Two major Sierran frontal faults transect the caldera, the Hilton Creek fault on the southeast and the

Hartley Springs fault on the northwest (Fig. 2). These faults, which show major pre- and postcaldera

displacement outside the caldera, terminate as major escarpments at the caldera margin and continue

within only as minor discontinuous splinters. This marked difference in intracaldera and extracaldera

displacement suggests that the subcaldera magma chamber tended to hydraulically dampen or absorb

intracaldera fault movements and that only recently has the chamber roof thickened sufficiently (as a

result of cooling and crystallization) to behave rigidly and transmit tectonic stresses through the roof block

(Bailey et al., 1976). An alternative, or possibly additional, explanation for the lack of significant structural

relief on intracaldera tectonic faults is that regional extension within the caldera has been accommodated

volumetrically by repeated injection of magma into the chamber or into fissures and faults (Bursik & Sieh,

1989).

Recent seismicity and ground deformation

Based on the 200,000 yr eruption periodicity of the moat rhyolite and on the age of the youngest 100

ka west moat domes, future eruptions from the residual Long Valley magma chamber would seem only a

remote possibility, not to be expected for another 100,000 yrs. However, an unusual succession of

earthquake swarms, including a sequence of magnitude-6 earthquakes in May 1980, and another of

magnitude-5 earthquakes in January 1983, accompanied by 50 cm uplift of the resurgent dome, suggests

that new magma has been injected into the chamber and possibly into the south moat ring-fracture zone

(Savage & Clark, 1982; Savage & Cockerham, 1984). Although the intensity of seismicity and the rate of

uplift and horizontal distention across the resurgent dome decreased since 1984, circumstantial structural

and geophysical evidence indicates the persistence of a residual magma chamber beneath the caldera as

well as the potential for future eruptions (Hill et al., 1985; Rundle and Hill, 1988). In the last few years

though, seismicity has increased again with several significant earthquake swarms in the area of the

caldera – these can be best appreciated through the USGS web site

http://quake.wr.usgs.gov/VOLCANOES/LongValley/Current.html.

Statistically, however, the more likely site for future eruptions in the regions is along the Mono-Inyo Craters

volcanic chain where eruptions have occurred most recently about 450, 700, 1200, and 1345-1469 A.D.

(Miller, 1985; Sieh & Bursik, 1986; Sieh, unpubl. data). The potential hazards associated with possible

future eruptions in the area have been outlined by Miller et al. (1982).

The route can be varied according to weather or other choices. The itinerary below provides an overview

on Day 1, including some of the youngest volcanism and travels as far north as Mono Lake. Day 2

focuses on the Bishop Tuff itself, and does not venture as far north, thus shortening the return journey to

Los Angeles.

Day 1

Travel north from Bishop up Sherwin Grade (the surface of the Bishop Tuff) and pull out beyond top of the

hill just before turn off to Toms Place. Take CARE crossing the road to the outcrop on the east side.

(7) Big Pumice Cut: Bishop Tuff-Sherwin Till contact. In this roadcut on the north side of US-395, the

basal, bedded pumice fall and overlying non-welded ash flows of the 0.73 Ma Bishop Tuff rest on the

weathered Sherwin Till (Sharp, 1968). The contrasting modes of emplacement of the pumice-fall and

ash-flow members of the Bishop Tuff are well displayed here--the eastward-dipping airborne pumice fall

mantling the initial sloping surface of the deeply weathered till, and the overlying fluidized ash flows

banking against the slope, with their nearly horizontal pumice swarms converging with the dipping pumice-

fall beds. Vertically cutting the tuff and till are several north-trending clastic dikes that are subparallel to

major Sierran frontal faults in the area. In the till, some of these dikes cut cleanly through large boulders

weathered to incoherent gruss. The source for the clastic dike fillings is a Pleistocene glacio-fluvial

deposit that overlies the tuff at the top of the roadcut; note how the dikes become finer-grained downward,

with only the upper parts containing pebbles and cobbles.

Follow 395 N. to Mammoth turnoff. Drive through Mammoth, turning right on rt 203, and right on

Mammoth Scenic Loop 0.9 miles later. Turn left to Inyo Craters (brown signpost) and follow track to

parking lot.

2. Inyo Craters: Phreatic explosion craters and deposits. The Inyo Craters are three northerly aligned,

phreatic explosion craters (Rinehart & Huber, 1965) on the south flank and summit of Deer Mountain, a

115 ka porphyritic rhyolite dome in the west moat of the caldera (Fig. 5). What is meant by phreatic? How

do we distinguish phreatic deposits from pyroclastic deposits? The two southernmost craters (informally

designated as north crater and south crater) are about 200 m in diameter, about 60 m deep, and contain

small lakes; the crater on the summit of Deer Mountain (informally referred to as summit crater) is smaller,

irregular in outline, breached on its south side, and dry. The lake in south crater is yellowish green,

suggesting the presence of suspended sulfur, but the water is cold (11°C) and other evidence of thermal

influx is lacking. The lake in north crater is colored brown with organic matter. Despite apparent

differences in morphology and vegetation, all three craters formed at nearly the same time, probably within

hours or days. They erupted in succession from north to south, as the light-colored deposits from summit

crater underlie the darkest ones of south crater--relations that are well exposed in the northeast wall of

north crater and on the east flank of Deer Mountain (Mastin, 1988). The light-colored debris around

summit crater is composed primarily of pulverized hornblende-biotite rhyolite of Deer Mountain dome,

whereas the darker debris around north and south craters consists largely of fragmented trachyandesite

derived from flows like those exposed in the walls of south crater. In the north wall of south crater,

exposed in succession above the trachyandesite flows, are: (1) trachyandesitic cinder from nearby vents

to the northeast and southeast (10 m), (2) pumiceous rhyolite tephra from the Inyo domes magmatic

eruptions centered a few kilometers to the north (1 m), and (3) coarse, crudely bedded phreatic explosion

deposits mainly from south crater (13 m). The latter, which extend as far as 1 km from the crater (Fig. 5),

consist mainly of trachybasalt-trachyandesite blocks in a grayish-brown, compact to semi-indurated, fine-

grained matrix; blocks and boulders of granite and metamorphic rocks, probably derived from subsurface

glacial till or possibly Sierran basement, also are included in the debris. A battered log incorporated in the

deposit of south crater yielded a radiocarbon age of 710±60 yrs B.P., which together with

dendrochronlogical data gave a calendar age of between 1340 and 1460 A.D. for the phreatic eruptions

(Wood, 1977a). How does radiocarbon dating work? Could you use it to date the Bishop Tuff?

The terrain around the craters is broken by many north-trending faults and fissures (Fig. 5), which

show up to 20 m displacement individually and define a graben 0.6 km wide and 2.5 km long. The graben

probably formed as a consequence of uplift and distension above a rising dike (Mastin & Pollard, 1988),

which probably generated the phreatic explosion craters. To test this hypothesis, Sandia National

Laboratories slant-drilled an 865 m scientific corehole (Fig. 6) eastward at an angle of 68° from a site on

the top of the fault scarp just west of the south crater (Eichelberger et al., 1988). It passed through a 320

m sequence of postcaldera trachyandesite and trachybasalt flows, about 50 m of gravel, 360 m of tuffs

and flows of the early rhyolite, 63 m of glacial till, and terminated in 30 m of Paleozoic quartzite presumed

to be Sierran basement rock. At a depth of about 600-650 m, within the early rhyolite section and directly

beneath the center of south crater, the hole penetrated an apparent vent breccia consisting primarily of

pulverized early rhyolite and postcaldera trachybasalt. Included within the breccia were small pumiceous

fragments of a distinctive high-silica rhyolite, apparently juvenile and considered to be the magma that

generated the phreatic explosions. Presumably a dike or conduit of this magma rising from depth

encountered water-saturated tuffs of the early rhyolite, causing explosive flashing of the water to steam,

which in turn reamed a vent through the overlying postcaldera trachybasalt sequence, producing the Inyo

explosion craters and surrounding phreatic deposits. Surprisingly, little or no juvenile rhyolite magma

reached the surface during the eruptions.

Return to the Scenic Loop road, and turn left, arriving at the junction with U.S. 395. Turn left, and at turn

left on the dirt road signposted to Hartley Springs Campground. 1.1 miles down the track fork left to

obsidian dome (signposted) and continue to end (1.3 miles) where you can park.

3. Glass Creek Dome: Commingled rhyolite lava. Directly south of Glass Creek is the steep, talus- and

block-mantled, north margin of Glass Creek Dome. The dome is composed of two distinctive rhyolites

(Bailey, 1984b; Sampson, 1987); (1) sparsely porphyritic-rhyolite, typically consisting of black obsidian or

vitrophyre and occurring mainly on the dome margins, and (2) light-gray, pumiceous, coarsely porphyritic,

hornblende-biotite-rhyolite, occurring mainly in the center. Between the margin and center is an

intervening zone where the two lava types have commingled, producing fascinating "marbled cake"

structures (Fig. 7) that illustrate the contrast in viscosity between the relatively fluid obsidian and the more

viscous porphyritic rhyolite. Petrologic and chemical studies (Sampson and Cameron, 1987) indicate that

the two contrasting rhyolites originated and evolved in separate chambers, possibly confirming a

suggestion (Bailey et al., 1976) that the sparsely and coarsely porphyritic types came from the Mono

Craters and Long Valley magma chambers, respectively, and that they commingled along an

interconnecting fissure just prior to eruption.

Turn south on US 395 N and turn left onto gravel road to Lookout Mountain. Keep right at first fork, taking

left fork at 1.1 miles after the turnoff.

4. Lookout Mountain: Overlook of Inyo Domes and western caldera moat. Lookout Mountain is a small

rhyolite stratovolcano with a summit crater 2 km in diameter, which forms the tree-covered depression

immediately west of the viewpoint. The volcano is in the northwest moat on the northern edge of the

resurgent dome, and, although not part of the resurgent dome, it is constructed of interstratified 0.69 Ma

flows and tuffs of the early rhyolite. The aphyric obsidian underfoot is typical of the early rhyolite.

The view from the summit provides an excellent 360 degree view of Caldera features (Fig 8). To the

west can be seen the west moat of the caldera, underlain mainly by 150-160 ka postcaldera trachybasalts

and trachyandestites, and the west and northwest wall of the caldera underlain by Sierran granitic and

metamorphic rocks and capped by Pliocene precaldera volcanic rocks. In the middle distance are the

barren, craggy, Holocene rhyolite dome-flows of the Inyo Craters chain, including from north to south:

Wilson Butte, Obsidian Dome, Deadman Creek Dome, and Glass Creek Dome which we will visit

tomorrow.

Return to U.S. 395 and turn right. In sight of Mono Lake, turn right to Benton. The (brown) signpost to

Panum Crater is just before you draw level with it (after turnoff to the landfill).

5. Panum Crater. Panum Crater (Fig. 9) is the northernmost and youngest vent of the North Mono

eruption (Sieh & Bursick, 1986). The stratigraphic sequence of deposits emplaced during this youngest

activity (Figs. 9, 10) includes (in ascending order): (1) a vent-clearing breccia, (2) four pyroclastic-flow and

-surge deposits varying considerably in character, (3) a tephra-ring deposit, and (4) a composite, crater-

filling dome. The sequence suggests two episodes of explosive activity from the same vent, each

followed by dome extrusion. The initial vent-clearing eruption produced an early tephra ring and crater,

from which early pyroclastic-flow and -surge deposits were erupted. An early dome rising into this crater

apparently collapsed or exploded laterally, breaching the northwest side of the tephra ring and sending a

block avalanche or coarse pyroclastic flow northwestward into Mono Lake (Wood, 1977b; Sieh & Bursik,

1986). This outburst possibly caused momentary decompression in the conduit and more pyroclastic

flows followed. Subsequent, less energetic, pyroclastic-fall eruptions partially healed the breach in the

northwest side of the tephra ring, and finally the current composite dome was extruded (Figs. 9, 10). Why

might you expect pyroclastic eruptions to be followed by lava extrusion? Panum Dome consists of three

subunits (Fig. 9b) extruded in succession in the east and southwest, the south, and finally the north. Both

the eastern and southwestern subunits have a thin mantle of tephra on them, suggesting that they are the

shouldered-aside remnants of the early dome that exploded and produced the northwest breach in the

crater rim and the block avalanche to the northwest. Extrusion of the southern and northern parts of

Panum Dome concluded the North Mono eruption episode. Reconstruction of the 1345 A.D. North Mono

vents suggests that they developed above a rising rhyolite dike that broke to the surface in northward

succession (Fig. 11).

Time permitting, continue on the Benton road and turn off to Mono Lake (Tufa Preserve). There is a $2

charge to visit the Reserve, and there are information boards and a labeled trail illustrating the geology

biology, ecology and history of Mono Lake.

6. Mono Lake State Nature Preserve: Tufa towers. The delicate, castellated, white towers along the

shore of Mono Lake are calcareous sublacustrine spring deposits formed where fresh-water springs

percolate through lake-bottom sediments and flow up through the saline water. Calcium in the fresh water

combines with carbonate of the saline lake water, resulting in precipitation, possibly catalyzed by algae, of

calcite or aragonite. The towers formed entirely underwater and are being exposed as the lake level

declines. Numerous signs along the well-marked trail through the tufa towers and along the lake shore

describe the many fascinating and unique features of the Mono Lake habitat, including its brine shrimp,

brine flies, bird life, and Indian and early settlement history.

Return to US 395. At the bottom of a hill just before re-entering the caldera, turn left onto Owens River

Road, which becomes gravel surface in a few miles. Stop to examine flow on N. side of road.

7. Owens River Road: Post caldera trachybasalt: On the north side of Owens River Road, the moat is

floored by a 15 km long, 108 ka, postcaldera trachybasalt flow that erupted from a vent in the west moat.

The dark-brownish crags marking the southern edge of this flow can be traced from the gap at Big Springs

just north of Lookout Mountain on the west to near the base of Glass Mountain on the east. The flow,

which shows no evidence of having flowed into water, indicates that the caldera lake was confined to the

eastern moat and may have been completely drained by this time. Just beyond the toe of the trachybasalt

flow, at the very base of Glass Mountain, is a small conical hill--a small, undated, postcaldera rhyodacite

dome that appears to have erupted subaqueously in Pleistocene Long Valley Lake in early postcaldera

time. The steep escarpment to the north is the northern caldera wall, underlain by Mesozoic plutonic and

Paleozoic metasedimentary rocks and capped with 3.5-2.2 Ma precaldera trachybasalts and

trachyandesites. On the rim to the northwest is Bald Mountain, composed of 3.4 Ma precaldera

rhyodacite domes and flows. On the far western skyline is the crest of the Sierra Nevada and somewhat

closer and lower the western caldera rim. On the northeast rim is Glass Mountain, a 2.1-0.8 Ma

precaldera rhyolite complex, the southwestern half of which is downfaulted and buried within the caldera.

The low, flat surface south of Glass Mountain is the southeast caldera rim and beyond is the Volcanic

Tableland, underlain by the Bishop Tuff. The high Sierran peaks on the southern skyline mark the south

rim of the caldera. The low forested hills in the south foreground are the resurgent dome, underlain by

730-650 ka tuffs and aphyric rhyolite flows of the early rhyolite. In the near southwest, the high knob rising

above the surface of the resurgent dome is a 520 ka lava dome of the moat rhylite, composed of coarsely

porphyritic hornblend-biotite rhyolite. On the sage-covered lower slopes of the resurgent dome, lake

terraces and strand lines of Pleistocene Long Valley Lake may be seen.

Continue on Owens River Road, crossing Hot Creek and a rhyolite flow before turning right. Keep right at

fork and reach Hot Springs parking lot.

8. Hot Creek Gorge: southeast moat rhyolite; hot-spring activity. Hot Creek is the lower reach of

Mammoth Creek where several vigorous hot springs issue from its bottom forming this popular bathing

area. Hot Creek Gorge has been incised into a 288 ka moat rhyolite lava flow that erupted from a vent on

the south shore of Pleistocene Long Valley Lake, which at that time was about 4 km to the south at an

elevation of 2195 m. The flow coursed northward into the lake, becoming entirely subaqueous in the

vicinity of the present gorge. The initially hot glassy rhyolite was pervasively hydrated by interaction with

the lake water and later was partly hydrothermally altered by local hot-spring activity. Outcrops of partly

altered, perlitized flow breccia are exposed along the paved trail leading to the bottom of the gorge. What

is perlite? Make sure you see it, and think about how it was formed.

The Hot Creek flow is a sparsely porphyritic, sanidine-augite-bearing rhyolite of the moat sequence.

Its chemistry and mineralogy differ from that of older, more coarsely porphyritic hornblende-biotite rhyolites

in the north and southeast moat, and its chemical fractionation pattern (relative to older flows) is similar to

that of Bishop Tuff. These petrographic and chemical changes at about 3000 ka possibly reflect thermal

rejuvenation of the Long Valley magma chamber brought about by a new influx of mafic magma into the

roots of the magmatic system, as shortly thereafter trachybasalts began erupting in the west moat (Bailey,

1984a). Renewed uplift of the resurgent dome probably accompanied this magmatic rejuvenation, as at

this time a sudden influx of coarse detritus derived from the resurgent dome built extensive flanking deltas

into the caldera lake. These deltaic sediments are intensely silicified and permeated with fossil hot

springs and fumaroles, suggesting that increased widespread hydrothermal activity also accompanied this

magmatic and structural activity.

Thermal springs issue from the stream banks all along Hot Creek Gorge, but the largest and hottest

springs are localized on two north-trending faults that bound a shallow, 1 km wide graben transecting the

Hot Creek flow and gorge (Fig. 2). The numerous smaller, cooler springs along the creek between the

two faults are produced by lateral flow of hot water from the faults through the basal breccia of the rhyolite

flow where mixing with cold meteoric water occurs. The several boiling pools in the gorge (93°C at this

elevation) commonly change in vigor and configuration in response to local eathquakes and are also

significantly affected by seasonal rise and fall of the water level in the creek. The white deposits lining the

pools are calcium carbonate (travertine). Numerous older travertine deposits surround extinct hot springs

at higher levels along the gorge.

Continue south/west. Turn left at the stop sign by the fish hatchery, and 0.5 miles later left on

U.S. 395, returning to Bishop

Day 2

From East Line/ US 395 junction drive 0.6 miles north to US Rt 6 - turn right. After 8.4 miles turn left

opposite Rudolph Rd on dirt track to pumice quarry (right fork as track heads west)

Stop 1a. Bishop Pumice Quarry: Bishop Tuff, basal pumice fall and distal ash flows. Exposed in this

quarry are 4 m of basal pumice fall and 4-6 m of distal ash flows of the Bishop Tuff. How can you

distinguish between the two? The pumice fall is roughly divisible into: (1) a lower third composed of

moderately well-bedded ash and fine pumice lapilli; (2) a middle third that is lighter-colored, less well

bedded and sorted, and coarser, with pumice lapilli up to 2 cm in diameter; and (3) an upper third that is

darker-colored, more distinctly bedded, with numerous fine-to-coarse alternations containing pumice lapilli

up to 4 cm in diameter. Discontinuous dark manganiferous(?) streaks are present in the upper third of the

deposit. The lack of distinct discontinuities in the lower two-thirds suggests continuous deposition over a

relatively short time, with activity becoming more variable or intermittent in the upper third. The tendency

for lithic clasts as well as pumice lapilli to increase upward in size suggests generally increasing eruption

intensity. Very subtle low-angle crossbedding together with the poor sorting in the lower part suggest

deposition from very dense, laterally drifting, eruption clouds. The overlying ashflows are generally

massive and poorly sorted, but the basal few centimeters consist almost entirely of fine ash, which locally

also shows faint low-angle cross-bedding suggestive of pyroclastic surge. The lowermost 2 m of the ash

flows contain many discontinuous swarms of relatively coarse pumice lapilli and blocks, indicating

complex multiflow deposition near this distal part of the formation.

Interbedded within the pumice-fall sequence at the north end of the quarry is a 1-2 m thick, unsorted,

tongue-shaped deposit with an irregular upper surface; it is buff-colored and consists of unusually coarse

pumice blocks and lapilli in an ashy matrix; it is a pyroclastic flow, correlative with extensive flow deposits

nearer to the caldera (Wilson and Hiildreth, 1997). What are the characteristics of this deposit that

suggest it is a pyroclastic flow - and how do these contrast with the pumice fall making up the rest of the

quarry walls?

At one or two localities in the quarry, low-angle reverse faults with as much as 1 m displacement can

be seen. The fault planes incline eastward with decreasing dip and become bedding-plane faults that are

difficult to trace laterally. In this structural setting, near the western boundary fault of the White Mountains

and on the eastern edge of the Volcanic Tableland, which is broken by numerous normal faults, reverse

faults seem anomalous. Possibly they are related to local compressional wedging or bending within

downfaulted blocks in an otherwise extensional region.

From pumice quarry, return to US 6 and turn right back towards Bishop. At 3.4 miles from the quarry, turn

right on Jean Blanc Road. At intersection with 5 Bridges continue through aggregate yard. At 6.8 miles

from stop 1, turn left on Chalk Bluff Road which parallels the river just to the south of the tuff. At 7.8 miles

reworked deposits are exposed on right hand side - probably from Glass Mountain eruptions predating the

Bishop Tuff.

1b. Glass Mountain Fluvio lacustrine deposits.

Stop and examine these deposits. How are they different from the pyroclastic deposits in the pumice

quarry? How were they deposited - and what is the evidence for this?

Continue to junction with US 395 (14.3 miles) turn right. At 20 miles turn right on Gorge Road, left at T

junction (top of the hill) and follow the pipeline. At 24 miles fork right to parking lot above Gorge.

2. Owens River Gorge: Section of Bishop Tuff; rosette jointing.

The Los Angeles Department of Water and Power (LADWP) upper powerhouse road descends the west

wall of Owens River Gorge and affords spectacular closeup views of a 150 m thick section of Bishop Tuff

on the east wall. At the same time, the road permits close-hand examination of roadcuts and outcrops

that show the vertical changes in density and texture of the entire tuff section.

On the east wall of the gorge, the Bishop Tuff displays a remarkable variety of columnar jointing. In

the upper half of the cliff the joint columns are remarkably well-formed five- to six-sided columns 1-3 m in

diameter. Locally these columns curve and converge downward toward common foci, forming joint

rosettes. These rosettes are the loci of large fossil fumaroles. What causes the rosette shape formed by

the columns? At the top of the gorge directly above many of these joint rosettes are fumarolic mounds like

those seen on the Bishop Tuff surface near Round Valley. The distribution of these mounds close to the

present gorge suggests that the volatiles responsible for their formation were derived in part from the

ancestral Owens River, which was overrun and vaporized by Bishop ash flows.

In the lower half of the gorge, in dark-gray, densely welded tuff, the joint columns are much larger,

10-20 m in diameter, and are relatively crudely formed as a consequence of slower cooling around more

widely spaced cooling centers. Locally, some of these large joint columns have secondary horizontal joint

columns developed perpendicular to their primary surface. Between the upper and lower tiers of joint

columns is a broadly undulating, ill-defined parting that marks the contact between two sub-cooling units

of the Bishop Tuff. This parting and another like it at the top of the gorge near the gate represent brief

hiatuses in deposition of the tuff and show that the tuff consists of multiple cooling units (Sheridan, 1967)

or at most 2-3 yrs (Snow & Und, 1988). The prominent crystal-rich parting near the gate is the contact

between the Tableland and Gorges cooling units (Hildreth, 1979); the Tableland (upper) unit contains two

pyroxenes and has Fe-Ti-oxide temperatures of 737-763°C, whereas the Gorges (lower) unit lacks

pyroxenes and has lower oxide temperatures of 725-736°C.

Exposures along the left (west) side of the powerhouse road display almost the entire range of

lithologic variation in the Bishop Tuff. Near the upper gate the rock is light-gray to pinkish or purplish,

porous, poorly to moderatly welded, vapor-phase tuff. Downward, the tuff grades through brown to dark

gray, becomes progressively less porous and more densely welded, and displays conspicuous eutaxitic

texture. In the most densely welded, darker-gray facies, eutaxitic texture is almost entirely obscured by

devitrification. At the first major left bend in the road, however, eutaxitic texture becomes more prominent

over a 5 m interval where scattered obsidian fiamme occur, suggesting a minor eruptive/cooling interval.

What is the difference between pumice and obsidian (such as that forming the fiamme) and why is the

former light-colored whereas obsidian is black? The tuff below this level is monotonously dark gray,

densely welded, and devitrified to the bottom of the gorge. The basal contact is not exposed in the gorge,

but the powerhouse foundation was excavated in dense eutaxitic vitrophyre, indicating that the base is not

far below. Why is the tuff welded here, but not at the pumice quarry where we stopped earlier?

Return to 395, cross it and continue N. to Lower Rock Creek. At the top of the hillpull off to left hand side

and walk north to gorge overlook.

3. Lower Rock Creek Gorge overlook: Zonation in the Bishop Tuff. Down-canyon on the east wall of the

gorge, the internal zonation of the Bishop Tuff is well displayed. Here, the tuff laps around a granitic high

on the Sierran frontal fault scarp and thickens southward toward Round Valley. Exposed from the base

upward are: (1) the lower nonwelded zone (in contact with the granite), (2) lower vitrophyre (dark-gray

layer dipping down-canyon), (3) central devitrified zone (brown zone gradually thickening down-canyon),

and (4) upper buff vapor-phase zone. The upper, vitric, nonwelded zone has been removed by erosion.

This is a good place to discuss the chemical properties of the Bishop tuff. Although loosely defined

as a rhyolite, there is distinct compositional zonation through the tuff from top to bottom (Fig 12a). The

chemical zonation is believed to reflect progressive tapping of a chemically zoned magma chamber (Fig

12b), such that the base of the tuff represents the earliest erupted material, and the top is the latest. The

top of the magma chamber is expected to be the most differentiated (highest SiO 2, and highest

incompatible element concentrations), and progressively less differentiated magma is tapped deeper into

the magma chamber (Hildreth, 1979; Cameron, 1984). What is responsible for the chemical zonation

believed to have existed in the Bishop Tuff magma chamber?

From here, continue N. to junction with 395, stopping at Toms Place for refreshments. Turn south and

return to Los Angeles. Time permitting, stop on Whitney Portal Road (turn right at Lone Pine) to view

Sierran granitoids and overview Owens Valley.

References

Bailey, R.A. (1984a) Chemical evolution and current state of the Long Valley magma chamber. In Hill,

D.P., Bailey, R.A. & Ryall, A.S. (eds.), Proceedings of Workshop XIX: Active tectonic and magmatic

processes beneath Long Valley caldera, eastern California: USGS Open-File Rept. 84-930, 24-40.

Bailey, R.A. (1984b) Introduction to the Late Cenozoic volcanism and tectonism of the Long Valley-Mono

Basin area, eastern California. In Lentz, J., Jr., ed., Guidebook for Western Geological Excursions.

Dept. Geol. Sci., MacKay School of Mines, Univ. Nevada, Reno, 2, 56-67.

Bailey, R.A. (1989) Geologic map of Long Valley caldera, Mono-Inyo Craters volcanic chain, and vicinity,

eastern California: USGS Misc. Invest. Ser. Map 1-1933, scale 1:62,500.

Bailey (19xx) Excursion 13B: Long Valley caldera and Mono-Inyo Craters volcanic chain, eastern

California. In Field Excursions to Volcanic Terranes in the Western United States, Bailey, R.A., Miller,

C.D. & Sieh, K. eds., New Mexico Bur. Mines Mineral Res. Mem. 47, 227-254.

Bailey, R.A., Dalrymple, G.B., & Lanphere, M.A. (1976) Volcanism, structure, and geochronology of Long

Valley caldera, Mono County, California. J. Geophys. Res. 81, 725-744.

Bursik, M.I. & Sieh, K.E. (1989) Range-front faulting and volcanism in the Mono Basin, eastern California.

J. Geophys. Res. (in press).

Cameron K.L. (1984) Bishop Tuff revisited: new rare earth element data consistent with crystal

fractionation. Science 224, 411-413.

Christensen, J.N. & DePaolo, D.J. (1993) Time scales of large volume silicic magma systems: Sr isotopic

systematics of phenocrysts and glass from the Bishop Tuff, Long Valley, California. Contr. Mineral.

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Christensen, J.N. & Halliday, A.N., 1996, Rb-Sr ages and Nd isotopic compositions of melt inclusions from

the Bihsop Tuff and the generation of silicic magma. Earth Planet. Sci. Lett. 144, 547-561.

Davies, G.R., Halliday, A.N., Mahood, G.A. & Hall, C.M. (1994) Isotopic constraints on the production

rates, crystallization histories and residence times of pre-caldera silicic magmas, Long Valley,

California. In 125, p. 17-37.

Eichelberger, J.C., Vogel, T.A., Younker, L.W., Miller, C.D., Heiken, G.H., & Wohletz, K.H. (1988)

Structure and stratigraphy beneath South Inyo Crater, Long Valley caldera, California. J. Geophys.

Res. 93, 13,208-13,220.

Gilbert, C.M. (1938) Welded tuff in eastern California. Geol. Soc. Am. Bull. 49, 1829-1862.

Halliday, A.N., Mahood, G.A., Holden, P., Metz, J.M., Dempster, T.J. & Davidson, J.P. (1989) Evidence

for long residence times of rhyolitic magma in the Long Valley magmatic system: the isotopic record in

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